Ancient Orogens and Modern Analogues (Geological Society, 2009)

Ancient Orogens and Modern Analogues
The Geological Society of London

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It is recommended that reference to all or part of this book should be made in one of the following ways:
MURPHY , J. B., KEPPIE , J. D. & HYNES , A. J. (eds) 2009. Ancient Orogens and Modern Analogues.
Geological Society, London, Special Publications, 327.
DOSTAL , J., KEPPIE , J. D. & FERRI , F. 2009. Extrusion of high-pressure Cache Creek rocks into the
Triassic Stikinia-Quesnellia arc of the Canadian Cordillera: implications for terrane analysis of ancient
orogens and paleogeography. In: MURPHY , J. B., KEPPIE , J. D. & HYNES , A. J. (eds) 2009. Ancient
Orogens and Modern Analogues. Geological Society, London, Special Publications, 327, 71 –87.
GEOLOGICAL SOCIETY SPECIAL PUBLICATION NO. 327
Ancient Orogens and Modern Analogues
EDITED BY
J. B. MURPHY
St Francis Xavier University, Canada
J. D. KEPPIE
Universidad Nacional Autonoma de Mexico, Mexico
and
A. J. HYNES
McGill University, Canada
2009
Published by
The Geological Society
London
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Contents
MURPHY , J. B., KEPPIE , J. D. & HYNES , A. J. Ancient orogens and modern analogues: an 1
introduction
Mesozoic –Cenozoic orogens

COLE , J. W. & SPINKS , K. D. Caldera volcanism and rift structure in the Taupo Volcanic Zone, 9
New Zealand
RAMOS , V. A. & FOLGUERA , A. Andean flat-slab subduction through time 31
MANDUJANO -VELAZQUEZ , J. J. & KEPPIE , J. D. Middle Miocene Chiapas fold and thrust belt of 55
Mexico: a result of collision of the Tehuantepec Transform/Ridge with the Middle America
Trench
DOSTAL , J., KEPPIE , J. D. & FERRI , F. Extrusion of high-pressure Cache Creek rocks into the 71
Triassic Stikinia–Quesnellia arc of the Canadian Cordillera: implications for terrane analysis of
ancient orogens and palaeogeography
STAMPFLI , G. M. & HOCHARD , C. Plate tectonics of the Alpine realm 89
JOHNSTON , S. T. & MAZZOLI , S. The Calabrian Orocline: buckling of a previously more linear 113
orogen
DILEK , Y. & SANDVOL , E. Seismic structure, crustal architecture and tectonic evolution of the 127
Anatolian –African Plate Boundary and the Cenozoic Orogenic Belts in the Eastern
Mediterranean Region
Palaeozoic/Neoproterozoic orogens
PUCHKOV , V. N. The evolution of the Uralian orogen 161
HOFMANN , M., LINNEMANN , U., GERDES , A., ULLRICH , B. & SCHAUER , M. Timing of dextral 197
strike-slip processes and basement exhumation in the Elbe Zone (Saxo-Thuringian Zone): the
final pulse of the Variscan Orogeny in the Bohemian Massif constrained by LA-SF-ICP-MS
U –Pb zircon data
PEREIRA , M. F., CHICHORRO , M., WILLIAMS , I. S., SILVA , J. B., FERNAŃDEZ , C., 215
DÍAZ - AZPÍROZ , M., APRAIZ , A. & CASTRO , A. Variscan intra-orogenic extensional tectonics in
the Ossa –Morena Zone (Évora –Aracena –Lora del Rı́o metamorphic belt, SW Iberian Massif):
SHRIMP zircon U – Th–Pb geochronology
NANCE , R. D., KEPPIE , J. D., MILLER , B. V., MURPHY , J. B. & DOSTAL , J. Palaeozoic 239
palaeogeography of Mexico: constraints from detrital zircon age data
VAN STAAL , C. R., WHALEN , J. B., VALVERDE -V AQUERO , P., ZAGOREVSKI , A. & 271
ROGERS , N. Pre-Carboniferous, episodic accretion-related, orogenesis along the Laurentian
margin of the northern Appalachians
MARTÍNEZ , S. S., ARENAS , R., FERNÁNDEZ -S UÁREZ , J. & JEFFRIES , T. E. From Rodinia to 317
Pangaea: ophiolites from NW Iberia as witness for a long-lived continental margin
MURPHY , J. B., GUTIÉRREZ - ALONSO , G., NANCE , R. D., FERNÁNDEZ - SUÁREZ , J., KEPPIE , 343
J. D., QUESADA , C., DOSTAL , J. & BRAID , J. A. Rheic Ocean mafic complexes: overview and
synthesis
vi CONTENTS
Proterozoic orogens
EVANS , D. A. D. The palaeomagnetically viable, long-lived and all-inclusive Rodinia 371
supercontinent reconstruction
RIVERS , T. The Grenville Province as a large hot long-duration collisional orogen – insights 405
from the spatial and thermal evolution of its orogenic fronts
OCCHIPINTI , S. A. & REDDY , S. M. Neoproterozoic reworking of the Palaeoproterozoic 445

Capricorn Orogen of Western Australia and implications for the amalgamation of Rodinia
CORRIGAN , D., PEHRSSON , S., WODICKA , N. & DE KEMP , E. The Palaeoproterozoic 457
Trans-Hudson Orogen: a prototype of modern accretionary processes
Index 481
Ancient orogens and modern analogues: an introduction
J. BRENDAN MURPHY1*, J. DUNCAN KEPPIE2 & ANDREW J. HYNES3
1
Department of Earth Sciences, St. Francis Xavier University, Antigonish,
Nova Scotia, Canada, B2G 2W5
2
Universidad Naçional Autónoma de México, México, D.F. 04510
3
Department of Earth and Planetary Sciences, McGill University,
Montreal, PQ, Canada, H3A 2A7
*Corresponding author (e-mail: bmurphy@stfx.ca)
Abstract: Plate-tectonics principles have been routinely applied to the study of Phanerozoic oro-
genic belts and, more controversially, to Precambrian orogens as far back as the Early Archaean.
Recent advances in a variety of fields have vastly improved our understanding of ancient orogenic
belts, so that realistic modern analogues can be entertained. This volume presents up-to-date synth-
eses of some classic modern and ancient orogenic belts as well as examples of some of the pro-
cesses responsible for their evolution.
Plate-tectonics principles provide a unifying that resolution of orogenic complexities has not
conceptual framework for understanding the evol- matched that of modern orogens. Detailed models
ution of modern oceanic lithosphere and Cenozoic that include the effects of, for example, delamina-
orogens (Dewey & Bird 1970). Since Tuzo Wilson tion, transform faults, subduction of oceanic and
(1966) first proposed that the evolution of the aseismic ridges, overriding of plumes and sub-
Appalachian– Caledonide orogen of eastern North duction erosion have rarely been considered in
America and western Europe records the birth and ancient orogens, although they have a demonstrably
death of an earlier ocean (now known as Iapetus), profound effect on the styles of Cenozoic orogens.
plate-tectonics principles have been routinely On the other hand, studies of deeply eroded
applied to the study of Phanerozoic orogenic belts ancient orogens provide insights into the hidden
(e.g. van Staal et al. 1998), and, more controver- roots of modern orogens, providing a complemen-
sially, to Precambrian orogens as far back as the tary perspective.
Early Archaean (Condie & Kröner 2008; Wyman
et al. 2008).
For pre-Jurassic orogenic belts, no seafloor- Modern orogens
spreading record is available. Furthermore, as we Subduction tectonics
go further back in time, faunal and palaeomagnetic
data become more sparse and difficult to apply. This volume treats examples of subduction-related
Tectonic models are therefore less well constrained tectonics from modern orogens, including the role
with increasing age. However, recent advances in of structure in determining the chemical compo-
geochronological and palaeomagnetic techniques, sition of volcanism, a description of flat-slab sub-
and conceptual advances in other fields such as duction, and the potential relationship between
geodynamics, have vastly improved the available ridge-trench collision and the formation of fold-
tools for the study of ancient orogenic belts, so that and-thrust belts.
realistic modern analogies can now be drawn with The Taupo Volcanic Zone (TVZ) of northern
some confidence. New Zealand is a classic modern example of an
This volume is an outgrowth of IGCP Project ensialic island arc that is undergoing extension.
453 (1999–2004) entitled ‘Uniformitarianism revis- The TVZ results from oblique subduction of the
ited: a comparison between modern and ancient Pacific plate beneath New Zealand’s North Island.
orogens’. The main goal of the project was to Magmatism is attributed to the rise of hot mantle
enhance our understanding of the causes and effects beneath thinned continental crust (c. 16 km, e.g.
of modern and ancient mountain belts, and how Harrison & White 2006) that generates crustal melts
these relationships have varied with time. Today, as well as melts from the decompressed mantle
most geoscientists apply these principles back to (Cole et al. 1998, 2005). In the northern and south-
the Early Proterozoic or Late Archaean. However, ern TVZ, the magmatism is predominantly andesi-
the state of preservation of these orogens is such tic, whereas in the central zone it is predominantly
From: MURPHY , J. B., KEPPIE , J. D. & HYNES , A. J. (eds) Ancient Orogens and Modern Analogues.
Geological Society, London, Special Publications, 327, 1–8.
DOI: 10.1144/SP327.1 0305-8719/09/$15.00 # The Geological Society of London 2009.
2 J. B. MURPHY ET AL.
rhyolitic. Spinks et al. (2005) show that the regions geometry, convergence rate, and the degree of coup-
with the greatest extension in the TVZ produce cal- ling between the subducting and overriding plates
deras up to 50 km wide that are dominated by felsic (e.g. McQuarrie 2002). The authors present
magmatism, whereas areas with the greatest trans- structural and stratigraphic sections across the
tension produce andesitic stratovolcanoes. The Middle Miocene Chiapas fold-and-thrust belt that
strain rate along local faults influences magma com- are constrained by seismic and well data, and they
position and eruption style (Paterson & Tobsich propose that the belt was formed as a result of ridge-
1992; Petford et al. 2000). In this volume, Cole & trench collision, thereby providing a potential
Spinks provide a comprehensive overview of the modern analogue to be considered in ancient belts.
structural and volcanological features of TVZ, and
include recent high precision radiometric data. Accretionary tectonics
Using these data, the authors draw attention to the
close correspondence between volcanism and struc- The Cordillera of North America are considered a
ture in TVZ, pointing out that vents are commonly classic example of an orogen dominated by terrane
aligned along local faults and that many calderas accretion and has resulted in the growth of North
have boundaries with a rectangular geometry that America by an average of about 500 km since the
reflect local fault patterns. They describe a model end of the Palaeozoic (e.g. Coney et al. 1980).
in which basalt rises into the lower crust causing Obducted vestiges of oceanic lithosphere are com-
partial melting of the lower crust. These crustal monly interpreted to reflect suture zones between
melts pond at the mid-crustal brittle-ductile tran- adjacent terranes. This interpretation is valid if the
sition zone, and rise into the upper crust exploiting subducted oceanic lithosphere is exhumed up the
structural weaknesses. subduction channel (e.g. Federico et al. 2007). In
The influence of subduction-zone geometry on this volume, however, Dostal et al. propose that
the tectonic style of orogenesis at convergent maran entire terrane (the Cache Creek terrane) in the
gins has long been recognized. Since Barazangi & Canadian Cordillera could represent subduction-
Isacks (1976, 1979) documented flat slab segments zone rocks that were extruded into the upper plate.
along the Andes continental margin, the causes and The Cache Creek terrane is a high-pressure
effects of flat-slab subduction have been a matter of blueschist-eclogitic terrane (Johnston & Borel
intense debate. Recent analysis of modern flat-slab 2007) that occurs between the Quesnellia and
subduction zones has drawn attention to their spatial Stikinia terranes. Middle–Upper Triassic arc mag-
and temporal correlation with subducting oceanic matic sequences in Quesnellia and Stikinia (Takla
plateaus (e.g. Gutscher et al. 2000; Ramos & Group) have very similar geochemical and isotopic
McNulty 2002). The modern Andean margin has characteristics that are typical of primitive arcs with
several flat-slab segments, up to 500 km wide, limited continental crust involvement. Extrusion of
with each correlated to subduction of anomalously the HP rocks into the upper plate indicates that
buoyant oceanic crust, represented by oceanic pla- rather than two terranes (Stikinia and Quesnellia),
teaus (Pilger 1981; Gutscher et al. 2000). In this the Takla Group represents one arc split at the site
volume, Ramos & Folguera identify and charac- of extrusion of the Cache Creek rocks. Dostal et al.
terize modern flat slab segments, incipient flat slab point out that existing models (post-Middle Jurassic
segments, and ancient flat slab segments, including duplication of the arc by strike-slip faulting, orocl-
an Early Permian example. They propose that inal or synformal folding) are inconsistent with
most of the Andes have experienced a stage of flat both palaeomagnetic and faunal data. Building on
subduction at some stage in their evolution. They the numerical models of Stöckhert & Gerya (2005)
also identify the characteristics of slab steepening, and Gerya & Stöckhert (2006), they propose an
which, beneath thick crust, results in delamination alternative model in which oblique eastward subduc-
and bimodal magmatism, but beneath thin crust tion and high-pressure metamorphism of the Cache
results in extensional tectonism and voluminous Creek accretionary prism and forearc was followed
mafic volcanism. by extrusion into the upper plate arc and exhumation
Fold-and-thrust belts are a common manifes- by the Middle Jurassic. In this scenario, the high-
tation of deformation in the overriding plate above pressure rocks occur between similar units and do
subduction zones (e.g. Price & Mountjoy 1970; not separate terranes or define an oceanic suture.
Allmendinger et al. 1997). Mandujano-Velazquez
& Keppie point out a disconnect between models Collisional tectonics
proposed to explain ancient fold-and-thrust belts
(which are commonly attributed to collisions Three contributions from the Mediterranean region
between continents or terranes), and Mesozoic – provide examples of the complexity and hetero-
Cenozoic belts which are typically interpreted to geneity of continent-continent collisional zones.
reflect changes in factors such as Benioff-zone Stampfli & Hochard provide detailed plate
ANCIENT OROGENS AND MODERN ANALOGUES 3
reconstructions (rather than traditional continental resulted in crustal thickening and plateau formation,
drift reconstructions) of the Mediterranean region and was followed by slab break-off which resulted
from late Triassic to Miocene. This time period inc- in extension. The mid-Miocene collision of Arabia
ludes the opening of the Neotethys and North with Eurasia at 13 Ma resulted in crustal shortening
Atlantic oceans, followed by the Pyrenean and and regionally extensive strike-slip fault systems
Alpine orogenies. The reconstructions describe the that accommodated the tectonic escape of the Ana-
evolution of several distinct tectonic domains and tolian plate. The authors compare this setting with
the relationships between them, and imply differen- that of the Tibetan plateau.
tial transport of thousands of kilometres for the
plates and the terranes within them. Other interesting
implications of the reconstructions are that: (i) the Ancient orogens
Neotethys and North Atlantic oceans were never Palaeozoic orogens
directly connected; (ii) a series of basins opened
and closed, due to subduction zone roll-back and In the absence of an ocean floor record, palaeocon-
back-arc extension; and (iii) the Alpine orogen is tinental reconstructions for the Palaeozoic Era are
an example of the juxtaposition of exotic tectonic not as precise as those available for the Mesozoic
elements. and Cenozoic. Nevertheless, since the landmark
The Calabrian orocline is one of the most papers of Wilson (1966) and Dewey (1969), plate-
distinctive structural elements in the Mediterranean tectonic principles have been routinely applied to
region and reflects the bending of the Apennine and Palaeozoic orogenic belts. Over the past twenty
Sicilian mountain chains. Although its status as an years, a combination of faunal, palaeomagnetic,
orocline is well accepted (e.g. Tapponnier 1977; lithostratigraphic, geochronological and petrologi-
Eldredge et al. 1985), the mechanisms responsible cal data have provided first-order constraints
for its development remain enigmatic. Oroclines, on palaeocontinental reconstructions, and although
as defined by Carey (1955, 1958), are arcuate oro- some important issues remain, a broad consensus
genic belts that were formerly more linear than they has emerged (e.g. McKerrow & Scotese 1990;
are today. They are commonly best established where Scotese 1997; Cocks & Fortey 1990; van Staal
palaeomagnetic data show a change of declination et al. 1998; Cocks & Torsvik 2002; Stampfli &
around the oroclinal bend (Eldredge et al. 1985). Borel 2002). Tectonic history of the Palaeozoic Era
The Calabrian orocline development previously has is dominated by the breakup of the supercontinent
been viewed as a thin-skinned, allochthonous struc- Pannotia (Powell 1995; Dalziel 1997; Cawood
ture (Eldredge et al. 1985), that may have involved et al. 2001) between 600–550 Ma, followed by
subduction zone roll-back (Kastens et al. 1988). the assembly and amalgamation of Pangaea by the
Johnston & Mazzoli present a geometric model Late Carboniferous –Permian. These reconstruc-
for development of the orocline that also accounts tions provide the backdrop for models for the evol-
for the coeval development of the Tyrrhenian Sea, ution of the Palaeozoic orogens presented herein.
a tract of oceanic lithosphere that began to form Puchkov reviews recent advances in the under-
about 7 to 9 million years ago. They propose that the standing of the Uralian orogen, which, he points
orocline formed by eastward buckling of an originally out, has many of the characteristics expected from
north-south continental ribbon that involved the the Wilson Cycle. Its evolution begins with the devel-
crust and some of the lithospheric mantle and was opment and destruction of the Palaeouralian ocean
mechanically facilitated by the presence of a sub- which involved collisions between Baltica, Siberia
duction zone along the length of its eastern margin. and Kazakhstania. Puchkov stresses that the Uralian
According to Dilek & Sandvol, the tectonic orogen also displays some very distinctive features
evolution of the Eastern Mediterranean in the late that preclude a simple geodynamic connection with
Mesozoic –Cenozoic is analogous to the develop- the Late Palaeozoic Variscide orogenic belts of
ment of the Himalayan –Indonesian orogen. Accord- western Europe as has been previously proposed
ing to their model, the tectonic evolution of the (e.g. International Committee of Tectonic Maps
Eastern Mediterranean was governed by collisions 1982). Late Cambrian–Early Ordovician continental
of Gondwana-derived continental terranes with rifting was followed by Middle Ordovician passive
Eurasia as well as by intra-oceanic subduction and margin development, Late Ordovician subduction,
closure of the intervening Neotethyan oceanic arc–continent collision in the Late Devonian–Early
basins. Jurassic –Cretaceous ophiolites, developed Carboniferous, and continent-continent collision
in the proto-arcs and forearcs of these subduction beginning in the mid-Carboniferous. Renewed and
zones, were emplaced during collision of the pas- relatively recent uplift is thought to be a far-field
sive margin with the trench, and partial subduction effect of the collision between India and Asia. The
of these margins yielded high-pressure metamor- Uralian orogen also has some rare and enigmatic
phic assemblages. Paleocene –Eocene collisions characteristics, such as the presence of a cold
isostatically equilibrated root, the abundance of that the evolution of the Middle American terranes
oceanic mafic complexes, the limited crustal shorten- of southern Mexico records the creation and des-
ing in the southern Urals, and the Silurian platinum- truction of a Palaeozoic ocean. Palaeozoic metasedi-
rich belt hosted by subduction-related layered mentary rocks occur in a number of these terranes,
plutons. including the Acatlán Complex of the Mixteca
The papers by Hofmann et al. and Pereira et al. terrane, which is the largest inlier of Palaeozoic
provide a detailed analysis of the late-stage tectonic rocks in Mexico and underlies an area almost the
evolution in two classical areas of the Late Palaeo- size of Belgium. Broad tectonostratigraphic simi-
zoic Variscan orogen (Bohemian Massif and larities to parts of the Appalachian-Ouachita orogen
Iberia, respectively). Global reconstructions show and close proximity to a c. 1 Ga basement (Oaxacan
that the Variscan orogeny was due to the destruction Complex) have led to models in which the Acatlán
of the Rheic Ocean, and was a key event in the for- Complex was interpreted in terms of Laurentia–
mation of Pangaea. Hofmann et al. present new geo- Gondwana (Amazonia) collision (e.g. Yañez et al.
chronological data that constrain the timing of the 1991). Ortega-Gutiérrez et al. (1999) proposed
final pulse of the Variscan orogeny in the northern that the complex represents a vestige of the Iapetus
(Saxo-Thuringian Zone) of the Bohemian Massif, suture formed during a Late Ordovician– Early
which they relate to closure of the Rheic Ocean Silurian collisional orogeny between Laurentia and
and terminal collision between Gondwana and Oaxaquia, a crustal block of Gondwanan affinity.
Laurussia. These data, together with field relation- More recently, Talavera-Mendoza et al. (2005)
ships, allow an understanding of the relationships proposed that the complex records suturing of both
between late-stage strike-slip motion, plutonism and peri-Laurentian and peri-Gondwanan arcs to Laur-
lateral extrusion, as well as the formation of post- entia during closure of both the Iapetus and Rheic
orogenic basins. oceans. In this volume, Nance et al. focus on the
The evolution of the Variscan orogen in SW provenance of the Palaeozoic metasedimentary
Iberia is widely recognized as a world-class example rocks in the Middle America terranes by compiling
of a transpressional orogen, and it is one of the few and reviewing available detrital zircon data. These
places in which an oceanic suture, reflecting the data confirm the Gondwanan affinity of the meta-
collision between Gondwana (Iberian massif) and sedimentary rocks, but indicate an origin on the
Laurentia (South Portuguese zone) is documented southern flank of the Rheic Ocean, rather than
(Quesada 1990; Quesada & Dallmeyer 1994). Iapetus. In this scenario, these rocks are interpreted
Pereira et al. focus on the post-collisional (Visean) to have been transferred to Laurentia during and
extensional phases of the origin, most specifically following the amalgamation of Pangaea.
the evolution in the footwall of several metamorphic Ever since the papers of Williams (1964, 1979),
core complexes, which occur within a 250 km long Wilson (1966) and Dewey (1969) were published,
belt of metamorphic rocks in the Ossa-Morena zone the Appalachian orogen has been a type area for
along the southwestern flank of the Iberian Massif. the application of modern plate tectonic concepts
They interpret new U –Pb SHRIMP data from zircon to Palaeozoic orogens. van Staal et al. present an
rims and monazites to reflect flushing of hydrother- overview paper that synthesizes recent interpret-
mal fluids along the detachment zones that bound ations of various parts of the northern Appalachians.
each of these core complexes. While accepting the The paper, which combines field and petrological
popular view that the syn-collisional style of defor- studies with comprehensive and precise geochro-
mation is compatible with regional sinistral trans- nology, shows that prior to the amalgamation of
pression (Quesada & Dallmeyer 1994), they also Pangaea, the development of the Appalachian oro-
point out that post-collisional sinistral shear could gen was typical of an accretionary orogen, ana-
provide the local regions of extension where the logous to the Mesozoic –Cenozoic evolution of the
core complexes developed. This extensional event western Pacific Ocean. The authors discuss four
is manifest at different crustal levels and is also episodes of orogenesis, each related to the accretion
held responsible for coeval basin formation and of a microcontinent, one peri-Laurentian (Dash-
magmatism. Taken together, the data presented by woods) during the Ordovician Taconic orogeny,
Pereira et al. provide important evidence for post- and three peri-Gondwanan terranes (Ganderia, Ava-
collisional intra-orogenic shear along the Gondwa- lonia and Meguma) which docked at various times
nan margin in the terminal phases of closure of the to the eastern margin of Laurentia between the
Rheic Ocean. Late Ordovician and Devonian. Sequential accre-
In recent years, there has been an increased tionary tectonics led to pulses of deformation and
understanding of the relationship between Palaeo- metamorphism.
zoic orogenesis in central America and coeval events Despite their obvious importance to the assem-
in the Appalachian– Caledonide orogen. The pio- bly of Pangaea, studies of the mafic complexes,
neering work of Ortega-Gutiérrez (1975) showed interpreted by most authors as ophiolites, formed
during the lifespan of the Rheic Ocean, have lagged crust that originated along the Laurussian margin,
behind the comprehensive studies of ophiolites but were thrust over Gondwana during Variscan
within the Iapetan realm. In this volume, there are orogenesis. The Carboniferous ophiolites may
two papers that provide such syntheses. Sanchez- have formed in a strike-slip regime within relict
Martinez et al. provide a detailed analysis of the ocean basins during closure of the Rheic Ocean.
range of ophiolites that occur in allochthonous com-
plexes in NW Iberia (Galicia). Murphy et al. provide Proterozoic orogens
an overview of mafic complexes in the Rheic Ocean
realm, from Mexico to Bohemia. Over the past 20 years, the existence of a super-
The allochthonous complexes of NW Iberia continent, Rodinia, and the corresponding global
within the Variscan suture zone provide one of ocean, Mirovoi (McMenamin & McMenamin
the best-preserved expressions of the collision 1990), between c. 1.1–0.75 Ga has gained accep-
between Laurussia and Gondwana (Arenas et al. tance (see Li et al. 2008 for a review). However,
1986, 2007). Sanchez-Martinez et al. show that the configuration of this supercontinent is highly
the ophiolites have a wide range in ages, tectonic controversial. Hoffman (1991) attributed global-
setting and subsequent tectonothermal evolution, scale orogenesis between 1.3 –1.0 Ga to the amalga-
precluding the simple interpretation that they mation of Rodinia, and the distribution of Late
reflect only closure of the Rheic Ocean during Pan- Neoproterozoic passive margins around Laurentia
gaean assembly. This study provides an to its c. 0.75 Ga breakup. Since that time, several
important example of the complexity of collisional authors have modified Hoffman’s original recon-
suture zones. The oldest ophiolite is Mesoprotero- struction, but its basic premise remains intact. In
zoic in age, has supra-subduction zone affinities, contrast to most reconstructions that are based on
and is thought to have been generated near the geological data, Evans produces a reconstruction
West African craton before the c. 1.1 Ga amalgama- of Rodinia that is built primarily upon the available
tion of the supercontinent Rodinia. An Early Cam- palaeomagnetic database, although geological rela-
brian ophiolite was probably generated in an arc tions are included as supporting evidence. The
setting, but a Late Cambrian ophiolite has MORB Evans reconstruction is a radical departure from
affinities and may have formed in the Iapetus- previous versions of Rodinia and is sure to generate
Tornquist ocean. Other Late Cambrian-Early Ordo- much discussion. In this reconstruction, Rodinia
vician ophiolites bodies have arc affinities, and was a long-lived supercontinent that contained all
together with coeval metamorphism, are thought major Precambrian continents (see e.g. Pisarevsky
to reflect the opening of the Rheic Ocean as a et al. 2003). Perhaps the most controversial features
back-arc basin. A Devonian ophiolite was formed are: (i) the lack of a collider for the Grenville orogen
during contraction of the Rheic Ocean, shortly pre- of eastern North America; (ii) the inverted position
ceding collision between Laurussia and Gondwana. of Australia and Antarctica relative to Baltica; and
Murphy et al. point out that mafic complexes (iii) the positioning of Amazonia, West Africa,
widely interpreted to represent vestiges of the and Rio de la Plata cratons close to western (rather
Rheic Ocean are widespread, from the Acatlán than eastern) Laurentia.
Complex in Mexico to the Bohemian Massif in The Grenville orogen is widely held to be
eastern Europe. Most of these complexes are either a classic Proterozoic example of a continent-
Late Cambrian-Early Ordovician or Late Palaeozoic continent collision (e.g. Ludden & Hynes 2000
in age. With the exception of those in NW Iberia, and references therein; Li et al. 2008). Rivers pro-
these Late Cambrian-Early Ordovician mafic com- vides a thorough summary of the thermal evolution
plexes are not ophiolites. Their geochemical and of the Grenville orogen exposed in Canada, which,
Sm-Nd isotopic signatures indicate that they are on the basis of global reconstructions, he interprets
rift-related continental tholeiites, derived from to have developed in late Mesoproterozoic to early
an enriched c. 1.0 Ga subcontinental lithospheric Neoproterozoic time by the collision between
mantle, and are associated with crustally-derived Laurentia and Amazonia during the assembly of
felsic volcanic rocks. They are interpreted to reflect Rodinia (Li et al. 2008; compare Evans, this volume).
magmatism along the Gondwanan margin during He points out that its characteristics such as width
the formation of the Rheic Ocean, and to have (.600 km), abundance of high grade metamorphic
remained along that margin as Avalonia and other rocks, and duration (100 My) is typical of large
peri-Gondwanan terranes drifted northward. The hot long-duration orogens as defined by Beaumont
Late Palaeozoic mafic complexes (Devonian and et al. (2006). Such orogens develop in orthogonal
Carboniferous), however, do preserve many of the collisional settings, and in ideal cases, the amount
characteristics of ophiolites. They are characterized of new crust added (by thrusting) is balanced by
by derivation from an anomalous ultra-depleted crust removed by erosion at orogenic fronts where
mantle, and may reflect narrow tracts of oceanic a mid-crustal channel of hot, low-viscosity rocks
becomes exposed. Rivers adopts this ‘channel-flow’ This special publication is an outcome of IGCP Project 453
model for the Grenville and also identifies features of the International Geoscience Programme of the Inter-
consistent with imbalances between crust added national Union of Geological Sciences and UNESCO.
and removed. We acknowledge the Natural Sciences and Engineering
Research Council, Canada (JBM, AH), the St. Francis
The paper by Occhipinti & Reddy provides a Xavier University Council for Research (JBM) and
link between the Palaeoproterozoic Capricorn Papiit and CONACyT grants (JDK). We thank all those
Orogen of Western Australia and its reactivation who organized meetings, field trips and special volumes
in the Neoproterozoic, as documented by precise in connection with this project. We also thank to Margarete
40
Ar/39Ar mica ages. The Capricorn Orogen com- Patzak (UNESCO) and Angharad Hills (Geological
prises rocks deformed and metamorphosed during Society of London) for their help and support. We are
several Palaeoproterozoic orogenic events rang- especially grateful to the following reviewers who gener-
ing from 2.0–1.6 Ga (e.g. Occhipinti et al. 2004). ously gave of their time and expertise: Valerio Acocella,
Together with the Archaean Pilbara and Yilgarn cra- Kevin Ansdell, Andrea Arnani, Sandra Barr, Pete Betts,
Myron (Pat) Bickford, Gilles Borel, Peter Cawood, David
tonic blocks, the Palaeoproterozoic Capricorn Corrigan, Brian Cousens, Allen Dennis, Richard Ernst,
Orogen form the West Australian Craton (Myers Mary Ford, John Gamble, Laurent Godin, Gabi Gutierrez-
1993). This craton was reworked during the Alonso, Stephen Johnston, Fred McDowell, Brendan
c. 1.0 Ga oblique collision of Western Australia McNulty, Damian Nance, Francisco Pereira, Sergei
with another continent (Kalahari or Greater India). Pisarevsky, Russell Pysklywec, Cecilio Quesada, Victor
The authors point out that although the evolution Ramos, Paul Robinson, Paul Ryan, Rob Strachan, Ben
of such reworked cratonic margins can be van der Pluijm, Rob Van Der Voo and several anonymous
complex, the timing of events affecting those reviewers.
margins may be recorded by low-temperature
events in the cratonic interior. These low- References
temperature ‘events’ are difficult to detect by struc-
tural or textural analysis, but precise 40Ar/39Ar A LLMENDINGER , R. W., J ORDAN , T. E., K AY , S. M. &
mineral ages can be utilized to document the I SACKS , B. L. 1997. The evolution of the Altiplano–
Puna Plateau of the Central Andes. Annual Review of
timing of reactivation and fluid flow events. The Earth and Planetary Science, 25, 139– 174.
authors further propose that the reactivation is a far- A NBAR , A. D. & K NOLL , A. H. 2002. Proterozoic ocean
field effect of the continental collisions and, more chemistry and evolution: A bioinorganic bridge?
generally, that the results of this study demonstrate Science, 297, 1137–1142.
that the evolution of orogens can be constrained A NSDELL , K. M. 2005. Tectonic evolution of the Manitoba-
by thermochronological studies in adjacent cratons. Saskatchewan segment of the Paleoproterozoic Trans-
The Capricorn orogen is one of several Paleo- Hudson Orogen, Canada. Canadian Journal of Earth
proterozoic collisional belts that formed during the Sciences, 42, 741– 759.
2.0–1.8 Ga amalgamation of supercontinent Nuna A RENAS , R., G IL I BARGUCHI , J. I. ET AL . 1986. Tecto-
noestratigraphic units in the complexes with mafic
(also known as Columbia, Bleeker 2003; Rogers & and related rocks of the NW of the Iberian Massif.
Santosh 2002). Corrigan et al. provide an up-to- Hercynica, II, 87–110.
date synthesis of the accretionary and collisional A RENAS , R., M ARTÍNEZ C ATALÁN , J. R. ET AL . 2007.
evolution of the largest of these belts, the Trans- Paleozoic ophiolites in the Variscan suture of Galicia
Hudson Orogen of the Canadian Shield. As noted (northwest Spain): distribution, characteristics and
by the authors, these orogenic events occurred meaning. In: H ATCHER , R. D., J R ., C ARLSON ,
during a fundamental interval in the evolution of M. P., M C B RIDE , J. H. & M ARTÍNEZ C ATALÁN ,
the Earth’s systems, with important changes in the J. R. (eds) 4-D Framework of Continental Crust.
ocean and atmosphere (e.g. Anbar & Knoll 2002), Geological Society of America Memoirs, 200,
425–444.
biosphere (e.g. Konhauser et al. 2002) as well as B ARAZANGI , M. & I SACKS , B. 1976. Spatial distribution
the growth of cratons into large stable continents of earthquakes and subduction of the Nazca plate
(e.g. Hoffman 1988). The Trans-Hudson orogen beneath South America. Geology, 4, 686– 692.
orogen is widely recognized as the oldest example B ARAZANGI , M. & I SACKS , B. 1979. Subduction of the
of a relatively complete Wilson-Cycle (e.g. Ansdell Nazca plate beneath Peru evidence from spatial distri-
2005; Corrigan et al. 2005; St-Onge et al. 2006), bution of earthquakes. Geophysical Journal of the
from c. 2.45– 1.92 Ga passive margin assemblages Royal Astronomic Society, 57, 537 –555.
deposited along Archean cratonic margins, to for- B EAUMONT , C., N GUYEN , M. H., J AMIESON , R. A. &
mation and accretion of c. 2.00– 1.88 Ga oceanic E LLIS , S. 2006. Crustal flow modes in large hot
orogens. In: L AW , R. D., S EARLE , M. P. & G ODIN , L.
and 1.88 –1.83 Ga continental arc systems, followed (eds) Channel Flow, Ductile Extrusion and Exhumation
by 1.83– 1.80 terminal collision. The paper by Cor- in Continental Collision Zones. Geological Society,
rigan et al. is an example of the complexity and London, Special Publications, 268, 91–145.
scale of Palaeoproterozoic orogenic processes, and B LEEKER , W. 2003. The late Archean record: a puzzle in
the type of geodynamic operative during this era. ca. 35 pieces. Lithos, 71, 99–134.
C AREY , S. W. 1955. The orocline concept in geotectonics. subduction: seismicity and tomographic constraints
Proceedings of the Royal Society of Tasmania, 89, from the Andean margin. Tectonics, 19, 814– 833.
255–258. H ARRISON , A. & W HITE , R. S. 2006. Lithospheric struc-
C AREY , S. W. 1958. The tectonic approach to continental ture of an active backarc basin: The Taupo Volcanic
drift. In: C AREY , S. W. (ed.) Continental Drift – A Zone, New Zealand. Geophysical Journal Inter-
Symposium. University Tasmania Press, Hobart, national, 167, 968–990.
Tasmania, 178–355. H OFFMAN , P. F. 1988. United plates of America, the birth
C AWOOD , P. A., M C C AUSLAND , P. J. A. & D UNNING , of a craton: Early Palaeozoic assembly and growth of
G. R. 2001. Opening Iapetus: constraints from the Laurentia. Annual Review of Earth and Planetary
Laurentian margin of Newfoundland. Geological Science, 16, 543– 603.
Society of America Bulletin, 113, 443– 453. H OFFMAN , P. F. 1991. Did the breakout of Laurentia turn
C OCKS , L. R. M. & F ORTEY , R. A. 1990. Biogeography of Gondwanaland inside-out? Science, 252, 1409–1412.
Ordovician and Silurian faunas. In: M C K ERROW , I NTERNATIONAL C OMMITTEE OF T ECTONIC M APS .
W. S. & S COTESE , C. R. (eds) Paleozoic Paleogeogra- 1982. Tectonics of Europe and Adjacent Areas. Varis-
phy and Biogeography. Geological Society, London, cides, epi-Paleozoic Platforms and Alpides. Explana-
Memoirs, 12, 97–104. tory note to the International tectonic map of Europe
C OCKS , L. R. M. & T ORSVIK , T. H. 2002. Earth geogra- and adjacent areas, scale 1:2 500 000. ‘Nauka’ Publish-
phy from 500 to 400 million years ago: a faunal and ing House.
palaeomagnetic review. Journal of the Geological J OHNSTON , S. T. & B OREL , G. D. 2007. The odyssey of
Society, London, 159, 631–644. the Cache Creek terrane, Canadian Cordillera: impli-
C OLE , J. W., B ROWN , S. J. A., B URT , R. M., B ERESFORD , cations for accretionary orogens, tectonic setting of
S. W. & W ILSON , C. J. N. 1998. Lithic types in Panthalassa, the Pacific superswell, and break-up of
ignimbrites as a guide to the evolution of a caldera Pangea. Earth and Planetary Science Letters, 253,
complex, Taupo volcanic centre, New Zealand. 415– 428.
Journal of Volcanology and Geothermal Research, K ASTENS , K., M ASCLE , J. ET AL . 1988. ODP Leg 107 in
80, 217–237. the Tyrrhenian Sea: insights into passive margin and
C OLE , J. W., M ILNER , D. M. & S PINKS , K. D. 2005. back-arc basin evolution. Geological Society of
Calderas and caldera structures: a review. Earth America Bulletin, 100, 1140– 1156.
Science Reviews, 69, 1 –26. K ONHAUSER , K. O., H AMADE , T., R AISWELL , R.,
C ONDIE , K. C. & K RÖNER , A. 2008. When did plate tec- M ORRIS , R. C., F ERRIS , F. G., S OUTHAM , G. &
tonics begin? Evidence from the geologic record. In: C ANFIELD , D. E. 2002. Could bacteria have formed
C ONDIE , K. C. & P EASE , V. (eds) When Did Plate the Precambrian banded iron formations? Geology,
Tectonics Begin on Planet Earth? Geological Society 30, 1079–1082.
of America, Special Papers, 440, 281–294. L I , Z.-X., B OGDANOVA , S. V. ET AL . 2008. Assembly,
C ONEY , P. J., J ONES , D. L. & M ONGER , J. W. H. 1980. configuration, and break-up history of Rodinia: a syn-
Cordilleran suspect terranes. Nature, 288, 329–333. thesis. Precambrian Research, 160, 179–210.
C ORRIGAN , D., H AJNAL , Z., N ÉMETH , B. & L UCAS , S. B. L UDDEN , J. & H YNES , A. J. (eds) 2000. The Abitibi–
2005. Tectonic framework of a Paleoproterozoic Grenville transect. Canadian Journal of Earth
arc-continent to continent-continent collisional zone, Sciences, 37, 115–516.
Trans-Hudson Orogen, from geological and seismic M C M ENAMIN , M. A. S. & M C M ENAMIN , D. L. S. 1990.
reflection studies. Canadian Journal of Earth Sciences, The Emergence of Animals: The Cambrian Break-
42, 421–434. through. Columbia University Press, New York.
D ALZIEL , I. W. D. 1997. Overview: Neoproterozoic- M C K ERROW , W. S. & S COTESE , C. R. (eds) 1990.
Paleozoic geography and tectonics: review, hypotheses Palaeozoic Palaeogeography and Biogeography.
and environmental speculations. Geological Society of Geological Society, London, Memoirs, 12, 1– 435.
America Bulletin, 109, 16– 42. M C Q UARRIE , N. 2002. The kinematic history of the
D EWEY , J. F. 1969. Evolution of the Appalachian- central Andean fold-thrust belt, Bolivia: implications
Caledonian orogen. Nature, 222, 124–129. for building a high plateau. Geological Society of
D EWEY , J. F. & B IRD , J. M. 1970. Mountain belts and the America Bulletin, 114, 950– 963.
new global tectonics. Journal of Geophysical Research, M YERS , J. S. 1993. Precambrian history of the West
75, 2625– 2647. Australian craton and adjacent orogens. Annual
E LDREDGE , S., B ACHTADSE , V. & VAN DER V OO , R. Review of Earth and Planetary Sciences, 21, 453–485.
1985. Paleomagnetism and the orocline hypothesis. O CCHIPINTI , S. A., S HEPPARD , S., P ASSCHIER , C.,
Tectonophysics, 119, 153–179. T YLER , I. M. & N ELSON , D. R. 2004. Palaeoprotero-
F EDERICO , L., C RISPINI , L., S CAMBELLURI , M. & zoic crustal accretion and collision in the southern
C APONNI , G. 2007. Ophiolite mélange zone records Capricorn Orogen: the Glenburgh Orogeny. Precam-
exhumation in a fossil subduction channel. Geology, brian Research, 128, 237– 255.
35, 499–502. O RTEGA -G UTIÉRREZ , F. 1975. The pre-Mesozoic Geology
G ERYA , T. & S TÖCKHERT , B. 2006. Two-dimensional of the Acatlán Area, South Mexico. PhD thesis, Univer-
numerical modeling of tectonic and metamorphic his- sity of Leeds, UK.
tories of active continental margins. International O RTEGA -G UTIÉRREZ , F., E LÍAS -H ERRERA , M., R EYES -
Journal of Earth Sciences, 95, 250 –274. S ALAS , M., M ACIAS -R OMO , C. & L ÓPEZ , R. 1999.
G UTSCHER , M. A., S PAKMAN , W., B IJWAARD , H. & Late Ordovician-Early Silurian continental collision
E NGDAHL , E. R. 2000. Geodynamic of flat orogeny in southern Mexico and its bearing on
Gondwana-Laurentia connections. Geology, 27, S TAMPFLI , G. M. & B OREL , G. D. 2002. A plate tectonic
719– 722. model for the Paleozoic and Mesozoic constrained by
P ATERSON , S. R. & T OBSICH , O. T. 1992. Rates of pro- dynamic plate boundaries and restored synthetic
cesses in magmatic arcs: implications for the timing oceanic isochrones. Earth and Planetary Science
and nature of pluton emplacement and wall rock Letters, 196, 17– 33.
deformation. Journal of Structural Geology, 14, S TÖCKHERT , B. & G ERYA , T. V. 2005. Pre-collisional
291– 300. high pressure metamorphism and nappe tectonics at
P ETFORD , N., C RUDEN , A. R., M C C AFFREY , K. J. W. & active continental margins: a numerical simulation.
V IGNERESSE , J. L. 2000. Granite magma formation, Terra Nova, 17, 102– 110.
transport and emplacement in the Earth’s crust. S T -O NGE , M. R., S EARLE , M. P. & W ODICKA , N. 2006.
Nature, 408, 669–673. Trans-Hudson Orogen of North America and Hima-
P ILGER , R. H. 1981. Plate reconstruction, aseismic ridges, laya-Karakoram-Tibetan Orogen of Asia: structural
and low angle subduction beneath the Andes. Geologi- and thermal characteristics of the lower and upper
cal Society of America Bulletin, 92, 448– 456. plates. Tectonics, 25, TC4006, doi: 10.1029/
P ISAREVSKY , S. A., W INGATE , M. T. D., P OWELL , 2005TC001907.
C. MC A., J OHNSON , S. & E VANS , D. A. D. 2003. T ALAVERA -M ENDOZA , O., R UIZ , J., G EHRELS , G.,
Models of Rodinia assembly and fragmentation. In: M EZA -F IGUEROA , D., V EGA -G RANILLO , R. &
Y OSHIDA , M., W INDLEY , B. & D ASGUPTA , S. (eds) C AMPA -U RANGA , M. F. 2005. U–Pb geochronology
Proterozoic East Gondwana: Supercontinent Assembly of the Acatlán Complex, and implications for the
and Breakup. Geological Society, London, Special Paleozoic paleogeographic and the tectonic evolution
Publications, 206, 35–55. of southern Mexico. Earth and Planetary Science
P OWELL , C. MC A. 1995. Are Neoproterozoic glacial Letters, 235, 682– 699.
deposits preserved on the margins of Laurentia T APPONNIER , P. 1977. Evolution tectonique du système
related to the fragmentation of two supercontinents? alpin en Méditerranée: poinçonnement et écrasement
Comment. Geology, 23, 1053–1054. rigide-plastique. Bulletin de la Société Géologique de
P RICE , R. A. & M OUNTJOY , E. W. 1970. Geologic France, 7, 437–460.
structure of the Canadian Rocky Mountains between VAN S TAAL , C. R., D EWEY , J. F., M AC N IOCAILL , C. &
Bow and Athabasca rivers-a progress report. Geo- M C K ERROW , W. S. 1998. The Cambrian-Silurian
logical Association of Canada Special Paper, 6, tectonic evolution of the Northern Appalachians
1 –166. and British Caledonides: history of a complex, west
Q UESADA , C. 1990. Precambrian terranes in the and southwest Pacific-type segment of Iapetus. In:
Iberian Variscan foldbelt. In: S TRACHAN , R. A. & B LUDELL , D. J. & S COTT , A. C. (eds) Lyell: The
T AYLOR , G. K. (eds) Avalonian and Cadomian Past is the Key to the Present. Geological Society,
Geology of the North Atlantic. Blackie and Son, London, Special Publications, 143, 199– 242.
Glasgow, 109–133. WILLIAMS, H. 1964. The Appalachians in Newfoundland –
Q UESADA , C. & D ALLMEYER , R. D. 1994. Tectonother- a two sided symmetrical system. American Journal of
mal evolution of the Badajoz–Córdoba shear zone Science, 262, 1137–1158.
(SW Iberia): characteristics and 40Ar/39Ar mineral W ILLIAMS , H. 1979. Appalachian Orogen in Canada.
age constraints. Tectonophysics, 231, 195– 213. Canadian Journal of Earth Sciences, 16, 792–807.
R AMOS , V. A. & M C N ULTY , B. (eds) 2002. Flat-slab sub- W ILSON , J. T. 1966. Did the Atlantic Ocean close and then
duction in the Andes. Journal of South American Earth re-open? Nature, 211, 676–681.
Sciences, 15, 1– 155. W YMAN , D. A., O’N EILL , C. & A YER , J. A. 2008. Evi-
R OGERS , J. J. W. & S ANTOSH , M. 2002. Configuration of dence for modern-style subduction to 3.1 Ga: a plateau-
Columbia, a Mesoproterozoic supercontinent. Gond- adakite-gold (diamond) association. In: C ONDIE ,
wana Research, 5, 5– 22. K. C. & P EASE , V. (eds) When Did Plate Tectonics
S COTESE , C. R. 1997. Continental Drift. 7th ed., PALEO- Begin on Planet Earth? Geological Society of
MAP Project, Arlington, Texas. America Special Paper, 440, 129 –148.
S PINKS , K. D., A COCELLA , V., C OLE , J. W. & B ASSETT , Y AÑEZ , P., R UIZ , J., P ATCHETT , P. J., O RTEGA -
K. N. 2005. Structural control of volcanism and caldera G UTIÉRREZ , F. & G EHRELS , G. 1991. Isotopic studies
development in the transtensional Taupo Volcanic of the Acatlán Complex, southern Mexico: implications
Zone, New Zealand. Journal of Volcanology and for Paleozoic North American tectonics. Geological
Geothermal Research, 144, 7 –22. Society of America Bulletin, 103, 817–828.
Caldera volcanism and rift structure in the Taupo Volcanic Zone,
New Zealand
J. W. COLE1* & K. D. SPINKS2
1
Department of Geological Sciences, University of Canterbury, Private Bag 4800,
Christchurch 8140, New Zealand
2
Mighty River Power Ltd, PO Box 445, Hamilton, New Zealand
*Corresponding author (e-mail: jim.cole@canterbury.ac.nz)
Abstract: The Taupo Volcanic Zone (TVZ) is an active continental volcanic arc/back-arc basin
in central North Island, New Zealand. It is the youngest area of volcanic activity that extends south-
wards from the Coromandel Volcanic Zone (CVZ), where andesitic volcanism began c. 18 Ma and
rhyolitic volcanism c. 10 Ma. It is an extensional basin (average c. 8 mm a21) with numerous, pre-
dominantly normal (dip .608) faults within the Taupo Rift, but with some strike-slip component.
TVZ can be divided into three parts. In the north (Whakatane Graben – Bay of Plenty) and south
(Tongariro volcanic centre) volcanism is predominantly andesitic, while in the central part it is
predominantly rhyolitic. This central area comprises eight caldera centres; the oldest of which
(Mangakino caldera; 1.62 –0.91 Ma) may be transitional between CVZ and TVZ. Kapenga
caldera (c. 700 ka) is completely buried by younger volcanics, but is probably a composite structure
with most recent subsidence related to volcano-tectonic processes. Of the remaining five caldera
centres, Rotorua, Ohakuri and Reporoa are all simple, sub-circular structures which collapsed
c. 240 ka, and are each associated with one ignimbrirte outflow sheet (Mamaku, Ohakuri and
Kaingaroa, respectively). Okataina and Taupo are caldera complexes with multiple ignimbrite
eruptions and phases of collapse. The three simple calderas are extra-rift, occurring outside the
main fault zone in the centre of the Taupo Rift system, while the two caldera complexes are
both intra-rift. There is a close relationship between volcanism and structure in TVZ, and many
of the structural caldera boundaries have rectangular geometry reflecting the fault pattern. Intrusion
of high-alumina basalts as dykes, parallel to the fault trend, may have had a strong influence in
causing rhyolitic eruptions in central TVZ.
The Taupo Volcanic Zone (TVZ) is the currently (Adams et al. 1994). Volcanism in CVZ was
active zone of calc-alkaline volcanism and intra-arc largely NNW-oriented and becomes progressively
rifting (Fig. 1), resulting from oblique subduction younger southwards (Fig. 3; Skinner 1986; Briggs
(c. 42 mm a21; De Mets et al. 1990) of the Pacific et al. 2005), culminating with eruptions in the
plate beneath the North Island, New Zealand Tauranga and Kaimai centres (Fig. 3) between 2.9
(Fig. 2), and is the most active rhyolitic system on and 1.9 Ma (Briggs et al. 2005). The earliest rhyoli-
Earth (Houghton et al. 1995). TVZ is a zone of tic centre considered part of TVZ is Mangakino
active rifting within continental lithosphere that (Fig. 3), with ages ranging from 1.62 –0.91 Ma, but
is extending at an average rate of c. 8 mm a21 this centre is on the western side of TVZ, and may
(Darby & Meertens 1995; Darby et al. 2000) in a itself be transitional between CVZ and TVZ.
NW–SE direction. Many geophysical papers (e.g. Stratford & Stern
Wilson et al. (1995) consider the earliest volcanic 2006) refer to the Central Volcanic Region (CVR).
activity was andesitic and began c. 2 Ma ago, with This is a wedge-shaped area comprising low-
rhyolitic volcanism commencing c. 1.6 Ma from density, low-velocity volcanics that occupy the
the central part of TVZ, accompanied by rifting eastern side of Coromandel, and TVZ (Fig. 3), as
from c. 1.1 Ma. The exact timing of initiation of defined by gravity and seismicity (Cole et al.
TVZ remains uncertain because much of the early 1995), and hence corresponds to the period of rhyo-
geological history is concealed by eruptives from litic volcanism in the last 10 Ma. Studies of volcanic
the largest (Whakamaru) eruption at c. 0.34 Ma. rocks in CVZ and TVZ suggest there is no obvious
Also, the relationship to the Coromandel Volcanic time break between volcanism in the two zones
Zone (CVZ) in both space and time is unclear (Fig. 3). If there is a complete transition, it may be
(Briggs et al. 2005). Andesitic volcanism in CVZ better to delineate the TVZ purely on visible struc-
began c. 18 Ma and rhyolitic volcanism c. 10 Ma ture, and to restrict it to volcanism between clearly
10 J. W. COLE & K. D. SPINKS
Fig. 1. Location and tectonic setting of the Taupo Volcanic Zone (TVZ) within North Island, New Zealand. Velocity of
the Pacific plate relative to the Australian plate after De Mets et al. (1990). AD, andesite-dominated northern and
southern part of TVZ; RD, rhyolite-dominated central part of TVZ; TF, Taranaki Fault, a long-lived structure last
active in the Miocene as a thrust fault (e.g. Knox 1982); NIDFB, North Island Dextral Fault Belt, a zone of dominantly
strike-slip faulting related to the rotation of the eastern North Island; HSDF, Hikurangi Subduction Deformation
Front, which marks the edge of subduction of the Pacific Plate. After Reyners et al. (2006).
defined NNE-trending faults, as indicated in gravity (e.g. Rogan 1982; Stern 1986; Bibby et al.
Figure 3. This would equate with the ‘Modern 1995), electrical modelling (e.g. Bibby et al. 2000),
TVZ’ of Wilson et al. (1995). earthquake tomography (e.g. Reyners et al. 2006;
This review considers current views on the 3D Sherburn et al. 2003, 2006), explosion seismology
structure of TVZ and discusses the location and (e.g. Harrison & White 2004; Stratford & Stern
nature of caldera volcanism in central TVZ. 2004, 2006) and magnetotelluric studies (Ogawa
et al. 1999).
Subsurface structure of TVZ The primary reason for TVZ existence is the
Wadati –Benioff zone, which descends westwards
Our understanding of what is beneath TVZ has from the Hikurangi Trough (Fig. 2) to a depth of at
largely come from geophysical studies, including least 300 km, and lies 80– 100 km beneath TVZ
CALDERA VOLCANISM AND RIFT STRUCTURE IN THE TAUPO VOLCANIC ZONE 11
Fig. 2. Schematic cross-section of the central North Island, showing crustal and upper mantle structure as determined
by seismology. From Cole (1984) with addition of new data from Stratford & Stern (2004).
(Reyners et al. 2006). This subduction zone has now Moho (see also Price et al. 2005). Instead there
been well defined by a dense seismograph deploy- is a distinct layer between c. 16 and 20 km depth,
ment (code named CNIPSE, for Central North which they interpret as a zone of mafic under-
Island Passive Seismic Experiment) which was plating, at the top of which is a very strong reflector.
carried out in 2001 (Reyners et al. 2006) concur- Rocks above 16 km are interpreted to be a mix of
rently with a number of other active and passive intruded igneous rock, volcanics and greywacke.
seismic experiments in the same region (Henrys Structure within the top 10 km is largely deter-
et al. 2003). It has been imaged as a high Vp mined from gravity data, earthquake tomography
(.8.5 km s21), low Vp/Vs feature, the latter prob- and electrical modelling. Gravity data indicates a
ably reflecting dehydration of the downgoing slab broad area of low gravity (c. 2350 mN/kg; Bibby
(Reyners et al. 2006). In contrast, the mantle et al. 1995; Stagpoole & Bibby 1999), with smaller
wedge is imaged as a low Vp, high Vp/Vs feature, areas of even lower gravity generally corresponding
with the lowest Vp (7.4 km s21) and highest Vp/ to the centres of rhyolitic volcanism (Rogan 1982).
Vs ratio (1.87) at about 65 km depth, possibly repre- These areas are also characterized by low-Vp
senting rising plumes of magma or hot upwelling, indicating low-density, volcaniclastic materials
decompressing mantle (Reyners et al. 2006). that have filled areas of caldera collapse (Sherburn
Material with Vp . 8.0 km s21 directly above the et al. 2003). There is a seismogenic zone at
dipping plane is interpreted as more viscous material c. 6.5 km that may represent the maximum depth
sinking into the wedge, which is largely insulating that brittle fracture may occur (Bibby et al. 1995;
the slab from the high-temperature part of the wedge. Bryan et al. 1999), and below this is probably
Explosion seismology also suggests a significant quartzo-feldspathic continental crust.
mantle reflector at c. 35 km beneath TVZ (Fig. 2) Additional information about the upper 2–5 km
within the mantle wedge (Stratford & Stern 2004), has come from extensive geothermal drilling (e.g.
but Harrison & White (2004) believe this simply Wood et al. 2001) and from studies of lithic
represents the base of the under-plated basaltic blocks incorporated in ignimbrites erupted during
crust. This may reflect a pool of magma within caldera-forming events (e.g. Brown et al. 1998a;
the mantle wedge. These authors also suggest the Cole et al. 1998).
upper part of the mantle is heterogeneous, with No indication has yet been found of molten or
c. 1–2% partial melt, and that the transition to the semi-molten magma bodies near the surface,
lower crust is broad without a sharp reflection although some plutons (e.g. Ngatamariki) are
Fig. 3. Map of northern and central North Island showing the Coromandel and Taupo volcanic zones, and the location
of the main rhyolitic calderas since 10 Ma. Modified from Briggs et al. (2005), with permission of New Zealand
Journal of Geology and Geophysics.
characterized by a high Vp anomaly (Sherburn et al. North Island Shear Belt (NISB) by Cole (1990)
2003). A rapid increase in conductivity occurs at and the North Island Dextral Fault Belt (NIDFB)
about 10 km depth, about 2 km below the base of by Cashman et al. (1992), Beanland (1995) and
the seismogenic zone, and well above the base of Beanland et al. (1998). These faults are probably
the quartzo-feldspathic crust, which suggests the closely related to rotation on the eastern side of
presence of a small fraction (,4%) of connected the North Island, but most strike-slip displacement
melt at this level (Heise et al. 2007). has been in the last 1–2 Ma, and the faults have
not contributed to major displacements within the
North Island (Nicol et al. 2006).
TVZ structure The inter-relationship between this belt and TVZ
is unclear. Some of the faults of the NIDFB termi-
Surface faulting in TVZ has largely been identified nate on the eastern side of TVZ, but there is no
from aerial photograph interpretation (e.g. Healy direct evidence that they continue beneath it.
et al. 1964; Nairn 1989, 2002), satellite imagery
(e.g. Oliver 1978; Cochrane & Tianfeng 1983;
Spinks 2005) supplemented by field observation Relationship between faulting
and interpretation (e.g. Rowland & Sibson 2001; and volcanism
Acocella et al. 2003; Spinks et al. 2005; Villamor
& Berryman 2006a, b; Nicol et al. 2006), infor- TVZ can be divided into three broad areas (Fig. 1):
mation from subsurface drilling (e.g. Grindley & (a) A wide (,50 km) central segment (from
Browne 1976; Wood et al. 2001) and seismic data Okataina to Taupo inclusive) dominated by rhyolitic
such as that from the ML 6.3 Edgecumbe earthquake volcanism including several calderas and caldera
of 2 March 1987 (Anderson & Webb 1989; Nairn & complexes; (b) lateral segments (c. 20 km wide) in
Beanland 1989). the north (Whakatane graben – Bay of Plenty); and
Faults (Fig. 4a) range from major topographic (c) south (Tongariro Volcanic Centre) dominated by
scarps with substantial displacements, such as the andesitic volcanism.
Paeroa Fault, to lineaments inferred to be the
morphologic expression of faults, but because of Whakatane graben – Bay of Plenty
the very high production rate of rhyolitic volcanism
(0.28 m3s21; Wilson et al. 1995; Wilson 1996), only This area is dominated by the andesite – dacite
the most recent are observed. The structure of volcanoes of Edgecumbe, the partially buried
TVZ is thus defined by a system of NNE-trending Manawahe, Motuhora (Whale Island) and White
faults that extend from Ruapehu in the south to the Island, although volcanic activity continues
Bay of Plenty coast with a 15 –20 km wide axial north for at least 150 km to the Whakatane sea-
rift zone, referred to as the Ruamoko Rift Zone by mount (Wright 1992; Gamble et al. 1993). Faults
Rowland & Sibson 2001 or more commonly the at the surface are hard to see onshore because of
Taupo Fault Belt (Grindley 1960) or Taupo Rift the high sedimentation rate in the graben, but in
(Villamor & Berryman 2006b; Acocella et al. 1987 the Edgecumbe earthquake created fault
2003; Spinks et al. 2005). The axial part of the rift traces which were slightly oblique to that of the
is roughly symmetric with a central graben axis graben (Beanland et al. 1990; Rowland & Sibson
but is partitioned along strike into a number of struc- 2001), and faulting is clear from geothermal drilling
tural domains or rift segments that are often coinci- logs (Wood et al. 2001). Interpretation of the
dent with major volcanic centres, for example faults by Nairn & Beanland (1989) suggest their
Okataina and Taupo caldera complexes (Spinks subsurface portions were block faults dipping
et al. 2005). at 45 + 108 on the SE side of the graben (signifi-
Dips of the faults within this zone appear to be cantly shallower than dips indicated in the central
steep (.608), and the presence of en echelon fault part of TVZ), with steeper antithetic faults on the
traces (Cole 1990) displaced streams (Villamor & NW side.
Berryman 2001), and the pitch of slip vectors (Aco- The area offshore is extending 13 + 6 mm a21,
cella et al. 2003) indicate some dextral strike-slip with up to 4.6 + 2.1 mm a21 of dextral motion, par-
component. Acocella et al. (2003) indicate that allel to the rift axis, and with the most intense fault-
b-values (angle between the direction perpendicular ing on the eastern side (Lamarche et al. 2006). Like
to the trend of the rift segment and extension direc- its onshore equivalent, this is a zone of subsidence
tion) vary from 08 to 258 (Fig. 4b). They suggest (Wright 1990) and thick sediment accumulation
that dextral shear is not confined to individual struc- (Wright 1993). Beyond the continental– oceanic
tures but appears to be distributed across TVZ. crust boundary (at 368450 S) the belt merges with
To the east of TVZ is a zone of primarily dextral the volcanic arc of the southern Kermadec ridge
strike-slip faults (Figs 1 & 5a) referred to as the (Parsons & Wright 1996).
Fig. 4. (a) Main fault structures within and to the east of TVZ, and locations of main vents for rhyolitic and
basaltic volcanism. (b) Rift segments of TVZ showing their extensional directions and the amount of shear (b) in
each segment. After Spinks et al. (2005).
Tongariro Volcanic Centre of Tongariro Volcanic Centre. Alternatively, in

these early vents we may be seeing relics of the
Volcanism here is dominated by andesite massifs of Coromandel structural grain.
Tongariro and Ruapehu, with older centres to the Villamor & Berryman (2006a) note that fault
north and west and associated satellite vents. initiation in the Tongariro Volcanic Centre is
Recent fault traces have a strong NNE trend with younger than to the north, indicating continued
a rift axis of 0278 (Rowland & Sibson 2001), and southward propagation with time. Extension
aligned vents have a similar orientation (Cole within the central area, referred to as the Mt.
1990). Earlier volcanic cones (e.g. Kakaramea and Ruapehu Graben by Villamor & Berryman
Pihanga) have a NW trend, and there is some evi- (2006b), is estimated at 2.3 + 1.2 mm a21. The
dence that the earliest vents on Tongariro had graben abruptly terminates at its southern end with
similar trends (Cole 1978). This may reflect the three sets of cross-cutting faults (Villamor &
importance of cross fractures early in the history Berryman 2006b).
Fig. 4. (Continued).
Central TVZ axis. Near Taupo, for example, a series of basalt

cones and phreatomagmatic deposits collectively
The central TVZ is structurally complex with a form the K Trig basalts (Cole 1972; Brown et al.
number of variable oriented and offset rift segments 1994). Vents for these basalts trend at 0238, parallel
(Spinks et al. 2005). At the surface it is overwhel- to the faults on the north side of Lake Taupo (Cole
mingly dominated by rhyolitic volcanism, manifest 1984). Further north, at Tarawera volcano, there
as a number of volcanic centres comprising clus- was an eruption on 10 June 1886 when basalt was
tered rhyolite lava domes, explosive vents and erupted from a series of en echelon dykes trending
large calderas surrounded and filled by lavas and between 0738 and 0808, within a fissure trending
the deposits of explosive eruptions, including large 0578 across the mountain parallel to the local fault
ignimbrites. Small amounts of high-alumina basalt trend (Nairn & Cole 1981). Most of the individual
(Cole 1972; Houghton et al. 1987) are also associ- dykes are ,2 m wide, and appear to have filled dila-
ated with these rhyolite centres. These were tional fractures without detectable horizontal or ver-
erupted along lineaments parallel to the local rift tical shear in the fissure walls.
Fig. 5. Distribution of calderas within TVZ. (a) Calderas shown by Wilson et al. (1995). (b) Distribution of calderas
within this paper. Dashed lines are maximum extent of calderas; solid lines are possible structural boundaries. After
Spinks et al. 2005.
Vents for rhyolite domes are also commonly (Leonard 2003). This is therefore only discussed
aligned parallel to the local faults, as at Maroa in the context of the Whakamaru Caldera.
(Leonard 2003) and in the Okataina centre (Nairn Our interpretation of the distribution of calderas
2002; Spinks et al. 2005). in TVZ is shown in Figure 5b.
Kapenga ‘Caldera’
Caldera volcanoes of central TVZ
The ‘Kapenga’ Caldera was postulated entirely on
Wilson et al. (1995) identified eight calderas in TVZ geophysical evidence (Rogan 1982; Wilson et al.
(Fig. 5a), expressed at the surface by clustering of 1984), and while several ignimbrites erupted
known or inferred vent locations and/or at depth between 0.89 and 0.71 Ma (Houghton et al. 1995),
by geophysically defined basement depressions. are attributed to a now buried structure, no more
As such, the surface expression of calderas at the recent deposits are unequivocally related to it.
surface in TVZ varies dramatically, and the mor- Gravley et al. (2007) suggest that much of
phology and structure of only four (Taupo, Oka- Kapenga’s current surface morphology is related
taina, Rotorua, Reporoa) is understood with any to collateral (non-eruptive) subsidence associated
level of precision (Spinks et al. 2005); other col- with the 240 ka rhyolitic eruptions from Rotorua
lapse structures, while recognized, are obscured by and Ohakuri (described below), which reactivated
younger pyroclastic units or largely destroyed by faulting associated with the older structure.
subsequent activity.
The oldest recognized caldera collapse structure Whakamaru Caldera and Maroa
is the Mangakino Caldera (1.6– 0.91 Ma), which is Dome Complex
attributed as the source of at least five major ignim-
brites and airfall deposits (Briggs et al. 1993) but is Whakamaru Caldera (Fig. 6) formed during the
little more than a shallow basin at the surface and is largest caldera-forming rhyolitic ignimbrite erup-
defined largely on gravity data (Rogan 1982; Wilson tion in TVZ history at c. 0.34 Ma, but has no clear
et al. 1984). It is outside TVZ, as defined in this topographic margin. It was proposed by Wilson
paper, and will not be discussed further. et al. (1986) on the basis of the thickness and distri-
Recent work within the earlier proposed Maroa bution of the Whakamaru-group ignimbrites,
Caldera does not indicate any separate caldera- exposed east and west of central TVZ (Fig. 6),
forming event or caldera structure associated with which constitute the largest eruptive episode in
emplacement of the Maroa Dome Complex TVZ history, with an exposed volume of at least
Fig. 6. Outflow sheets of the Whakamaru ignimbrite group and probable location of the Whakamaru caldera complex
(dashed line). Redrawn from Brown et al. (1998b).
1000 km3. Several episodes of collapse occurred at faulting likely relates to continued movement
Whakamaru Caldera, with dome emplacement along the margins of the Whakamaru Caldera.
between collapse events (Brown et al. 1998b). Maroa Dome Complex (location shown in
Other caldera boundaries are defined on drillhole Fig. 5a) is an accumulation of simple and composite
thickness of Whakamaru ignimbrite, which is par- silicic lava domes (Leonard 2003). Domes are
ticularly thick inside the structure from observations strongly aligned along NNE trends and extensively
of geothermal drilling at Wairakei, Mokai and faulted along the same lineation, indicating the
Rotokawa (Brown 1994). same regional structure is responsible for control-
The proposed Whakamaru Caldera is defined to ling vent locations and subsequent deformation.
the west by the Western Dome Belt (WDB), a Vent and fault lineations are sub-parallel with fault-
32 km long curvilinear chain of simple and com- ing in the Kapenga graben to the NE.
pound silicic domes inferred from field evidence
to post-date the Whakamaru ignimbrite eruptions Rotorua Caldera
(Wilson et al. 1986; Brown et al. 1998b), a result
of post-collapse volcanism localized along the Rotorua Caldera is perhaps the most conspicuous
western caldera margin. Eruption of the domes has caldera in TVZ, accentuated the local topography
clearly been controlled by a north– south-trending and by its large, sub-circular caldera lake, Lake
fault system, which has subsequently ruptured to Rotorua (Fig. 7). The caldera is located 15–20 km
displace the domes. Domes immediately east of NW of the junction between Okataina and
the faults are probably younger features. Given Kapenga rift axes, and formed during and immedi-
that the domes and faults are aligned along a linea- ately following the eruption of the 0.24 + 11 Ma
tion oblique to the regional trend, if the lavas of the Mamaku Ignimbrite (Milner et al. 2002; date from
WDB are post-collapse features (Fig. 6), then Leonard 2003), with a minimum eruption volume
Fig. 7. Rotorua Caldera showing main structures derived from analyses of DEM data; pre-caldera and post-caldera
lavas and distribution of Mamaku ignimbrite outflow sheet. Grid coordinates shown in the New Zealand map grid
(NZMG).
(including intra-caldera ignimbrite) of 145 km3 eruptive phase from the Mamaku magma system
DRE (Milner et al. 2003). Some authors have pro- (Milner et al. 2002). Smaller rhyolite domes are geo-
posed earlier events at Rotorua Caldera to account chemically distinct and thought to be much younger.
for older ignimbrites in the area (Wood 1992), or Stratigraphic evidence outlined by Milner et al.
that Rotorua is not a caldera at all (Hunt 1992), (2002) indicates that caldera collapse occurred
but a detailed study by Milner et al. (2002, 2003) throughout the eruption and emplacement of the
strongly indicates Rotorua is a single-event Mamaku Ignimbrite during a single eruptive
caldera, and the source of the Mamaku Ignimbrite. episode. Milner et al. (2002) describe asymmetric
A number of pre-caldera rhyolite domes are caldera collapse that was deepest in the SW of the
exposed in the vicinity of Rotorua Caldera caldera, with a component of downsag expressed
(Fig. 7). Milner et al. (2002) showed that those on in the overlying Mamaku Ignimbrite. Mamaku
the rim of the caldera are geochemically distinct Ignimbrite geochemistry indicates the eruption of
from each other and from the Mamaku magma a single, compositionally zoned magma reservoir,
system. The post-caldera Ngongotaha and Puke- represented by three petrogenetically related
hangi rhyolite dome complexes, and lavas exposed pumice types. An andesitic juvenile component in
at Kawaha Point are geochemically similar to the upper parts of the Mamaku Ignimbrite is thought
Mamaku Ignimbrite system and may reflect a final to reflect a discrete magma injected into the residual
silicic chamber and tapped during later phases of the Rotorua (Figs 5b & 8), which is considered the
eruption during advanced stages of caldera collapse source of the c. 240 ka Ohakuri pyroclastic deposits
(Milner et al. 2003). (Gravley et al. 2006, 2007). It overlies the northern
Rotorua is a simple sub-circular caldera with margin of the older Whakamaru Caldera and lies
dimensions of c. 20 16 km. The caldera floor is adjacent to the axis of the modern Taupo Fault
dominated by the c. 9 km diameter caldera lake Belt. It is difficult to define either the topographic
and the youthful morphologies of the post-caldera or structural caldera margins precisely because of
rhyolite dome complexes. The caldera is character- subsequent deposition of volcaniclastic sediments,
ized by a North –South elongate negative residual but its presence is inferred from the distribution of
gravity anomaly to the west and SW of Lake air fall deposits, distribution of and transport direc-
Rotorua, including Rotorua city and the post-caldera tions within the Ohakuri Ignimbrite, and from
rhyolite dome of Ngongotaha (Rogan 1982; Hunt geophysical data.
1992; Davy & Caldwell 1998). The surrounding Stratigraphic evidence presented by Gravley
gravity contours are not specifically concentric to et al. (2007) suggests that the Mamaku Ignimbrite
the caldera margin, defining an asymmetric rise in and the Ohakuri pyroclastic deposits were erupted
basement towards the NE and NW caldera in rapid succession, from the Rotorua and
margins. The basement gradient is shallowest Ohakuri Calderas respectively. They suggest that
towards the east and steepest around the southwes- the most feasible way of achieving this linkage
tern margin. In the NW of the caldera several low is through volcano-tectonic processes, in which con-
scarps sub-parallel to the caldera margin are temporaneous faulting triggered almost concurrent
located within the area defined by Milner et al. events 25 km apart.
(2002) as having deformed by downsag into the
caldera during collapse. Reporoa Caldera
Ohakuri Caldera Reporoa Caldera (Nairn et al. 1994) is located at the
northern end of the Taupo-Reporoa depression,
The Ohakuri Caldera is a newly recognized struc- c. 15 km east of the Kapenga segment axis
ture in the Whangapoa basin, 25 km SSW of (Fig. 9), and is the source of the 0.23 + 0.01 Ma
Fig. 8. Ohakuri Caldera and distribution of the Ohakuri ignimbrite (darker ornament). From Gravley et al. (2007).
Fig. 9. Reporoa Caldera showing main structures derived from analyses of DEM data; pre-caldera and post-caldera
lavas and distribution of Kaingaroa ignimbrite outflow sheet. Symbols and grid as in Figure 7.
(Houghton et al. 1995) Kaingaroa Ignimbrite, with a The caldera has a small but distinctive negative
total eruptive volume of c. 100 km3 (Nairn et al. gravity anomaly (Nairn et al. 1994; Stagpoole
1994; Beresford & Cole 2000). Kaingaroa Ignim- 1994; Stagpoole & Bibby 1999) and a low Vp
brite extends radially for 20 –30 km beyond the anomaly consistent with low density, low Vp
caldera, mostly to the east of Reporoa Caldera caldera fill (Sherburn et al. 2003). The gravity
where it caps the Kaingaroa Plateau (Fig. 9). anomaly corresponds well with the topographic
Pre-caldera volcanism in the Reporoa area com- expression of the caldera, with a gentle and largely
prises rhyolite lavas unrelated to the Kaingaroa open western margin and a steep eastern margin,
magma system or the formation of Reporoa consistent with asymmetric collapse (Beresford &
Caldera (Beresford & Cole 2000) and older ignim- Cole 2000). Nairn et al. (1994) interpret post-
brites from caldera sources to the west (Wilson caldera rhyolite domes and a buried dome
et al. 1986; Ritchie 1996; Beresford et al. 2000). complex to have erupted along fractures related
Minor (,2 km3) post-caldera rhyolite domes are to the caldera rim and a supposed inner caldera
geochemically and isotopically distinct from the ring fault.
Kaingaroa magma system. Lithic componentry Reporoa is a morphologically simple sub-
data for the Kaingaroa Ignimbrite presented by circular caldera (Fig. 9), with approximate dimen-
Beresford & Cole (2000) suggest multiple stages sions of 11 13 km and well-preserved 250 m
in the eruption event: (1) an initial single vent high collapse scarps along the northern boundary.
phase; (2) a multiple vent or ring fracture phase on The north–south long axis of the caldera (eccentri-
the eastern side with asymmetric caldera collapse city E ¼ 0.81) is oblique to the regional trend
leading to eastward-directed pyroclastic flows; and of faults to the west; in the east a NE-trending
(3) piston collapse accompanied by radially directed fault scarp merges with the north–south trending
pyroclastic flows. eastern caldera margin. The flat-floored caldera
has a well-defined topographic margin in the north and the 65 ka Rotoiti eruption (Nairn 1981, 1989;
and east, but is open to the west and south. The date from Houghton et al. 1995), and modified
caldera margin is neither dissected by younger by substantial intra-caldera rhyolite volcanism
faults nor does it truncate older structures. Minor (e.g. Jurado-Chichay & Walker 2000; Nairn 1989,
lineaments within the caldera may record modern 2002). Older (.300 ka) collapse events are
subsidence or reflect the lacustrine history of likely, but potential deposits and precise collapse
the basin. margins are obscured and/or overprinted by
subsequent activity. Magmatic volume estimates
Okataina Caldera Complex for the Matahina and Rotoiti events (including
intra-caldera estimates) are 150 km3 (Bailey &
Okataina Caldera Complex (OCC) is a complex of Carr 1994) and 120 km3 (Froggatt & Lowe 1990),
overlapping and nested collapse structures, largely respectively; other eruptives from within and adja-
filled by the products of post-caldera rhyolite cent to OCC account for at least 150 km3 (e.g.
volcanism (Fig. 10). The composite structure is the Nairn 1989; Froggatt & Lowe 1990; Bellamy
result of two main collapse events associated with 1991; Jurado-Chichay & Walker 2000). The
the 0.28 + 0.01 Ma Matahina Ignimbrite (Bailey Matahina and Rotoiti ignimbrites extend predomi-
& Carr 1994; date from Houghton et al. 1995) nantly to the east and north, and appear related to
Fig. 10. Okataina Caldera complex showing main structures derived from analyses of DEM data, pre-caldera and
post-caldera lavas and distribution of Matahina and Rotoiti ignimbrite outflow sheets. Symbols and grid as in Figure 7.
overlapping but distinct sources in the southern and the structure but gravity highs separate the structure
northern parts of the caldera complex respectively from basement highs to the east and Rotorua
(Nairn 1989). Caldera to the west. Topographic embayments in
A number of rhyolite dome lavas scattered the caldera complex margin are also peripheral to
around the rim of the caldera complex record vol- the main basement depression indicated by gravity
canism in the Okataina area predating the first col- and Vp data.
lapse event (Nairn 1989, 2002). No precise dates The topographic margin at Okataina (Nairn
exist for these lavas, but Bowyer (2001) showed 2002) is variably manifest as scalloped slump
that they are chemically distinct and relate to dis- scars in pre-caldera rhyolite domes and ignimbrites,
crete magma batches. Geochemical variation is eroded caldera walls, rectilinear fault scarps coinci-
small however, and no significant variability exists dent with regional faulting (Fig. 10), and in the SW
between pre-caldera lavas adjacent to the Okataina by the steep margins of post-caldera constructional
and Rotorua calderas; this would imply that these rhyolite domes. As such, the composite topographic
lavas are not related specifically to a Rotorua or margin defines a depression considerably modified
Okataina ‘volcanic centre’. Only one pre-caldera from the original multiple collapse structure. Dis-
magma batch has a similar chemistry with the tinct embayments occur on each side of OCC
Matahina magma system (e.g. Nairn 1981; where it is intersected by regional faulting of the
Bowyer 2001). axial rift within the Okataina transfer zone. These
Lavas and associated pyroclastics erupted are contiguous with two intra-caldera dome com-
between the major caldera-forming events are pre- plexes forming two overlapping linear vent zones,
dominantly exposed to the SW of the caldera which transect the caldera complex as the lateral
where it intersects the Kapenga axial rift segment continuation of the adjacent rift segment axes. The
and relate to multiple magma batches (Bellamy boundaries of individual collapse events are
1991; Bowyer 2001). Following the caldera collapse complex and largely overprinted by subsequent vol-
event associated with the eruption of the Rotoiti canism and tectonism, but caldera reconstructions
Pyroclastics, a major phase of explosive volcanism suggest the major collapses are centred on the
ensued from sources within the caldera complex axes of the intersecting rift segments. Davy &
prior to the development of the two large rhyolite Bibby (2005) indicate from a seismic reflection
lava massifs that currently fill the caldera. Two survey that Mamaku Ignimbrite dips ,68E across
main magma types were erupted during the explo- the western side of Lake Tarawera, and then
sive phase, generating fourteen eruptive episodes becomes cut off, suggesting a caldera structural
(Smith et al. 2002). During more recent effusive boundary at this point. Lakes at Okataina exhibit a
activity, multiple magmas were often involved moat pattern where they have ponded between the
with a single eruptive episode. The two documented topographic rim of the caldera and caldera-filling
caldera-forming events at Okataina are spatially post-caldera constructional volcanism. Lakes
overlapping but are significantly separated tem- filling earlier manifestations of the structure may
porally (.200 ka), reflecting geochemically distinct have been much larger.
magma systems (Burt et al. 1998). Volcanism in the
Okataina area from the earliest to the most recent Taupo Caldera Complex
eruptives, including the caldera forming events,
therefore records the eruption of multiple discrete Taupo Caldera Complex (Fig. 11) has been fre-
magma chambers rather than the progressive quently active in the past c. 65 ka (Wilson et al.
tapping of a single large chamber. 1986; Houghton et al. 1995), while its poorly con-
A distinct large negative residual gravity strained early eruptive history indicates activity
anomaly (Rogan 1982; Davy & Caldwell 1998) over c. 300 ka (Wilson et al. 1986; Cole et al.
and low Vp anomaly (Sherburn et al. 2003) define 1998). The caldera-forming Oruanui eruption at
a north– south elongated depression consistent with 26.5 ka (calibrated; Wilson 1993) generated a
the mapped Okataina Caldera margin and filled c. 430 km3 fall deposit, a 320 km3 bulk volume non-
with a large volume of low Vp, low density, volcani- welded density current deposits (mostly ignimbrite)
clastic sediment. A clear low Vp anomaly at 4 km and c. 420 km3 of caldera-fill material erupted,
effectively corresponds to a minimum depth extent equivalent to c. 530 km3 of magma (Wilson 2001).
of the collapse structure (Sherburn et al. 2003). The Oruanui event is thus largely responsible for
Gravity data presented by Nairn (2002) indicates a the modern caldera morphology. Wilson (1993)
north–south elongate negative gravity anomaly has identified 28 separate eruptions since the
roughly concentric to the topographic margin, and Oruanui eruption; the most recent and largest of
centred beneath Tarawera Volcanic Complex and these, the caldera modifying 35 km3 Taupo Ignim-
the southern part of Haroharo Volcanic Complex brite eruption, occurred about 1800 years ago from
(Fig. 10). The contours open on the west side of vents near the Horomatangi Reefs in the eastern
Fig. 11. Taupo Caldera complex showing main structures derived from analyses of DEM data, pre-caldera and
post-caldera lavas and distribution of Oruanui and Taupo ignimbrite outflow sheets. Symbols and grid as in Figure 7.
part of the lake (Wilson & Walker 1985; Smith & dome complex in this area during the post-Oruanui
Houghton 1995). phase is suggested by lithic componentry data for
The early history (.65 ka) of volcanism in the the Taupo Ignimbrite (Cole et al. 1998) with its
vicinity of modern Lake Taupo is represented likely destruction during the Taupo eruption.
mainly by domes and limited pyroclastics scattered Lithic componentry analysis of both the Oruanui
around the lake (Sutton et al. 1995). Two ignim- and Taupo ignimbrites identifies different lithic
brites exposed on the margin of the caldera are com- suites, interpreted by Cole et al. (1998) as reflecting
monly attributed to a Taupo source (Sutton et al. dissimilar sub-caldera geology beneath mutually
1995; Cole et al. 1998), although their limited exclusive collapse structures.
extent means their relationship to the current Petrological studies (e.g. Sutton et al. 1995,
caldera complex is ambiguous. Pre-caldera rhyolite 2000; Charlier et al. 2005) show a complex mag-
lavas form a series of headlands along the northern matic system, involving the stepwise appearance
caldera margin and to the SW of the caldera, while of compositionally distinct magma batches with
the large caldera-filling domes and flows character- short crustal-residence times. Eruptives prior to
istic of OCC are noticeably absent. Vents for the the Oruanui eruption form distinct compositional
post-Oruanui explosive eruptions are inferred by and spatial groups, while a large isotopically
isopach data to be concentrated along a NE-trending homogeneous magma body was generated prior to
lineation (Fig. 11) in the eastern part of modern the Oruanui caldera-forming eruption (Sutton
Lake Taupo (Wilson 1993). Construction of a et al. 1995). Some of the pre-Oruanui domes
exposed on the northern caldera margin, and wide- presented by Davy & Caldwell (1998) all suggest
spread tephras erupted between 65 ka and the that a line between Karangahape Cliffs and
Oruanui eruption, are the same composition as the Motutere Point marks a major structural boundary
Oruanui magma, and thus record the coalescence perpendicular to the trend of TVZ.
of a large magma chamber. Distinct magmas were The southern part of Lake Taupo occupies a
erupted during the same period from different NE-trending fault-bounded depression intersecting
areas around Taupo, suggesting that magma the southern caldera margin; the western margin
batches may have been erupted during the Oruanui of this structure is a continuation of fault systems
eruption. Eruptives of the post-Oruanui sequence dissecting the Tongariro Volcanic Centre to the
form four temporally grouped magma types that south. A prominent structural feature is the diver-
are compositionally distinct from the Oruanui gence in fault trend (c. 208) to the south and north
magma. The youngest magma, associated with the of the caldera complex. This bend effectively
Taupo eruption, represents the largest homogeneous forms the boundary between the Tongariro and
magma accumulation in the post-Oruanui sequence Taupo South Rift segments. Vents for pre-caldera
(Sutton et al. 1995). lavas to the south and north of the caldera
A large trapezoidal-shaped negative Bouguer complex lie along NE-trending lineaments; the
gravity anomaly, the most intense negative gravity vents for post-Oruanui eruptions mostly occur
anomaly in TVZ (Davy & Caldwell 1998), is docu- along the eastern edge of the caldera complex
mented over the northern part of Lake Taupo. (Wilson 1993).
The gravity anomaly is consistent with a caldera col-
lapse structure elongate NW–SE, perpendicular to Discussion
the axial rift zone in this segment. The gravity
data do not facilitate identification of individual col- The characteristics of TVZ calderas in the last
lapse structures, and Davy & Caldwell (1998) con- 300 ka (Taupo, Okataina, Rotorua, Ohakuri and
sider the structures are nested, with the Taupo Reporoa) clearly indicate the complex eruptive and
eruption producing additional subsidence in the structural history of Okataina and Taupo compared
NE part of the modern lake. Geophysical data also to Rotorua, Ohakuri and Reporoa. These data show
demonstrate differential subsidence towards the that these caldera structures can be divided into
caldera in the southern part of the lake, and a two groups: (1) extra-rift calderas (Reporoa and
NW–SE-trending structural boundary marking the Rotorua) are simple, relatively small, sub-circular,
southern caldera margin (Davy & Caldwell 1998). monogenetic structures; and (2) intra-rift caldera
Seismic reflection, gravity and magnetic data complexes (Okataina and Taupo) are large, multiple
Fig. 12. Schematic cross-section of TVZ indicating likely features within the crust and upper mantle. After Hiess
et al. (2007).
collapse structures, with rectangular geometries, References

clear coupling to regional structure, and broadly
homogeneous magmas. This distinction demon- A COCELLA , V., S PINKS , K. D., C OLE , J. W. & N ICOL , A.
2003. Oblique back arc rifting of Taupo Volcanic
strates the role of active regional tectonics in influen-
Zone, New Zealand. Tectonics, 22, 1045, doi:
cing the location, structure and development of 0.1029/2002TC001447.
caldera systems within a rift zone. Ohakuri is a A DAMS , C. J., G RAHAM , I. J., S EWARD , D. & S KINNER ,
simple structure, but is intra-rift. D. N. B. 1994. Geochronological and geochemical
With an increasing number of high precision evolution of late Cenozoic volcanism in the Coroman-
radiometric dates from ignimbrites and rhyolite del Peninsula, New Zealand. New Zealand Journal of
domes (e.g. Houghton et al. 1995; Leonard 2003), Geology and Geophysics, 37, 359–379.
it has become evident that caldera-forming events A NDERSON , H. & W EBB , T. 1989. The rupture process of
are not uniformly spread in time and space across the 1987 Edgecumbe Earthquake, New Zealand.
New Zealand Journal of Geology and Geophysics,
TVZ. The earliest caldera-forming events occurred
32, 43– 52.
in the most central part of TVZ and the currently B AILEY , R. A. & C ARR , R. G. 1994. Physical geology and
active areas are to the north and south, at Okataina eruptive history of the Matahina Ignimbrite, Taupo
and Taupo. Equally, as shown by the above authors, Volcanic Zone, North Island, New Zealand. New
there are periods of more intense caldera-forming Zealand Journal of Geology and Geophysics, 37,
activity (e.g. 0.89–0.71 Ma; 0.34 –0.32 Ma; 0.28 – 319– 344.
0.24 Ma). This may well relate to variability in B EANLAND , S. 1995. The North Island Dextral Fault
rate of extension. It must be remembered however Belt, Hikurangi subduction margin, New Zealand.
that rhyolite dome/tephra eruptions occur through- PhD thesis, Victoria University of Wellington,
New Zealand.
out the history of TVZ, and in any one area there
B EANLAND , S., B LICK , G. H. & D ARBY , D. J. 1990.
may be activity over a considerable period of time Normal faulting in a back-arc: geological and geodetic
(as at Kapenga). The caldera-forming event is there- characteristics of the 1987 Edgecumbe earthquake,
fore just one episode in the activity, so using the New Zealand. Journal of Geophysical Research, 95,
terms ‘pre’- and ‘post’-caldera volcanism needs to 4693– 4707.
be used with caution. B EANLAND , S., M ELHUISH , A., N ICOL , A. & R AVENS , J.
Magmatism and geological structure are closely 1998. Structure and deformational history of the
related in the TVZ, as shown by the common align- inner forearc region, Hikurangi subduction margin,
ment of vents along local faults. This is true of many New Zealand. New Zealand Journal of Geology and
Geophysics, 41, 325– 342.
rhyolite domes and of high-alumina basalt fissure
B ELLAMY , S. 1991. Some studies of the Te Wairoa ignim-
eruptions (Cole 1972; Nairn & Cole 1981; brites and the associated volcanic geology of the
Houghton et al. 1987). Spinks et al. (2005) also SW Okataina volcanic centre, Taupo Volcanic
show that there is a correlation between areas of Zone. MSc thesis, University of Waikato, Hamilton,
maximum extension, particularly Okataina and New Zealand.
Taupo, and eruptive volume. A combination of B ERESFORD , S. W. & C OLE , J. W. 2000. Kaingaroa
these observations suggests that basalt rises into Ignimbrite, Taupo Volcanic Zone, New Zealand:
the lower crust, providing high heat flow. This is evidence for asymmetric caldera subsidence of the
likely to cause partial melting of parts of the lower Reporoa Caldera. New Zealand Journal of Geology
and Geophysics, 43, 471–481.
crust, composed of metasediments and igneous
B ERESFORD , S. W., C OLE , J. W. & W EAVER , S. D. 2000.
rocks (Fig. 12), which rise and pond at the Weak chemical and mineralogical zonation in
ductile/brittle transition zone in the mid-crust. At the Kaingaroa Ignimbrite, Taupo Volcanic Zone,
this point they become strongly influenced by the New Zealand. New Zealand Journal of Geology and
many faults in central TVZ, and so erupt at the Geophysics, 43, 639– 680.
surface along linear fault zones. The greatest B IBBY , H. M., C ALDWELL , T. G., D AVEY , F. J. & W EBB ,
volume of magma will occur where the greatest T. H. 1995. Geophysical evidence of the structure
amount of extension takes place, and this in turn of the Taupo Volcanic Zone and its hydrothermal cir-
will create the greatest volume of eruptives, and culation. Journal of Volcanology and Geothermal
Research, 68, 29– 58.
the largest and most complex calderas. The shapes
B IBBY , H. M., R ISK , G. F., C ALDWELL , T. G., O GAWA ,
of these calderas will also be strongly influenced Y., T AKAKURA , S. & U CHIDA , T. 2000. Deep electri-
by the local faults. cal structure beneath the Taupo Volcanic Zone and the
source of geothermal heat. Proceedings of the 22nd
Early drafts of the paper have been read by H. Bibby and Geothermal Workshop, 81–86.
D. Gravley, and their comments are much appreciated. B OWYER , D. A. 2001. Petrologic, geochemical, and isoto-
The paper was reviewed by J. Gamble and V. Acocella, pic evolution of rhyolite lavas from the Okataina,
and their helpful and constructive suggestions have Rotorua, and Kapenga volcanic centres, Taupo Volca-
considerably improved the paper. Assistance with figure nic Zone, New Zealand. PhD thesis, University of
preparation by S. Hampton is also acknowledged. Waikato, Hamilton, New Zealand.
B RIGGS , R. M., G IFFORD , M. G., M OYLE , A. R., C OLE , J. W. 1978. Andesites of the Tongariro Volcanic
T AYLOR , S. R., N ORMAN , M. D., H OUGHTON , B. F. Centre, New Zealand. Journal of Volcanology and
& W ILSON , C. J. N. 1993. Geochemical zoning and Geothermal Research, 3, 121– 153.
eruptive mixing in ignimbrites from Mangakino C OLE , J. W. 1984. Taupo-Rotorua Depression – an ensia-
volcano, Taupo Volcanic Zone, New Zealand. lic marginal basin of North Island, New Zealand. In:
Journal of Volcanology and Geothermal Research, K OKELAAR , P. & H OWELLS , M. F. (eds) Marginal
56, 175–203. Basin Geology: volcanic and associated sedimentary
B RIGGS , R. M., H OUGHTON , B. F., M C W ILLIAMS , M. & and tectonic processes in modern and ancient marginal
W ILSON , C. J. N. 2005. 40Ar/39Ar ages of silicic rocks basins. Geological Society, London, Special Publi-
in the Tauranga-Kaimai area, New Zealand: dating the cations, 16, 109–120.
transition between volcanism in the Coromandel Arc C OLE , J. W. 1990. Structural control and origin of volcan-
and the Taupo Volcanic Zone. New Zealand Journal ism in the Taupo volcanic zone, New Zealand. Bulletin
of Geology and Geophysics, 48, 459–469. of Volcanology, 52, 445–459.
B ROWN , S. J. A. 1994. Geology and geochemistry of the C OLE , J. W., D ARBY , D. J. & S TERN , T. A. 1995. Taupo
Wkahamaru-group ignimbrites, and associated rhyo- Volcanic Zone and Central Volcanic Region: backarc
lite domes, Taupo Volcanic Zone, New Zealand. PhD Structures of North Island, New Zealand. In:
thesis, University of Canterbury, Christchurch, New T AYLOR , B. (ed.) Backarc Basins: Tectonics and
Zealand. Magmatism. Plenum Press, New York, 1 –27.
B ROWN , S. J. A., S MITH , R. T. & C OLE , J. W. 1994. Com- C OLE , J. W., B ROWN , S. J. W., B URT , R. M.,
positional and textural characteristics of the strombo- B ERESFORD , S. W. & W ILSON , C. J. N. 1998. Lithic
lian and surtseyan K-Trig basalts, Taupo Volcanic types in ignimbrites as a guide to the evolution of
Centre, New Zealand: implications for eruption a caldera complex, Taupo volcanic centre, New
dynamics. New Zealand Journal of Geology and Geo- Zealand. Journal of Volcanology and Geothermal
physics, 37, 113– 126. Research, 80, 217– 237.
B ROWN , S. J. A., B URT , R. M., C OLE , J. W., K RIPPNER , D ARBY , D. J. & M EERTENS , C. M. 1995. Terrestrial and
S. J. P., P RICE , R. C. & C ARTWRIGHT , I. 1998a. Plu- GPS measurements of deformation across the Taupo
tonic lithics in ignimbrites of the Taupo Volcanic back-arc and Hikurangi fore-arc regions in New
Zone, New Zealand; sources and conditions of crystal- Zealand Journal of Geophysical Research, 100,
lisation. Chemical Geology, 148, 21– 41. 8221– 8232.
B ROWN , S. J. A., C OLE , J. W., W ILSON , C. J. N. & D ARBY , D. J., H ODGKINSON , K. M. & B LICK , G. H.
W OODEN , J. 1998b. Geochemistry of the Whakamaru 2000. Geodetic measurement of deformation in the
group ignimbrites, Taupo Volcanic Zone, New Taupo Volcanic Zone, New Zealand: the north Taupo
Zealand; evidence for non-sequential tapping of a network revisited. New Zealand Journal of Geology
large silicic magma system. Journal of Volcanology and Geophysics, 43, 157 –170.
and Geothermal Research, 84, 1 –37. D AVY , B. W. & B IBBY , H. 2005. Seismic reflection
B RYAN , C. J., S HERBURN , S., B IBBY , H. M., B ANNIS- imaging of the Haroharo Caldera boundary beneath
TER , S. C. & H URST , A. W. 1999. Shallow seismicity Lake Tarawera, Okataina Volcanic Centre, New
of the central Taupo Volcanic Zone, New Zealand: its Zealand. New Zealand Journal of Geology and Geo-
distribution and nature. New Zealand Journal of physics, 48, 153 –166.
Geology and Geophysics, 42, 533– 542. D AVY , B. W. & C ALDWELL , T. G. 1998. Gravity, mag-
B URT , R. M., B ROWN , S. J. A., C OLE , J. W., S HELLEY , D. netic and seismic surveys of the caldera complex,
& W AIGHT , T. E. 1998. Glass-bearing plutonic frag- Lake Taupo, North Island, New Zealand. Journal of
ments from ignimbrites of the Okataina volcanic Volcanology and Geothermal Research, 81, 69–89.
complex, Taupo Volcanic Zone, New Zealand; D E M ETS , C., G ORDON , R. G., A RGUS , D. F. & S TEIN , S.
remnants of a partially molten intrusion associated 1990. Current plate motions. Geophysical Journal
with preceding eruptions. Journal of Volcanology International, 101, 425 –478.
and Geothermal Research, 84, 209 –337. F ROGGATT , P. C. & L OWE , D. J. 1990. A review of late
C ASHMAN , S. M., K ELSEY , H. M., E RDMAN , C. F., quaternary silicic and some other tephra formations
C UTTEN , H. N. C. & B ERRYMAN , K. 1992. Strain par- from New Zealand; their stratigraphy, nomenclature,
titioning between structural domains in the forearc of distribution, volume, and age. New Zealand Journal
the Hikurangi subduction zone. Tectonics, 11, of Geology and Geophysics, 33, 89–109.
242– 257. G AMBLE , J. A., W RIGHT , I. C. & B AKER , J. A. 1993. Sea-
C HARLIER , B. L. A., W ILSON , C. J. N., L OWENSTERN , floor geology and petrology in the oceanic to continen-
J. B., B LAKE , S., VAN C ALSTEREN , P. W. & tal transition zone of the Kermadec-Havre-Taupo
D AVIDSON , J. P. 2005. Magma generation at a large Volcanic Zone Arc System, New Zealand. New
hyperactive silicic volcano (Taupo, New Zealand) Zealand Journal of Geology and Geophysics, 36,
revealed by U–Th and U–Pb systematics in zircons. 417–435.
Journal of Petrology, 46, 3 –32. G RAVLEY , D. M., W ILSON , C. J. N., R OSENBERG , M. D.
C OCHRANE , G. R. & T IANFENG , W. 1983. Interpretation & L EONARD , G. S. 2006. The nature and age of
of structural characteristics of the Taupo Volcanic Ohakuri Formation and Ohakuri Group rocks in
Zone, New Zealand, from Landsat imagery. Inter- surface exposures and geothermal drillhole sequences
national Journal of Remote Sensing, 4, 111– 128. in the central Taupo Volcanic Zone, New Zealand.
C OLE , J. W. 1972. High-alumina basalts of Taupo Volca- New Zealand Journal of Geology and Geophysics,
nic Zone, New Zealand. Lithos, 6, 53–64. 49, 305– 308.
G RAVLEY , D. M., W ILSON , C. J. N., L EONARD , C. S. & PhD thesis, University of Canterbury, Christchurch,
C OLE , J. W. 2007. Double Trouble: paired ignimbrite New Zealand.
eruptions and collateral subsidence in the Taupo Vol- M ILNER , D. M., C OLE , J. W. & W OOD , C. P. 2002.
canic Zone, New Zealand. Geological Society of Asymmetric, multiple-block collapse at Rotorua
America Bulletin, 119, 18– 30. Caldera, Taupo Volcanic Zone, New Zealand. Bulletin
G RINDLEY , G. W. 1960. Sheet 8 – Taupo. Geological map of Volcanology, 64, 134– 149.
of New Zealand 1:250:000. Department of Scientific M ILNER , D. M., C OLE , J. W. & W OOD , C. P. 2003.
and Industrial Research, Wellington. Mamaku Ignimbrite: a caldera-forming ignimbrite
G RINDLEY , G. W. & B ROWNE , P. R. L. 1976. Structural erupted from a compositionally zoned magma
and hydrological factors controlling the permeabilities chamber in Taupo Volcanic Zone, New Zealand.
of some hot-water systems. Proceedings of the 2nd UN Journal of Volcanology and Geothermal Research,
Symposium on Development and Use of Geothermal 122, 243–264.
Resources, 377– 386. N AIRN , I. A. 1981. Some studies of the geology, volcanic
H ARRISON , A. J. & W HITE , R. S. 2004. Crustal structure history and geothermal resources of the Okataina
of the Taupo Volcanic Zone, New Zealand: stretching volcanic centre, Taupo Volcanic Zone, New Zealand.
and igneous intrusion. Geophysical Research Letters, PhD thesis, Victoria University, Wellington, New
31, 1– 4. doi: 10.1029/2004GL019885. Zealand.
H EALY , J., S CHOFIELD , J. C. & T HOMPSON , B. N. 1964. N AIRN , I. A. 1989. Sheet V16AC – Mount Tarawera, geo-
Sheet 5 – Rotorua. Geological map of New Zealand logical map of New Zealand, 1:50,000. Department of
1:250,000. Department of Scientific and Industrial Scientific and Industrial Research, Wellington, New
Research, Wellington. Zealand.
H EISE , W., B IBBY , H. M., C ALDWELL , T. G., N AIRN , I. A. 2002. Geology of the Okataina Volcanic
B ANNISTER , S. C., O GAWA , Y., T AKAKURA , S. & Centre. Institute of Geological and Nuclear Sciences
U CHIDA , T. 2007. Melt distribution beneath a young Geological Map, 25, 156p þ map.
continental rift: the Taupo Volcanic Zone, New N AIRN , I. A. & B EANLAND , S. 1989. Geological setting
Zealand. Geophysical Research Letters, 34; L14313. of the 1987 Edgecumbe earthquake, New Zealand.
doi: 10.1029/2007GL029629. New Zealand Journal of Geology and Geophysics, 32,
H ENRYS , S. H., R EYNERS , M. & B IBBY , H. 2003. Explor- 1–13.
ing the plate boundary structure of the North Island, N AIRN , I. A. & C OLE , J. W. 1981. Basalt dykes in the 1886
New Zealand. EOS, Transaction of the American Geo- Tarawera Rift. New Zealand Journal of Geology and
physical Union, 84, 289– 295. Geophysics, 24, 585– 592.
H IESS , J., C OLE , J. W. & S PINKS , K. D. 2007. High- N AIRN , I. A., W OOD , C. P. & B AILEY , R. A. 1994. The
alumina basalts of the Taupo Volcanic Zone, New Reporoa Caldera, Taupo Volcanic Zone: source of
Zealand: influence of the crust and crustal structure. the Kaingaroa Ignimbrites. Bulletin of Volcanology,
New Zealand Journal of Geology and Geophysics, 56, 529 –537.
50, 327–342. N ICOL , A., M AZENGARB , C., C HANIER , F., R AIT , G.,
H OUGHTON , B. F., W ILSON , C. J. N., L LOYD , E. F., U RUSKI , C. & W ALLACE , L. 2006. Large (.50 km)
G AMBLE , J. A. & K OKELAAR , B. P. 1987. A post Oligocene strike-slip displacements in the
catalogue of basaltic deposits within Taupo Volcanic North Island; Fact or Fiction. Abstract, Geosciences
Zone. New Zealand Geological Survey Record, 18, ’06 – Our planet, our future. Geological Society
95– 101. of New Zealand Miscellaneous Publication,
H OUGHTON , B. F., W ILSON , C. J. N., M C W ILLIAMS , 122A, 59.
M. O., L ANPHERE , M. A., W EAVER , S. D., B RIGGS , O GAWA , Y., B IBBY , H. M. ET AL . 1999. Wide-band mag-
R. M. & P RINGLE , M. S. 1995. Chronology and netotelluric measurements across the Taupo Volcanic
dynamics of a large silicic magmatic system, Central Zone, New Zealand – preliminary results. Geophysical
Taupo Volcanic Zone, New Zealand. Geology, 23, Research Letters, 26, 3673–3676.
13– 16. O LIVER , P. J. 1978. Seismotectonic, structural, volca-
H UNT , T. M. 1992. Gravity studies in the Rotorua area, nological and geomorphological study of New
New Zealand. Geothermics, 21, 65–74. Zealand. In: E LLIS , P. J. ET AL . (eds) ‘Landsat II
J URADO -C HICHAY , Z. & W ALKER , G. P. L. 2000. over New Zealand, Monitoring our Resources from
Stratigraphy and dispersal of the Mangaone Subgroup Space’. Department of Scientific and Industrial
pyroclastic deposits, Okataina Volcanic Centre, New Research Bulletin, 221, 298.
Zealand. Journal of Volcanology and Geothermal P ARSONS , L. M. & W RIGHT , I. C. 1996. The Lau-Havre-
Research, 104, 319 –383. Taupo back-arc basin: a southward propagating,
K NOX , G. J. 1982. Taranaki Basin, structural style and tec- multi-stage evolution from rifting to spreading. Tecto-
tonic setting. New Zealand Journal of Geology and nophysics, 263, 1– 22.
Geophysics, 25, 125–140. P RICE , R. C., G AMBLE , J. A., S MITH , I. E. M., E GGINS ,
L AMARCHE , G., B ARNES , P. & B ULL , J. M. 2006. S. M. & W RIGHT , I. C. 2005. An integrated model
Faulting and extension rate over the last 20,000 years for the temporal evolution of andesites and rhyolites
in the offshore Whakatane Graben, New Zealand con- in New Zealand’s North Island. Journal of Volcanol-
tinental shelf. Tectonics, 25, TC4005, doi: 10.1029/ ogy and Geothermal Research, 140, 1– 24.
2005TC001886. R EYNERS , M., E BERHART -P HILLIPS , D., S TUART , G. &
L EONARD , G. S. 2003. The evolution of Maroa Volcanic N ISHIMURA , Y. 2006. Imaging subduction from the
Centre, Taupo Volcanic Zone, New Zealand. trench to 300 km depth beneath the central North
Island, New Zealand, with Vp and Vp/Vs. Geophysi- S TRATFORD , W. R. & S TERN , T. A. 2006. Crust and upper
cal Journal International, 165, 565– 583. mantle structure of a continental backarc: central North
R ITCHIE , A. 1996. Volcanic geology of the Waiotapu Island, New Zealand. Geophysical Journal Inter-
and related ignimbrites, Taupo Volcanic Zone. MSc national, 166, 469 –484.
thesis, University of Canterbury, Christchurch, New S UTTON , A. N., B LAKE , S. & W ILSON , C. J. N. 1995. An
Zealand. outline geochemistry of rhyolite eruptives from Taupo
R OGAN , M. 1982. A geophysical study of the Taupo Vol- volcanic centre, New Zealand. Journal of Volcanology
canic Zone, New Zealand. Journal of Geophysical and Geothermal Research, 68, 153– 175.
Research, 87, 4073– 4088. S UTTON , A. N., B LAKE , S., W ILSON , C. J. N. & C HAR-
R OWLAND , J. & S IBSON , R. H. 2001. Extensional fault LIER , B. L. A. 2000. Late Quaternary evolution of a
kinematics within the Taupo Volcanic Zone, New hyperactive rhyolite magmatic system: taupo volcanic
Zealand: soft-linked segmentation of a continental centre, New Zealand. Journal of the Geological
rift system. New Zealand Journal of Geology and Geo- Society, London, 157, 537– 552.
physics, 44, 271– 283. V ILLAMOR , P. & B ERRYMAN , K. 2001. A late Quaternary
S HERBURN , S., B ANNISTER , S. & B IBBY , H. 2003. extension rate in the Taupo Volcanic Zone, New
Seismic velocity structure of the central Taupo Volca- Zealand, derived from fault slip rate. New Zealand
nic Zone, New Zealand, from local earthquake tom- Journal of Geology and Geophysics, 44, 243–269.
ography. Journal of Volcanology and Geothermal V ILLAMOR , P. & B ERRYMAN , K. 2006a. Late Quaternary
Research, 122, 69–88. geometry and kinematics of faults at the southern ter-
S HERBURN , S., R EYNERS , M., E BERHART -P HILLIPS , D. mination of the Taupo Volcanic Zone, New Zealand.
& B ANNISTER , S. 2006. Seismic velocity structure of New Zealand Journal of Geology and Geophysics,
the Taupo Volcanic Zone. Abstract, Geosciences ’06 – 49, 1 –21.
Our planet, our future. Geological Society of New V ILLAMOR , P. & B ERRYMAN , K. 2006b. Evolution of the
Zealand Miscellaneous Publication, 122A, 76. southern termination of the Taupo Rift, New Zealand.
S KINNER , D. N. B. 1986. Neogene volcanism of the New Zealand Journal of Geology and Geophysics, 49,
Hauraki Volcanic region. In: S MITH , I. E. M. (ed.) 23–37.
Late Cenozoic volcanism in New Zealand. Royal W ILSON , C. J. N. 1993. Stratigraphy, chronology, styles
Society of New Zealand Bulletin, 23, 21– 47. and dynamics of late Quaternary eruptions from
S MITH , R. T. & H OUGHTON , B. F. 1995. Vent migration Taupo volcano, New Zealand. Philosophical Trans-
and changing eruptive style during the 1800 ka actions of the Royal Society of London A, 343,
Taupo eruption: new evidence from the Hatepe and 205–306.
Rotongaio phreatoplinian ashes. Bulletin of Volcanol- W ILSON , C. J. N. 1996. Taupo’s atypical arc. Nature, 379,
ogy, 57, 432– 439. 27–28.
S MITH , V. C., S HANE , P. & S MITH , I. E. M. 2002. W ILSON , C. J. N. 2001. The 26.5 ka Oruanui eruption,
Tephrostratigraphy and geochemical fingerprinting of New Zealand: an introduction and overview. Journal
the Mangaone Subgroup tephra beds, Okataina Volca- of Volcanology and Geothermal Research, 112,
nic Centre, New Zealand. New Zealand Journal of 133–174.
Geology and Geophysics, 45, 207– 219. W ILSON , C. J. N. & W ALKER , G. P. L. 1985. The Taupo
S PINKS , K. D. 2005. Rift Architecture and Caldera Vol- eruption, New Zealand. Philosophical Transactions of
canism in the Taupo Volcanic Zone, New Zealand. the Royal Society of London, Series A, 314, 199– 228.
PhD thesis, University of Canterbury, Christchurch, W ILSON , C. J. N., R OGAN , M. A., S MITH , I. E. M.,
New Zealand. N ORTHEY , D. J., N AIRN , I. E. & H OUGHTON , B. F.
S PINKS , K. D., A COCELLA , V., C OLE , J. W. & B ASSETT , 1984. Caldera volcanoes of the Taupo Volcanic
K. N. 2005. Structural control on volcanism and Zone, New Zealand. Journal of Geophysical Research,
caldera development in the transtensional Taupo Vol- 89, 8463–8484.
canic Zone, New Zealand. Journal of Volcanology W ILSON , C. J. N, H OUGHTON , B. F. & L LOYD , E. F.
and Geothermal Research, 144, 7– 22. 1986. Volcanic history and evolution of the Maroa–-
S TAGPOOLE , V. M. 1994. Interpretation of refraction Taupo area, Central North Island. In: S MITH , I. E. M.
seismic and gravity data across the eastern margin of (ed.) Late Cenozoic volcanism in New Zealand.
the Taupo Volcanic Zone. Geothermics, 23, 501–510. Royal Society of New Zealand Bulletin, 23, 194– 223.
S TAGPOOLE , V. M. & B IBBY , H. M. 1999. Residual W ILSON , C. J. N., H OUGHTON , B. F., M C W ILLIAMS ,
gravity anomaly map of the Taupo Volcanic Zone, M. O., L ANPHERE , M. A., W EAVER , S. D. &
New Zealand, 1:250 000, version 1.0. Institute of Geo- B RIGGS , R. M. 1995. Volcanic and structural evolution
logical & Nuclear Sciences Geophysical Map, 13. of Taupo Volcanic Zone, New Zealand: a review.
Institute of Geological & Nuclear Sciences Limited, Journal of Volcanology and Geothermal Research,
Lower Hutt, New Zealand. 68, 1 –28.
S TERN , T. A. 1986. Geophysical studies of the upper crust W OOD , C. P. 1992. Geology of the Rotorua Geothermal
within the Central Volcanic region. In: S MITH , I. E. M. System. Geothermics, 21, 25– 41.
(ed.) Late Cenozoic volcanism in New Zealand. Bulle- W OOD , C. P., B RAITHWAITE , R. L. & R OSENBERG ,
tin of the Royal Society of New Zealand, 23, 92–111. M. D. 2001. Basement structure, lithology and per-
S TRATFORD , W. R. & S TERN , T. A. 2004. Strong seismic meability at Kawerau and Ohaaki geothermal fields,
reflections and melts in the mantle of a continental New Zealand. Geothermics, 30, 461– 481.
back-arc basin. Geophysical Research Letters, 31, W RIGHT , I. C. 1990. Late Quaternary faulting of the off-
L06622. shore Whakatane Graben, Taupo Volcanic Zone,
New Zealand. New Zealand Journal of Geology and W RIGHT , I. C. 1993. Southern Havre Trough-Bay
Geophysics, 33, 245–256. of Plenty (New Zealand): structure and seismic
W RIGHT , I. C. 1992. Shallow structure and active tecton- stratigraphy of an active back-arc basin complex.
ism of an offshore continental back-arc spreading Chapter 11. In: B ALANCE , P. (ed.) South Pacific
system: the Taupo Volcanic Zone, New Zealand. Sedimentary Basins of the World 2. Elsevier,
Marine Geology, 103, 287–309. Amsterdam, 195–211.
Andean flat-slab subduction through time
VICTOR A. RAMOS & ANDRÉS FOLGUERA*
Laboratorio de Tectónica Andina, Universidad de Buenos Aires – CONICET
*Corresponding author (e-mail: andes@gl.fcen.uba.ar)
Abstract: The analysis of magmatic distribution, basin formation, tectonic evolution and
structural styles of different segments of the Andes shows that most of the Andes have experienced
a stage of flat subduction. Evidence is presented here for a wide range of regions throughout the
Andes, including the three present flat-slab segments (Pampean, Peruvian, Bucaramanga), three
incipient flat-slab segments (‘Carnegie’, Guañacos, ‘Tehuantepec’), three older and no longer
active Cenozoic flat-slab segments (Altiplano, Puna, Payenia), and an inferred Palaeozoic flat-
slab segment (Early Permian ‘San Rafael’). Based on the present characteristics of the Pampean
flat slab, combined with the Peruvian and Bucaramanga segments, a pattern of geological processes
can be attributed to slab shallowing and steepening. This pattern permits recognition of other older
Cenozoic subhorizontal subduction zones throughout the Andes. Based on crustal thickness, two
different settings of slab steepening are proposed. Slab steepening under thick crust leads to dela-
mination, basaltic underplating, lower crustal melting, extension and widespread rhyolitic volcan-
ism, as seen in the caldera formation and huge ignimbritic fields of the Altiplano and Puna
segments. On the other hand, when steepening affects thin crust, extension and extensive
within-plate basaltic flows reach the surface, forming large volcanic provinces, such as Payenia
in the southern Andes. This last case has very limited crustal melt along the axial part of the
Andean roots, which shows incipient delamination. Based on these cases, a Palaeozoic flat slab
is proposed with its subsequent steepening and widespread rhyolitic volcanism. The geological
evolution of the Andes indicates that shallowing and steepening of the subduction zone are thus
frequent processes which can be recognized throughout the entire system.
Introduction The objective of the present study is to charac-

terize geological processes linked to shallowing
The pioneer work of Barazangi & Isacks (1976, and steepening of the subduction zones and their
1979) described the first two well documented seg- geological consequences. We aim to characterize
ments along the Andes without late Cenozoic arc these parameters along the Andes in order to be
magmatism and adscribed them to flat-slab subduc- able to identify palaeo flat slab segments during
tion (Fig. 1). This cold subduction was associated the Phanerozoic. Based on these premises, three
with a subhorizontal Benioff zone identified in the palaeo flat slabs were identified in Cenozoic times.
retroarc area that was characterized by large and fre- Even further, it is speculated that a late Palaeozoic
quent intracrustal earthquakes driven by important flat slab could have developed in the Central
basement shortening. As a result, important foreland Andes. These new data enhance the importance of
basement uplifts took place in late Cenozoic times flat-slab subduction through time, and indicate that
giving rise to the present Sierras Pampeanas it is not an anomalous feature of the present-day
(González Bonorino 1950; Jordan et al. 1983a, b). margin, but has been an important feature of the
Another detailed seismotectonic study in the north- geological record and its frequency is higher
ern Andes recognized a flat-slab segment in the than expected.
northern Colombian Andes with similar character-
istics (Pennington 1981). Present flat-slab subduction segments
Multidisciplinary research performed during
the last two decades, mainly based on seismological Seismological data clearly show that there are three
and geological data on the continents, and oceano- distinct segments with horizontal subduction along
graphic studies in the adjacent areas, depict the the Andean margin: the Bucaramanga, Peruvian and
present setting of these three segments, where Pampean segments (Gütscher et al. 2000; Ramos
shallowing of the Benioff zone was closely related 1999a). There is also a striking transition to a sub-
to collision of aseismic ridges (Pilger 1981, 1984). horizontal subduction in the Ecuadorian Andes
However, it was only recently that geological (Gütscher et al. 1999a) that will be described to
evidence was obtained along the Andes showed show the initial geological processes linked to the
steepening of past subhorizontal subduction. beginning of shallowing. These segments will be
Geological Society, London, Special Publications, 327, 31– 54.
32 V. A. RAMOS & A. FOLGUERA
Fig. 1. Present flat-slab segments along the Andes (modified from Barazangi & Isacks 1976; Pennington 1981;
Ramos 1999a; Gütscher et al. 2000).
ANDEAN FLAT-SLAB SUBDUCTION THROUGH TIME 33
described from south to north, in order to move from the Mercedario (6850 m) and the La Ramada
better known segments to less known settings. (6400 m) among others, correspond to tectonically
uplifted areas with Miocene to Late Palaeozoic
Pampean flat-slab segment rocks above 6000 m (Ramos et al. 1996a). The
description of the geological evidence will encom-
This segment was one of the first where systematic pass the magmatic, sedimentological and structural
data were collected to reconstruct the tectonic history (Fig. 2), later linked to the oceanic features
history associated with flat-slab subduction in the associated with the shallowing.
Andes (Isacks et al. 1982; Jordan et al. 1983a, b).
The segment is located between 278 and 338300 S Magmatic evidence. The recognition of volcanic
latitude along the Pampean foreland. The highest gaps in the Quaternary volcanic arc of the Andes
segment of the Main Andes coincides with the emphasized the presence of cold subduction that
central part of the Pampean flat slab, where moun- coincides with the flat-slab segments (Barazanghi &
tain peaks, such as the Aconcagua (6967 m a.s.l.), Isacks 1976). Subsequent studies were able to
Fig. 2. Pampean flat-slab segment with indication of isobaths to the Nazca oceanic plate based on Cahill & Isacks
(1992) (compare with the Benioff geometry proposed by Pardo et al. 2002 and Alvarado et al. 2005a, b); main
basement uplifts of Sierras Pampeanas (Jordan et al. 1989), and location of the Precordillera fold and thrust belt
(Ramos et al. 2002).
recognize that the gap had existed since Late than 10 000 m of continental fluvial deposits, such
Miocene times (Jordan et al. 1983b). Detailed petro- as in the Neogene depocentre of Sierra de Los
graphic studies performed in the late Cenozoic arc Colorados at 298S lat. (Ramos 1999b).
show that the geochemical signature changes in The beginning of the broken foreland
the main arc through time (Kay et al. 1987), and stage coincided with the eastward advance of the
that the arc expanded towards the foreland region shallowing of the subducted slab beneath the
(Kay & Gordillo 1994). The geochemistry shows retroarc area. Sedimentological studies show
that the La/Yb ratios increased in Early to Late that the Early Miocene foreland basin was canniba-
Miocene arc rocks, at the same time that crustal lized during the Miocene, with the largest subsi-
stacking thickened the crust (Kay et al. 1991; Kay & dence rates experienced during the Middle
Mpodozis 2002; Litvak et al. 2007). Miocene inception of the Pampean flat-slab at
Geochronological data show that the main ande- these latitudes (Fig. 4).
sitic arc was active from 22–8.6 Ma (Fig. 3), Some basin remnants in the western interior
although volumes of erupted magmas were drasti- areas between the Frontal Cordillera and the
cally reduced throughout this time (Ramos et al. Precordillera, such as the Iglesia Valley basin,
1996a). A minor late rhyolitic eruption of Vacas were reactivated as piggy-back basins by out-of-
Heladas Ignimbrites at 7.67 Ma was the last activity sequence thrusts (Beer et al. 1990; Zapata &
in the area (Ramos et al. 1989). Subsequent hydro- Allmendinger 1996). There is also a great variation
thermal mineralization was widespread along the in the timing of deformation when the sedimentary
segment in El Indio, Valle del Cura and Maricunga record is compared from north to south. Synoro-
mineral districts (Mpodozis et al. 1995; Kay & genic deposition gets younger to the east and to
Mpodozis 2001). The latest activity east of the the south (Vergés et al. 2001), when comparing
previous main arc was the eruption of the Cerro de time of deposition along the Rı́o San Juan and
Vidrio rhyolitic dome dated at 2.0 + 0.2 Ma (Ar– Jachal further north. The same trend is regionally
Ar in glass) by Bissig et al. (2002) in Valle del observed along the entire segment (Jordan et al.
Cura. Both rhyolitic episodes are interpreted as 2001; Ramos 1999b).
minor melts of the crust.
The expansion of the arc magmatism is first Tectonic history. The timing of shortening in the
associated with a second dehydration front. At the Principal and Frontal cordilleras and the Precordil-
latitude of the Aconcagua for example, the main lera show some striking relations when analyzed
Middle Miocene arc was characterized by large in conjunction with: (1) the shortening rates of
volumes of andesites and dacites in the Principal the fold and thrust belts; (2) the propagation of the
Cordillera, whereas in the Precordillera at orogenic front; (3) the subsidence rate of the
c. 130 km east of the main arc, small volcanic adjacent foreland basin; and (4) the uplift of
centres and subvolcanic bodies were emplaced in Sierras Pampeanas (Fig. 5). The shortening of
Paramillos and Cerro Colorado (Kay et al. 1991). this fold-and-thrust belt was concentrated in a
The main arc, as well as the second volcanic front, thin-skinned belt within the Principal Cordillera
shifted eastward. The shifting and subsequent cessa- prior to the shallowing. This period recorded a
tion of the magmatic arc simultaneously moved shortening rate of 5.5–5.75 mm/a, and a slow
from west to east, and from north to south, ending propagation rate of 2.5 mm/a of the thrust or
at 5 + 0.5 Ma west of Sierra de Aconquija orogenic mountain front. The propagation rate
(278200 S lat.), 4.7 + 0.3 Ma at the Pocho volcanic increased to 13.3 mm/a soon after the beginning
field (318300 S lat.), and 1.9 + 0.2 Ma in Sierra del of shallowing, while the shortening was reduced to
Morro at 338100 S lat. (Ramos et al. 2002). 3.6 mm/a. This change from thin to thick skinned
shortening is also reflected in the subsidence rate
Sedimentary evolution. Several retroarc foreland of the foreland basin (Fig. 5).
basins were formed along the flat-slab segment This data – when compared with the tectonic
(Jordan 1984). Besides the general Andean trend evolution of the adjacent oceanic region – show
of east migration of the synorogenic depocentres close time and space relationships between collision
recorded from Late Cretaceous to Neogene times of the Juan Fernández aseismic ridge against the
through the entire Andes (Ramos 1999b), the flat- margin and the beginning of the shallowing of the
slab segment superimposed a special character. subducted slab (Yañez et al. 2001). The south and
East migration of the foreland system is linked to eastward shifting of the magmatic arc, the time of
fragmentation of the foreland basement (Jordan deformation and basin evolution accompany the
et al. 1989). Detailed magnetostratigraphic studies migration of the Juan Fernández ridge along and
show that subsidence rates were exceptional beneath the upper plate, as clearly demonstrated by
during the broken foreland stage (Reynolds et al. Pilger (1984), Gütscher et al. (2000) and Kay &
1990). Locally, some depocentres recorded more Mpodozis (2002). The most active neotectonic
Fig. 3. Evolution of arc magmatism through time in the Pampean flat-slab: (a) Representative ages after Ramos et al.
(2002) with indication of the isobath of 200 km depth corresponding to the oceanic slab; (b) Cross-section at crustal
scale showing the expansion and migration of the main volcanic centers during the shallowing of the oceanic slab.
Main elevations in the High Andes not related to the Quaternary volcanoes are also indicated.
Fig. 4. Subsidence rates in the proximal, intermediate and distal areas of the Bermejo broken foreland basin, with
indication of the beginning of flat-slab subduction at these latitudes (modified from Ramos 1999b). Seismostratigraphic
data after Reynolds et al. (1990).
area corresponds to the Pie de Palo uplift, an area of by Dorbath et al. (1986, 1991). This survey demon-
high intracrustal seismicity (Kadinsky-Cade strated that the subduction zone starts under the
et al. 1985; Regnier et al. 1992) and a western trench with a 308 dip until approximately 100 km
Sierras Pampeanas block where an average uplift depth (Fig. 6), where it becomes horizontal
rate of 1.0 mm/a during the last 3 Ma has been beneath the Eastern Cordillera and the Subandean
observed (Ramos et al. 2002; Siame et al. 2006a). zone (Dorbath et al. 1991). Pilger (1984) showed
Pie de Palo is just above the track of the Juan the kinematics between the Nazca Ridge collision
Fernández ridge, as indicated by the coincidence and the shallowing of the central Peru segment.
between high density of earthquake epicentres and This region was examined again by Gütscher et al.
the projection of the oceanic feature (Kirby et al. (1999b), who challenged the previous proposal
1996), and is located where the ridge is presently and instead of the collision of an aseismic ridge pro-
shallowing the subducting slab. posed that the large Peruvian flat-slab segment was
the result of the Nazca Ridge and the Inca Plateau
Peruvian flat-slab segment subduction. Precise timing of the Nazca Ridge col-
lision, and constraints in the length of the ridge,
This segment is encompassed between the Gulf of support that collision started at c. 11.2 Ma at about
Guayaquil at 58S and Arequipa at 148S latitudes. 118S, moving later to the present position, as
It has been described by Barazanghi & Isacks depicted by Hampel (2002). The segment north of
(1976, 1979) based on global data of the ISC catalo- this latitude requires a collision of a plateau or
gue, and with more precision using local networks other oceanic feature.
Fig. 5. The Aconcagua fold and thrust belt in the Central Andes at 328S latitude with variations on shortening and
propagation rates through time (after Ramos et al. 1996b and Hilley et al. 2004) and the subsidence rates in the foreland
basin after Irigoyen et al. (2002).
The Peruvian flat-slab segment shares many zircon ages as young as 6 Ma. The magmatic lull
common features with the Pampean flat slab. The following Nazca Ridge subduction began at the
second highest part of the Andes coincides with end of the Miocene. Most of the emplacement of
the Cordillera Blanca, with mountains such as the Cordillera Blanca Batholith and coeval ignim-
Huascarán (6778 m a.s.l.), which is only 110 m brites took place during the southern sweep of the
lower than the Aconcagua massif, and other Late Nazca Ridge (Aleman 2006).
Miocene granitic peaks over 6000 m. The Cordillera Neotectonics in the forearc where the Nazca
Blanca is in the central part of an important Ridge intersects the trench are described by
basement high, which includes the Marañón Macharé et al. (1986). Further support includes
Massif further to the east in the Eastern Cordillera. active tectonics and uplift in the foreland region in
The Cordillera de Marañón is a basement uplift the Fitzcarrald arch in the Subandean region,
that exposed middle crustal rocks very similar in where the aseismic ridge is being presently sub-
composition and metamorphic degree to the ducted. Evidence consists of a radial drainage
Sierras Pampeanas. The Peruvian segment also network and deformation of Pliocene–Recent
coincides with an area of no-arc volcanism, at fluvial deposits on both sides of this structural
least since latest Miocene times. Radiometric ages high (Espurt et al. 2007). Both forearc and foreland
document several Cenozoic pulses of eastward geology, together with the distribution of late
magmatic migration (Aleman 2006). The cessation Cenozoic arc volcanoes, highlight the relationships
(c. 12 Ma) of magmatism in the northern part of between aseismic ridge subduction, active uplift
the flat slab correlates with the complete subduction and cessation of magmatism.
of the Inca Plateau and the arrival of the Nazca
Ridge. As in the Pampean flat slab, the cessation Bucaramanga segment
of the main magmatic activity in the volcanic arc
is followed by the emplacement of minor crustal The early proposal of Pennington (1981), based on
melts of acidic composition. An example are the limited seismological data, showed a shallow sub-
granites of Cordillera Blanca where McNulty et al. duction zone beneath northern Colombia. This fact
(1998) and Giovanni et al. (2006) reported U– Pb has been confirmed by the seismological studies of
Fig. 6. (a) General features of the Peruvian flat slab based on Hampel (2002) and Aleman (2006). See the coincidence
between the projection of the Nazca Ridge into the foreland and the uplift of the Fitzcarrald arch and associated alluvial
fan (Espurt et al. 2007). (b) Geometry of Benioff– Wadatti zone beneath central Peru at 14–128S latitude (based on
Dorbarth et al. 1991).
Gütscher et al. (2000) in the Northern Andes north microplate and South America, after the middle
of 58N, and the analysis made by Corredor (2003), Miocene collision of the Chocó block (12 –13 Ma,
who shows the shallow subduction produced by Duque Caro 1990).
the recent subduction of the Caribbean plate The cessation of the late Cenozoic magmatic arc
beneath the Northern Andes. The dense concen- north of Cerro Bravo and Nevados de Ruiz (Mendez
tration of intracrustal earthquakes of the Bucara- Fajury 1989), as well as the intense widespread
manga nest (Fig. 7) is associated with basement neotectonic intracrustal activity, is better explained
deformation and uplift of the Eastern Cordillera, by the flat slab model of Gütscher et al. (2000).
characteristic of flat-slab subduction. However, an Their regional seismic tomography depicts a cold
alternative hypothesis was advanced by Taboada mantle and lower crust in this segment.
et al. (2000), where most of this intraplate The latest volcanic activity is exposed near
deformation at Bucaramanga was explained as a Boyacá in the retroarc region in the northern part
palaeo-Benioff zone associated with an old but of the Eastern Cordillera. The Tunja and Paipa
still active subduction between the Panamá volcanoes, among others, are associated with
Fig. 7. Seismic activity and main morphostructural units of the Bucaramanga flat slab based on Dimate et al. (2003).
Volcanic arc based on Mendez Fajury (1989) and retroarc volcanoes based on Cepeda (2004).
pyroclastic flows that range in age from 2.5–2.1 Ma wedge (Bourdon et al. 2003), and in the expansion
for the oldest eruption, to pyroclastic flows younger of the volcanic arc that coincides with the projection
than 1.5 Ma (Cepeda et al. 2004). This dominant of the Carnegie ridge, an aseismic oceanic ridge that
Pliocene–Quaternary explosive volcanic activity is now obliquely colliding against the margin
has a high-K rhyolitic composition that resembles (Gütscher et al. 1999a). The forearc crust is over-
the last magmatic activity, characterized by the thickened only in the segment where the Carnegie
rhyolitic dome described by Bissig et al. (2002) in ridge (Fig. 1) is colliding against the margin, as
the Pampean flat slab. The rhyolitic composition demonstrated by wide-angle seismic data recently
of both areas, the residual thermal fields, and the collected offshore (Gailler et al. 2007). This col-
mechanism of emplacement are very similar in lision is also related to the abnormal present uplift
both regions (Cepeda et al. 2004). This volcanic of the Cordillera Real and the Subandean block
activity is better explained by the shallowing of that controls the Pastaza alluvial megafan (Bés de
the subducted Bucaramanga Pacific slab than by Berc et al. 2005). Uplift rates during the Pleistocene
the inception of a new volcanic arc, as the result of 1.37– 1.4 cm/a, associated with an exhumation
of the subduction zone that is being developed of the late Cenozoic alluvial plain of 500 m, are
from the Caribbean margin of Colombia and closely linked to the Carnegie ridge collision
Venezuela (Audemard & Audemard 2001). (Christophoul et al. 2002; Baby et al. 2004). Impor-
Therefore, the flat slab hypothesis explains the tant intracrustal seismic activity is related to the
active present uplift of the northern segment of the basement structure of the Cutucu high. This uplift
Eastern Cordillera, the cessation of arc magmatism, may correlate with the fission track data for the
the neotectonic features associated with the tectonic Cordillera Real that shows more than 9 km uplift
inversion of previous rifts (Sarmiento-Rojas et al. in late Cenozoic times (Spikings et al. 2001).
2006), the large intracrustal seismicity (Dimate Another segment with incipient evidence of
et al. 2003) and the complex latest Cenozoic struc- shallowing is the Guañacos segment, located
ture of the Pie-de-monte Llanero (Martı́nez 2006). between 368 and 388300 S latitudes. It is character-
ized by strong neotectonic and intracrustal activity
in both: i) the forearc region at the Nahuel Buta
Incipient flat-slab subduction segments Cordillera and offshore Cretaceous –Paleogene
Arauco Basin (368300 –378300 S; Melnick et al.
One of the best lines of evidence of early-stage 2006a) and ii) the western retroarc zone at the
shallowing is documented inland of the collision Guañacos fold and thrust belt (368–388S; Folguera
of the Carnegie aseismic ridge (Gütscher 1999a). et al. 2004a). The two sectors correspond to
The volcanic arc of Colombia is composed of a ancient deformed belts that have been suddenly
line of individual volcanoes from Cerro Bravo at reactivated in Late Pliocene to Quaternary times.
58N to Cumbal at 28300 N latitude (Fig. 7). South of The offshore Arauco Basin, which was previously
the border with Ecuador, it changes to a complex uplifted in the Late Cretaceous, as indicated by
volcanic arc system, which is expanded towards fission track ages (Glodny et al. 2007), has been
the foreland. Active volcanoes are emplaced on shortened since 3.6 Ma at a rate of 0.8 mm a21 as
the Western Cordillera, the Inter-Andean Valley, the an eastward vergent fold and thrust belt. On the
Eastern Cordillera (or Cordillera Real) and in the other hand, recent neotectonics characterized the
Subandean zone across 120 km from the volcanic Guañacos fold and thrust belt, which was a Palaeo-
arc front. gene basin inverted during Late Miocene times. The
Individual volcanoes, such as the Cayambé and Pleistocene magmatic arc has migrated about 30 km
Quimsacocha volcanoes, show a trend from old to the east in this segment regarding the Pliocene
calc-alkalic volcanic rocks to a more recent new volcanic front. Petrological studies performed in
edifice with a typical adakitic signature (Beate the Cenozoic arc at these latitudes show crustal
et al. 2001; Samaniego et al. 2002). Although the thickening and subduction erosion, both processes
origin of this adakitic signal was early ascribed to consistent with shallowing of the subduction zone
slab melting, this has been questioned with their (Kay et al. 2005). Gravimetric studies show that
formation being attributed to melting of thickened the 368 –388300 S segment is characterized by a
continental crust or forearc subduction erosion long wavelength residual gravimetric anomaly that
(Ramos 2004). Both processes, crustal thickening can only be explained (see density model in
and forearc crustal erosion, are consistent with flat Alasonati Tašárová 2007; Hackney et al. 2006) by
subduction (Kay & Mpodozis 2002). the shallowing by 108 of the subduction angle of
The variation in the dip of the subducted slab has the Nazca subducted plate. Therefore, the anoma-
been addressed by the change in petrological lous concentration of crustal earthquakes linked to
characteristics, such as the depth of generation and unusual neotectonic activity in a 200 km wide sub-
degree of partial melting in the asthenospheric ducted segment, may indicate incipient shallowing
at the transition between the Central and Patagonian segments: (1) rapid cessation of the magmatic arc
Andes since Late Pliocene times. between 45 and 35 Ma; (2) widespread deformation
Another segment with an incipient flat slab is and crustal thickening in the Eastern Cordillera; (3)
the Transmexican volcanic belt in central Mexico, the tectonothermal Zongo San Gabán effect that per-
which is related to the collision of the Tehuantepec vasively resets the Ar– Ar ages along 450 km, over-
aseismic ridge, although a different mechanism for printing Permian and Triassic metamorphic rocks
uplift has been proposed (Ferrari 2006). A detailed with a cryptic 38 Ma age; and (4) no igneous
analysis of this segment is outside the scope of rocks of this age are known in this segment. This
this paper. effect was interpreted as the result of heat advection
by fluids at 38 Ma that predated the activity of the
sub-Andean fold and thrust belt (Farrar et al.
Past flat-slab subduction segments 1988). These processes were explained by a shal-
lowing of the subducted slab that became subhori-
There is a strong correlation between the segment zontal at about c. 35 Ma and lasted until c. 25 Ma.
with current arc volcanism in the Central Andes The steepening of the subduction zone was
(see central volcanic zone in Fig. 1) and the area evidenced by widespread bimodal volcanism
of past flat-slab subduction extending from southern where rhyolites and basalts cover a wide area. As
Peru to northern Argentina (Fig. 8). A summary of a result, great volumes of rhyolites up to 530 km3
the geological processes involved in the changes were spread on the present Altiplano and western
from normal to flat, and from flat to normal subduc- slope of Eastern Cordillera between 26 and 22 Ma
tion, will be discussed updating the proposal of (Sandeman et al. 1995). During flat subduction the
James & Sacks (1999) (also see Sebrier et al. 1988). overlying lithosphere is hydrated by dewatering of
the flat slab (James & Sacks 1999). Consequent
Altiplano flat-slab segment of Southern Peru steepening and expansion of the mantle wedge con-
trolled the flow of hot asthenosphere and melting
A period of flat-slab subduction was recorded in of the hydrated lithosphere beneath the Altiplano
southern Peru and northern Bolivia, between 148 and Eastern cordilleras. Volcanic arc retreat is
and 208S latitudes (James & Sacks 1999). The reflected by the shifting to the trench of the Tacaza
evidence was similar to the previous described arc between 29 and 15 Ma, the Upper Barroso arc
Fig. 8. Segments that recorded flat subduction in Late Eocene to Early Miocene times that correspond to the present
Central Volcanic Zone (based on James & Sacks 1999 and Kay et al. 1999).
(10–6 Ma) and Lower Barroso in the last 3 Ma to foreland basin (De Celles & Horton 2003) enabled
meet the present frontal arc during the Pleistocene. the onset of the deformation in the Subandean
The main points of these processes are the weak- region to be constrained to after 10 Ma.
ening of the lithosphere during steepening of the Again, the same processes indicate that strong
subduction, delamination of the lithosphere and deformation in the axial part of the Puna and
part of the lower crust (Kay & Kay 1993), and the Eastern cordilleras were related to shallowing of
collapse of the crust to form the Subandean fold the subduction zone, while steepening produced
and thrust belt. For further details, see James & important hot asthenospheric flow, which in contact
Sacks (1999) and Kay et al. (1999). with the hydrated lithosphere (Oncken et al. 2006),
led to important crustal and lithospheric delamina-
Puna flat-slab segment of southern Bolivia– tion. As a result, huge rhyolitic calderas and ignim-
britic fields are associated with thermal uplift and
northern Argentina the consequent horizontal collapse and weakening
This trend of shallowing progressed to the south, of the crust with the deformation of the Subandean
where another period of flat subduction was recog- belt (Isacks 1988; Kay et al. 1999; Beck & Zandt
nized between 208 and 248S (Kay et al. 1999). 2002; Garzione et al. 2006).
The shallowing took place between 18 and 12 Ma,
as recognized by the cessation of the magmatism, Payenia segment
crustal shortening and deformation of the southern
Altiplano and northern Puna. Precise timing of Arc related rocks were emplaced more than 550 km
the deformation established by Allmendinger et al. away from the trench during Late Miocene times,
(1997), Baby et al. (1995) and Oncken et al. from 348300 to 378450 S (Fig. 9), suggesting
(2006), together with the palaeogeography of the shallow subduction processes at that time (Kay
Fig. 9. Expansion of the magmatic arc during the Middle to Late Miocene showing the location of exhumed andesitic to
dacitic arc rocks on the San Rafael block. Subsequent extensional structures, within plate basaltic flows, and huge
rhyolitic calderas and ignimbritic flows along the main Andes suggest steepening of the subducted slab.
2001; Kay et al. 2006a, b). Intermediate positions of dacites and andesites emplaced in the San Rafael
the arc are located on the eastern slope of the Andes block between 13 and 4 Ma.
near the drainage divide area (Nullo et al. 2002) to In addition, the main phase of deformation in the
the east of the Late Oligocene arc, emplaced eastern section of the Malargüe fold and thrust belt
mainly on the western Andean slope. Easternmost at these latitudes has been constrained to 13–10 Ma
centres were emplaced over the San Rafael block, (Giambiagi et al. 2008), which indicates a genetic
a basement block that cannibalized the distal relationship between the initial phase of arc expan-
section of the Rio Grande foreland basin. The sion, uplift of the main Andes, sedimentation in the
uplift of this block was associated with the foreland adjacent foreland basin, and the breaking of the
migration of the Malargüe fold and thrust belt to the foreland area.
east (Kozlowski et al. 1993; Manceda & Figueroa During latest Miocene –Early Pliocene times,
1995). The San Rafael block was exhumed in Late this compressional crustal stage changed to an exten-
Miocene times (Dessanti 1956; González Dı́az sional regime with the development of extensional
1964; Polanski 1964; Yrigoyen 1993, 1994). The troughs across the area that had previously recorded
Middle Miocene age assigned to the synorogenic arc expansion (Fig. 10) (Bermúdez et al. 1993;
sequences at the San Rafael block (Soria 1984; Melnick et al. 2006b; Folguera et al. 2008). Arc
Marshall et al. 1986) points to a Late Miocene dynamics were characterized during this period by
exhumation that coincides with the age of the fast retreat to the present position on the western
Fig. 10. Distribution of Upper Palaeozoic magmatic rocks and deformation in the southern Central Andes (based
on Caminos 1979; Ramos et al. 1988; Varela et al. 1993; Mpodozis & Ramos 1989; Mpodozis & Kay 1990). Maximum
expansion of arc volcanic rocks in the Early Permian was followed by subsequent extensional regime associated
with the Choiyoi volcanic province.
flank of the Andes. Extensional deformation is arc expansion; and (4) resetting of remanent magne-
associated at depth with crustal attenuation as well tization in the area suggests abnormal lithospheric
as anomalous sublithospheric heating inferred by heating that preceeds eruption of intraplate melts.
teleseismic and tomographic analysis (Gilbert These facts point to a flat subduction cycle in
et al. 2006; Yuan et al. 2006). Gravimetric studies Early Permian times, followed by slab steepening
show high positive residual anomalies with areas and consequent orogenic collapse in the Late
submitted to extension, inferring an area of continu- Permian to Early Triassic, as proposed by Martı́nez
ous asthenospheric upwelling in coincidence with et al. (2006).
the area of previous arc expansion (Folguera et al.
2007a). This extensional setting hosted rhyolitic Sedimentary evolution. Late Carboniferous– Early
associations derived from crustal melts at the Permian 7000–8000 m thick marine to non-marine
highest collapsed sector of the Andes in the Las sequences are hosted along the eastern slope of the
Loicas trough (Fig. 9) (Hildreth et al. 1984, 1991, Principal Cordillera of Mendoza and San Juan
1999), whereas in foreland sectors it was associated (Fig. 10). Those are locally covering a Late Proter-
with poorly differentiated mantle derived products ozoic basement indicating an important erosional
(González Dı́az 1972; Rossello et al. 2002; Kay hiatus prior to their deposition. The broad area
et al. 2006b). uplifted in the Main Andes was the source of these
These two contrasting stages of deformation and sequences, which are characterized by coarsening-
arc dynamics, which occurred during the last 15 Ma up cycles. This episode of mountain building, known
between 348300 and 378450 S, point to a scenario as the San Rafael orogenic phase (280 –270 Ma:
in which progressive shallow subduction from Azcuy & Caminos 1987; Llambı́as et al. 1993;
15–5 Ma was followed by sudden slab steepening Cortés & Kleiman 1999), ended in the Lower
during the last 4 Ma, associated with the partial col- Permian with an important angular unconformity.
lapse of the orogen at these latitudes. From west to east these sequences were gathered
in the Loma de los Morteritos and El Plata for-
mations, with palynomorphs indicative of a Late
Palaeozoic flat-slab subduction segment? Carboniferous to Early Permian age. These units,
located on the eastern slope of the Frontal Cordil-
Palaeozoic deformations exhumed by Andean lera, formed the maximum depocentre of the Late
events through the Southern Central Andes have Palaeozoic in the region (Polanski 1958; Caminos
been connected to collisional episodes (Ramos 1965; Folguera et al. 2004b). To the east, Late
et al. 1984; Ramos 2004). Early Permian defor- Palaeozoic thicknesses fall in the western Precordil-
mations of the San Rafael tectonic phase have also lera region (Fig. 10), where several coarsening-up
been related to collision of an unidentified X tectonostratigraphic units do not reach 500 m.
terrane (Mpodozis & Kay 1990). This deformation These sequences, as determined by invertebrate
exhibits some peculiarities in the foreland sedimen- and palynomorph associations (Ottone 1987), are
tation and is associated with a phase of orogenic col- coeval with the magmatic rocks and the structural
lapse that led Mpodozis & Kay (1990) to propose deformation of the region.
structural instabilities after orogenic development. This main depocentre of several thousand metres
The analysis of the late Palaeozoic orogenies in flanked the Early Permian belt of deformation, and
other areas of Gondwana led Cawood & Buchan pinch out to the platform area. The foreland basin
(2007) to argue that deformation is not always started with shore sediments over which deltaic
related to a collisional event. Furthermore, in this bodies and turbiditic lobes prograded, ending with
segment of the Andes, little attention has been braided fluvial systems (Heredia et al. 2002). More-
paid to coeval arc dynamics, which constitutes a over, the dominance of westward palaeocurrents
direct indicator of Benioff zone variations through and lithoclasts of crystalline basement indicate
time. The Early Permian San Rafael tectonic phase that the basement may have been exhumed east of
is associated with unique processes that resemble the Early Permian orogenic front, potentially as an
more those of Andean tectonics than those occurred incipient Sierras Pampeanas system, similar to the
in Palaeozoic times at these latitudes (Ramos & present setting of the Pampean flat slab (Fig. 2).
Folguera 2007): (1) arc related volcanic assem- Lower Permian mesosiliceous lavas are part of
blages cover diachronically Early Permian com- the basal section of the Choiyoi Group. The upper
pressive deformation features; (2) Early Permian part of this unit accumulated either in the Frontal
arc abnormally expanded to extend through the Cordilleran or Precordilleran areas in a contrasting
entire region and probably its front shifted to the tectonic regime when compared to the basal
east; (3) extensional processes followed the main member. As revealed by the structural style of the
phase of orogenic building and intraplate rhyolitic Andean fold and thrust belt at these latitudes, the
sequences were erupted through the area of previous main basement thrusts are the result of tectonic
inversion of extensional faults that controlled the Main Andes that correspond to sutures formed as
main depocentres of the Choiyoi Group (Cristallini a result of Late Proterozoic to Early Palaeozoic
& Ramos 2000; Rodrı́guez Fernández et al. 1997). terrane amalgamation (Ramos 1988). These
sutures were reactivated with important strike-slip
Magmatic evidence. Several studies have pointed displacements in the Late Palaeozoic. The dominant
out that Lower Permian calc-alkaline series, right lateral displacements were caused by the
gathered with different names in the southern Pre- oblique convergence of the subducting Pacific
cordilleran region, have unconformably covered (Panthalassa) oceanic plate (Rapalini & Vilas
the San Rafael unconformity in the Frontal Cordil- 1991), which originated several deep transtensional
lera (Coira & Koukharsky 1976; Vilas & Valencio depocentres (Fernández Seveso et al. 1993;
1982; Cortés 1985; Kay et al. 1989; Rapalini & Fernández Seveso & Tankard 1995). The depocen-
Vilas 1991; Sato & Llambı́as 1993; Sotarello et al. tres are associated with alkaline eruptions typical
2005). In addition, other isolated minor volcanic of extensional intraplate settings (Koukharsky
bodies with similar chemical patterns and Early et al. 2001; Ramos et al. 2002), found in the
Permian age have been found to the east up to Paganzo Basin. Fernández Seveso et al. (1993)
250 km away from their westernmost position discuss the relation between Early Permian com-
(Fig. 10), on the Precordillera and Sierras Pampea- pressive thrusting in the western Andean sector
nas domains (Rubinstein & Koukharsky 1995; and transtension at the eastern foreland area. They
Castro de Machuca et al. 2007). The magmatic arc propose that the origin of the extension could have
was located mainly westward of the Frontal Cordil- been related to breaking up of the foreland basement
lera during the Carboniferous (Hervé et al. 1987), due to crustal downwarping as found in modern ana-
which implies a strong eastward shifting and expan- logues. This transtension in the Paganzo Basin
sion from the Late Carboniferous to the Early would be a passive response in the foreland area to
Permian (Rodrı́gez Blanco 2004). Early Permian orogenic loading of the San Rafael thrust wedge.
sequences are in turn separated by an erosional An alternative hypothesis would be to consider a
hiatus from an extensive intraplate rhyolitic associ- high partitioned subduction system where displace-
ation of the Choiyoi Group of Late Permian to Early ments perpendicular to the trench would have been
Triassic age (Rapalini & Vilas 1991). On geochem- absorbed in the San Rafael fold and thrust belt;
ical grounds, the plutonic and volcanic rocks of the lateral displacements imposed by oblique conver-
Choiyoi Group define a large within plate volcanic gence between plates would have been concentrated
province (Kay et al. 1989; Mpodozis & Ramos and localized in ancient lithospheric boundaries
1989) that covers important sectors of the Main (Rapalini & Vilas 1991; Fernández Seveso et al.
Andes and Precordillera regions (Fig. 10). The 1993). In this context, high oblique convergence
area of Early Permian arc expansion coincides and strong coupling associated with shallow sub-
with a phase of extensional collapse with peak duction would be the condition for the development
igneous activity around 260 –240 Ma at these lati- of a high strain partitioned subduction regime
tudes (Martı́nez 2004). during Late Carboniferous –Early Permian times.
Arc expansion, stacking of the western sector
Tectonic history. A wide volcanic arc, in excess of of the fold and thrust belt during San Rafael
200 km, developed in Late Carboniferous–Early tectonic phase, formation of foreland basins, and
Permian times and has been exhumed along the transtensional to transpressional reactivation of
Pampean flat slab zone. The volcanic sequences Proterozoic-early Palaeozoic sutures in the foreland
are interfingered in the west with a 7000–8000 area, ended in the tectonic wedge collapse. As a
thick turbiditic to deltaic succession whose eastern- result, a multitude of rift systems were filled by
most section is preserved at the eastern Frontal the Choiyoi Group.
Cordillera (Fig. 10). Towards the east, the volcanic Rotation of half grabens produced erosional
rocks were emplaced over folded and thrust unconformities that separate Early Permian volca-
sequences deformed during the San Rafael orogenic nics from the rest of the late Palaeozoic sequence.
phase. The sedimentary depocentre, characterized This zone of orogenic collapse coincides with the
by the stacking of coarsening-up cycles, was area of arc expansion and San Rafael orogenic com-
affected by the Early Permian deformation. This pressional deformations, suggesting a common
basin was formed during the arc expansion stage, mechanism. Therefore, slab steepening and conse-
with its subsidence controlled by orogenic loading. quent asthenospheric injection in the broadened
It experienced rapid thinning towards the east in asthenospheric wedge, after shallow subduction,
the present eastern Precordillera. are the mechanisms proposed for the origin of
Regional analysis of the Late Carboniferous to the anomalously voluminous rhyolitic magmas of
Early Permian tectonics shows some striking facts. the Choiyoi Group and its extensional tectonic
There are major crustal anisotropies east of the control. As a result, delamination of the lower
crust took place after thickening and eclogitization and the Gondwana margin. The larger area and
during the San Rafael compressive phase. Sublitho- elevations up to 5250 m reached by the Sierras
spheric heating due to slab steepening explains the Pampeanas in the Sierra de Aconquija could be
massive crustal melting, as the lower crust was related to the more segmented nature of the base-
directly in contact with the rising asthenospheric ment with several sutures and ophiolitic belts reacti-
flux (Martı́nez et al. 2006). vated first as extensional faults during the opening
of the South Atlantic, and later, as a thrust during
the shallowing of the oceanic slab (Ramos et al.
Normal to flat-slab transition 2002). Other segments such as the Bucaramanga
Several examples of different ages and several are related to the reactivation and uplift of the
distinct segments of the Andes show that the Eastern Cordillera of Colombia by tectonic inver-
transit from normal subduction to flat-slab subduc- sion of extensional faults, partially coinciding with
tion is associated with a series of events: sutures (Cortés et al. 2006; Ramos & Moreno
2006). Even in a small segment as the Payenia flat-
Migration of the volcanic front and expansion of the slab, the uplift of the San Rafael Block coincided
arc magmatism. It is important to note that with the maximum expansion of the arc. There is a
migration of the arc is indicated by the location of close relationship between arc migration, thermal
the largest volume of magmatic rocks; although weakening of the crust and basement uplift
magmatism in the previous setting may last for (James & Sacks 1999; Ramos et al. 2002) during
several million years, but with insignificant volu- the shallowing of the oceanic slab. Some pervasive
mes. Such migration involves a decreasing volume tectonothermal effects, such as the Zongo –San
of magmatic rocks that parallel the decline of Gabán (Farrar et al. 1988) and the San Rafael effect
dehydration in the subducted slab. This migration (Rapalini & Astini 2005), are associated with
can be correlated to crustal weakening of the this stage.
foreland and subsequent faulting. Geochemical sig- Basin subsidence. The increase in subsidence has a
nature of these magmas changes with the distance to clear relationship with the approximation of the
the trench as well as the depth of generation (Kay & thrust front, as shown in several Subandean basins
Mpodozis 2002). Final products may be as far as (Irigoyen et al. 2002; Jordan 1995). However, the
600 km from the trench, as in the Bucaramanga subsidence achieves a critical collapse when the
segment (Jaramillo & Rojas 2003; Cepeda et al. basement is broken and maximum thicknesses
2004), and up to 750 km in the Pampean flat slab are obtained. This is seen in the Pampean flat slab,
(Kay & Gordillo 1994). where more than 10 000 m of sediments in the
Uplift of the Main Andes. Tectonic uplift is well synorogenic deposits of the Bermejo foreland
documented in the Peruvian segment and in the basin have been reported by Ramos et al. (2002).
Pampean flat slab, where the Cordillera Blanca There are incomplete records in other segments,
and the High Cordillera of Mendoza and San Juan but De Celles & Horton (2003) described several
encompass the highest sectors of the Andes with thousand metres in the Oligocene and Early
the Huascarán (6778 m) and the Aconcagua Miocene of the Altiplano. The Payenia segment
(6967 m) mountains. The main difference between nicely depicts the migration and cannibalization of
these two segments is that the Peruvian one registers the previous basins until the broken foreland stage
some extensional collapse of the Cordillera Blanca is reached.
(Siame et al. 2006b), while the Aconcagua shows
no evidence of extension (Ramos et al. 1996b). Flat to normal slab transition
This could imply that extension is more related to
slab buoyancy from ridge subduction of the pre- On the other hand, the processes related to the tran-
thickened continental crust, as proposed by sition from flat-slab to normal subduction are less
McNulty & Farber (2002), than to orogenic collapse well known, but have interesting characteristics:
in the sense of Dewey (1988).
Rhyolitic flare-up. One of the first results of stee-
Broken foreland. Although the Sierras Pampeanas pening of the subducted oceanic slab is the presence
is one of the most typical features of the Pampean of large crustal melts that are suddenly erupted over
flat slab (Jordan et al. 1983a, b), most other areas the flat-slab area in thick continental crust (Kay
have recorded basement uplifts. The Peruvian et al. 1999). Recent studies demonstrate that these
segment is characterized by the Marañón Massif large lower crustal melts are associated with litho-
(3400 m a.s.l.), a basement uplift of the Eastern Cor- spheric removal, sinking of the eclogitized lower
dillera produced in Late Miocene times almost crust, and crustal delamination, as earlier proposed
along the suture between an allochthonous terrane by Kay & Kay (1993).
Fig. 11. Segments that have experienced shallowing of the subduction zone during Cenozoic times along the Andes.
Note the almost continuous outline of flat-slabs.
Thermal uplift. This effect is a direct consequence 408 and 438S has experienced some shallowing
of the lithospheric removal (Isacks 1988), although during late Paleogene times (de Ignacio et al.
it has only been well documented in the Altiplano– 2001). There is no obvious trend or wave of shallow-
Puna segment (Whitman et al. 1996; Allmendinger ing, except among the Altiplano, Puna and Pampean
et al. 1997). Different geophysical tools have been segments, where there is some defined younging to
used to confirm this evidence (see review in the south. The other segments, at the present level of
Oncken et al. 2006). Evidence of thermal uplift knowledge, show a random inception.
has not been documented in other segments.
Reduced uplift in a thermal weakened area has Funding for this research was provided by grants ANPCYT
been recently proposed in the Payenia segment PICT 14144, CONICET PIP 5965 and UBACyT 160.
with reduced geophysical datasets by Folguera The authors are grateful to C. Mpodozis (Sipetrol, Chile)
and S. M. Kay (Cornell University, USA) for many years
et al. (2007b).
of fruitful discussions on these topics, as well as to the
researchers of Laboratorio de Tectónica Andina (Univer-
Extensional regime. The onset of the steepening of sity of Buenos Aires). The critical reviews of
the subducted slab in some areas is associated with B. McNulty and P. Cawood are greatly appreciated.
the vertical collapse by extension of the previous
contracted structures. This is seen in the Payenia
segment, where the pre-Miocene peneplain, uplifted
in the Late Miocene, is segmented by normal faults References
(Ramos & Folguera 2005). Although the Puna has A LASONATI T AŠÁROVÁ , Z. 2007. Towards understand-
evidence of Pliocene extensional faulting that has ing the lithospheric structure of the southern Chilean
been interpreted in different ways (Allmendinger subduction zone (368S–428S) and its role in the
et al. 1997), it is important here to note that exten- gravity field. Geophysical Journal International, doi:
sion occurs immediately after the thermal uplift of 10.1111/j.1365-246X.2007.03466.
the area. A LEMAN , A. M. 2006. The Peruvian flat-slab. Backbone of
the Americas, Asociación Geológica Argentina – Geo-
logical Society of America Symposium, Abstract with
Intense deformation shifted to the foreland. The Programs, p. 17, Mendoza.
best example of migration of deformation that post- A LLMENDINGER , R. W., J ORDAN , T. E., K AY , S. M. &
dates thermal uplift and some extension in the axial I SACKS , B. L. 1997. The evolution of the Altiplano-
area, is the formation of the southern Subandean Puna Plateau of the Central Andes. Annual Reviews
fold and thrust belt. The spatial and temporal Earth Planetary Sciences, 25, 139– 174.
relationships are clearly seen in southern Bolivia A LVARADO , P., C ASTRO DE M ACHUCA , B. & B ECK , S.
(Beck & Zandt 2002). In some other segments, 2005a. Comparative seismic and petrographic crustal
this relationship is not evident, although in the study between the Western and Eastern Sierras
Pampeanas region (318S). Revista de la Asociación
Payenia flat slab, the Guañacos fold and thrust belt
Geológica Argentina, 60, 787– 796.
was developed after the emplacement of calderas A LVARADO , P., B ECK , S., Z ANDT , G., A RAUJO , M. &
and rhyolitic domes as well as the San Rafael T RIEP , E. 2005b. Crustal deformation in the south-
block. This belt along the axis of the Andean Cordil- central Andes backarc terranes as viewed from
lera has evidence of neotectonic activity (Folguera regional broad-band seismic waveform modeling.
et al. 2004a). Geophysical Journal International, 163, 580 –598.
A UDEMARD , F. E. & A UDEMARD , F. A. 2001. Structure
Widespread mafic within plate floods. The segment of the Mérida Andes, Venezuela: relation with the
with thin crust, even after gentle shortening, shows South America– Caribbean geodynamic interaction.
Tectonophysics, 345, 299– 327.
an important basaltic flood linked to the inception
A ZCUY , C. & C AMINOS , R. 1987. Diastrofismo. In:
of the steepening. These basaltic floods, in the A RCHANGELSKY , S. (ed.) El sistema Carbonı́fero en
Payenia segment, indicate a mantle-derived poorly- la República Argentina. Academia Nacional de
evolved magma of mafic composition and within Ciencias, Córdoba, 239–251.
plate signature (Kay et al. 2006a, b). Acidic rocks B ABY , P., M ORETTI , I. ET AL . 1995. Petroleum system of
of Pliocene to Quaternary age in this area are scarce the northern and central Bolivian Subandean zone. In:
and are mainly small crustal melts as in Cerro T ANKARD , A. J., S UÁREZ S ORUCO , R. & W ELSINK ,
Peceño, in the San Rafael Block. H. J. (eds) Petroleum basins of South America.
In conclusion, it is interesting to show that when American Association of Petroleum Geologists,
Memoir, 62, 445– 458.
the present and past Cenozoic segments that had
B ABY , P., R IVADENEIRA , M. & B ARRAGÁN , R. 2004.
experienced flat-slab subduction are posted along La Cuenca de Oriente: Geologı́a y Petróleo. Travaux
the Andes (Fig. 11), an almost continuous belt of de l’Institut Francais d’Etudes Andines, 144, 1– 195.
flat-slabs is outlined. The area that does not show B ARAZANGI , M. & I SACKS , B. 1976. Spatial distribution
evidence is Patagonia, although some studies pos- of earthquakes and subduction of the Nazca plate
tulate that the northern Patagonian massif between beneath South America. Geology, 4, 686– 692.
B ARAZANGI , M. & I SACKS , B. 1979. Subduction of the foreland basin: the Ecuadorian Oriente Basin from
Nazca plate beneath Peru evidence from spatial distri- Eocene to Oligocene times. Tectonophysics, 345,
bution of earthquakes. Geophysical Journal of Royal 281– 298.
Astronomic Society, 57, 537– 555. C EPEDA , N. 2004. Estratigrafı́a de las vulcanitas asocia-
B EATE , B., M ONZIER , M., S PIKINGS , R., C OTTEN , J., das al Volcán de Paipa, municipios de Paipa y Tuta,
S ILVA , J., B OURDON , E. & E ISSEN , J.-P. 2001. Mio- Departamento de Boyacá, Colombia. PhD thesis,
Pliocene adakite generation related to flat subduction Universidad Nacional de Colombia, (unpublished),
in southern Ecuador: the Quimsacocha volcanic Bogotá.
center. Earth and Planetary Science Letters, 192, C EPEDA , H., P ARDO , N. & J ARAMILLO , J. M. 2004. The
561–570. Paipa Volcano, S.A. International Association of
B ECK , S. L. & Z ANDT , G. 2002. The nature of orogenic Volcanology and Chemistry of the Earth Interior.
crust in the central Andes. Journal of Geophysical IAVCEI International meeting 2004, Poster Session,
Research, 107, doi: 10.1029/2000JB000124. Pucón.
B ÈS DE BERC , S., S OULA , J. C., B ABY , P., S OURIS , M., C OIRA , B. & K OUKHARSKY , M. 1976. Efusividad tardı́o-
C HRISTOPHOUL , F. & R OSERO , J. 2005. Geomorphic hercı́nica en el borde oriental de la cordillera Frontal,
evidence of active deformation and uplift in a modern zona del arroyo del Tigre, provincia de Mendoza,
continental wedge-top-foredeep transition: example of República Argentina. Congreso Geológico Chileno,
the eastern Ecuadorian Andes. Tectonophysics, 399, Abstracts, 1, 105–123.
351–380. C ORREDOR , F. 2003. Seismic strain rates and distributed
B EER , J. A., A LLMENDINGER , R. W., F IGUEROA , D. E. & continental deformation in the northern Andes and
J ORDAN , T. 1990. Seismic stratigraphy of a Neogene three-dimensional seismotectonics of northwestern
Piggyback basin, Argentina. American Association South America. Tectonophysics, 372, 147– 166.
Petroleum Geologist, Bulletin, 74, 1183– 1202. C ORTÉS , J. 1985. Vulcanitas y sedimentitas lacustres en la
B ERMÚDEZ , A., D ELPINO , D., F REY , F. & S AAL , A. base del Grupo Choiyoi al sur de l estancia de Tambil-
1993. Los basaltos de retroarco extraandinos. In: los, Provincia de Mendoza. República Argentina. IV
R AMOS , V. A. (ed.) Geologı́a y Recursos Naturales Congreso Geológico Chileno, Abstracts, 1, 89– 108.
de Mendoza. XII8 Congreso Geológico Argentino y C ORTÉS , J. & K LEIMAN , L. 1999. La orogenia Sanrafaé-
II8 Congreso de Exploración de Hidrocarburos, Rela- lica en los Andes de Mendoza. 148 Congreso Geológio
torio, 161–172, Buenos Aires. Argentino, Abstracts, 1, 31.
B ISSIG , T., C LARK , A. H. & L EE , J. K. W. 2002. Cerro de C ORTÉS , M., C OLLETTA , B. & A NGELIER , J. 2006. Struc-
Vidrio rhyolitic dome: evidence for Late Pliocene ture and tectonics of the central segment of the Eastern
volcanism in the central Andean flat-slab region, Cordillera of Colombia. Journal of South American
Lama-Veladero district, 298200 S, San Juan Province, Earth Sciences, 21, 437–465.
Argentina. Journal of South American Earth Sciences, C RISTALLINI , E. & R AMOS , V. A. 2000. Thick-skinned
15, 571–576. and thin-skinned thrusting in the La Ramada fold and
B OURDON , E., E ISSEN , J.-P., G UTSCHER , M.-A., thrust belt: crustal evolution of the High Andes of
M ONZIER , M., H ALL , M. L. & C OTTEN , J. 2003. Mag- San Juan, Argentina (328S). Tectonophysics, 317,
matic response to early aseismic ridge subduction: the 205– 235.
Ecuadorian margin case (South America). Earth and D ESSANTI , R. 1956. Hoja Cerro Diamante 1 sheet
Planetary Science Letters, 205, 123– 138. 1: 250,000. Provincia de Mendoza. Servicio Nacional
C AHILL , T. & I SACKS , B. L. 1992. Seismicity and the Minero Geológico, Boletı́n, 85, 1 –79.
shape of the subducted Nazca plate. Journal of Geo- D E C ELLES , P. G. & H ORTON , B. 2003. Early to middle
physical Research, 97, 17503–17529. Tertiary foreland basin development and the history
C AMINOS , R. 1965. Geologı́a de la vertiente oriental del of Andean crustal shortening in Bolivia. Geological
Cordón del Plata, Cordillera Frontal de Mendoza. Society of America, Bulletin, 115, 58–77.
Revista de la Asociación Geológica Argentina, 20, D E I GNACIO , C., L ÓPEZ , I., O YARZUN , R. & M ÁRQUEZ ,
351–392. A. 2001. The northern Patagonia Somuncura plateau
C AMINOS , R. 1979. Cordillera frontal. In: T URNER , basalts: a product of slab-induced, shallow astheno-
J. C. M. (ed.) Segundo Simposio de Geologı́a Regional spheric upwelling? Terra Nova, 13, 117– 121.
Argentina. Academia Nacional de Ciencias, Córdoba, D EWEY , J. F. 1988. Extensional collapse of orogens.
1, 397–453. Tectonics, 7, 1123–1139.
C ASTRO DE M ACHUCA , B., C ONTE -G RAND , A., D IMATE , C., R IVERA , L. ET AL . 2003. The 19 January
M EISSL , E., P ONTORIERO , S., S UMAY , C. & 1995 Tauramena (Colombia) earthquake: geometry
M ORATA , D. 2007. Manifestaciones del magmatismo and stress regime. Tectonophysics, 363, 159 –180.
neopaleozoico (Carbonı́fero superior-Pérmico) en la D ORBATH , C., D ORBATH , L., C ISTERNAS , A.,
sierra de la Huerta, sierras Pampeanas occidentales, D EVERCHERE , J., D IAMENT , M., O COLA , L. &
Provincia de San Juan. Caracterización petrológica M ORALES , M. 1986. On crustal seismicity of the
de los pórfidos Marayes Viejo y El Arriero. Revista de Amazonian foothill of the central Peruvian Andes.
la Asociación Geológica Argentina, 62, 447–459. Geophysical Research Letters, 13, 1023– 1026.
C AWOOD , P. A. & B UCHAN , C. 2007. Linking accretion- D ORBATH , L., D ORBATH , C., J IMENEZ , E. & R IVERA , L.
ary orogenesis with supercontinent assembly. Earth- 1991. Seismicity and tectonic deformation in
Science Reviews, 82, 217– 256. the Eastern Cordillera and the sub-Andean zone of
C HRISTOPHOUL , F., B ABY , P. & D ÁVILA , C. 2002. Central Perú. Journal of South American Earth
Stratigraphic responses to a major tectonic event in a Sciences, 4, 13–24.
D UQUE C ARO , H. 1990. The Chocó block in the north- mantle lithosphere. Earth and Planetary Science
western corner of South America: structural, tectonos- Letters, 241, 543– 556.
traigraphic and paleogeographic implications. Journal G IAMBIAGI , L., B ECHIS , F., G ARCÍA , V. & C LARK , A.
of South American Earth Sciences, 3, 71–84. 2008. Temporal and spatial relationships of
E SPURT , N., B ABY , P. ET AL . 2007. How does the Nazca thick- and thin-skinned deformation: a case study
Ridge subduction influence the modern Amazonian from the Malargüe fold and thrust belt, Southern
foreland basin? Geology, 35, 515–518. Central Andes. Tectonophysics, Special Issue ISAG,
F ARRAR , E., C LARK , A. H., K ONTAK , D. J. & 459(1– 4), 123 –139.
A RCHIBAL , D. A. 1988. Zongo-San Gabán zone: G ILBERT , H., B ECK , S. & Z ANDT , G. 2006. Lithospheric
eocene foreland boundary of the Central Andean and upper mantle structure of central Chile and
orogen, northwest Bolivia and southeast Peru. Argentina. Geophysical Journal International, 165,
Geology, 16, 55–58. 383, doi: 10.1111/j.1365-246X.2006.02867.x.
F ERNÁNDEZ S EVESO , F. & T ANKARD , A. 1995. G LODNY , J., G RÂFE , K., E CHTLER , H. & R OSENAU , M.
Tectonics and stratigraphy of the Late Paleozoic 2007. Mesozoic to Quaternary continental margin
Paganzo basin of Western Argentina and its regional dynamics in South-Central Chile (36–428S): the
implications. In: T ANKARD , A. J., S UÁREZ apatite and zircon fission track perspective. Inter-
S ORUCO , R. & W ELSINK , H. J. (eds) Petroleum national Journal of Earth Sciences (Geologische
Basins of South America. American Association of Rundschau), doi: 10.1007/s00531-007-0203-1.
Petroleum Geologists, Memoir, 62, 285– 301. G ONZÁLEZ B ONORINO , F. 1950. Algunos problemas
F ERNÁNDEZ S EVESO , F., P ÉREZ , M., B RISSON , I. & geológicos de las Sierras Pampeanas. Revista de la
A LVAREZ , L. 1993. Sequence stratigraphy and Asociación Geológica Argentina, 5, 81– 110.
tectonic analysis of the Paganzo basin, Western G ONZÁLEZ D ÍAZ , E. F. 1964. Rasgos geológicos y evolu-
Argentina. 128 International Congress on the ción geomorfológica de la Hoja 27 d (San Rafael) y
Carboniferous-Permian systems (Buenos Aires), zona occidental vecina (Provincia de Mendoza).
Comptes Rendus, 2, 223 –260. Revista de la Asociación Geológica Argentina, 19,
F ERRARI , L. 2006. Laramide and Neogene shallow 151–188.
subduction in Mexico: constraints and contrasts. G ONZÁLEZ D ÍAZ , E. F. 1972. Descripción geológica de la
Backbone of the Americas, Asociación Geológica Hoja 30d Payún-Matrú, Provincia de Mendoza. Direc-
Argentina – Geological Society of America Sym- ción Nacional de Geologı́a y Minerı́a, Boletı́n, 130,
posium, Abstract with Programs, p. 17, Mendoza. 1– 88.
F OLGUERA , A., R AMOS , V. A., H ERMANNS , R. L. & G ÜTSCHER , R., M ALAVIEILLE , J., L ALLEMEND , S. &
N ARANJO , J. 2004a. Neotectonics in the foothills of C OLLOT , J. Y. 1999a. Tectonic segmentation of the
the southernmost central Andes (378– 388S): evidence North Andean margin: impact of the Carnegie Ridge
of strike-slip displacement along the Antiñir-Copahue Collision. Earth and Planetary Science Letters, 168,
fault zone. Tectonics, 23, TC5008, doi: 10.1029/ 255–270.
2003TC001533. G ÜTSCHER , M. A., O LIVET , J. L., A SLANIAN , D.,
F OLGUERA , A., E TCHEVERRÍA , M. ET AL . 2004b. Hoja E ISSEN , J. P. & M AURY , R. 1999b. The ‘lost Inca
Geológica Potrerillos 3369–15. Servicio Geológico Plateau’: cause of flat subduction beneath Peru?
Minero Argentino, Boletı́n, 301, 1 –142. Earth and Planetary Science Letters, 171, 335– 341.
F OLGUERA , A. & I NTROCASO , A. 2007a. Crustal attenu- G ÜTSCHER , M. A., S PAKMAN , W., B IJWAARD , H. &
ation in the Southern Andean retroarc determined E NGDAHL , E. R. 2000. Geodynamic of flat subduction:
from gravimetric studies (388–398300 S): The Lonco- seismicity and tomographic constraints from the
Luán astenospheric anomaly. Tectonophysics, 439, Andean margin. Tectonics, 19, 814–833.
129– 147, doi: 10.1016/j.tecto.2007.04.001. H ACKNEY , R., E CHTLER , H. ET AL . 2006. The segmented
F OLGUERA , A., A LASONATI T AŠÁROVÁ , S., H ESE , F., overriding plate and coupling at the South-Central
H ACKNEY , R., G OTZE , H., S CHMIDT , S. & R AMOS , Chilean margin (368–428S). In: O NCKEN , O.,
V. A. 2007b. Crustal attenuation along the Andean C HONG , G. ET AL . (eds) The Andes – Active Subduc-
retroarc zone 368– 398S from gravimetric análisis. tion Orogeny. Frontiers in Earth Sciences Series, 1,
208 Colloquium on Latinoamerican Earth Sciences, Springer, Berlin.
Abstracts, 38–41, Kiel. H AMPEL , A. 2002. The migration history of the Nazca
F OLGUERA , A., B OTTESI , G., Z APATA , T. & R AMOS , ridge along the Peruvian active margin: a
V. A. 2008. Crustal collapse in the Andean back-arc re-evaluation. Earth and Planetary Science Letters,
since 2 Ma: Tromen volcanic plateau, Southern 203, 665 –679.
Central Andes (368400 – 378300 S). Tectonophysics H EREDIA , N., R ODRÍGUEZ F ERNÁNDEZ , L. R., G ALLAS-
(Special Issue on Andean Geodynamics), 459(1– 4), TEGUI , G., B USQUETS , P. & C OLOMBO , F. 2002.
140– 160, doi: 10.1016/j.tecto.2007.12.013. Geological setting of the Argentine Frontal Cor-
G AILLER , A., C HARVIS , P. & F LUEH , E. R. 2007. Seg- dillera in the flat-slab segment (308000 –318300 S
mentation of the Nazca and South American plates latitude). Journal of South American Earth Sciences,
along the Ecuador subduction zone from wide angle 15, 79–99.
seismic profiles. Earth and Planetary Science Letters, H ERVÉ , F., G ODOY , E., P ARADA , M., R AMOS , V. A.,
260, 444– 464. R APELA , C., M PODOZIS , C. & D AVIDSON , J. 1987.
G ARZIONE , C. N., M OLNAR , P., L IBARKIN , J. C. & A general view on the Chilean Argentinian Andes,
M AC F ADDEN , B. J. 2006. Rapid late Miocene rise of with emphasis on their early history. In: M ONGER , J.
the Bolivian Altiplano: evidence for removal of & F RANCHETEAU , J. (eds) Circum-Pacific orogenic
belts and evolution of the Pacific Basin. American Pampeanas, Argentina. Journal of South American
Geophysical Union, Geoynamics Series, 18, 97–113. Earth Sciences, 2, 207–222.
H ILDRETH , W., G RUNDER , A. & D RAKE , R. 1984. The J ORDAN , T. E., S CHULENEGGER , F. & C ARDOZO , N.
Loma Seca Tuff and the Calabozos Caldera: a major 2001. Unsteady and spatially variable evolution of
ash-flow and caldera complex in the southern Andes the Neogene Andean Bermejo Foreland basin,
of Central Chile. Geological Society of America, Argentina. Journal of South Americn Earth Sciences,
Bulletin, 95, 45–54. 14, 775 –798.
H ILDRETH , W., D RAKE , R., G ODOY , E. & M UNIZAGA , F. K ADINSKY -C ADE , K., R EILINGER , R. & I SACKS , B.
1991. Bobadilla caldera and 1.1 Ma ignimbrite 1985. Surface deformation associated with the
at Laguna del Maule, Southern Chile. VI8 Congreso November 23, 1977, Caucete, Argentina, earthquake
Geológico Chileno Actas, 62– 63. sequence. Journal of Geophysical Research, 90,
H ILDRETH , W., F IERSTEIN , J., G ODOY , E., D RAKE , R. & 12 691– 12 700.
S INGER , B. 1999. The Puelche volcanic field: exten- K AY , R. W. & K AY , S. M. 1993. Delamination and dela-
sive Pleistocene rhyolite lava flows in the Andes of mination magmatism. Tectonophysics, 219, 177 –189.
Central Chile. Revista Geológica de Chile, 26, K AY , S. 2001. Tertiary to recent magmatism and tectonics
275–309. of the Neuquén basin between 36805 0 and 388S
H ILLEY , G. E., S TRECKER , M. R. & R AMOS , V. A. 2004. latitude. Repsol-YPF Buenos Aires. Unpublished
Growth and erosion of fold-and-thrust belts with Report, 77 p.
application to the Aconcagua fold-and-thrust belt, K AY , S. M. & G ORDILLO , C. E. 1994. Pocho volcanic
Argentina. Journal of Geophysical Research 109, rocks and the melting of depleted continental litho-
B01410, doi: 10.1029/2002JB002282. sphere above a shallowly dipping subduction zone in
I RIGOYEN , M. V., B UCHAN , K. L., V ILLENEUVE , M. E. & the Central Andes. Contributions to Mineralogy and
B ROWN , R. L. 2002. Cronologı́a y significado tectó- Petrology, 117, 25–44.
nico de los estratos sinorogénicos neógenos aflorantes K AY , S. M. & M PODOZIS , C. 2001. Central Andean ore
en la región de Cacheuta-Tupungato, provincia de deposits linked to evolving shallow subduction
Mendoza. Revista de la Asociación Geológica systems and thickening crust. GSA Today, 11(3), 4– 9.
Argentina, 57, 3– 18. K AY , S. M. & M PODOZIS , C. 2002. Magmatism as a probe
I SACKS , B. 1988. Uplift of the Central Andean plateau and to the Neogene shallowing of the Nazca plate beneath
bending of the Bolivian orocline. Journal Geophysical the modern Chilean flat-slab. Journal of South
Research, 93, 3211–3231. American Earth Sciences, 15, 39– 57.
I SACKS , B., J ORDAN , T., A LLMENDINGER , R. & R AMOS , K AY , S. M., M AKSAEV , V., M OSCOSO , R., M PODOZIS ,
V. A. 1982. La segmentación tectónica de los Andes C. & N ASI , C. 1987. Probing the evolving Andean
Centrales y su relación con la placa de Nazca subduc- lithosphere: Mid-late Tertiary magmatism in Chile
tada. V8 Congreso Latinoamericano de Geologı́a, (298–308300 S) over the modern zone of subhorizontal
Actas, 3, 587–606. subduction. Journal Geophysical Research, 92(B7),
J AMES , D. E. & S ACKS , S. 1999. Cenozoic formation of 6173– 6189.
the Central Andes: a geophysical perspective. In: K AY , S., R AMOS , V. A., M PODOZIS , C. & S RUOGA , P.
S KINNER , B. ET AL . (eds) Geology and Mineral Depos- 1989. Late Paleozoic to Jurassic silicic magmatism at
its of Central Andes. Society of Economic Geology, the Gondwana margin: analogy to the Middle Protero-
Special Publication, 7, 1 –25. zoic in North America? Geology, 17, 324– 328.
J ARAMILLO , J. M. & R OJAS , P. 2003. Depósitos de K AY , S. M., M PODOZIS , C., R AMOS , V. A. & M UNI-
cenizas volcánicas en la Cordillera Oriental. 98 Con- ZAGA , F. 1991. Magma source variations for mid to
greso Colombiano de Geologı́a, Resúmenes, p. 150, late Tertiary volcanic rocks erupted over a shallowing
Medellı́n. subduction zone and through a thickening crust in the
J ORDAN , T. E. 1984. Cuencas, volcanismo y acortamien- Main Andean Cordillera (28–338S). In: H ARMON ,
tos cenozoicos, Argentina, Bolivia y Chile, 208288 R. S. & R APELA , C. (eds) Andean Magmatism and
lat. S. 98 Congreso Geológico Argentino, Actas, 2, its Tectonic Setting. Geological Society of America,
7–24. Special Paper, 265, 113–137.
J ORDAN , T. E. 1995. Retroarc foreland and related basins. K AY , S. M., M PODOZIS , C. & C OIRA , B. 1999. Neogene
In: S PERA , C. & I NGERSOLL , R. V. (eds) Tectonics of magmatism, tectonism, and mineral deposits of
Sedimentary Basins. Blackwell Scientific, Cambridge, the Central Andes (228S to 338S). In: S KINNER , B.
MA, 331–362. ET AL . (eds) Geology and Mineral Deposits of Central
J ORDAN , T. E., I SACKS , B., R AMOS , V. A. & Andes. Society of Economic Geology, Special Publi-
A LLMENDINGER , R. W. 1983a. Mountain cation, 7, 27– 59.
building in the Central Andes. Episodes, 1983(3), K AY , S. M., G ODOY , E. & K URTZ , A. 2005. Episodic arc
20– 26. migration, crustal thickening, subduction erosion, and
J ORDAN , T. E., I SACKS , B. L., A LLMENDINGER , R. W., magmatism in the south-central Andes. Bulletin of
B REWER , J. A., R AMOS , V. A. & A NDO , C. J. the Geological Society of America, 117, 67–88.
1983b. Andean tectonics related to geometry of sub- K AY , S. M., M ANCILLA , O. & C OPELAND , P. 2006a.
ducted Nazca plate. Geological Society of America, Evolution of the Backarc Chachahuén volcanic
Bulletin, 94, 341–361. complex at 378S latitude over a transient Miocene
J ORDAN , T., Z EITLER , P., R AMOS , V. A. & G LEADOW , shallow subduction zone under the Neuquén Basin.
A. J. W. 1989. Thermochronometric data on the In: K AY , S. M. & R AMOS , V. A. (eds) Evolution of
development of the basement peneplain in the Sierras an Andean Margin: A Tectonic and Magmatic View
from the Andes to the Neuquén Basin (358– 398S lat.). changes in Paleo-Benioff zone. Backbone of the
Geological Society of America, Special Paper, 407, Americas, Asociación Geológica Argentina –
215– 246. Geological Society of America Symposium, Abstract
K AY , S. M., B URNS , M. & C OPELAND , P. 2006b. Upper with Programs. Mendoza, 1– 77.
Cretaceous to Holocene Magmatism over the M ARTINEZ , J. A. 2006. Structural evolution of the Llanos
Neuquén basin: evidence for transient shallowing of foothills, Eastern Cordillera, Colombia. Journal of
the subduction zone under the Neuquén Andes (368S South American Earth Sciences, 21, 510–520.
to 388S latitude). In: K AY , S. M. & R AMOS , V. A. M C N ULTY , B. A. & F ARBER , D. 2002. Active detachment
(eds) Evolution of an Andean Margin: A Tectonic faulting above the Peruvian flat-slab. Geology, 30,
and Magmatic View from the Andes to the Neuquén 567–570.
Basin (358– 398S lat.). Geological Society of M C N ULTY , B. A., F ARBER , D. L., W ALLACE , G. S.,
America, Special Paper, 407, 19– 60. L OPEZ , R. & P ALACIOS , O. 1998. Role of plate kin-
K IRBY , S., E NGDAHL , E. R. & D ENLINGER , R. 1996. ematics and plate-slip-vector partitioning in continen-
Intermediate-depth intraslab earthquakes and arc tal magmatic arcs; evidence from the Cordillera
volcanism as physical expressions of crustal and Blanca, Peru. Geology, 26, 827 –830.
uppermost mantle metamorphism in subducting M ELNICK , D., B OOKHAGEN , B., E CHTLER , H. &
slabs. In: B EBOUT , G. E., S CHOLL , D. W., K IRBY , S TREICKER , M. 2006a. Coastal deformation and
S. H. & P LATT , J. P. (eds) Subduction Top to great subduction earthquakes, Isla Santa Marı́a, Chile
Bottom. Geophysical Monograph Series, 96, 195–214. (378S). Geological Society of America, Bulletin, 118,
K OUKHARSKY , M., T ASSINARI , C., B RODTKORB , M. & 1463– 1480.
L EAL , P. 2001. Basaltos del Neopaleozoico-Triásico M ELNICK , D., C HARLET , F., E CHTLER , H. & B ATIST , M.
temprano? en las sierras del norte de Córdoba y de 2006b. Incipient axial collapse of the main cordillera
Ambargasta, Sierras Pampeanas Orientales: Petrograı́a and strain partitioning gradient between the central
y edades K/Ar. Revista de la Asociación Geológica and Patagonian Andes, Lago Laja, Chile. Tectonics,
Argentina, 56, 400 –403. 25, TC 5004, doi: 10.1029/2005TC001918: 1 –22.
K OZLOWSKI , E., M ANCEDA , R. & R AMOS , V. A. 1993. M ÉNDEZ F AJURY , R. A. 1989. Catálogo de volcanes
Estructura. In: R AMOS , V. A. (ed.) Geologı́a y Recur- activos de Colombia. Boletı́n Geológico Ingeominas,
sos Naturales de Mendoza. XII8 Congreso Geológico 30(3), 1 –75.
Argentino y II8 Congreso de Exploración de Hidrocar- M PODOZIS , C. & K AY , S. 1990. Provincias magmáticas
buros, Relatorio, 235–256, Buenos Aires. ácidas y evolución tectónica del Gondwana: Andes
L ITVAK , V. D., P OMA , S. & M AHLBURG K AY , S. 2007. Chilenos (288 –318S). Revista Geológica de Chile,
Paleogene and Neogene magmatism in the Valle del 17, 153– 180.
Cura region: new perspective on the evolution of the M PODOZIS , C. & R AMOS , V. A. 1989. The Andes of Chile
Pampean flat-slab, San Juan province, Argentina. and Argentina. In: E RICKSEN , G. E., C AÑAS P INO-
Journal of South American Earth Sciences, 24, CHET , M. T. & R EINEMUD , J. A. (eds) Geology of
117– 137. the Andes and its relation to Hydrocarbon and
L LAMBÍAS , E., K LEIMAN , L. & S ALVARREDI , J. 1993. Mineral Resources. Circumpacific Council for
El magmatismo gondwánico. In: R AMOS , V. A. (ed.) Energy and Mineral Resources, Earth Sciences
Geologı́a y Recursos Naturales de Mendoza. XII8 Series, 11, 59–90.
Congreso Geológico Argentino y II8 Congreso de M PODOZIS , C., C ORNEJO , P., K AY , S. M. & T ITTLER , A.
Exploración de Hidrocarburos, Relatorio, 53– 64, 1995. La franja de Maricunga: sı́ntesis de la evolución
Buenos Aires. del frente volcánico Oligoceno–Mioceno de la zona
M ACHARÉ , J., S EBRIER , M., H UAMÁN , D. & M ERCER , sur de los Andes Centrales. Revista Geológica de
J. L. 1986. Tectónica cenozoica de la margen peruana. Chile, 22, 273– 313.
Boletı́n de la Sociedad Geológica del Perú, 76, 45– 77. N ULLO , F., S TEPHENS , G., O TAMENDI , J. & B ALDAUF ,
M ANCEDA , R. & F IGUEROA , D. 1995. Inversion of the P. 2002. El volcanismo del Terciario superior del sur
Mesozoic Neuquén rift in the Malargüe fold-thrust de Mendoza. Revista de la Asociación Geológica
belt, Mendoza, Argentina. In: T ANKARD , A. J., Argentina, 57, 119–132.
S UÁREZ , R. & W ELSINK , H. J. (eds) Petroleum O NCKEN , O., C HONG , G. ET AL . (eds) 2006. The Andes –
Basins of South America. American Association of Active Subduction Orogeny. Frontiers in Earth
Petroleum Geologists, Memoir, 62, 369– 382. Sciences Series, Springer, Berlin.
M ARSHALL , L., D RAKE , R. & C URTISS , G. 1986. O TTONE , E. G. 1987. Estudios Bioestratigráficos y
40
K– 40Ar calibration of Late Miocene-Pliocene Paleoambientales de la Formación Santa Máxima,
Mammal-bearing Hayquerı́as and Tunuyán For- Paleozoico Superior, provincia de Mendoza, Repúb-
mations, Mendoza Province, Argentina. Journal of lica Argentina. PhD thesis, Universidad de Buenos
Paleontology, 60, 448 –457. Aires.
M ARTÍNEZ , A. 2004. Secuencias volcánicas permo- P ARDO , M., C OMTE , D. & M ONFRET , T. 2002. Seismo-
triásicas de los cordones del Portillo y del Plata, Cor- tectonic and stress distribution in the central Chile sub-
dillera Frontal, Mendoza: su interpretación tectónica. duction zone. Journal of South American Earth
PhD thesis, Universidad de Buenos Aires. Sciences, 15, 11–22.
M ARTÍNEZ , A., R ODRÍGUEZ B LANCO , L. & R AMOS , P ENNINGTON , W. D. 1981. Subduction of the Eastern
V. A. 2006. Permo-Triassic magmatism of the Panama Basin and seismotectonics of Northwest
Choiyoi Group in the cordillera Frontal of Mendoza, South America. Journal of Geophysical Research,
Argentina: geological variations associated with 86, 10753–10770.
P ILGER , R. H. 1981. Plate reconstructions, aseismic R AMOS , V. A., C RISTALLINI , E. & P ÉREZ , D. J. 2002.
ridges, and low angle subduction beneath the The Pampean flat-slab of the Central Andes. Journal
Andes. Geological Society of America, Bulletin, 92, of South American Earth Sciences, 15, 59–78.
448–456. R APALINI , A. E. & A STINI , R. A. 2005. La remagnetiza-
P ILGER , R. H. 1984. Cenozoic plate kinematics, subduc- ción sanrafaélica de la Precordillera en el Pérmico:
tion and magmatism: South American Andes. nuevas evidencias. Revista de la Asociación Geológica
Journal of the Geological Society, London, 141, Argentina, 60, 290– 300.
793–802. R APALINI , A. & V ILAS , J. 1991. Tectonic rotations in the
P OLANSKI , J. 1958. El bloque varı́scico de la Cordillera Late Paleozoic continental margin of southern South
Frontal de Mendoza. Revista de la Asociación Geoló- America determined and dated by paleomagnetism.
gica Argentina, 12, 165 –196. Geophysical Journal International, 107, 333–351.
P OLANSKI , J. 1964. Descripción geológica de la Hoja 26c, R EGNIER , M., C HATELAIN , J. L., S MALLEY , R., M ING
La Tosca (Prov. de Mendoza). Dirección Nacional de C HIU , J., I SACKS , B. L. & P UEBLA , N. 1992.
Geologı́a y Minerı́a, Boletı́n, 101, 1 –86. Seismotectonic of the Sierra Pie de Palo, a basement
R AMOS , V. A. 1988. Tectonics of the Late Proterozoic – block uplift in the Andean foreland, Argentina.
Early Paleozoic: a collisional history of Southern Bulletin of Seismological Society of America, 82,
South America. Episodes, 11(3), 168– 174. 2549– 2571.
R AMOS , V. A. 1999a. Plate tectonic setting of the Andean R EYNOLDS , J. H., J ORDAN , T. E., J OHNSON , N. M.,
Cordillera. Episodes, 22(3), 183–190. D AMANTI , J. F. & T ABBUTT , K. D. 1990. Neogene
R AMOS , V. A. 1999b. Los depósitos terciarios sinorogén- deformation on the flat-subduction segment of the
icos de la región andina. In: C AMINOS , R. (ed.) Geolo- Argentine-Chilean Andes: magneteostratigraphic con-
gı́a Argentina. Instituto de Geologı́a y Recursos straints from Las Juntas, La Rioja Province, Argentina.
Minerales, Anales, 29, 651–682. Geological Society of America, Bulletin, 102,
R AMOS , V. A. 2004. Cuyania, an exotic block to Gond- 1607– 1622.
wana: review of a historical success and the present R ODRÍGUEZ B LANCO , L. 2004. Geologı́a, estructura y
problems. Gondwana Research, 7, 1009–1026. volcanismo permotriásico entre los los Cerros
R AMOS , V. A. & F OLGUERA , A. 2005. Tectonic evolution Puntudo y Colorado, San Juan. Trabajo Final de
of the Andes of Neuquén: constraints derived from the Licenciatura, Universidad de Buenos Aires, unpub-
magmatic arc and foreland deformation. In: V EIGA , G. lished, Buenos Aires.
ET AL . (eds) The Neuquén Basin: A case study in R ODRÍGUEZ F ERNÁNDEZ , L., H EREDIA , N., E SPINA , R.
sequence stratigraphy and basin dynamics. Geological & C EGARRA , M. 1997. Estratigrafı́a y estructura de
Society, London, Special Publications, 252, 15–35. los Andes Centrales Argentinos entre los 308 y 318
R AMOS , V. A. & F OLGUERA , A. 2007. Evidence of de latitud sur. Acta Geológica Hispánica, 32, 51–75.
flat and steep subduction on the Andes: geological R OSSELLO , E., C OBBOLD , P., D IRAISON , M. & A RNAUD ,
processes and tectonic inferences. 208 Colloquium on N. 2002. Auca Mahuida (Neuquén Basin, Argentina): a
Latinoamerican Earth Sciences, Abstracts, 17– 18, Quaternary shield volcano on a hydrocarbon-
Kiel. producing substrate. 58 International Symposium on
R AMOS , V. A. & M ORENO , M. 2006. Tectonic evolution Andean Geodynamics, Extended Abstracts, 549 –552.
of the Colombian Andes, Editorial. Journal of South R UBINSTEIN , N. & K OUKHARSKY , M. 1995. Edades K/
American Earth Sciences, 21, 319– 321. Ar del volcanismo neopaleozoico en la Precordillera
R AMOS , V. A., J ORDAN , T. E., A LLMENDINGER , R. W., noroccidental sanjuanina (308S; 698030 W). Revista de
K AY , S. M., C ORTÉS , J. M. & P ALMA , M. A. 1984. la Asociación Geológica Argentina, 50, 270– 272.
Chilenia: Un terreno alóctono en la evolución paleo- S AMANIEGO , P., M ARTIN , H., R OBIN , C. & M ONZIER ,
zoica de los Andes Centrales. 98 Congreso Geológico M. 2002. Transition from calc-alkalic to adakitic mag-
Argentino, Actas, 2, 84– 106. matism at Cayambe volcano, Ecuador: insights into
R AMOS , V., M UNIZAGA , F. & M ARÍN , G. 1988. Las slab melts andmantle wedge interactions. Geology,
riolitas neopaleozoicas de la sierra de la Huerta (pro- 30, 967 –990.
vincia de San Juan): evidencias de una metalogénesis S ANDEMAN , H. A., C LARK , A. H. & F ARRAR , E. 1995.
aurı́fera gondwánica en Sierras Pampeanas. 38 Con- An integrated tectono-magmatic model for the evol-
greso Nacional de Geologı́a Económica, Actas, 1, ution of the Southern Peruvian Andes (138– 208S)
149–159. since 55 Ma. International Geology Review, 37,
R AMOS , V. A., K AY , S. M., P AGE , R. N. & M UNIZAGA , F. 1039– 1073.
1989. La Ignimbrita Vacas Heladas y el cese del volca- S ARMIENTO -R OJAS , L. F., V AN W ESS , J. D. & C LOE-
nismo en el Valle del Cura, provincia de San Juan. TINGH , S. 2006. Mesozoic transtensional basin
Revista de la Asociación Geológica Argentina, 44, history of the Eastern Cordillera, Colombian Andes:
336–352. inferences from tectonic models. Journal of South
R AMOS , V. A., A GUIRRE U RRETA , M. B. ET AL . 1996a. American Earth Sciences, 21, 383 –411.
Geologı́a de la Región del Aconcagua, Provincias de S ATO , A. & L LAMBÍAS , E. 1993. El Grupo Choiyoi, pro-
San Juan y Mendoza. Dirección Nacional del Servicio vincia de San Juan: equivalente efusivo del Batolito de
Geológico, Anales, 24, 1– 510. Colangüil. 128 Congreso Geológico Argentino y 28
R AMOS , V. A., C EGARRA , M. & C RISTALLINI , E. 1996b. Congreso de Exploración de Hidrocarburos, Actas,
Cenozoic tectonics of the High Andes of west-central 4, 156– 165.
Argentina, (308 –368S latitude). Tectonophysics, 259, S ÉBRIER , M., M ERCIER , J., M ACHARÉ , D., B ONNOT , J.
185–200. & C ABRERA Y , J. L. BLANC . 1988. The state of
stress in an overriding plate situated above a flat-slab: tectonic evolution of the Andean Precordillera at
the Andes of central Peru. Tectonics, 7, 895–928. 318S, Argentina. Journal of South American Earth
S IAME , L. L., B ELLIER , O. & S EBRIER , M. 2006a. Active Sciences, 14, 735– 750.
tectonics in the Argentine Precordillera and Western V ILAS , J. & V ALENCIO , D. 1982. Implicancias geo-
Sierras Pampeanas. Revista de la Asociación Geoló- dinámicas de los resultados paleomagnéticos de forma-
gica Argentina, 61, 604– 619. ciones asignadas al Paleozoico tardı́o-Mesozoico
S IAME , L. L., S ÉBRIER , M., B ELLIER , O. & B OURLES , D. temprano del centro-oeste Argentino. 58 Congreso
2006b. Can cosmic ray exposure dating reveal the Latinoamericano de Geologı́a, Actas, 3, 743–758.
normal faulting activity of the Cordillera Blanca W HITMAN , D., I SACKS , B. & K AY , S. M. 1996. Litho-
Fault, Peru? Revista de la Asociación Geológica spheric structure and along-strike segmentation of the
Argentina, 61, 536 –544. central Andean plateau. Tectonophysics, 259, 29–40.
S ORIA , M. F. 1984. Vertebrados fósiles y edad de la For- Y AÑEZ , G., R ANERO , G. R., VON H UENE , R. & D IAZ , J.
mación Aisol, provincia de Mendoza. Revista de la 2001. Magnetic anomaly interpretation across a
Asociación Geológica Argentina, 38, 299– 306. segment of the southern Central Andes (32– 348S):
S OTARELLO , G., B ELVIDERI , I., DE M ACHUCA , B., implications on the role of the Juan Fernández ridge
F ERNÁNDEZ , G., M ACHUCA , E. & M C G UINTY , W. in the tectonic evolution of the margin during the
2005. Avances en el conocimiento y exploración del Upper Tertiary. Journal of Geophysical Research,
depósito epitermal de oro y plata ‘Casposo’, Departa- 106, 6325–6345.
mento de Calingasta, Provincia de San Juan. 88 Con- Y RIGOYEN , M. 1993. Los depósitos sinorogénicos terciar-
greso Argentino de Geologı́a Económica, Actas, ios. In: R AMOS , V. A. (ed.) Geologı́a y Recursos Nat-
299– 305. urales de Mendoza. XII8 Congreso Geológico
S PIKINGS , R. A., W INKLER , W. & S EWARD , D. 2001. Argentino y II8 Congreso de Exploración de Hidrocar-
Along-strike variations in the thermal and tectonic buros, Relatorio, 123– 148, Buenos Aires.
response of the continental Ecuadorian Andes to the Y RIGOYEN , M. 1994. Revisión estratigráfica del Neógeno
collision with heterogeneous oceanic crust. Earth and de las Huayquerı́as de Mendoza septentrional,
Planetary Science Letters, 186, 57–73. Argentina. Ameghiniana, 31, 125– 138.
T ABOADA , A., R IVERA , L. A. ET AL . 2000. Geodynamics Y UAN , X., A SCH , G. ET AL . 2006. Deep seismic images of
of the northern Andes: subductions and intracontinen- the Southern Andes. In: K AY , S. M. & R AMOS , V. A.
tal deformation (Colombia). Tectonics, 19, 787–813. (eds) Late Cretaceous to Recent Magmatism and Tec-
V ARELA , R., C INGOLANI , C., D ALLA S ALDA , L., tonism of the Southern Andean Margin at the Latitude
A RAGÓN , E. & T EIXEIRA , W. 1993. Las monzodiori- of the Neuquen Basin (36–398S). Geological Society
tas y monzogabros de Cacheuta, Mendoza: edad, of America, Special Paper, 407, 61– 72.
petrografı́a e implicancias tectónicas. 128 Congreso Z APATA , T. R. & A LLMENDINGER , R. W. 1996. The
Geológico Argentino y 28 Congreso de Exploración thrust front zone of the Precordillera thrust belt,
de Hidrocarburos, Actas, 4, 75– 80. Argentina: a thick-skinned triangle zone. American
V ERGÉS , J., R AMOS , E., S EWARD , D., B USQUETS , P. & Association of Petroleum Geologists, Bulletin, 80,
C OLOMBO , F. 2001. Miocene sedimentary and 359–381.
Middle Miocene Chiapas fold and thrust belt of Mexico: a result
of collision of the Tehuantepec Transform/Ridge with the
Middle America Trench
J. J. MANDUJANO-VELAZQUEZ1 & J. DUNCAN KEPPIE2*
1
Subdirección de Exploración y Producción, Instituto Mexicano del Petróleo. Avenida Lázaro
Cárdenas Norte 152, San Bartolo Atepehuacán, Delegación Gustavo A. Madero,
C. P. 07730, México D. F.
2
Departamento de Geologı́a Regional, Instituto de Geologı́a, Universidad Nacional
Autónoma de México, C. P. 04510, México D. F.
*Corresponding author (e-mail: keppie@glinx.com)
Abstract: The Middle Miocene, thin-skinned, Chiapas fold-and-thrust belt (Gulf of Mexico–
southeastern Mexico– Belize) consists of WNW-trending folds and thrusts, and East– West sinis-
tral transcurrent faults resulting from N608E shortening. Balanced cross-sections indicate that
shortening varies from 48% (SW) to c. 8% (NE) with a total shortening of 106 km, and that
thrusts merge into a basal décollement in the Callovian salt horizon. The Middle Miocene age
of the deformation is synchronous with collision of the Tehuantepec Transform/Ridge with the
Middle America Trench off Chiapas. The presently exposed Tehuantepec Transform/Ridge
varies from a transform fault across which the age of the oceanic crust changes producing a step
(down to the east) to a ridge resulting from compression following a change in plate motion and
a series of seamounts. On the other hand, the earthquake data show that the part of the Tehuantepec
Transform/Ridge subducted during the past 5 Ma is a step with no accompanying ridge. Whereas
collision of a ridge segment with the trench is inferred to be responsible for the 13– 11 Ma defor-
mation in the upper plate, its termination at 11 Ma suggests an along-strike transition to a step.
Collision of the Tehuantepec Transform/Ridge also appears to have terminated arc magmatism
along the Pacific coast of Chiapas. The similarity between the petroleum-producing, Cantarell
structure in the Sonda de Campeche and the buried foldbelt in the Sierra de Chiapas suggests
there is considerable further hydrocarbon potential.
Introduction better understanding of modern analogs, such as the

Chiapas foldbelt, which forms the topic of this paper.
The Middle Miocene, Chiapas–Tabasco–Campeche The genesis of the Chiapas foldbelt has been
fold-and-thrust belt (Chiapas foldbelt hereafter) in linked to the inferred eastward relative movement
Mexico (Fig. 1) is remarkable because it formed of the Chortis block (mainly Honduras) from a pos-
during a restricted time interval (c. 2.5 Ma), is of ition off southwestern Mexico during the last
restricted size (300 600 km) and is host to some 45 Ma (Ross & Scotese 1988; Pindell et al. 1988,
of the largest petroleum resources in the world. Its 2006; Schaaf et al. 1995). The contrast in duration
origin is enigmatic. Origins inferred for fold-and- of these events (c. 2.5 versus 45 Ma) poses a
thrust belts elsewhere in the world vary depending problem for this model. Alternatively, Keppie &
on their age. The older ones are related to colli- Moran-Zenteno (2005) proposed that the fold-and-
sions, such as continent– continent, arc –continent thrust belt resulted from collision of the Tehuan-
and terrane –continent collisions. On the other tepec Ridge with the Middle America Trench. To
hand, younger fold-and-thrust belts bordering sub- resolve the origin of the Chiapas foldbelt, we present
ducting oceanic plate containing little or no continen- an analysis of seismic reflections data from two
tal lithosphere have been attributed to other factors, NE –SW composite sections across the Sonda de
such as a shallowing of the dip of the Benioff zone, Campeche and the Sierra de Chiapas (Fig. 1) that
an increased convergence rate and absolute plate document the thin-skinned nature of the deforma-
motion, a change in the convergence vector, overrid- tion, whereas well data indicate its restricted dur-
ing of an active plume or seamount chain, an increase ation. This information is then considered in a plate
in friction between the overriding and subducting tectonic context. The final integration of all the infor-
plates, and the thickness of the upper continental mation suggests that collision of the Tehuantepec
plate (Sobolev & Babeyko 2005). Recognition of Ridge with the trench may have induced short-lived
these genetic factors in ancient orogens requires a deformation of the overriding plate.
56 J. J. MANDUJANO-VELAZQUEZ & J. D. KEPPIE
Fig. 1. Map of eastern Mexico showing the locations of the Chiapas foldbelt, the reflection seismic sections, the
Tehuantepec Transform/Ridge, the Chiapas massif and various other tectonic and geographical elements.
Geological setting Trench that marks the boundary with the northward
subducting Cocos Plate (Fig. 1). The Chiapas fold-
The NW-trending Chiapas foldbelt lies close to the belt is bound to the south by the Motagua Fault
southern margin of the North American Plate and is Zone that forms the northern border of the Chortis
approximately parallel to the Middle America block, to the SW by the Permo-Triassic Chiapas
MIDDLE MIOCENE CHIAPAS FOLD AND THRUST BELT OF MEXICO 57
Massif and the Gulf of Tehuantepec, to the west by form the regional topseal for the reservoir rocks.
the Miocene–Recent Trans-Mexican Volcanic Belt, The contact beneath the Middle Miocene is gener-
and to the north and east it is unconformably over- ally an unconformity that locally rests upon sedi-
lain by Late Neogene sediments. Where it is ments as old as Eocene. This unconformity marks
exposed in southern Mexico, Guatemala and a period of deformation that produced the Chiapas
Belize, the foldbelt consists of a series Z-shaped, foldbelt. The overlying Middle Miocene to Holo-
NW- to E-trending folds and thrusts that verge cene sediments consist of shale, sandstone, lime-
both to the NE and to the SW associated with stone and conglomerate deposited in environments
common E– W striking, sinistral transcurrent faults that range from basin through shelf, alluvial, fluvial
(de la Rosa et al. 1989) (Fig. 1). The foldbelt has to lacustrine.
been traced northwestwards beneath the Late Development of the Chiapas foldbelt at
Neogene rocks into the Sonda de Campeche and c. 13–11 Ma coincides with a beginning of a
the well studied NE-trending Reforma– Akal uplift, hiatus in arc magmatism defined by Damon &
where the Cantarell structure hosts major petroleum Montesinos (1978: Cenozoic Igneous in Fig. 1).
resources. The Reforma –Akal uplift is bounded by Between c. 17.8 and 12.4 Ma, arc magmatism in
the Comalcalco and Macuspana basins that formed the Chiapas massif and eastern Oaxaca State
during Neogene extensional growth faulting associ- formed a western continuation of the Central Amer-
ated with salt tectonics (Ricoy 1989; Galloway et al. ican Volcanic Arc (Damon & Montesinos 1978;
1991). In southeastern Mexico, constrains on the Keppie et al. 2009). However, between c. 10 and
age of the fold is defined by an unconformity 2.8 Ma, a hiatus in arc magmatism produced a
between folded Latest Cretaceous to Early Middle magmatic arc gap between the eastern end of the
Miocene (12–15 + 1 Ma: K –Ar ages on biotite: Trans-Mexican Volcanic Belt and the western end
Ferrusquı́a 1996) and Upper Miocene –Quaternary of the Central American Volcanic Arc in Guate-
sediments and volcanic rocks (c. 11–0 Ma) mala. Sporadic alkalic–adakitic arc magmatism in
(Fig. 2b). Synchronous deformation in the Sonda the gap was reinitiated at 2.8 Ma within the
de Campeche is similarly bracketed between Chiapas foldbelt in the NW-trending Chiapanecan
folded Jurassic – Middle Miocene (Serravallian: volcanic arc, which lies c. 200 km north of the
13.65 –11.61 Ma: Gradstein et al. 2004) rocks and present Pacific coast (Mora et al. 2007) (Fig. 1).
unconformably overlying Late Miocene (Tortonian: Offshore in the Cocos Plate, the Tehuantepec
11.61 –7.25 Ma; Gradstein et al. 2004)–Recent Transform Fault juxtaposes contrasting ages of the
sediments (Mitra et al. 2005; Mandujano-Velazquez ocean floor due to a combination of mid-ocean
& Keppie 2006) (Fig. 2c). In contrast to the defor- ridge offset and different rates of spreading north
mation recorded in the Chiapas foldbelt, contem- and south of the transform (Manea et al. 2005). It
poraneous Late Cretaceous and Tertiary rocks in is also coincident with a topographic ridge (hence
the Gulf of Tehuantepec are not folded; however its alternate name, the Tehuantepec Ridge) resulting
there are two unconformities: between the Eocene from a combination of: (i) compression produced by
and Early Miocene, and between the Early and a change in the relative plate motions at 13 Ma
Late Miocene (Sanchez-Barreda 1981; Keppie & (Manea et al. 2005); and (ii) the presence of a sea-
Morán-Zenteno 2005) (Fig. 2a). mount chain (Keppie & Morán-Zenteno 2005). As
The rocks deformed in the Chiapas foldbelt its character varies along its length, we hereafter
range in age from Early Jurassic (possibly Triassic) use the term, Tehuantepec Transform/Ridge. How-
to Middle Miocene. The oldest rocks consist of allu- ever, the inferred location of the subducted Tehuan-
vial fan and fluvial red beds deposited in grabens tepec Transform/Ridge beneath Mexico over the
and half grabens bordered by normal faults formed past 15 Ma varies between authors. Keppie & Morán-
during opening of the Gulf of Mexico. Red beds Zenteno (2005) infer that, following a Middle
overlain by Middle–Upper Jurassic evaporites Miocene plate reorganization, the pre-13 Ma old
(including salt) pass upwards into Upper Jurassic segment of the Tehuantepec Transform/Ridge
and Middle Cretaceous platformal carbonates, rotated counterclockwise about a pole in the mouth
and shales and sandstone of Upper Cretaceous – of the Gulf of California from an ENE –North
Palaeogene age. In the Gulf of Tehuantepec, trend. Rotation would have resulted in a westward
conglomerates with tonalite, dacite and basalt migration of its intersection with the Middle
fragments occur. The Upper Jurassic rocks formed America Trench. Manea & Manea (2006) projected
the source of most of the hydrocarbons (Mitra the Tehuantepec Ridge northeastwards beneath
et al. 2005). The Cretaceous –Palaeocene boundary Mexico, omitting the sharp bend in the transform
is marked by carbonate breccias derived from the that formed during the Middle Miocene plate reor-
Chixhulub impact crater (Grajales et al. 2000), ganization: this results in an easterly migration of
which forms the main reservoir for hydrocarbons the Tehuantepec/Middle America Trench intersec-
in the Sonda de Campeche. Palaeocene shales tion in the last 13 Ma. Earthquake data indicate
Fig. 2. Representative stratigraphic columns of the: (a) Tehuantepec Platform (from Sanchez-Barreda 1981, well
Salina Cruz-1); (b) Sierra de Chiapas; (c) Sonda de Campeche. Abbreviations: E, Early; M, Middle; L, Late.
that instead of a ridge, a N –S step in the Benioff data from seven wells in the Sierra de Chiapas and
Zone extends northwards beneath Mexico from ten in the Sonda de Campeche (Figs 3 & 5). Two
the exposed Tehuantepec Transform/Ridge at the composite NE-SW PEMEX reflection seismic sec-
Middle America Trench (Bravo et al. 2004). The tions, roughly perpendicular to the trend of the
dip of the Benioff zone changes from west to east folds, were compiled: Section A –B is 97.75 km
across the step from c. 118N to c. 408N (Pardo & long across the Sierra de Chiapas, and the other
Suárez 1995; Rebollar et al. 1999a). Using the con- (C –D) is 172 km in length across the Sonda de
vergence rate of 67 km/Ma (Pindell et al. 1988), the Campeche (Figs 4 & 6).
step-like geometry extends back to at least 5 Ma.
These observations are consistent with the location Sierra de Chiapas
of the Tehuantepec Transform/Ridge inferred by
Keppie & Morán-Zenteno (2005); however, it high- In the Sierra de Chiapas, the deepest wells (#11 &
lights its changing nature along strike. 17) reached Kimmeridgian and Tithonian (Upper
Jurassic) carbonates. The Cretaceous is typically
dolomite in the Early Cretaceous passing upwards
Interpretation of reflection seismic data into mixed carbonate and dolomite in the Middle
Cretaceous, and marl and breccia in the Upper Cre-
Stratigraphic and structural interpretation was sup- taceous (up to 2142 m) (Fig. 3). The Tertiary is
ported by lithological and chronostratigraphical characterized by clastic rocks (shale, sandstone
MIDDLE MIOCENE CHIAPAS FOLD AND THRUST BELT OF MEXICO
Fig. 3. Chronostratigraphic correlation of wells in the Sierra de Chiapas (Fig. 4). Abbreviations: E, Early; M, Middle; L, Late; Jk, Kimmeridgian; Jt, Tithonian; K, Cretaceous;
P, Palaeocene; Eo, Eocene; O, Oligocene; M, Miocene; Ind, undifferentiated.
59
60
J. J. MANDUJANO-VELAZQUEZ & J. D. KEPPIE
Fig. 4. Sierra de Chiapas. (a) Seismic reflection data and interpretation; (b) Geological interpretation; (c) balanced cross-section (from PEMEX, Region Sur del Activo
Regional de Exploración, Coordinación de Plays Establecidos: Medina et al. 1997).
MIDDLE MIOCENE CHIAPAS FOLD AND THRUST BELT OF MEXICO
Fig. 5. Chronostratigraphic correlation of wells in the Sonda de Campeche section (Fig. 6). Abbreviations: E, Early; M, Middle; L, Late; Jo, Oxfordian; Jk, Kimmeridgian;
Jt, Tithonian; K, Cretaceous; P, Palaeocene; Eo, Eocene; O, Oligocene; M, Miocene; and Quater– Plioc, Quaternary–Pliocene.
61
62
J. J. MANDUJANO-VELAZQUEZ & J. D. KEPPIE
Fig. 6. Sonda de Campeche. (a) Seismic reflection data and interpretation, (b) Geological interpretation (modified after PEMEX, Region Marina del Activo Regional de
Exploración, Grupo de Cuencas y Sistemas Petroleros: Ortega & Nolasco 2005) (c) Balanced cross-section.
and conglomerate) with minor limonite and bento- Early Miocene consist of shale, carbonate shales,
nite horizons. Unconformities are present at various marl, sandstone and bentonite with rare carbonate.
levels: at the base of the Palaeocene (well #12), at Middle Miocene sandstone and shale unconform-
the base of the Eocene (wells #14 and 15), at the ably overlying the Early Miocene range from
base of the Early Miocene (well #14) and at the 1302 m in the SW to 175 m in the NE. In contrast,
base of the Pliocene/Quaternary (wells #14, 16 the Late Miocene lithologies are thickest in the
and 17). In places, the Eocene rests directly upon NE, tapering to 175 m in the SW. The Pliocene –
the Cretaceous (well #14). The unconformities Quaternary sediments are 750 m thick.
lead to considerable variation in thickness: Palaeo- The structures observed in the reflection seismic
cene 0–800 m, Eocene up to 1175 m, Oligocene section along line C– D may be divided into two
0–185 m, Early Miocene 550 m, Middle Miocene types: (i) folds and thrusts in rocks older than
.250 m and Pliocene– Quaternary .3000 m. Middle Miocene; and (ii) listric normal faults that
The Chiapas fold-and-thrust belt is characterized cut part or all of the stratigraphic section (Fig. 6).
by NE verging folds and thrusts with trends, varying Most of the anticlines have salt cores and are
from East –West through NW–SE to NNW–SSE. associated with ramps in underlying NE-vergent
Sinistral transcurrent faults strike East – West thrusts that root into a décollement in the salt
and, rarely, NW–SE (Fig. 1). The reflection horizon. In the hanging wall of many of these
seismic section (A– B) reveals numerous folds and NE-vergent thrusts are small SW-vergent thrusts.
associated thrusts. Thrust duplexes are floored by The largest and best documented of the NE-vergent
a décollement zone beneath the Upper Jurassic, thrusts is the Sihil Thrust, which ramps upwards to
which is inferred to be located in the Middle-Upper the east and places Upper Jurassic rocks upon
Jurassic salt horizon (Fig. 4). Towards the south- Early Miocene sediments in the Cantarell anticline
western end of the section, the thrusts appear to (Mitra et al. 2005). This Sihil thrust is truncated by
have been folded during back-thrusting. Two the Middle Miocene unconformity. Two SW-
normal faults downthrown to the south are present vergent thrusts occur below the Sihil Thrust on the
in the middle of the section (faults 7 and 11). northern limb of the Cantarell anticline.
The northern half of the section shows that all of Listric normal faults are common in the upper
these structures are unconformably overlain by part of the section, where some delimit graben and
Miocene –Present sediments that overstep and half graben located over older synclines (Fig. 6).
overlap towards the south. Whereas some of these faults are truncated by Holo-
A balanced section was produced using standard cene sediments, others extend to the surface indicat-
techniques (Marshak & Woodward 1988) (Fig. 4c), ing continuing activity. The variations in the
and reveals that the present 97.75 km section was thickness of the Middle Miocene–Present units is
originally 148.33 km in length, that is a shortening especially marked in the half graben on the south
of 50.68 km (¼34.1%). The shortening increases side of the Cantarell anticline and shows that the
from 6% in the NE to 48% in the SW. bounding faults were active throughout their depo-
sition. Listric normal faults at the northeastern end
Sonda de Campeche of the section cut only Mesozoic and Palaeogene
rocks and are interpreted as being associated with
In the Sonda de Campeche, the deepest wells the boundary between the Macuspana Basin and
reached Callovian–Oxfordian salt, anhydrite, sand- the Yucatan Platform.
stone, mudstone and shale (Fig. 5). Salt is visible in In order to calculate the amount of shortening, a
the seismic reflection data along the base of the balanced section was produced in a two-stage
section, and as diapers that, in the southwestern process using standard procedures (Marshak &
part, penetrate up through the entire stratigraphic Woodward 1988). First, extension during listric
sequence. The salt diapirs have associated anticli- normal faulting was removed, and then the thrusts
nes. Salt also occurs as lenses in the Palaeogene were moved back to connect displaced units and
and Miocene (Fig. 6). The Kimmeridgian is rep- unfolding the strata (Fig. 6c). In this way we deter-
resented by clastic rocks overlain by carbonate mined that the present distance of 157.5 km between
rocks, anhydrite and dolomite, which continues C and the Pin Line near D was originally 170.8 km,
upwards into the 150–400 m-thick Tithonian sedi- that is 13.3 km (¼ 7.8%) of shortening. The magni-
ments. The Cretaceous consists of 315–751 m-thick tude of shortening is similar to that at the northeast-
carbonates that pass upwards into shaly carbonate ern end of the Chiapas section (A –B) and indicates
and marl rocks; however, towards the northeastern that the Chiapas foldbelt continues 200 km to the
end of the section the total thickness of the Cretac- NE beneath the Pliocene–Quaternary cover.
eous increases to 1420 m where thick meteorite The similarity in the structure of the Chiapas
impact breccias occur. The 395 –1294 m-thick foldbelt both onland in the Sierra de Chiapas, and
Palaeogene, 475 m-thick Eocene and 2085 m-thick offshore in the Sonda de Campeche, suggests that
there should be a similar potential for hydrocarbons. retreat of the trench over the last c. 23 Ma of
However, in the exposed part of the belt, the 1 km/Ma off Mexico and 0.9 km/Ma off Guate-
absence of an adequate topseal and apparent lack mala. Within error, this is similar to the average
of maturation of organic matter may account for rate calculated from the data recorded in DSDP
the present lack of known reserves. On the other leg 67, hole #496 at 90.88W (Fig. 7); however,
hand, eastward projection of the foldbelt from detailed data from this hole indicate that the subsi-
the Sonda de Campeche suggests considerable dence rate was variable and yields a faster landward
unexplored potential beneath the Middle Miocene migration rate of c. 3 km/Ma between 18 and
unconformity. 11 Ma. This faster rate of trench migration mirrors
the faster rate at c. 22–13 Ma calculated for hole
#493 (Clift & Vannucchi 2004). These legs are
Evidence for subduction erosion located west (leg #66 Acapulco, México) and east
(leg 84, San José, Guatemala) of the Gulf of Tehuan-
Subduction erosion along the Mexican and Guate- tepec, respectively, and it assumed that similar
malan Pacific margin appears to have occurred subduction rates occurred south of the Gulf of
over the past 20 Ma (Clift & Vannucchi 2004), Tehuantepec. These results suggest that the period
and so may be a factor in development of the of subduction erosion predated development of the
Chiapas foldbelt: this is investigated now. Estimates Chiapas foldbelt, and thus subduction erosion is
of subduction erosion are based on the subsidence probably not a factor in its genesis.
history of the forearc using fossils as depth indi-
cators; subsidence has been related to landward
migration of the trench as a result of frontal subduc- Tectonic interpretation
tion erosion (see Clift & Vannucchi 2004 for the
methodology, which we follow herein). Depths of The orientation of the finite strain ellipse in the
deposition has been documented by the Deep Sea Chiapas foldbelt may be deduced from the orien-
Drilling Project (DSDP leg #66, holes #489 and tation of the faults and folds. Assuming little or no
#493 at c. 998W) and off Guatemala (DSDP leg rotation of the E-W sinistral faults, constrain the
84, hole #569 at 90.848W). Clift and Vannucchi’s shortening direction to N608E (Fig. 1). This direc-
(2004) calculations give average rates of landward tion is c. 228 clockwise of the N388E relative plate
Fig. 7. Subsidence Evolution: (a) Mexican margin, DSDP 493, [(----------) Clift & Vannucchi 2004],
(2 .. 2 .. 2 .. 2 ..2) this paper); Guatemala margin: DSDP 569 (Clift & Vannucchi 2004), DSDP 496 (this
paper); (b) Subduction-Erosion of the Guatemala Margin. Abbreviations: E, Early; M, Middle; L, Late. N. B. The
palaeobatymetric curves calculated in this work (Mexican margin 493 and Guatemala margin 569) were constructed
based on the benthic foraminifers reported in these wells, which were drilled by the Deep Sea Drilling Project.
motion between the Cocos and North American associated with slightly oblique convergence
plates at 10 –20 Ma (Pindell et al. 1988; Engebretson (Jiang et al. 2001). The total shortening across the
et al. 1985) (Fig. 1). The difference is explicable Chiapas foldbelt (106 km) may be calculated by
in terms of the general difference between strain adding the 51.7 km of shortening along A –B, the
axes in finite strain versus infinitesimal strain 13.3 km of shortening along C –D, the c. 30 km of
(approximately parallel to stress directions) shortening between A and the Chiapas massif
Fig. 8. Tectonic evolution of Chiapas (a) before, (b) during and (c) after development of the Chiapas and Campeche
fold-and-thrust.
assuming 50% shortening (similar to the 48% short- magmatic arc, the dip of the Benioff zone before
ening at the southwestern end of A –B) and 11 km 13 Ma may be calculated as 318 (Fig. 8a).
between B and C assuming a 7.8% shortening (as The dip of the Benioff zone during development
in C –D). Seismic data indicates no shortening of the Chiapas foldbelt cannot be estimated directly.
across the Gulf of Tehuantepec (Sánchez-Barreda However, its present dip west of the Tehuantepec
1981) and none has been reported across the Ridge is c. 118 (Pardo & Súarez 1995; Bravo et al.
Chiapas massif. This is analogous to that observed 2004). We assume a dip of 118 was present in the
in California where thrusts east of the Sierra Middle Miocene (Fig. 8b).
Nevada batholith merge into a basal décollement The present c. 398 dip of the Benioff zone
that passes beneath a more rigid block under the beneath Chiapas has been derived by Bravo et al.
Sierra Nevada batholith and the Great Valley (Fay (2004) using earthquake data (Fig. 8c). The slab
et al. 2006). In order to restore the position of beneath the Central American Volcanic Arc in Gua-
the Middle America Trench prior to 18 Ma, the temala to the south dips at 458 (Bevis & Isacks 1984;
20 km loss to subduction erosion between 18 and Burbach et al. 1984) and decreases to 118 west of the
c. 11 Ma must be added to the 106 km of shortening Tehuantepec Transform/Ridge (Pardo & Suárez
accommodated by the Chiapas fold and thrust belt. 1995). These changes along strike to the south
Thus the Middle America Trench at 18 Ma lay increase in slab dip can be interpreted as resulting
126 km SSW of its present location off Chiapas. from the progressive westward migration of the
Assuming the present arc-trench width of 150 km Tehuantepec Transform/Ridge, which would have
is applicable to the Early Miocene arc (17.8– resulted in a gradual decrease in slab dip of the
12.4 Ma) and 90 km for the slab depth beneath the Benioff zone before passage of the ridge followed
Fig. 9. Reconstructions showing the locations of the volcanic arcs, the Chiapas Massif, the Middle America Trench and
the Tehuantepec Transform/Ridge (modified from Keppie & Morán-Zenteno 2005) at 13– 0 Ma.
by an increase in slab dip after the ridge had passed the Middle America Trench migrated westwards
through (Fig. 9). In our model, we apply these dips along the Chortis block and Chiapas margins from
to the last 10 Ma. c. 15 –12 Ma to 12– 0 Ma (Fig. 9), respectively.
The Middle Miocene age of the Chiapas foldbelt Our plate reconstructions imply that the intersection
is synchronous with: (a) the 12.5 –11 Ma, Middle would have migrated west through time. If collision
Miocene plate reorganization during which the of the Tehuantepec ridge with the trench was
Pacific-Cocos pole of rotation moved c. 600 km responsible for formation of the Chiapas fold-and-
southwards along the Pacific coast resulting in a thrust belt, then thrust belt development should
decrease in convergence rate between the Cocos also be diachronous. Although our data are
and North American plates (Fig. 1) (Mammerickx restricted to the western part of the foldbelt, such
& Klitgord 1982); (b) collision of the Tehuantepec diachronism is not apparent because the Chiapas
Transform/Ridge with the Middle America foldbelt does not appear to continue into the
Trench off Chiapas near the Guatemala-Mexico Chortis block (Donnelly et al. 1990) and defor-
border (Keppie & Morán-Zenteno 2005); (c) the mation appears to have been of short duration
start of an hiatus in arc magmatism in Chiapas; (13 –11 Ma). The reason for initiation and termin-
and (d) the end of a period of increased subduction ation of the deformation may perhaps be related to
erosion. Subduction erosion affected the whole the variation in the geometry of the Tehuantepec
Mexican, Guatemalan, Nicaraguan and Costa Transform/Ridge. Thus, when a ridge was being
Rican margins suggesting that there is no causal subducted it induced deformation in the upper
link with either the Chiapas foldbelt or collision of plate. However, when the transform was only a
the Tehuantepec Transform/Ridge. The decrease in step with no accompanying ridge, no deformation
the convergence rate during the Middle Miocene occurred in the upper plate: this certainly applies
would not appear to be a factor in producing the to the period 5 Ma – Present because the earthquake
Chiapas foldbelt. data indicate that the geometry of the subducted
The difference in age of the subducted oceanic Tehuantepec Transform/Ridge is a step with no
plate across the Tehuantepec Transform/Ridge accompanying ridge (Bravo et al. 2004).
is c. 10 Ma at the Middle America Trench and In conclusion, the origin of the Chiapas foldbelt
produces a step down to the east in the sea-floor as a result of subduction of a transform fault where it
elevation of c. 1100 m (Manea et al. 2005). A is accompanied by a compressional ridge and sea-
similar N –S step in the top of the subducted mount chain indicates the importance of the topo-
Cocos Plate, documented from earthquake data, graphy of the subducting plate in the genesis of
shows an increase from c. 25 km beneath the coast modern and ancient fold-and-thrust belts. This
of the Gulf of Tehuantepec to c. 100 km at 188N may find application to short-lived episodes of
(Bravo et al. 2004). This indicates that the offset deformation in the long pre-collisional history of
in ages across the Tehuantepec Transform/Ridge orogenic belts.
results in an increase in the dip of the Benioff
zone from 118 on the western side to c. 308 on the We are grateful to A. O. Pérez, A. A. López, J. H. Garcı́a of
eastern side. Slab dip of the Cocos plate increases PEMEX and A. A. P. Luna of Instituto Mexicano del Pet-
eastwards to 458 at the Guatemalan border. Earth- róleo for allowing us to use the seismic data Chiapas and
the Sonda de Campeche. We thank M. E. V. Meneses for
quake data indicates that the dip of the Benioff
a constructive review of the structural model and
zone gradually increases westwards beneath the R. Strachan and S. T. Johnston for comments on the manu-
Trans-Mexican Volcanic Belt (Pardo & Suárez script that helped in revision of the paper.
1995). The geometry of the Benioff zone across
the subducted Tehuantepec Transform/Ridge
with its step down to the east is similar to that
of central South America where a step down References
to the south occurs across the subducted Juan B EVIS , M. & I SACKS , B. L. 1984. Hypocentral trend
Fernandez Ridge (Ramos & Aleman 2000). A surface analysis: probing the geometry of Benioff
decrease in the Benioff zone dip is accompanied zones. Journal of Geophysical Research, 89,
by an increase in shortening across the Andes that 6153– 6170.
produced a fold-and-thrust belt (Ramos et al. B RAVO , H., R EBOLLAR , C. J., U RIBE , A. & J IMÉNEZ , O.
2002), thereby providing an analogy for the 2004. Geometry and state of stress of the Wadati-
Benioff zone in the Gulf of Tehuantepec, Mexico.
Chiapas foldbelt.
Journal of Geophysical Research, 109, doi: 10.1029/
Assuming that a similar geometry in the Benioff 2003JB002854, 1 –14.
zone across the Tehuantepec Transform/Ridge has B URBACH , G. B., F ROHLICH , C., P ENNINGTON , W. D. &
existed since 15 Ma, the reconstructions of Keppie M ATUMOTO , T. 1984. Seismicity and Tectonics of the
& Morán-Zenteno (2005) indicate that the intersec- Subducted Cocos Plate. Journal of Geophysical
tion between Tehuantepec Transform/Ridge and Research, 89, 7719– 7735.
C LIFT , P. D. & V ANNUCCHI , P. 2004. Controls on tectonic present. Journal of Geophysical Research, 87,
accretion versus erosion in subduction zones: impli- 6751– 6759.
cations for the origin and recycling of the continental M ANDUJANO -V ELAZQUEZ , J. J. & K EPPIE , D. J. 2006.
crust. Review of Geophysics, 42, 1– 31. Cylindrical and conical fold geometries in the Cantar-
D AMON , P. E. & M ONTESINOS , E. 1978. Late Cenozoic ell structure, Southern Gulf of Mexico: implications
volcanism and metallogenesis over an active Benioff for hydrocarbon exploration. Journal of Petroleum
zone in Chiapas, Mexico. Arizona Geological Society Geology, 29, 215 –226.
Digest, 11, 155– 168. M ANEA , M., M ANEA , V. C., F ERRARI , L., C OSTOGLO-
D E LA R OSA , J. L., E BOLI , M. A., D ÁVILA , S. M., DOV , V. & B ANDY , W. L. 2005. Tectonic evolution
D ENGO , G., Y AMAZAKI , M. F. & B ALINAS , G. R. of the Tehuantepec Ridge. Earth and Planetary
1989. Geologı́a del Estado de Chiapas: Subdirección Science Letters, 238, 64–77.
de Construcción y departamento de Geologı́a de la M ANEA , V. C. & M ANEA , M. 2006. Origin of the modern
Comisión Federal de Electricidad, superintendencia Chiapanecan Volcanic arc in southern México inferred
de estudio Zona Sureste. Tuxtla Gutiérrez, Chiapas, from thermal models. Geological Society of America,
México, 1– 89. Special Paper, 412, 27–38.
D ONNELLY , T. W., H ORNE , G. S., F INCH , R. C. & M ARSHAK , S. & W OODWARD , N. 1988. Introduction to
L ÓPEZ -R AMOS , E. 1990. Northern Central America; cross-section balancing. In: M ARSHAK , S. & M ITRA ,
the Maya and Chortis blocks. In: D ENGO , G. & C ASE , G. (eds) Basic Methods of Structural Geology. Prentice
J. E. (eds) The Caribbean region: the geology Hall, Englewood Cliffs, 303–332.
of North America. Geological Society of America, M EDINA , F. U., H AM , W. J. M., N AMSON , J. & A LCÁN-
37–76. TARA , G. J. R. 1997. Proyecto Simojovel. Petróleos
E NGEBRETSON , A. C., C OX , A. & G ORDON , R. G. 1985. Mexicanos (PEMEX), Región Sur del Activo Regional
Relative motions between oceanic and continental de Exploración, Coordinación de Plays Establecidos,
plates in the Pacific Basin. Geology Society of (unpublished report).
America, Special Paper, 206, 1– 59. M ITRA , S., C ORREA , F. G., H ERNÁNDEZ , G. J. &
F AY , N., H UMPHREYS , E. D. & C OBLENTZ , D. D. 2006. M URILLO , A. A. 2005. Three-dimensional structural
The Pacific-North America plate boundary in the model of the Cantarell and Sihil structures, Campeche
western United States: dynamics of the Sierra Nevada Bay, Mexico. American Association of Petroleum
block and implications for western US tectonics. Geo- Geology Bulletin, 89, 1 –26.
logical Society of America. Abstracts with programs M ORA , J. C., J AIMES , V. M. C., G ARDUÑO , M. V. H.,
Speciality Meeting No. 2, p. 109. Paper No. 14–8. L AYER , P. W., P OMPA , M. V. & G ODINEZ , M. L.
F ERRUSQUÍA , V. 1996. Contribución al conocimiento 2007. Geology and geochemistry characteristics of
geológico de Chiapas, el área Ixtapa-Soyaló. Universi- the Chiapanecan Volcanic Arc (central area), Chiapas
dad Nacional Autónoma de México. Boletı́n del Insti- Mexico. Journal of Volcanology and Geothermal
tuto de Geologı́a, 109, 1 –130. Research, doi: 10.1016/j.jvolgeores.2006.12.009.
G ALLOWAY , W. E., B EBOUT , D. G., F ISHER , W. L., O RTEGA , V. & N OLASCO , J. 2005. Modelo geológico de
D UNLAP , J. B., J R , C ABRERA , C. R., L UGO , R. J. E. la Sonda de Campeche. Petróleos Mexicanos
& S COTT , T. M. 1991. Cenozoic, In: S ALVADOR , A. (PEMEX), Región Marina del Activo Regional de
(ed.) The Gulf of Mexico Basin. The Geology of Exploración, Grupo de Cuencas y Sistemas Petroleros,
North America, J, 245– 324. unpublishing report.
G RADSTEIN , F. M., O GG , J. G., S MITH , A. G., B LEEKER , P ARDO , M. & S UÁREZ , G. 1995. Shape of the subducted
W. & L OURENS , L. J. 2004. A New Geologic Time Rivera and Cocos plates in southern Mexico: seismic
Scale, with special reference to Precambrian and and tectonic implications. Journal of Geophysical
Neogene. Episodes, 27, 83– 100. Research, 100, 12 357–12 373.
G RAJALES , N. J. M., C EDILLO , P. E., R OSALES , D. C. & P INDELL , J. L., C ANDE , S. C. W., P ITMAN , W. C.,
M ORÁN , C. D. 2000. Chicxulub impact: the origin of R OWLEY , D. B., D EWEY , J. F., L EBRECQUE , J. &
reservoir and seal facies in the southeastern Mexico H AXBY , W. 1988. A plate-kinematic framework for
oil fields. Geology, 28, 307 –310. models of Caribbean evolution. Tectonophysics, 155,
J IANG , D., L IN , S. & W ILLIAMS , P. F. 2001. Deformation 121–138.
path in high strain zones, with reference to slip parti- P INDELL , J. L., L ORCAN , K., W ALTER , M., K LAUS , S. &
tioning in transpressional plate-boundary regimes. G RENVILLE , D. 2006. Foundations of Gulf of Mexico
Journal of Structural Geology, 23, 991– 1005. and Caribbean evolution: eight controversies resolved.
K EPPIE , J. D. & M ORÁN -Z ENTENO , D. J. 2005. Tectonic Geologica Acta, 4, 303–341.
implications of alternative Cenozoic reconstructions R AMOS , V. A. & A LEMAN , A. 2000. Tectonic evolution
for southern Mexico and the Chortis Block. Inter- of the Andes. In: C ORDANI , U. G., M ILANI , E.,
national Geology Review, 47, 473– 491. T HOMAZ , F. A. & C AMPOS , D. A. (eds) Tectonic evol-
K EPPIE , J. D., M ORÁN -Z ENTENO , D. J., M ARTINY , B. & ution of South America. 31st International Geological
G ONZÁLEZ -T ORRES , E. 2009. Synchronous 29– Congress, Rio de Janeiro, Brazil. Geological Society
19 Ma arc hiatus, exhumation and subduction of of America, Colorado, 635–685.
forearc in southwestern Mexico. In: J AMES , K. H., R AMOS , V. A., C RISTALLINI , E. O. & P ÉREZ , D. J.
L ORENTE , M. A. & P INDELL , J. L. (eds) The Origin 2002. The Pampean flat-slab of the central Andes.
and Evolution of the Caribbean Plate. Geological Journal of South American Earth Sciences, 15, 59–78.
Society, London, Special Publications, 328, 169–179. R EBOLLAR , C. J., E SPÍNDOLA , V. H., U RIBE , A.,
M AMMERICKX , J. & K LITGORD , K. D. 1982. Northern M ENDOZA , A. & P ÉREZ , V. A. 1999a. Distribution
East Pacific Rise: evolution from 25 m.y. to the of stresses and geometry of the Wadati– Benioff zone
under Chiapas, Mexico. Geophysical International, 38, southern Mexico. PhD thesis, University of Texas,
95– 106. Austin, Texas.
R ICOY , U. J. 1989. Tertiary Terrigenous Depositional S CHAAF , P., M ORAN , Z. D., H ERNÁNDEZ , B. M., S OLÍS ,
Systems of the Mexican Isthmus Basins. PhD thesis, P. G., T OLSON , G. & K ÖHLER , H. 1995. Paleogene
University of Texas, Austin, Texas. continental margin truncation in southwestern
R OSS , M. I. & S COTESE , C. R. 1988. A hierarchical tec- Mexico: geochronological evidence. Tectonics, 14,
tonic model of the Gulf of Mexico and Caribbean 1339– 1350.
regions. Tectonophysics, 135, 139– 168. S OBOLEV , S. V. & B ABEYKO , A. Y. 2005. What drives
S ÁNCHEZ -B ARREDA , L. A. 1981. Geologic evolution of orogeny in the Andes? Geology, 33, doi: 10.1130/
the continental margin of the Gulf of Tehuantepec in G21557AR.1, 617–620.
Extrusion of high-pressure Cache Creek rocks into the Triassic
Stikinia –Quesnellia arc of the Canadian Cordillera: implications for
terrane analysis of ancient orogens and palaeogeography
JAROSLAV DOSTAL1*, J. DUNCAN KEPPIE2 & FILIPPO FERRI3
1
Department of Geology, Saint Mary’s University, Halifax, Nova Scotia B3H 3C3, Canada
2
Departamento de Geologı́a Regional, Instituto de Geologia, Universidad Nacional
Autonoma de Mexico, 04510 Mexico, D.F., Mexico
3
British Columbia Ministry of Energy and Mines, 1810 Blanchard Street, Victoria,
British Columbia V8W 9N3, Canada
*Corresponding author (e-mail: jdostal@smu.ca)
Abstract: The volcanic Triassic Takla Group constitutes a significant part of Stikinia and Quesnel-
lia, two major terranes of the Canadian Cordillera that are separated by high-pressure rocks of the
Cache Creek terrane containing Asian fauna. The geochemical and isotopic characteristics of the
Takla Group in Quesnellia and Stikinia are similar, that is, tholeiitic basalts characterized by low
abundances of strongly incompatible trace elements, negative Nb anomalies, þ6 to þ8 1Nd
values, the low initial Sr isotopic ratios, and relatively horizontal chondrite-normalized heavy
REE patterns, all features typical of relatively primitive arcs with little or no continental crust invol-
vement. These similarities have led to several geometric models: post-Middle Jurassic duplication
by strike-slip faulting, and oroclinal or synformal folding. However, they are all inconsistent with
either palaeomagnetic or faunal data, and the presence of a Triassic overstep sequence, which indi-
cates amalgamation c. 50 ma before emplacement of the youngest oceanic rocks of the Cache Creek
terrane. An alternative model is proposed: oblique eastward subduction of the Cache Creek accre-
tionary prism and fore-arc producing high-pressure metamorphism, followed by extrusion into the
arc and exhumation by the Middle Jurassic. This model implies that these high-pressure rocks, rather
than marking an oceanic suture between disparate arc terranes, support a para-autochthonous origin.
The Canadian Cordillera is composed of various There is currently considerable debate about the
tectono-stratigraphic terranes that were accreted to palaeogeography of the northern Cordillera between
the western margin of Laurentia (North America) 100–55 Ma (e.g. Cowan et al. 1997; Colpron et al.
before the early Cenozoic (Colpron et al. 2006, 2006). Some palaeomagnetic data suggest that a
2007). The geology of the northern Cordillera has large region of the Cordilleran hinterland (Baja,
recently been synthesized by Colpron et al. (2006, British Columbia) made a 2100 + 700 km south-
2007), who divided it into five first-order tectonic ward excursion from its present latitude relative to
entities (Fig. 1): (1) ancestral North America Laurentia during the Late Cretaceous, returning
(Laurentia), which includes the western craton northward during the Palaeocene (Enkin 2006).
margin and overlying miogeocline and fringing On the other hand, other such data suggest that the
para-autochthonous terranes; (2) allochthonous Yukon-Tanana terrane has been para-autochthonous
pericratonic terranes (Intermontane terranes, which since 215 Ma (Late Triassic: Norian) and that, in the
includes the Yukon-Tanana terrane, presumed base- interval 100–55 Ma, the Intermontane terranes
ment to the Quesnel and Stikine arc terranes rotated 35 + 148 clockwise as they moved
addressed in this paper) that are separated from 915 + 775 km northwards relative to Laurentia,
Laurentia by the discontinuous, Upper Devonian- followed by a further 16 + 68 clockwise rotation
Permian, Slide Mountain back-arc terrane; (3) in the Cenozoic (Symons et al. 2005). However,
Insular and Northern Alaska terranes, which about half of the northward relative motion may
evolved in the Arctic during the Palaeozoic; (4) be accounted for by the 430 km dextral displace-
oceanic and accretionary complexes (including the ment along the Tintina Fault (Gabrielse et al.
oceanic Cache Creek terrane that presently lies 2006), with an additional 130–150 km on the
between Quesnellia and Stikinia, two major terranes Fraser-Pinchi Fault system (Coleman & Parrish
of the Canadian Cordillera); and (5) Late accreted 1991; Enkin et al. 2006). The latter data are consist-
arc and accretionary terranes that form the western ent with other Upper Triassic palaeomagnetic data
and southern fringe of the Cordillera. that indicate little or no relative motion between
72 J. DOSTAL ET AL.
Fig. 1. Simplified terrane map of the Canadian Cordillera (after Colpron et al. 2007), showing the location of previous
geochemical studies and Figure 3. Inset shows terrane grouping and tectonic realms.
the Intermontane terranes and Laurentia (Irving & and compare these with similar data from Takla
Wynne 1992). These data are also consistent with Group rocks located in a more outboard terrane
the recognition of a Triassic overstep sequence (Stikinia). All of these rocks were originally
that ties the Intermontane terranes to North mapped as one unit, the Takla Group, based on
America. Thus, provenance of Late –Middle Devo- lithological and age similarities (Church 1975;
nian detrital zircons in Lower Triassic (Smithian) Monger 1977; Monger & Church 1977), even
rocks of the Selwyn basin (overlying the Palaeozoic though they are separated by the Mississippian –
miogeocline) can only be found in the Yukon- Lower Jurassic high-pressure (HP) rocks of the
Tanana terrane that underlies the Quesnel and Cache Creek terrane and a major dextral fault
Stikine terranes (Beranek & Mortensen 2006). (Pinchi-Fraser Fault: Fig. 1). These observations
These conclusions are inconsistent with the sugges- have led to a wide variety of tectonic models,
tion that Stikinia, Quesnellia and Cache Creek are including the development of one arc that was
exotic, accreting to one another in western Pantha- subsequently either offset by strike-slip motions
lassa at 230 Ma, and then traveling c. 10 000 km (e.g. Wernicke & Klepacki 1988; Beck 1991,
across Panthalassa to accrete with western North 1992; Irving et al. 1996), bent into an orocline
America by c. 150 Ma (e.g. Johnston & Borel (Nelson & Mihalynuk 1993; Mihalynuk et al.
2007, and references therein). 1994; Nelson et al. 2006), or synformal folding of
In this paper we present geochemical data for the Cache Creek klippe that was thrust over the arc
Middle–Upper Triassic volcanic arc rocks of the during the Middle Jurassic (Samson et al. 1991;
Takla Group in an inboard terrane (Quesnellia) Gehrels et al. 1991).
EXTRUSION OF HIGH-PRESSURE CACHE CREEK ROCKS 73
The Takla Group volumetrically forms signifi- 1991; Ferri & Melville 1994; Pantaleyev et al.
cant portions of both Quesnellia and Stikinia 1996; MacIntyre et al. 2001; MacIntyre 2006;
(Fig. 1). Correlative rocks of the Takla Group to Beatty et al. 2006; Breitsprecher et al. 2007). Both
the south include the Nicola Group in Quesnellia sequences consist of arc-related, mafic-intermediate
(Mortimer 1987) and to the north the Stuhini Group pyroclastic rocks, massive flows and epiclastic
of Stikinia (Mortimer 1986). Whereas igneous geo- rocks, and lie on or laterally interfinger with
chemical data are available at several localities in argillites, limestones and minor volcanic-derived
Quesnellia (Fig. 1: Mortimer 1987; Pantaleyev epiclastic rocks containing fauna with Laurentian
et al. 1996; Nelson & Bellefontaine 1996), geochem- affinities (Carter et al. 1992; Stanley & Senowbari-
ical data is only available in central Stikinia (Dostal Daryn 1999). The Permo-Carboniferous rocks of
et al. 1999; MacIntyre et al. 2001). Thus, the purpose Quesnellia have been correlated with similar rocks
of this paper is to present more extensive geochem- (Finlayson and Klinkit units) of the Yukon–Tanana
ical data on the volcanic rocks of the Quesnel terrane (Simard et al. 2003), which unconformably
Takla Group from the Germansen Landing-Manson overlie Neoproterozoic-Lower Devonian distal
Creek and Aiken Lake area in north–central British Laurentian margin rocks of the Snowcap assemblage
Columbia and compare them with Stikine Takla (Colpron et al. 2006, 2007). Similarly, the Stikine
volcanic rocks from the McConnel Creek map assemblage is inferred to overlie the Yukon–
area, just to the west (Fig. 1). Tanana terrane (Monger & Struik 2006).
Geochemical data presented in this paper indi-
cate that the Takla Group volcanic rocks in both Slide Mountain terrane
Quesnellia and Stikinia were erupted in an almost
identical volcanic arc setting and represent one arc Uppermost Devonian –Permian rocks of the Slide
terrane rather than two. Geochemical variations indi- Mountain terrane (Fig. 2; ultramafic rocks, oceanic
cate a west-facing arc above an east-dipping subduc- basalts, radiolarian cherts, argillites and carbonates:
tion zone. Accretion of the Intermontane terranes to Sylvester Allochthon, Nina Creek Group, Slide
western Laurentia by late Lower Triassic, combined Mountain Group, Fennell Group: Ferri 1997;
with palaeomagnetic and faunal province data Nelson et al. 2006; Colpron et al. 2006, 2007)
suggesting little or no lateral offset relative to Laur- reflect cold water conditions in the Permian
entia, pose difficult problems for several models. We (Orchard 2006). These rocks are inferred to
propose an alternative model in which the Cache have formed in a back-arc basin behind the
Creek rocks were subducted and extruded into the Permo-Carboniferous arc in Quesnellia-Stikinia-
arc during the Upper Triassic –Lower Jurassic. Yukon-Tanana that collapsed in Late Permian–
Early Triassic times. Overthrusting of Stikinia
over the Slide Mountain terrane is recorded by
Geological setting clasts of Upper Permian Stikine arc rocks and eclo-
Quesnellia and Stikinia gites in Upper Permian conglomerates deposited on
the Slide Mountain units (Murphy et al. 2006). A
Quesnellia is separated from autochthonous-para- lower Norian sandy limestone, which is part of a
autochthonous Laurentian margin rocks by the dis- sequence lying unconformably upon Slide Moun-
continuous, ophiolitic Slide Mountain terrane tain, contains Carboniferous detrital zircons inferred
(Fig. 1). Furthermore, Quesnellia and Stikinia are to have been derived from the Intermontane terranes
separated by a belt of HP ophiolitic Cache Creek (Beranek & Mortensen 2006).
rocks that cuts diagonally across them in southern
British Columbia (Fig. 1). Thus, although Quesnellia Cache Creek terrane
and Stikinia are typically c. 300 km wide, the width
of Quesnellia tapers out in northern British In southern British Columbia, the Cache Creek
Columbia as the width of Stikinia tapers to zero kilo- terrane cuts diagonally across Stikinia-Quesnellia
metres in southern British Columbia. The upper (Fig. 1). It is composed of a series of thrust sheets
Palaeozoic– Lower Mesozoic geological record of cut by strike-slip faults that are made up of several
Quesnellia and Stikinia consists of (Fig. 2): (a) assemblages consisting of Pennsylvanian–Early
Mississippian–Permian, volcano-sedimentary, arc- Jurassic (Pliensbachian) slates, siltstones, grey-
related and pericratonic rocks (Stikine assemblage, wackes, cherts and carbonates (Fig. 2) (Struik
Asitka Group, Harper Ranch Group, Lay Range et al. 2001) containing Tethyan fauna (Cordey
Assemblage: Hart 1997; Colpron et al. 2006; Roots et al. 1987; Stanley 1994; Orchard et al. 2001).
et al. 2006; Beatty et al. 2006); and (b) Middle– These are accompanied by Permian and Upper
Upper Triassic Takla, Stuhini, and Nicola groups Triassic oceanic island basalts (OIB) with some
unconformably overlain by Lower –Middle Jurassic transitional to N –type mid-oceanic ridge basalts
Hazelton Group (Mortimer 1987; Monger et al. (N –MORB), Upper Permian primitive arc lavas
74 J. DOSTAL ET AL.
Fig. 2. Stratigraphic columns of Intermontane terranes and North American para-autochthonous and autochthonous
margin (sources given in text). W ¼ pebbles; and indicate transcurrent motion on the Pinchi-Fraser fault, away
and towards reader, respectively. Heavy dashed line indicates the base of the Triassic overstep sequence.
(Sitlika assemblage) and undated, supra-subduction schist and metavolcanic rocks (Figs 2 & 3). Like
zone, ultramafic rocks (Schiarizza & Massey 2000; the Stikine Takla Group, the QTG is separated from
Tardy et al. 2001). Some of these rocks underwent the overlying Lower to Middle Jurassic volcano-
HP blueschist and eclogite facies metamorphism sedimentary sequences by a regional unconformity
(200 –300 8C at 8–10 kb ¼ 28– 35 km, and 450– (Monger et al. 1991; Lang et al. 1995; Nelson &
565 8C at 12 –13 kb ¼ 42–46 km, respectively: Bellefontaine 1996). Ferri & Melville (1994)
Ghent et al. 1993, 1996) with white micas yielding divided the QTG into two units, the Slate Creek
Upper Triassic ages (221 + 2–224 + 2 Ma: Ghent Succession and the Plughat Mountain Succession.
et al. 1996). Slightly older white mica ages have The lower Slate Creek Succession is composed
been recorded in the Bridge River blueschists that mainly of fine-grained clastic sedimentary rocks con-
occur along the southern margin of Stikinia taining conodont-bearing limestones that yielded a
(Archibald et al. 1991). Similar rocks occur in Middle (Ladinian) to Late (Carnian) Triassic age.
northern British Columbia (English & Johnston The upper part of the unit interfingers with the
2005), where unroofing of some of the HP rocks in contemporaneous and dominant Plughat Mountain
the earliest Jurassic is tightly constrained. Here fos- Succession, which consists of mafic volcanic and
siliferous Pliensbachian –Toarcian cherts (c. 196– volcaniclastic rocks with minor amount of limestone
176 Ma) underwent blueschist metamorphism (at and shale. Comparisons with similar volcanic rocks
200 8C and 5 kb ¼ 17.5 km depth) that yielded to the south (Bailey 1988; Panteleyev & Hancock
white mica ages of 174 + 2 Ma, which was the 1989) led Ferri & Melville (1994) to conclude that
source for HP eclogite and rare blueschist detritus the Plughat Mountain Succession is of Norian age
that was deposited in the adjacent Whitehorse (Late Triassic). This age was later confirmed by
Trough before c. 171 Ma in front of westward fossil collections (Ferri et al. 2001a, b) and also by
vergent, Lower-Middle Jurassic thrusts (Mihalynuk palaeontological data for QTG (Stanley & Nelson
et al. 2004; English & Johnston 2005). The eastern 1996; Stanley & Senowari-Daryan 1999) collected
margin of the Cache Creek terrane is generally south of the study area.
defined by steeply dipping faults with c. 140 km
Tertiary dextral displacements, which are generally Takla Group in Stikinia
inferred to have been superimposed on easterly
The Takla Group in Stikinia (STG) is situated on the
vergent thrusts (Price & Monger 2003). Restoration
western side of the Pinchi Fault. It unconformably
of 140 km dextral offset places the Bridge River
terrane along-strike of the Cache Creek terrane. overlies the Lower Permian Asitka Group (Fig. 3),
which is made up of cherts, limestones and felsic
All of these terranes, Stikinia, Cache Creek, Ques-
volcanic rocks, and is in fault contact with the
nellia, Slide Mountain and para-autochthonous–
Cache Creek terrane (Fig. 1). Like QTG, STG is
autochthonous terranes, are unconformably overlain
composed of basaltic lava flows, volcaniclastic
by an upper Lower–Upper Triassic overstep
sequence that provides a depositional link between rocks and argillites deposited in subaerial and sub-
marine environments (Dostal et al. 1999). Fossil
them (Fig. 2) (Beranek & Mortensen 2006).
collections from the STG indicate an Upper Triassic
age in the late Carnian to early Norian stages
Local Geology (Monger & Church 1977; Monger et al. 1991).
Monger & Church (1977) divided the STG into
In north –central British Columbia, Upper Triassic three formations. The lowermost unit (Dewar For-
volcanic and sedimentary rocks in both Quesnellia mation) contains mainly tuffaceous siltstones and
and Stikinia terranes have been assigned to the argillites. The overlying Savage Mountain For-
Takla Group and are juxtaposed along the dextral mation, up to 3000 m thick, is made up of pillowed
Pinchi Fault, which passes southwards into the and massive basalt, breccia and pyroclastic rocks.
Cache Creek terrane (Fig. 1). The uppermost unit of the STG is the Moosevale
Formation, which is comprised of up to 1800 m
Takla Group in Quesnellia thick of marine and nonmarine, mafic to interme-
The Takla Group in Quesnellia (QTG) in the study diate, volcaniclastic rocks with a few lava flows.
area (Fig. 3; Germansen Landing-Manson Creek Dostal et al. (1999) documented that the lava
and Aiken Lake areas) was described by Ferri & flows from the Savage Mountain and Moosevale
Melville (1994), who estimated the total thickness formations have overlapping compositions.
of the group to be at least 5 km. The QTG overlies
and is in fault contact with the late Palaeozoic Petrography
(Mississippian to Permian) Lay Range Assemblage
(Ferri 1997), which contains internally imbricated Volcaniclastic rocks volumetrically dominate the
fault slices composed mainly of phyllite, chlorite Pughat Mountain Succession, and most interbedded
76 J. DOSTAL ET AL.
Fig. 3. Generalized geological map of Quesnellia and the westernmost part of Ancestral North America between
Manson Creek and the Ingenika River (modified after Ferri & Melville 1994 and Ferri et al. 2001a, b). The Pinchi
Fault marks the terrane boundary of Stikinia (west) and Quesnellia (east). QTG. samples are from the area between
Germansen Landing and 1268W. STG samples (see Dostal et al. 1999) are from the area of Stikinia west of the
Pinchi Fault. The inset is a map of British Columbia showing the location of the map area.
lava flows of the QTG are distinctly porphyritic groundmass. Some samples exhibit a pilotaxitic
basaltic rocks with clinopyroxene and plagioclase texture characterized by subparallel orientation of
phenocrysts. The phenocrysts are commonly the plagioclase microcrystals (Ferri & Melville
2–5 mm in size, but may reach 1 cm in length. In 1994).
addition, amphibole and rare olivine are also Plagioclase and clinopyroxene phenocrysts
present in some rocks. The groundmass is composed are typically euhedral to subhedral and the latter
of clinopyroxene, plagioclase, altered glass and often display a glomeroporphyric texture. Pyroxene
minor opaques, usually replaced by secondary min- is augite with compositional range of Wo41 – 46
erals. Accessory opaque minerals (Fe –Ti oxides) En40 – 46 Fs10 – 15 with Al2O3 ranging from 1.4–
are present both as microphenocrysts and in the 4.3 wt.% and with low TiO2 (,1.0 wt.%). The
composition of QTG clinopyroxene is comparable Newfoundland. The concentrations of Rb and Sr

to that from the STG basaltic rocks (Dostal et al. were obtained by X-ray fluorescence. A description
1999). Relatively rare, light-brown amphiboles of the analytical technique was given by Kerr et al.
occur as elongate euhedral to subhedral pheno- (1995). Measured 143Nd/144Nd values were nor-
crysts (2–4 mm in length) in plagioclase-phyric malized to a natural 146Nd/144Nd ratio of 0.7219.
lavas with minor clinopyroxene (,10%). Clasts in Replicate analyses of La Jolla standard yielded
143
the volcaniclastic rocks are mostly .1 cm in size, Nd/144Nd ¼ 0.511849 + 9. Replicate runs for
and have the same petrographic and mineralogical the NBS 987 standard gave 87Sr/86Sr ¼
characteristics as the lava flows. 0.710250 + 11. Initial Nd and Sr isotopic ratios
All the lava flows were affected by very low- and 1Nd values were calculated assuming an age of
grade metamorphism: prehnite, pumpellyite and 220 Ma for the Takla Group (Monger & Church
calcite are the most common index minerals in 1977).
QTG rocks. Prehnite and pumpellyite are abundant Like for the STG rocks (Dostal et al. 1999),
in both the groundmass and amygdules, whereas low-grade metamorphism of the QTG did not
calcite frequently fills veinlets and amygdules. significantly modify the whole-rock chemical
Other secondary minerals include chlorite, which composition. Evaluation of petrogenesis and tec-
is abundant in the groundmass, and sericite (or saus- tonic setting of the mafic rocks is based mainly on
suritization) replacing plagioclase. trace elements (e.g. high-field-strength elements
The petrographic characteristics of the mafic [HFSE] and REE), which are considered to be
rocks of the STG are similar to those of QTG. relatively ‘immobile’ in hydrothermal fluids
They both contain distinct coarse, platy plagioclase (Winchester & Floyd 1976).
and (or) clinopyroxene phenocrysts, although
the STG rocks from the type area lack amphibole.
The STG rocks were metamorphosed to zeolite Geochemistry
and prehnite –pumpelleyite facies (Monger 1977;
Dostal et al. 1999). The QTG volcanic rocks in the the Germansen
Landing –Manson Creek and Aiken Lake areas
are mainly basaltic rocks with SiO2 ,56 wt.%
Analytical methods and alteration (LOI-free), although most fall within a narrow
SiO2 range of 48–53 wt.% (Figs 4 & 5). Their
Samples were selected from a suite of specimens Mg# values (Mg/Mg þ FeT) are variable, ranging
collected during the detailed mapping of the between 0.75 and 0.35. The low contents of tran-
Germansen Landing –Manson Creek and Aiken sition elements, such as Ni and Cr in many
Lake areas (Fig. 3) by the British Columbia Geo- samples, relative to primitive mantle melts (BVSP
logical Survey, in particular by Ferri & Melville 1981), suggest that the rocks underwent extensive
(1994) and Ferri et al. (2001a, b). In addition, for fractionation. All the samples have .400 ppm Ni,
comparative purposes, we have augmented data a lower limit for melts in equilibrium with mantle
from our previous study (Dostal et al. 1999) by lherzolite. Ni and Cr show positive correlations
additional analyses of some samples of the Stikine with Mg#, indicating the early crystallization of
Takla Group. The STG samples are from the olivine and clinopyroxene. The QTG rocks display
McConnel Creek map sheet just west of the QTG a tholeiitic trend on the SiO2 versus FeOT/MgO
study area (Figs 1 & 3). plot (Fig. 5). Their low abundances of TiO2
The major and partial trace element (Cr, Ni, Sc, (0.6– 1 wt.%), HFSE and other incompatible trace
V, Rb, Ba and Sr) analyses of the whole rock elements (Table 1) are also typical of island arc
samples were done by X-ray fluorescence spec- tholeiitic suites (Pearce & Peate 1995).
trometer (Table 1). The additional trace element Using trace element abundances, the QTG volca-
analyses for rare earth elements (REE), Th, Hf, Zr, nic rocks can be separated into two groups, which
Nb and Y were performed by inductively coupled both occur throughout the studied area. The
plasma-mass spectrometry at the Memorial first group displays nearly horizontal chondrite-
University of Newfoundland (Table 1). The normalized REE patterns (Fig. 6) with (La/Yb)n
method and quality of the data were described by c. 1–2 and (La/Sm)n c. 0.75–1.25, accompanied
Longerich et al. (1990). by low La (,6 ppm), Zr (,60 ppm) and SiO2
Eight whole-rock QTG basalts were selected for (,50.5 wt.%). The rocks of the second group
Nd and Sr isotopic analyses (Table 2). Samarium have light REE-enriched patterns (Fig. 6),
and Nd abundances and Nd and Sr isotope ratios with (La/Yb)n c. 3–5 and (La/Sm)n c. 1.5–2.5.
were determined by isotope dilution mass spec- These rocks also have higher abundances of La
trometry in the laboratory of the Department of (7– 17 ppm), Zr (60–130 ppm) and SiO2
Earth Sciences at the Memorial University of (.51.5 wt.%), and higher Zr/Y ratios. However,
Table 1. Chemical composition of Quesnel Takla Group volcanic rocks
78
Sample (2 –6) (5 –4) (19 –14) (15–7) (23 –18) (11 – 10) (14 – 3) (38 – 1) (38 – 5) (38 – 12) (1 – 2) (10– 9) (7 – 9) (23 – 7)
Group 1 1 1 1 1 1 2 2 2 2 2 2 2 2
SiO2 (%) 44.71 48.96 48.87 46.78 48.37 46.84 50.2 49.69 50.57 48.97 52.03 50.82 49.9 50.62
TiO2 0.82 0.77 0.84 0.56 0.87 0.89 0.92 0.7 0.82 0.89 0.74 0.63 0.91 0.64
Al2O3 14.03 13.92 14.53 12.24 16.54 14.08 13.44 16.82 14.55 12.92 16.76 11.38 15.16 16.33
Fe2OT3 8.76 11.29 10.53 9.01 9.85 11.1 9.32 9.01 10.06 9.56 8.85 10.94 10.8 8.83
MnO 0.18 0.16 0.19 0.18 0.16 0.18 0.16 0.16 0.17 0.14 0.14 0.21 0.19 0.21
MgO 7.37 6.41 7.58 13.76 6.72 8.89 5.13 2.85 7.23 3.26 3.53 8.57 6.12 3.86
CaO 12.81 11.25 10.55 9.68 8.95 12 5.98 4.89 7.96 9.31 7.16 11.72 7.55 6.33
Na2O 2.82 2.47 3.05 0.94 3.36 2.43 3.16 4.14 2.55 2.23 2.45 1.03 3.49 3.32
K2O 0.43 1.97 0.87 1.34 1.56 0.48 2.81 4.7 2.53 2.1 1.54 2.74 1.97 4.76
P2O5 0.15 0.24 0.19 0.19 0.17 0.15 0.31 0.75 0.32 0.27 0.33 0.36 0.58 0.71
LOI 7.06 1.69 2.22 5.08 2.67 2.68 7.7 5.34 2.31 9.45 5.82 1.28 3.02 3.96
Total 99.14 99.13 99.42 99.76 99.22 99.72 99.13 99.05 99.07 99.1 99.35 99.68 99.69 99.57
Mg# 0.62 0.53 0.59 0.75 0.57 0.61 0.52 0.39 0.59 0.4 0.44 0.46 0.39 0.36
J. DOSTAL ET AL.
Cr(ppm) 260 168 229 801 90 324 130 49 222 98 39 380 120 62
Ni 104 48 83 141 78 97 61 28 86 28 31 89 42 27
Sc 41 39 40 29 41 63 30 14 32 26 22 36 32 12
V 274 329 317 199 302 338 281 308 271 240 141 260 282 217
Rb 11 73 20 21 14 10 51 158 58 43 18 41 54 90
Ba 188 520 392 410 796 213 937 794 643 671 664 1063 338 755
Sr 240 482 284 652 338 305 374 843 565 318 1338 581 399 879
Nb 3.42 1.26 1.75 0.89 1.2 1.43 7.6 2.83 6.76 8 7.33 7 6.5 1.9
Hf 1.64 1.19 1.04 1.05 1.41 1.23 2.32 1.8 2.56 2.64 2.87 0.87 1.85 1.34
Zr 52 44 44 49 51 50 93 56 90 86 121 59 110 56
Y 16 19 19 17 20 20 20 16 21 19 26 12 14 9
Th 0.78 0.93 0.62 0.54 1.34 0.88 2.49 1.64 2.54 1.95 4.78 1.34 2.32 1.22
La 4.96 5.5 3.6 2.54 3.7 4.19 13.07 11.47 13.32 10.94 15.56 6.51 14.15 9.09
Ce 11.68 12.6 9.55 6.64 9.26 10.4 26.48 21.84 26.57 24.68 32.42 14.14 28.44 17.68
Pr 1.71 1.85 1.49 1.01 1.4 1.57 3.23 2.79 3.26 3.19 3.96 1.96 3.48 2.25
Nd 8.49 9.14 7.9 5.48 7.43 7.49 13.87 12.27 14.12 14.16 15.29 9.01 14.52 9.7
Sm 2.44 2.75 2.4 1.83 2.29 2.28 3.39 3.01 3.43 3.44 3.62 2.45 3.45 2.3
Eu 0.91 0.9 0.79 0.59 0.74 0.83 1.06 1.02 1.08 1.08 1.1 0.77 1.08 0.8
Gd 3.31 3.87 3.3 2.18 3.06 2.91 4.4 3.48 4.23 4.08 3.9 3.16 4 2.81
Tb 0.49 0.51 0.5 0.34 0.47 0.44 0.6 0.49 0.57 0.55 0.56 0.4 0.56 0.38
Dy 3.41 3.26 3.15 2.49 3.25 2.87 3.49 2.98 3.52 3.32 3.83 2.41 3.61 2.32
Ho 0.69 0.68 0.64 0.46 0.63 0.56 0.75 0.57 0.74 0.69 0.78 0.51 0.69 0.47
Er 2.05 1.95 1.89 1.4 1.92 1.67 2.06 1.72 2.13 1.94 2.32 1.35 2.02 1.39
Tm 0.3 0.29 0.25 0.2 0.27 0.23 0.32 0.25 0.34 0.27 0.32 0.21 0.27 0.19
Yb 1.98 1.85 1.7 1.32 1.78 1.55 2.09 1.58 2.14 1.73 2.17 1.29 1.83 1.29
Lu 0.28 0.25 0.24 0.19 0.26 0.21 0.33 0.24 0.31 0.25 0.34 0.18 0.27 0.19
Notes: Fe2OT3 ¼ total Fe as Fe2O3; Mg# ¼ Mg/(Mg þ FeT); FeT ¼ total Fe.
Table 2. Nd and Sr isotopic composition of basaltic rocks of the Quesnel Takla Group
143
Sample No. Group Nd(ppm) Sm(ppm) Nd/144Ndm 147
Sm/144Nd 1Nd 143
Nd/144Ndi Rb(ppm) Sr(ppm) 87
Sr/86Srm 87
Sr/86Sri
19–14 1 8.08 2.42 0.512978 (10) 0.1812 7.07 0.512717 20 284 0.704290 (12) 0.703653
23–18 1 7.45 2.31 0.512951 (12) 0.1479 7.48 0.512738 14 338 0.704191 (10) 0.703816
11–10 1 7.52 2.32 0.512942 (11) 0.1478 7.31 0.512729 10 305 0.703952 (11) 0.703655
10–9 2 8.99 2.38 0.512898 (11) 0.1604 6.09 0.512667 41 581 0.704152 (15) 0.703513
EXTRUSION OF HIGH-PRESSURE CACHE CREEK ROCKS

23–7 2 10.91 2.66 0.512964 (14) 0.1473 7.76 0.512752 90 879 0.704618 (10) 0.703691
7–9 2 14.27 3.41 0.512946 (12) 0.1445 7.47 0.512738 54 399 0.704923 (9) 0.703698
14–3 2 13.91 3.4 0.512931 (10) 0.1462 7.14 0.51272 51 374 0.705824 (10) 0.703622
38–5 2 14.2 3.44 0.512978 (17) 0.1465 8.04 0.512767 58 565 0.704367 (10) 0.703438
Note: 143Nd/144Ndm and 87Sr/86Srm are measured Nd and Sr isotopic ratios, respectively, 143Nd/144Ndi and 87Sr/86Sri are initial Nd and Sr isotopic ratios, respectively, and 1Nd is the fractional difference
between the 143Nd/144Nd of rock and the bulk earth at the time of crystallization. 1Nd, 143Nd/144Ndi and 87Sr/86Sri assume an age of 220 Ma for QTG rocks. Values of 1Nd were calculated using
modern 143Nd/144NdCHUR ¼ 0.512638 and 147Sm/144NdCHUR ¼ 0.1967. Concentrations of Nd and Sm were determined by isotope dilution, and those of Rb and Sr by X-ray fluorescence.
Precision of concentrations of Nd and Sm is +1%.
is after Miyashiro (1974).
The line separating the calc-alkaline and tholeiitic fields
rocks of the Quesnel (a) and Stikine (b) Takla Group.
Fig. 5. FeOT/MgO versus SiO2 (wt.%) for the volcanic
Th/La, Ba/La, Tb/Lu (Fig. 7) and Nb/La ratios.

and many trace elements including TiO2 as well as
both groups have comparable abundances of major
desite; Bas, basanite; Trach, trachyte; Neph, nephelinite.

subalkaline basalt; AB, alkaline basalt; TrAn, trachyan-
Quesnel (a) and Stikine (b) Takla Group. Sub-AB,
Winchester & Floyd (1977) for the volcanic rocks from
Fig. 4. Zr/TiO2 versus SiO2 (wt.%) diagrams of
horizontal heavy REE patterns of the QTG rocks
ses that led to an enrichment of LILE. The nearly
mantle source modified through subduction proces-
relative to light REE, require derivation from a
elements (LILE), accompanied by depletion of Nb
Their slight enrichment in large-ion-lithophile
primitive arc environment (Pearce & Peate 1995).
characteristic of rocks emplaced in a relatively
element diagrams (Fig. 8). These patterns are
negative Nb anomalies on mantle-normalized trace
crustal contamination. Both subgroups display
groups suggest that the differences are not due to
cator of upper crustal contamination, in the two
Similar values of the Th/La ratios, a sensitive indi-
79
80 J. DOSTAL ET AL.
argue against a significant role for garnet in their

genesis and suggest mantle melting in the stabi-
lity field of spinel peridotite at a depth of less
than 60 km (White et al. 1992). The differences
between the two QTG groups cannot be readily
explained by either fractional crystallization, vari-
able degrees of partial melting or crustal conta-
mination, but are consistent with a heterogeneous
mantle source.
The QTG basalts have positive initial 1Nd values
(þ6 to þ8) and the low initial Sr isotopic ratios
(0.7034–0.7038; Fig. 9; Table 2) with the two
QTG groups having overlapping isotopic character-
istics. Differences in trace element distributions
between the two basalt groups with similar isotope
ratios suggest that the subduction-related trace
element enrichment of a mantle source occurred
Fig. 6. Chondrite-normalized REE patterns of the

basaltic rocks of the Takla Group. Normalizing values
after Sun & McDonough (1989). (a) QTG basalt group 1;
(b) QTG basalt group 2; (c) range of STG basalts (after
Dostal et al. 1999).
Fig. 8. Primitive mantle-normalized trace element

Fig. 7. Variations of Lan/Ybn versus Tbn/Lun patters of the Takla Group rocks. Normalizing values
(n-chondrite-normalized) in the QTG (þ) and STG (o) after Sun & McDonough (1989). (a) QTG basalt group 1;
basalts. Composition of STG rocks is from Dostal et al. (b) QTG basalt group 2; (c) range of STG basalts (after
(1999). Dostal et al. 1999).
shortly prior to melting so that the metasomatic Stikinia and suggest that they represent a single
process had not produced isotopic enrichment. island arc. This arc must have been in a peri-
Similar high positive 1Nd values are present in Laurentian location given the presence of the Triassic
some modern island arc tholeiites, such as those overstep sequence extending from the Yukon-
from the Mariana arc in the Pacific, where they are Tanana terrane to western margin of Laurentia
usually attributed to subduction of juvenile material (Beranek & Mortensen 2007).
(basalts with little sediment) with no involvement
of old continental crust (Pearce & Peate 1995). Arc polarity
In southernmost British Columbia, the Quesnellian
Comparison of STG and QTG segment has a width of c. 270 km and is bounded
The Takla Group in Stikinia and Quesnellia were on its western side by HP rocks of the Cache
originally mapped as a single unit (Lord 1948; Creek and Bridge River terranes: the Stikinian
Monger 1977). These two units show many strati- segment is absent (Fig. 1). Based upon variations
graphic similarities (Monger 1977; Monger & in the geochemistry of volcanic rocks in the
Church 1977; Ferri & Melville 1994; Dostal et al. Nicola Group (low-K calc-alkaline in the west chan-
1999). They are of the same age and both are com- ging to shoshonitic in the east), Mortimer (1987)
posed of mafic volcanic flows, volcanogenic sand- inferred an east-dipping Benioff zone beneath the
stones and argillites. Both Takla units show an Quesnellian segment. This is consistent with the
evolution from the lower marine assemblages to inference that the Bridge River rocks represent an
upper subaerial sections, and contain similar fossil accretionary prism (Mortimer 1986; Garver &
associations (Orchard 2006). Volcanic assemblages Scott 1995) separated from the Nicola arc by the
are characterized by clinopyroxene- and feldspar- Cadwallader –Methow terrane (fore-arc volcanic
phyric basalts with clinopyroxene typically and sedimentary rocks containing some HP rocks:
forming large phenocrysts. Gabrielse & Yorath 1992; Garver & Scott 1995;
In addition to petrographic and stratigraphic Monger & Sturik 2006).
commonalities, QTG and STG rocks have similar Similar shoshonitic, Upper Triassic rocks may
chemical (major, trace and REE elements and Nd be traced northwards along the eastern side of
and Sr isotopes) compositions (Figs 4– 8). Both the Quesnellian arc segment throughout British
suites are dominated by basaltic rocks with geo- Columbia (Garnett 1978; Pantaleyev et al. 1996;
chemical characteristics of island arc tholeiites. Nelson & Bellefontaine 1996; Nelson & Friedman
Emplacement of both QTG and STG in an island 2004). In central British Columbia, the Upper
arc environment is supported by the dominance of Triassic volcanic rocks of the Takla Group in the
basaltic over andesitic rocks and the association of Stikinian segment are less alkaline than those to
volcanic rocks with shallow marine to subaerial the east (Nelson & Bellefontaine 1996; Dostal
sediments (Dostal et al. 1999). The compositional et al. 1999; MacIntyre et al. 2001; Nelson &
similarities of QTG and STG basaltic rocks with Friedman 2004). Furthermore, the present width
some modern island arc tholeiites suggest a similar of the arc is similar to that in the southernmost
origin; that is, by melting of an upper mantle that Quesnellian segment. However, here the Quesnel-
was modified by subduction processes. Both the lian and Stikinian arc, segments are generally
STG and QTG have overlapping Nd and Sr isotopic separated by the Cache Creek HP rocks (Fig. 1).
data, high positive 1Nd, low initial Sr ratios (Fig. 9),
and their flat hREE patterns (Fig. 6), all of which are Evaluation of existing models
consistent with an island arc setting and derivation
from a similar spinel peridotite mantle source. Existing hypotheses proposed to explain the present
Dostal et al. (1999) inferred that the Takla Group distribution of Upper Triassic arc rocks separated
can be correlated with the other Upper Triassic vol- by the HP accretionary complex of the Cache
canic units in both Stikinia and Quesnellia, in par- Creek terrane (strike-slip duplication and oroclinal
ticular with the Stuhini Group of Stikinia and the bending) both require relative latitudinal displace-
Nicola Group in Quesnellia (Fig. 2) (Mortimer ments between the two arc segments. The strike-slip
1986, 1987). All these rocks are of similar age and duplication implies .1500 km northward motion of
exhibit lithological and chemical similarities. the Stikinian segment relative to the Quesnellian
segment (Wernicke & Klepacki 1988; Beck 1992).
On the other hand, the oroclinal model requires
Discussion and Conclusions that the Stikinian segment lay c. 1500 km north of
the Quesnellian segment in the Upper Triassic
The geochemical data presented here support corre- and was bent into an orocline during the Lower-
lations between the Takla Group in Quesnellia and Middle Jurassic with final closure by 187–174 Ma
82 J. DOSTAL ET AL.
Fig. 9. 1Nd versus initial 87Sr/86Sr ratio for the volcanic rocks of the Takla Group from Quesnellia (þ) and Stikinia
(o: after Dostal et al. 1999) showing also for comparison the composition of modern oceanic basalts (Zindler & Hart
1986), island-arc volcanic rocks [IAV; dashed line; sediment-contaminated (e.g., Sunda arc) and uncontaminated
(e.g., Aleutians, the Marianas) after Samson et al. 1989]. Position of mantle end-member components (DM, HIMU,
EM I and EM II) is after Zindler & Hart (1986).
(Mihalynuk et al. 1994). However, palaeomagnetic Creek HP rocks were extruded into the upper plate
data from volcanic rocks of the Upper Triassic rather than thrust over it (Fig. 10). Most extrusion
Stuhini and Lower Jurassic Hazelton groups of the models for the emplacement of HP rocks involve
Stikinian segment show zero to minimal southerly subduction of one continental block beneath
offsets relative to Laurentia: c. 300 –700 km and another followed by collision, slab breakoff
0–500 km, respectively (Irving & Wynne 1992). and extrusion back up the subduction channel
Similarly, Upper Triassic Nicola Group rocks (e.g. Ernst et al. 1997; Burov et al. 2001; Hynes
yielded a palaeolatitude consistent with its present 2002; Liou et al. 2004): in this case the HP rocks
location relative to Laurentia (Irving & Wynne mark an oceanic suture between separate terranes.
1992). This implies negligible offset between the However, HP rocks extruded into the upper plate
Stikinian and Quesnellian segments since Triassic above an active subduction zone have recently
times, a result that is consistent with recent Lower been discovered (Keppie et al. 2008), and have
Jurassic palaeomagnetic data from either side of been recorded in numerical models (Stöckhert &
the Fraser– Pinchi Fault that limits dextral move- Gerya 2005; Gerya & Stöckhert 2006). In these
ments to ,38 (Enkin et al. 2006). Thus, these data cases, the HP rocks neither separate disparate ter-
are inconsistent with both the strike-slip and ranes nor mark an oceanic suture; instead, they
oroclinal models. occur between similar units. Protoliths of such HP
rocks may include a wide range of different environ-
Extrusion model ments, including beheaded seamounts (subducting
plate) and forearc-arc material removed from the
Although an intra-arc rift interpretation for the upper plate by subduction erosion. Such protoliths
Cache Creek terrane may satisfy the palaeomagnetic have been recorded in the Cache Creek HP rocks
constraints, it cannot account for the prolonged (Tardy et al. 2001; Struik et al. 2001; English &
subduction indicated by the range of HP rocks Johnston 2005). We propose that the Cache Creek
(.50 Ma) and Carboniferous–Jurassic Asian fauna HP rocks represent trench complex and fore-arc
that implies a wide ocean, possibly as wide as material removed from the leading edge of the
3500 km using a reasonable translation rate of upper plate by subduction erosion that was taken
7 cm a21 (Gordon 1998). The geometrical enigma down to depths up to 50 km before being extruded
can be resolved if one considers that the Cache into the arc in the upper plate (Fig. 10). Whereas
Fig. 10. Block diagram showing the inferred relationship between HP rocks of the Cache Creek/Bridge River
terrane that have been extruded into the volcanic arc rocks of the Stikinia-Quesnellia during the Late Triassic and
Early-Middle Jurassic. The HP rocks are inferred to represent parts of the subducting slab (e.g. decapitated seamounts
and associated reef sediments), and parts of the fore-arc that were subducted to depths of 17–46 km before being
extruded into the upper plate: the cross-cutting relationship between the HP rocks and the Stikine-Quesnel arc is inferred
to have resulted from oblique convergence between the Palaeo-Pacific and North American plates.
subduction erosion is likely in a convergent arc, reached by some of the Cache Creek rocks within
extrusion into the upper plate may be aided by the Stikine-Quesnel arc is puzzling because one
extension of the arc. Such extrusion may be been might expect extruded HP rocks to have also come
further facilitated by trench rollback. from .90 km. However, an analogous situation
Preservation of blueschists requires relatively occurs along the southern Caribbean plate boundary
rapid exhumation, and the numerical models of in northeastern Venezuela, where oblique subduc-
Gerya & Stöckhert (2006) show that a complete tion has been related to extension of the arc and
cycle from the surface through 50– 100 km depth oblique –lateral extrusion of HP rocks (Lallemant
and back to near the surface can occur in c. 20 Ma. & Guth 1990). Such a model is consistent with the
Geochronological data for the southern Cache obliquity of the Cache Creek HP rocks relative to
Creek rocks suggest some of the HP metamorphism the Stikine– Quesnel arc (Fig. 1), and may indicate
took place at c. 223 Ma, although some rocks dextral plate convergence during the Late Triassic
(Sowchea succession: greenschist-lower amphibo- and Early Jurassic. Such a model is also consistent
lite facies) containing Pliensbachian –Toarcian radi- with the south to north transition from a normal
olarian (c. 183 + 2 Ma) were also involved in the sequence of HP trench complex-forearc– arc in
process (Struik et al. 2001), indicating post-183 Ma southernmost British Columbia and northernmost
metamorphism. In the northern Cache Creek, rocks USA through an abnormal forearc –HP –forearc
with similar-aged radiolarian underwent blueschist between 49 –508N to arc– HP–arc in the rest of
metamorphism dated at 174 + 2 Ma before being British Columbia. Thus, a model involving oblique
unroofed and eroded into rocks at least 171 Ma old extrusion of the HP Cache Creek rocks into the
(Mihalynuk et al. 2004). Thus the HP metamorphism Stikine–Quesnel arc appears to overcome most of
appears to have lasted at least c. 50 Ma. These data the apparently contradictory data and explains the
are consistent with the continuity of synchronous enigmatic relationships reported for the Intermon-
arc magmatism (with a minor hiatus in the latest tane Belt of British Columbia. A test for this model
Triassic-earliest Jurassic in Stikinia and Quesnellia) involves the nature of the faults bounding the HP
throughout the Late Triassic and Early Jurassic. rocks: thrust below and listric normal above.
Given that arc magmatism generally requires a Although a thrust lower boundary is evident in exist-
depth of .90 km for the Benioff zone (Tatsumi & ing mapping, the upper boundary is obscured by
Eggins 1995), the maximum depth of c. 17 km subsequent, Cenozoic vertical faulting; however,
84 J. DOSTAL ET AL.
further examination of this boundary may reveal (NTS 105I/13). In: E MOND , D. S., B RADSHAW ,
the nature of the Middle Jurassic boundary. G. D., L EWIS , L. L. & W ESTON , L. H. (eds) Yukon
HP rocks in ancient orogens are generally Geological Survey, 79– 91.
inferred to mark an oceanic suture between terranes B ERANEK , L. P. & M ORTENSEN , J. K. 2007. Investigating
a Triassic overlap assemblage in Yukon: On-going
that originated on opposite sides of an ocean basin. field studies and preliminary detrital zircon age data.
However, the proposal that the HP Cache Creek In: E MOND , D. S., L EWIS , L. L. & W ESTON , L. H.
rocks were extruded into the overriding plate (eds) Yukon Exploration and Geology 2006. Yukon
implies Quesnellia and Stikinia are one terrane, Geological Survey, 83– 92.
not two. It also leads to simpler palaeogeographic B REITSPRECHER , K., S COATES , J. S., A NDERSON , R. G.
reconstructions, in which subduction occurred & W EIS , D. 2007. Geochemistry of Mesozoic intru-
along an active ocean –continent boundary, not sions, Quesnel and Stikine Terranes (NTS 082; 092;
involving collision. Such a model has implications 093), South-central British Columbia: Preliminary
for other ancient orogens, for it is now imperative Characterization of Sampled Suites. British Columbia
Ministry of Energy, Mines and Petroleum Resources,
to distinguish between HP rocks extruded up along Geological Fieldwork 2006, 2007-1, 247–257.
the subduction channel (Ernst et al. 1997) from B UROV , E., J OLIVET , L., L E P OURHIET , L. & P OLIAKOV ,
those extruded into the overriding plate (Keppie A. 2001. A thermomechanical model of exhumation of
et al. 2008; this paper). Criteria include distinct high pressure (HP) and ultra-high pressure (UHP)
versus similar ‘terranes’, and a precollisional metamorphic rocks in Alpine-type collision belts. Tec-
versus collisional origin for HP belts. tonophysics, 342, 113– 136.
BVSP 1981. Basaltic Volcanism on the Terrestrial
The study was supported by Natural Sciences and Engin- Planets. Pergamon Press, New York, 1–1286.
eering Research Council of Canada grant to Dostal and C ARTER , E. S., O RCHARD , M. J., R OSS , C. A., R OSS ,
by the British Columbia Geological Survey Branch. We J. R. P., S MITH , P. L. & T PPER , H. W. 1992. Paleonto-
are grateful to R. Corney for cheerful technical assistance logical signatures of terranes. In: G ABRIELSE , H. &
and to V. Owen, B. Murphy, B. Cousens and R. Strachan Y ORATH , C. J. (eds) Geology of the Cordilleran
for their constructive reviews. Orogen in Canada. (The Geology of North America).
Geological Society of America. G-2, 28–38.
References C HURCH , B. N. 1975. Geology of the Sustut Area. British
Columbia Department of Mines and Petroleum
A RCHIBALD , D. A., S CHIARIZZA , P. & G ARVER , J. I. Resources, Geology, Exploration and Mining in
1991. 40Ar/39Ar Evidence for the Age of Igneous and British Columbia, 1974, 305–309.
Metamorphic Events in the Bridge River and Shulaps C OLEMAN , M. E. & P ARRISH , R. R. 1991. Eocene dextral
Complexes, Southwestern British Columbia. British strike-slip and extensional faulting in the Bridge River
Columbia Ministry of Energy, Mines and Petroleum Terrane, southwest British Columbia. Tectonics, 10,
Resources, Geological Fieldwork 1990, 1991-1, 1222– 1238.
75–83. C OLPRON , M., N ELSON , J. L. & M URPHY , D. C. 2006. A
B AILEY , D. G. 1988. Geology of the Central Quesnel Belt, tectonostratigraphic framework for the pericratonic
Hydraulic, South-central British Columbia (93A/12). terranes of the northern Canadian Cordillera. In:
British Columbia Ministry of Energy, Mines and C OLPRON , M. & N ELSON , J. L. (eds) Paleozoic Evol-
Petroleum Resources, Geological Fieldwork 1987, ution and Metallogeny of Pericratonic Terranes at the
1988-1, 147–153. Ancient Pacific Margin of North America, Canadian
B EATTY , T. W., O RCHARD , M. J. & M USTARD , P. S. and Alaskan Cordillera. Geological Association of
2006. Geology and tectonic history of the Quesnel Canada Special Paper, 45, 1 –23.
terrane in the area of Kamloops, British Columbia. C OLPRON , M., N ELSON , J. L. & M URPHY , D. C. 2007.
In: C OLPRON , M. & N ELSON , J. L. (eds) Paleozoic Northern Cordilleran terranes and their interactions
Evolution and Metallogeny of Pericratonic Terranes through time. GSA Today, 17, 4 –10.
at the Ancient Pacific Margin of North America, C ORDEY , F., M ORTIMER , N., D EWEVER , P. & M ONGER ,
Canadian and Alaskan Cordillera. Geological Associ- J. W. H. 1987. Significance of Jurassic radiolarians
ation of Canada Special Paper, 45, 483–504. from the Cache Creek terrane, British Columbia.
B ECK , M. E., J R . 1991. Case for northward transport Geology, 15, 1151–1154.
of Baja and coastal southern California: paleomagnetic C OWAN , D. S., B RANDON , M. T. & G ARVER , J. I. 1997.
data, analysis, and alternatives. Geology, 19, 506–509. Geological tests of hypotheses for large coastwise dis-
B ECK , M. E., J R . 1992. Tectonic significance of paleo- placements; a critique illustrated by the Baja British
magnetic results for the western conterminous United Columbia controversy. American Journal of Science,
States. In: B URCHFIEL , B. C., L IPMAN , P. W. & 297, 117 –173.
Z OBACK , M. L. (eds). The Cordilleran Orogen: D OSTAL , J., G ALE , V. & C HURCH , B. N. 1999. Upper
Conterminous US. (The Geology of North America). Triassic Takla Group volcanic rocks, Stkinine
Geological Society of America, G-3, 683– 697. Terrane, north-central British Columbia: geochemis-
B ERANEK , L. P. & M ORTENSEN , J. K. 2006. Triassic try, petrogenesis and tectonic implications. Canadian
overlap assemblages in the northern Cordillera: Pre- Journal of Earth Sciences, 36, 1483– 1494.
liminary results from the type section of the Jones E NGLISH , J. M. & J OHNSTON , S. T. 2005. Collisional
Lake Formation, Yukon and Northwest territories orogenesis in the northern Canadian Cordillera:
implications for Cordilleran crustal structure, ophiolite detrital zircons from a continental margin assemblage
emplacement, continental growth, and the terrane in the northern Coast Mountains, southeastern
hypothesis. Earth and Planetary Science Letters, 232, Alaska. Canadian Journal of Earth Sciences, 28,
333–344. 1285– 1300.
E NKIN , R. J. 2006. Paleomagnetism and the case for G ERYA , T. & S TÖCKHERT , B. 2006. Two-dimensional
Baja British Columbia. In: H AGGART , J. W., numerical modeling of tectonic and metamorphic
M ONGER , J. W. H. & E NKIN , R. J. (eds) Paleogeo- histories of active continental margins. International
graphy of the North American Cordillera: Evidence Journal of Earth Sciences, 95, 250– 274.
For and Against Large-scale Displacements. Geologi- G HENT , E. D., S TOUT , M. Z. & E RDMER , P. 1993.
cal Association of Canada Special Paper, 46, 233– 253. Pressure-temperature evolution of lawsonite-bearing
E NKIN , R. J., M AHONEY , J. B. & B AKER , J. 2006. Paleo- eclogites, Pinchi Lake, British Columbia. Journal of
magnetic signature of the Silverquick/Powell Creek Metamorphic Geology, 11, 279– 290.
succession, south-central British Columbia: reaffirma- G HENT , E. D., E RDMER , P., A RCHIBALD , A. & S TOUT ,
tion of Late Cretaceous large-scale terrane translation. M. 1996. Pressure-temperature and tectonic evolution
In: H AGGART , J. W., M ONGER , J. W. H. & E NKIN , of Triassic lawsonite-aragonit blueschists from Pinchi
R. J. (eds) Paleogeography of the North American Lake, British Columbia. Canadian Journal of Earth
Cordillera: Evidence For and Against Large-scale Sciences, 33, 800–810.
Displacements. Geological Association of Canada G ORDON , R. G. 1998. The plate tectonic approximation:
Special Paper, 46, 201– 220. plate nonrigidity, diffuse plate boundaries, and global
E RNST , W. G., M ARAYAMA , S. & W ALLIS , S. 1997. reconstructions. Annual Reviews in Earth and Plane-
Buoyancy-driven, rapid exhumation of ultrahigh- tary Sciences, 26, 615– 642.
pressure metamorphosed continental crust. Proceed- H ART , C. J. R. 1997. A Transect Across Northern Stikinia:
ings of the National Academy of Science USA, 94, Geology of the Northern Whitehorse Map-area,
9532–9537. Southern Yukon Territory (105D/13-16). Exploration
F ERRI , F. 1997. Nina Creek Group and Lay Range Assem- and Geological Services Division, Yukon Region,
blage, north-central British Columbia: remnants of late Indian and Northern Affairs Canada, Bulletin, 8,
Paleozoic oceanic and arc terranes. Canadian Journal 1– 122.
of Earth Sciences, 3, 854–874. H YNES , A. 2002. Encouraging the extrusion of deep-
F ERRI , F. & M ELVILLE , D. M. 1994. Bedrock Geology of crustal rocks in collisional zones. Mineralogical
the Germansen Landing – Manson Creek Area, British Magazine, 66, 5 –24.
Columbia. British Columbia Ministry of Energy, I RVING , E. & W YNNE , P. J. 1992. Paleomagnetism:
Mines and Petroleum Resources, Bulletin, 91, 147. review and tectonic implications. In: G ABRIELSE , H.
F ERRI , F., D UDKA , S., R EES , C., M ELDRUM , D. & & Y ORATH , C. J. (eds) Geology of the Cordilleran
W ILLSON , M. 2001a. Geology of the Uslika Lake Orogen in Canada. (The Geology of North America).
Area, North-Central British Columbia. British Geological Society of America, G-4, 61–86.
Columbia Ministry of Energy, Mines and Petroleum I RVING , E., W YNNE , P. J., T HORKELSON , D. J. &
Resources, Geoscience Map 2001–4. S CHIARIZZA , P. 1996. Large (1000 to 4000 km) north-
F ERRI , F., D UDKA , S., R EES , C. & M ELDRUM , D. 2001b. ward movements of tectonic domains in the northern
Geology of the Aiken Lake area, North-Central British Cordillera, 83 to 45 Ma. Journal of Geophysical
Columbia. British Columbia Ministry of Energy, Research, 101, 17 901–17 916.
Mines and Petroleum Resources, Geoscience Map J OHNSTON , S. T. & B OREL , G. D. 2007. The odyssey of
2001–10. the Cache Creek terrane, Canadian Cordillera: impli-
G ABRIELSE , H. & Y ORATH , C. J. (eds) 1992. Geology cations for accretionary orogens, tectonic setting of
of the Cordilleran Orogen in Canada. Geology of Panthalassa, the Pacific superswell, and break-up of
Canada No. 4, Geological Survey of Canada, Pangea. Earth and Planetary Science Letters, 253,
Ottawa, 1–844. 415– 428.
G ABRIELSE , H., M URPHY , D. C. & M ORTENSEN , J. K. K EPPIE , J. D., D OSTAL , J., M URPHY , J. B. & N ANCE ,
2006. Cretaceous and Cenozoic dextral orogen-parallel R. D. 2008. Synthesis and tectonic interpretation of
displacements, magmatism and paleogeography, north- the westernmost Paleozoic Variscan orogen in
central Canadian Cordillera. In: H AGGART , J. W., southern Mexico: from rifted Rheic margin to active
M ONGER , J. W. H. & E NKIN , R. J. (eds) Paleogeogra- Pacific margin. Tectonophysics, http://dx.doi.org/10.
phy of the North American Cordillera: Evidence For 1016/j.tecto.2008.01.012.
and Against Large-scale Displacements. Geological K ERR , A., J ENNER , G. A. & F RYER , B. J. 1995. Sm-Nd
Association of Canada Special Paper, 46, 255–276. isotopic geochemistry of Precambrian to Paleozoic
G ARNETT , J. A. 1978. Geology and Mineral Occurrences granitoid suites and the deep-crustal structure of the
of the Southern Hogem Batholith. British Columbia southeast margin of the Newfoundland Appalachians.
Ministry of Mines and Petroleum Resources, Bulletin, Canadian Journal of Earth Sciences, 32, 224–245.
70, 1– 75. L ALLEMANT , H. G. A. & G UTH , L. R. 1990. Role of
G ARVER , J. I. & S COTT , T. J. 1995. Trace elements in extensional tectonics in exhumation of eclogites and
shale as indicators of crustal provenance and terrane blueschists in an oblique subduction setting: northeast-
accretion in the southern Cordillera. Geological ern Venezuela. Geology, 18, 950–953.
Society of America Bulletin, 107, 440– 453. L ANG , J. R., L UECK , B., M ORTENSEN , J. K., R USSEL , J. K.,
G EHRELS , G. E., M C C LELLAND , W. C., S AMSON , S. D. S TANLEY , C. R. & T HOMPSON , J. F. H. 1995. Triassic-
& P ATCHETT , P. J. 1991. U– Pb geochronology of Jurassic silica-undersaturated and silica-saturated
86 J. DOSTAL ET AL.
alkalic intrusions in the Cordillera of British Columbia. Columbia. Canadian Journal of Earth Sciences, 24,
Geology, 23, 451– 454. 2521– 2536.
L IOU , J. G., T SUJIMORI , T., Z HANG , R. Y., K ATAYAMA , M URPHY , D. C., M ORTENSEN , J. K., P IERCEY , S. J.,
I. & M ARUYAMA , S. 2004. Global UHP meta- O RCHARD , M. J. & G EHRELS , G. E. 2006. Mid-
morphism and continental subduction/collision: the Paleozoic to early Mesozoic tectonostratigraphic
Himalayan model. International Geology Review, evolution of Yukon-Tanana and Slide Mountain ter-
46, 1– 17. ranes and affiliated overlap assemblages, Finlayson
L ONGERICH , H. P., J ENNER , G. A., F RYER , B. J. & Lake massive sulphide district, southeastern Yukon.
J ACKSON , S. E. 1990. Inductively coupled plasma- In: C OLPRON , M. & N ELSON , J. L. (eds) Paleozoic
mass spectrometric analysis of geological samples: a Evolution and Metallogeny of Pericratonic Terranes
critical evaluation based on case studies. Chemical at the Ancient Pacific Margin of North America,
Geology, 83, 105– 118. Canadian and Alaskan Cordillera. Geological
L ORD , C. S. 1948. McConnell Creek Map-area, Cassiar Association of Canada Special Paper, 45, 75–105.
District, British Columbia. Geological Survey of N ELSON , J. L. & B ELLEFONTAINE , K. A. 1996. The
Canada, Memoir, 251, 1– 72. geology and mineral deposits of north-central Quesnel-
M AC I NTYRE , D. G. 2006. Geology and Mineral Deposits lia; Tezzeron Lake to Discovery Creek, Central British
of the Skeena Arch, West-Central British Columbia: A Columbia. British Columbia Geological Survey
Geoscience BC Digital Data Compilation Project. Branch Bulletin, 99, 112.
British Columbia Ministry of Energy, Mines and N ELSON , J. L. & F RIEDMAN , R. 2004. Superimposed
Petroleum Resources, Geological Fieldwork 2005, Quesnel (late Paleozoic-Jurassic and Yukon-Tanana
2006-1, 303–312. (Devonian-Mississippian) arc assemblages, Cassiar
M AC I NTYRE , D. G., V ILLENEUVE , M. E. & S CHIAR- Mountains, northern British Columbia: field, U– Pb,
IZZA , P. 2001. Timing and tectonic setting of Stikine and igneous petrochemical evidence. Canadian
Terrane magmatism, Babine-Takla lakes area, central Journal of Earth Sciences, 41, 1201– 1235.
British Columbia. Canadian Journal of Earth N ELSON , J. L. & M IHALYNUK , M. 1993. Cache Creek
Sciences, 38, 579–601. ocean: closure or enclosure. Geology, 21, 173– 176.
M IHALYNUK , M. G., N ELSON , J. & D IAKOW , L. J. 1994. N ELSON , J. L., C OLPRON , M., P IERCEY , S. J.,
Cache Creek terrane entrapment: Oroclinal paradox D USEL -B ACON , C., M URPHY , D. C. & R OOTS , C.
within the Canadian Cordillera. Tectonics, 13, F. 2006. Paleozoic tectonic evolution and metallogenic
575– 595. evolution of the pericratonic terranes in Yukon, north-
M IHALYNUK , M. G., E RDMER , P., G HENT , E. D., ern British Columbia and eastern Alaska. In:
C ORDEY , F., A RCHIBALD , D. A., F RIEDMAN , R. M. C OLPRON , M. & N ELSON , J. L. (eds) Paleozoic Evol-
& J OHANNSON , G. G. 2004. Coherent French Range ution and Metallogeny of Pericratonic Terranes at the
Blueschist: Subduction to Exhumation in ,2.5 m.y.? Ancient Pacific Margin of North America, Canadian
Geological Society of America Bulletin, 116, and Alaskan Cordillera. Geological Association of
910– 922. Canada Special Paper, 45, 323– 360.
M IYASHIRO , A. 1974. Volcanic rock series in island arcs O RCHARD , M. J. 2006. Late Paleozoic and Triassic
and active continental margins. American Journal of conodont faunas of Yukon and northern British
Sciences, 274, 321– 357. Columbia and implications for the evolution of the
M ONGER , J. W. H. 1977. The Triassic Takla Group in Yukon-Tanada terrane. In: C OLPRON , M. &
McConnell Creek Map Area, North-Central British N ELSON , J. L. (eds) Paleozoic Evolution and Metallo-
Columbia. Geological Survey of Canada, Paper, geny of Pericratonic Terranes at the Ancient Pacific
76–29, 145. Margin of North America, Canadian and Alaskan Cor-
M ONGER , J. W. H. & C HURCH , B. N. 1977. Revised stra- dillera. Geological Association of Canada Special
tigraphy of the Takla Group, north-central British Paper, 45, 229–260.
Columbia. Canadian Journal of Earth Sciences, 14, O RCHARD , M. J., S TRUIK , L. C., R UI , L., B AMBER , B.,
318– 326. M AMET , B., S ANO , H. & T AYLOR , H. 2001. Paleonto-
M ONGER , J. W. H. & S TRUIK , L. C. 2006. Chilliwack logical and biogeographical constraints on the Carbon-
terrane: a slice of Stikinia? A tale of terrane transfer. iferous to Jurassic Cache Creek terrane in central
In: C OLRON , M. & N ELSON , J. L. (eds) Paleozoic British Columbia. Canadian Journal of Earth
Evolution and Metallogeny of Pericratonic Terranes Sciences, 38, 551– 578.
at the Ancient Pacific Margin of North America, P ANTELEYEV , A. & H ANCOCK , K. D. 1989. Quesnel
Canadian and Alaskan Cordillera. Geological Associ- Mineral Belt: Summary of the Geology of the Beaver
ation of Canada Special Paper, 45, 351–368. Creek-Horselly Map Area. British Columbia Geologi-
M ONGER , J. W. H., W HEELER , J. O. ET AL . 1991. Cordil- cal Survey Branch, Geological Fieldwork 1988, 1989-
leran terranes. In: G ABRIELSE , H. & Y ORATH , C. J. 1, 159–166.
(eds) Geology of the Cordilleran Orogen in Canada. P ANTELEYEV , A., B AILEY , D. G., B LOODGOOD , M. A. &
(The Geology of North America). Geological Society H ANCOCK , K. D. 1996. Geology and Mineral Deposits
of America, G-2, 281– 327. of the Quesnel River-Horsefly Map Area, Central
M ORTIMER , N. 1986. Late Triassic arc-related potassic Quesnel Trough. British Columbia Ministry of
igneous rocks in the North American Cordillera. Employment and Investment, Bulletin, 97, 156.
Geology, 14, 1035– 1038. P EARCE , J. A. & P EATE , D. W. 1995. Tectonic impli-
M ORTIMER , N. 1987. The Nicola Group: Late Triassic and cations of the composition of volcanic arc magmas.
Early Jurassic subduction-related volcanism in British Annual Review of Earth Sciences, 23, 251– 285.
P RICE , R. & M ONGER , J. 2003. A transect of the southern S TANLEY , G. D., J R . & S ENOWBARI -D ARYN , B. 1999.
Canadian Cordillera from Calgary to Vancouver. 2003 Upper Triassic reef fauna from Quesnel terrane,
Annual Meeting of GAC/MAC Calgary, fieldguide, 16. Central British Columbia, Canada. Journal of Paleon-
R OOTS , C. F., N ELSON , J. L., S IMARD , R.-L. & H ARMS , tology, 73, 787– 802.
T. A. 2006. Continental fragments, mid-Paleozoic arcs S TÖCKHERT , B. & G ERYA , T. V. 2005. Pre-collisional
and overlapping late Paleozoic arc and Triassic sedi- high pressure metamorphism and nappe tectonics at
mentary strata in the Yukon-Tanana terrane of northern active continental margins: a numerical simulation.
British Columbia and southern Yukon. In: C OLPRON , Terra Nova, 17, 102 –110.
M. & N ELSON , J. L. (eds) Paleozoic Evolution and S TRUIK , L. C., S CHIARIZZA , P. ET AL . 2001. Imbricate
Metallogeny of Pericratonic Terranes at the Ancient architecture of the upper Paleozoic to Jurassic
Pacific Margin of North America, Canadian and oceanic Cache Creek Terrane, central British Colum-
Alaskan Cordillera. Geological Association of bia. Canadian Journal of Earth Sciences, 38, 495–514.
Canada Special Paper, 45, 153– 177. S UN , S. S. & M C D ONOUGH , W. F. 1989. Chemical and
S AMSON , S. D., M C C LELLAND , W. C., P ATCHETT , P. J., isotopic systematics of oceanic basalts: implications
G EHRELS , G. E. & A NDERSON , R. G. 1989. Evidence for mantle composition and processes. In: S AUNDERS ,
from neodymium isotopes for mantle contributions to A. D. & N ORRY , M. J. (eds) Magmatism in the Ocean
Phanerozoic crustal genesis in the Canadian Cordillera. Basins. Geological Society, London, Special Publi-
Nature, 337, 705– 709. cations, 42, 313– 345.
S AMSON , S. D., P ATCHETT , P. J., M C C LELLAND , W. C. S YMONS , D. T. A., H ARRIS , M. J., M C C AUSLAND ,
& G EHRELS , G. E. 1991. Nd isotopic characterization P. J. A., B LACKBURN , W. H. & H ART , C. J. R.
of metamorphic rocks in the Coast Mountains, Alaskan 2005. Mesozoic-Cenozoic paleomagnetism of the
and Canadian Cordillera: ancient crust bounded by Intermontane and Yukon-Tanana terranes, Canadian
juvenile terranes. Tectonics, 10, 770– 780. Cordillera. Canadian Journal of Earth Sciences, 42,
S CHIARIZZA , P. & M ASSEY , N. 2000. Volcanic facies and 1163– 1185.
geochemistry of the Sitlika assemblage, central British T ARDY , M., L APIERRE , H., S TRUIK , L. C., B OSCH , D. &
Columbia: part of a primitive intra-oceanic arc in B RUNET , P. 2001. The influence of mantle plume in
western Cache Creek terrane. Geological Society of the genesis of the Cache Creek oceanic igneous
America Abstracts with Programs, Cordilleran rocks: implications for the geodynamic evolution of
Section, 32(16), A66. the inner accreted terranes of the Canadian Cordillera.
S IMARD , R.-L., D OSTAL , J. & R OOTS , C. F. 2003. Devel- Canadian Journal of Earth Sciences, 38, 515–534.
opment of late Paleozoic volcanic arcs in the Canadian T ATSUMI , Y. & E GGINS , S. 1995. Subduction Zone Mag-
Cordillera: an example from the Klinkit Group, north- matism. Blackwell Publishing, Massachusetts, USA.
ern British Columbia and southern Yukon. Canadian W ERNICKE , B. & K LEPACKI , D. W. 1988. Escape hypoth-
Journal of Earth Sciences, 40, 907 –924. esis for the Stikine block. Geology, 16, 461– 464.
S TANLEY , G. D. J. 1994. Late Paleozoic and early Meso- W HITE , R. S., M C K ENZIE , D. & O’N IONS , R. K. 1992.
zoic reef-building organizations and paleogeography: Oceanic crustal thickness from seismic measurements
the Tethyan-North American connection. Courier and REE inversions. Journal of Geophysical Research,
Forschungsinstitut Seckenberg, 172, 69– 75. 97, 19683– 19715.
S TANLEY , G. D., J R . & N ELSON , J. L. 1996. New Investi- W INCHESTER , J. A. & F LOYD , P. A. 1977. Geochemical
gations on Eaglenest Mountain, Northern Quesnel discrimination of different magma series and their
Terrane: An Upper Triassic Reef Facies in the Takla differentiation products using immobile elements.
Group, Central British Columbia (93N/IE). British Chemical Geology, 20, 325–343.
Columbia, Ministry of Energy, Mines and Petroleum Z INDLER , A. & H ART , S. R. 1986. Chemical geody-
Resources, Geological Survey Branch, Geological namics. Annual Review of Earth and Planetary
Fieldwork 1995, 1996-1, 127– 135. Sciences, 14, 493–571.
Plate tectonics of the Alpine realm
GÉRARD M. STAMPFLI* & CYRIL HOCHARD
Institute of Geology and Palaeontology, University of Lausanne, Anthropole,
CH 1015 Lausanne, Switzerland
*Corresponding author (e-mail: gerard.stampfli@unil.ch)
Abstract: New field data on the East Mediterranean domain suggest that this oceanic basin
belonged to the larger Neotethyan oceanic system that opened in Permian times. A Greater
Apulia domain existed in Mesozoic times, including the autochthonous units of Greece and SW
Turkey. It also included a united Adria and Apulia microplate since Early Jurassic times. This
key information implies that a new post-Variscan continental fit for the western Tethyan area is
necessary, where the relationships between the Adriatic, Apulian and Iberian plates are defined
with greater confidence. To construct a reliable palinspastic model of the Alpine realm, plate
tectonic constraints must be taken into consideration in order to assess the magnitude of lateral dis-
placements. For most of the plates and their different terranes, differential transport on the scale of
thousands of kilometres can be demonstrated. This plate tectonic framework allows a better
geodynamic scenario for the formation of the Alpine chain to be proposed, where the western
and eastern transects have experienced contrasting geological evolutions. The eastern Alps–
Carpathians domain evolved from the north-directed roll-back of the Maliac– Meliata slab and
translation of the Meliata suture and Austroalpine domain into the Alpine domain. In the
western Alps, the changing African plate boundary in space and time defined the interaction
between the Iberian– Briançonnais plate and the Austroalpine accretionary wedge.
Introduction For a long time, the Tethys was considered as a

large and single oceanic space, mostly of Mesozoic
To construct a plate tectonic scenario for the age, located between Gondwana and Eurasia.
Alpine domain is not an easy task, as the wealth of Already in the 1940s and 1950s, a distinction
information on these regions cannot be grasped, between a Palaeo- and a Neo-Tethys appeared (see
nor properly cited, in a single publication. Therefore references in Sengör 1985; Stampfli & Kozur
here we try to improve on our former attempts 2006) and it was recognized that the latter com-
at reconstructing one of the most complicated prised marine Permian and younger strata,
areas of the planet (Fig. 1), by applying plate whereas the former opened during the Early Palaeo-
tectonic concepts as much as possible, and by zoic. Stöcklin (1968, 1974), following extensive
working on a large enough scale to integrate the field work in Iran, gave a formal definition of
Alps into the geodynamic framework of the these two large oceanic entities, the Neotethys
western Tethyan area, and also integrating infor- becoming a Permian to Cretaceous peri-Gondwanan
mation from the Atlantic domain, such as magnetic ocean (whose suture was between Iran and Arabia),
anomalies (Fig. 2). whereas the Palaeotethys suture was located just
Our approach is totally ‘non-fixist’ – meaning north of Iran, thus between the Cimmerian terranes
that if two terranes are now juxtaposed they most (sensu Sengör 1979) and Eurasia. In that sense, the
likely were not so before, and we would like to ded- Palaeotethys separates the Variscan domain from
icate this paper to H. Schardt who, more than a late Palaeozoic Gondwana-derived terranes.
century ago (e.g. Schardt 1889, 1898, 1900, 1907), Besides the two large Palaeotethyan and Neo-
proposed that the Briançonnais domain of the tethyan oceanic domains (one replacing the other
western Prealps was an exotic terrane. Through his during the Triassic) many oceanic back-arc type
1900 paper ‘encore les régions exotiques’ (again oceans opened just north of the Palaeotethys
the exotic regions) one can see that his ‘non-fixist’ subduction zone. They are sometimes erroneously
ideas were strongly rejected at that epoch, we considered as Neotethyan because of their Triassic
expect present geologists will be more open- to Jurassic age, but most of these had no direct
minded. Schardt’s proposal was finally proven connection (either geographical or geological)
correct, and the Briançonnais terrane is a good with the peri-Gondwanan Neotethys ocean, and
example of margin duplication in an orogen should therefore be called by their local names
(Frisch 1979; Stampfli 1993; Stampfli et al. 2002). (e.g. Meliata, Küre, Maliac, Pindos, Huglu,
90 G. M. STAMPFLI & C. HOCHARD
Fig. 1. Western Tethys terrane map. Line with triangle marks the limit between the African and Eurasian plates,
grey areas are terranes implied in the Alpine orogenic events, darker grey terranes represent fragments of the former
Iberian plate. Black area represents the Vardar suture zone, black dash line the Izmir Ankara suture zone. Thick lines
with ticks represent the passive margin of respectively the East Mediterranean basin and the Gulf of Biscay.
Vardar) (Stampfli & Kozur 2006). During the ophiolitic belts found in the Tethyan realm,
break-up of Pangaea, another relatively long, if whereas older oceanic domains totally disappeared
not wide, oceanic domain appeared, consisting of without leaving large remnants of their sea floors
the Central Atlantic and its eastern extension in (Stampfli & Borel 2002).
the Alpine –Carpathian domain. This Jurassic
ocean was named ‘Alpine Tethys’ (Favre & Stamp-
fli 1992), in order to mark the difference between The East Mediterranean Neotethys
this relatively northerly ocean (opening in the Var- connection
iscan domain) and the peri-Gondwanan Neotethys.
This Alpine Tethys is made of oceanic segments In the light of the large amount of new data provided
represented by the Ligurian, Piemont and Penni- by the CROP Atlas (Finetti 2005), the geodynamic
nic/Vahic ophiolitic sequences (e.g. Lagabrielle evolution of the Adria and Apulia micro-continents
et al. 1984; Liati et al. 2005). has been recently redefined (Stampfli 2005). One
The resulting picture of the Tethys realm in Jur- of the main issues is the age of the East
assic time is, therefore, quite complex, and made of Mediterranean –Ionian sea basin, and the nature of
numerous small oceans and a large peri-Gondwanan the sea floor in this area. This, in turn, influences
Neotethys. Further complexity arose during the con- Mesozoic continental re-assembly models.
vergence stages, as many of these oceanic realms This problem was already reviewed in some
gave birth to new back-arc basins, especially detail in previous publications (Stampfli 2000;
around Turkey and Iran (Moix et al. 2008). These Stampfli et al. 2001b); depending on the authors,
are, in most cases, at the origin of the many the East Mediterranean –Ionian sea basin is regarded
PLATE TECTONICS OF THE ALPINE REALM 91
Fig. 2. Wander path of some terranes involved in the Alpine orogen. Central Iberia, Apulia, Adria, the western part
of the Briançonnais and the North Calcareous Alps are shown in dark grey in their 220 Ma position, and in light
grey in their present-day position. Their related velocities are shown in Table 1.
as already opening in the Late Palaeozoic (Vai complicated scheme should be envisaged when con-
1994) or as late as the Cretaceous (e.g. Dercourt sidering that Apulia belonged for a while to the
et al. 1985, 1993). Most people would regard drifting Cimmerian continent, whose rotation point
this ocean as opening in the Late Triassic or Early changed in time. Therefore, a slightly different
Jurassic (e.g. Garfunkel & Derin 1984; Sengör placing and rotation of Apulia is shown in the
et al. 1984; Robertson et al. 1996) and therefore following reconstructions.
possibly related to the Alpine Tethys–Central
Atlantic opening.
A new interpretation (Stampfli et al. 2001b), The Apulia – Adria problem and
showing that the East Mediterranean domain corre- the Late Triassic fit
sponded to a westward extension of the Permian
Neotethys, is supported by a large volume of geo- The late Triassic reconstruction shown in Figure 3
logical and geophysical data. We proposed a incorporates all the terranes/geological elements
Middle to Late Permian onset of sea-floor spreading shown on the present-day map (Fig. 1). The main
in the eastern Mediterranean basin, concomitant problem regarding this model is the position of
with the opening of the Neotethys eastward, and Adria and Apulia. The present-day shape of these
the northward drift of the Cimmerian continents two terranes (present-day Italy) cannot be fitted
since late Early Permian. This model also implies between Iberia and Africa without some important
a late closure of the Palaeotethys (Middle to Late deformation in classical fits and remains a major
Triassic) on a Mediterranean transect (Stampfli problem at the centre of the western Tethyan
et al. 2003; Stampfli & Kozur 2006). realm (e.g. Wortmann et al. 2001; Schettino &
Thus, the original position of Apulia with respect Scotese 2002).
to Africa can be well determined by closing the The continuity between the active subduction
Ionian sea by a c. 408 rotation of Apulia around a zone under Greece and the outer Dinarides
point located north of Tunisia (Finetti 2005). This (Wortel & Spakman 1992; de Jonge et al. 1994)
represents the total rotation of Apulia from Late shows that there is a possible plate limit between
Carboniferous to Late Triassic, but a more Apulia s.l. and the autochthonous terrane of
92
Table 1. Velocities and distances to rotation poles of peri-Alpine terranes (see the paths in Fig. 2). Ages are expressed in Ma, and velocities correspond to the
centroids velocities and are expressed in cm/a
Initial age Final age Adria Apulia Central Iberia North Calcarous Alps West Briançonnais
Velocity Distance to Velocity Distance to Velocity Distance to Velocity Distance to Velocity Distance to
G. M. STAMPFLI & C. HOCHARD

(cm/a) pole (8) (cm/a) pole (8) (cm/a) pole (8) (cm/a) pole (8) (cm/a) pole (8)
220 200 0.26 27.71 0.44 31.92 0.16 32.58 0.52 27.48 0.00 2.58
200 180 0.75 28.15 0.84 31.79 0.00 2.86 0.16 16.86 0.00 2.58
180 165 0.70 36.57 1.53 29.68 0.76 33.57 0.64 52.10 0.00 42.54
165 155 1.33 25.06 1.53 29.18 0.75 36.80 0.89 16.44 0.00 42.54
155 142 2.16 34.30 2.35 37.65 2.26 38.86 1.23 8.11 0.52 14.30
142 131 1.38 32.06 1.51 35.08 0.62 5.46 0.98 40.73 0.58 11.03
131 121 1.08 29.94 0.97 26.68 1.48 40.58 1.26 59.34 1.18 30.71
121 112 2.33 26.24 2.73 31.14 2.24 25.06 0.89 9.15 1.92 21.30
112 103 1.36 10.96 2.00 16.27 1.37 11.04 1.60 17.50 0.70 5.63
103 95 2.89 12.92 4.05 18.25 2.29 10.20 1.00 13.78 1.73 7.69
95 84 1.98 28.64 2.28 33.39 1.10 15.35 1.83 51.40 1.73 24.65
84 70 1.63 17.62 2.07 22.58 0.59 6.29 0.96 12.29 1.13 21.55
70 57 0.25 66.92 0.24 62.41 0.27 83.39 0.23 5.04 0.00 44.67
57 48 0.73 35.00 0.75 35.71 0.71 23.91 2.25 19.27 0.41 18.71
48 40 1.45 23.15 1.69 27.32 0.37 6.64 1.18 6.09 1.18 8.02
40 33 1.42 21.43 1.69 25.83 0.17 6.09 0.58 3.47 0.88 1.74
33 20 1.45 27.12 1.65 31.19 0.07 9.93 0.50 6.74 0.55 2.75
20 0 0.74 36.68 0.76 37.89 0.00 40.57 0.27 4.69 0.40 1.58
Greece. However, we regard this feature as recent were tried and many resulted in aborted rifts.
and as having no bearing on the fact that a Greater This pervasive Jurassic rifting finally gave birth,
Apulia plate existed in Mesozoic times, in which in the Alps, to passive margins, flanked by aborted
all the autochthonous portion of Greece was rifts that became rim basins (e.g. Subbriançonnais,
included (PIM, Tor), as well as the Bey –Daglari Helvetic-Dauphinois domains, Lombardian basin,
(Bdg) of SW Turkey, representing the northern Subbetic basin, Magura basin), and a narrow
margin of the East Mediterranean basin (Moix oceanic strip dominated by mantle denudation
et al. 2008). The apparent present plate limit (Stampfli & Marchant 1997; Rampone & Piccardo
(between Ap and PIM, Fig. 1) comes from the fact 2000). This can be regarded as forced rifting
that the Hellenic orogenic/accretionary wedge is through an already thinned lithosphere, and
oblique with respect to former palaeogeographic effectively all this rifting took place under water,
domains. It is therefore still colliding with Apulia with little isostatic/thermal rebound of the rift
on an Albanides transect while still subducting the shoulders.
East Mediterranean sea floor on a Greek transect. Thus, the final pre-collisional length and geome-
The CROP seismic lines through the Adriatic try of the Adriatic plate would have been established
domain (Finetti 2005) clearly show that there is no by Middle–Late Jurassic times only (Figs 4 & 5).
major tectonic accident cutting Italy into two Deformation of this plate interior during Alpine
units, at least since the Jurassic. Thus, Italy must times also took place, but only at a small scale as
have reached its present configuration between the shown by the seismic data, and despite the fact
Triassic and Middle Jurassic. Palaeomagnetic data that most of its border was strongly involved in sub-
show that the Apulian plate s.l. (Italy) suffered duction and/or crustal shortening.
relatively little rotation in regard to Africa since The larger Apulia domain had already gone
the Triassic (e.g. Channell 1992, 1996; Muttoni through major rifting phases in the Permian and
et al. 2001). Triassic, even leading to the opening of small
Our basic hypothesis is that the Apulian part of oceans between Greater Apulia and the Austroal-
Italy was definitely an African promontory pine (AA) domain. This finally led to the subduction
(Argnani 2002) from Middle Triassic to Recent of large amounts of continental crust material due to
times, without much displacement with respect to the negative buoyancy of former rift zones and atte-
Africa, and that the Adriatic and Apulian micro- nuated margins, and regions on which large-scale
plates were welded in an Eocimmerian collision ophiolite obduction had taken place. On a Hellinides
phase during the Middle–Late Triassic, when both transect this was calculated to be in the order of
units became part of the African plate. 900 km of subducted sub-upper-crust continental
The need to cut Italy into two microplates comes lithosphere (van Hinsbergen et al. 2005), and
from the fact that in a Triassic Pangaean reconstruc- from our reconstructions a similar amount can be
tion, there is very little room to insert the present calculated through a Dinarides –Balkan transect
form of Italy in its proper place. In order to solve (e.g. compare Figs 3 & 12).
this dilemma, we had to reconsider the fit of Iberia
with Europe, as well as the size and position of the Construction of the models
Alboran domain microplates. We made a much
tighter fit of these elements with Europe, following The reconstructions shown in Figures 3 to 12 are
similar previous proposals (Srivastava & Tapscott based on a tight pre-Pangaea break-up Permian fit
1986; Srivastava et al. 1990). Still, there was not as explained above. From the Early Jurassic
enough room to put the entire present length of onward they are based on magnetic anomalies
Italy in its proper position. The conclusion was from the Central Atlantic. Plate tectonic concepts
that Italy was shorter in Late Triassic times than it have been systematically applied to our palinspatic
is now (or already in the Late Jurassic). A few models of the western Tethys, moving away from
hundred kilometres were gained through major pure continental drift models, not constrained by
phases of rifting affecting Greater Apulia since the plate limits, to produce a model which is increas-
Triassic and related to the break-up of Pangaea ingly self-constrained. In this approach, first
and the opening of the Alpine Tethys [Alpine Lom- explained in Stampfli & Borel 2002, inter-dependent
bardian basin in Italy and Ionian rift system affect- reconstructions are created from the past to the
ing the Hellenic domain (Stampfli 2005)]. present. Except during collisions, plates are moved
The reason for widespread Jurassic rifting affect- step by step, as single rigid entities. The only evol-
ing Italy was that for the Pangaean break-up to ving elements are the plate boundaries, which are
succeed, the Atlantic rift system had to join an preserved and follow a consistent geodynamic
active plate limit located far to the east in the Neo- evolution through time and an interconnected
tethyan domain. Thus, all possible ways to break network through space. Hence, lithospheric plates
through the Alpine– Mediterranean lithosphere are constructed by adding to, or removing oceanic
Fig. 3. 220– 200 Ma. Late Triassic. As the Palaeotethys (PaT) subduction reached its final stage, slab roll-back along its
northern margin accelerated and was marked by the opening of successive oceanic back-arc basins before the final
closure of Palaeotethys in late Carnian times. Extension affected southwestern Eurasia and north Africa. A possible link
with the north Atlantic rift system existed through the Pyrenean rift. The East Mediterranean part of Neotethys ceased
spreading; on its northern margin, Greater Apulia (LNg, Ap, PIM, Tor, Bdg, Tau) represents at this time the westernmost
part of the Cimmerian terranes, detached now from the Iranian blocks (SS). Triassic-Jurassic boundary. At this time
Palaeotethys was completely closed. The Central Atlantic rift widened but had difficulty finding a way to link with a
plate limit to the east. The closure of Palaeotethys south of the Küre basin generated the southward subduction of the
latter. Slab roll-back, both in Küre and the Neotethys, allowed the opening of the Izmir-Ankara (IzAn) ocean. See
Figure 1 for abbreviation. All reconstructions are in spherical equidistant projection, centred 20E20N. Symbols:
1: passive margin; 2: magnetic or synthetic anomalies; 3: seamount; 4: intraoceanic subduction; 5: mid-ocean ridge;
6: active margin; 7: active rift; 8: inactive rift (basin); 9: collision zone; 10: thrust; 11: suture. Oceanic lithosphere in
black. Abbreviations for Figures 3 to 12: AA, Austro– Alpine; Abr, Abruzzi; ACy, Attica–Cyclades; Adr, Adria; Ana,
Anatolides; And, Andrusov; AnT, Antalya; APr, Algero– Provençal; Ap, Apulia; Apu, Apuseni; Bal, Baleares; Bdg,
Beydaglari; BDu, Bosnia–Durmitor; Bet, Betic; Big, Biga; Bri, Briançonnais; BS, Black Sea; Buc, Bucovinian; Bud,
Budva; Cal, Calabria; Car, Carnic; CLu, Campania–Lucania; Cor, Corsica; Dac, Dacides; Dal, Dalmatian; Dan,
Danubian; Dau, Dauphinois; DoN, Dobrogea North; DoS, Dobrogea South; ECa, East-Carpathian; EPt, East-Pontides;
Era, Eratostene; GCa, Great Caucasus; Get, Getic; Gos, Gosau; GTP, Gavrovo-Tripolitza-Pindos; Hat, Hatay; Hel,
material (symbolized by synthetic isochrones) from The Jurassic ocean: Alpine Tethys, Central
major continents and terranes. Atlantic and Vardar (Figs 3, 4 & 5)
In recent years we changed our tools and moved
into GIS softwares and built a geodynamic database The results of field-work on the Canary Islands and
to support the reconstructions, and the model was, in Morocco (Favre et al. 1991; Favre & Stampfli
and still is, extended to the whole globe (Hochard 1992; Steiner et al. 1998) indicate that the onset of
2008). An example of this new approach can be sea-floor spreading in the northern part of the
found in Ferrari et al. (2008). Central Atlantic occurred in the Toarcian. Similar
Most geodynamic/geological constraints and subsidence patterns between this region and the
data used for the reconstructions can be found in Lombardian basin (Stampfli 2000) led us to
our previous publications (Stampfli 2000; Stampfli propose a direct connection between these areas.
et al. 2001a, b; Stampfli & Borel 2002, 2004). A The Lombardian basin aborted in Middle Jurassic
new global terrane, or rather geological elements times (Bertotti et al. 1993) as it could not link up
map, was established following the approach with the nearby oceanic Meliata-Maliac domains
outlined in Ferrari et al. (2008). Part of this map is whose already cold lithosphere was rheologically
presented in Figure 1, where only the elements dis- considerably stronger than the surrounding conti-
placed during the Alpine orogenic event are shown. nental areas. Therefore, the Alpine Tethys rift
The present-day outlines of the elements should opened to the north of the Meliata basin, separating
be regarded as geographic markers, their shape Adria and the Austro-Carpathian domain from
having little to do with their original shape, gener- Europe (Bernoulli 1981). Thermal subsidence of
ally most larger, excepted for elements having areas flanking the Alpine Tethys commenced in
gone through large-scale Tertiary extension, such the Aalenian in the west (Briançonnais margin:
has Corsica or Sardinia, or the Cyclades elements. Stampfli & Marchant 1997; Stampfli et al. 1998,
Due to the GIS database, many geodynamic/ 2002) and in the Bajocian further to the east
kinematic information can be derived from the (Helvetic and Austroalpine margin: Froitzheim &
reconstructions; one of the most interesting is Manatschal 1996; Bill et al. 1997). The Alpine
the velocity map of the moving terranes, and their Tethys ocean spreading was considerably delayed
wander paths in respect to a fixed Europe with respect to the Central Atlantic; very slow
(Fig. 2), which can be established here with great spreading gave birth to a limited amount of
precision. This quantitative information can then oceanic crust, the oceanic area being dominated by
be confronted to major structural patterns and continental mantle denudation. A larger transform
kinematic indicators in the field; they show the Maghrebide ocean linked the central Atlantic and
changing direction and speed of convergence in the Alpine Tethys, and was also characterized by
space and time. delays in thermal subsidence (e.g. Rif area, Favre
1995; Stampfli 2000).
Within the Alpine domain, there is a fundamental
The geodynamic evolution of the larger difference between the AA –Carpathian and
Alpine area Western Alps systems. The AA –Carpathian evol-
ution was rooted in the dynamics of the Triassic
We shall review some of the major steps of the back-arc basins located to the south (Meliata–
peri-Alpine evolution used to constrain the recon- Maliac domain). These back-arc basins were
structions; Palaeotethys ocean evolution was shortened in conjunction with the opening of the
recently reviewed in Stampfli and Kozur (2006) Central Atlantic and rotation of Africa with respect
and is not repeated here. We also do not discuss to Europe. Subsequent slab roll-back of the subduct-
alternative models, quite numerous, and not often ing Küre, Maliac –Meliata oceanic lithosphere
taking into consideration plate tectonics principles, induced opening of the Vardar back-arc ocean,
but always very useful for the geological constraints which by Late Jurassic times had completely
they offer. replaced the pre-existing oceanic basins (Fig. 5).
Fig. 3. (Continued) Helvetic; Ion, Ionian; Ist, Istanbul; Kab, Kabylies; Kar, Karst; LNg, Lagonegro; Lig, Ligurian
ocean; Lom, Lombardian; Lyc, Lycian; Mag, Magura; Men, Menderes; Moe, Moesia; NCA, North Calcareous Alps;
NDo, North Dobrogea; Pan, Panormides; Pel, Pelagonia; Pen, Penninic, Vahic ocean; Pie, Piemontais ocean; PIM,
Paxi-Ionian-Mani; Pyr, Pyrenean ocean; Rho, Rhodope; Rif, Rif; SA, South Alpine; Sak, Sakarya; Sar, Sardinia; SCr,
South Crimea; SDz, Shatsky– Dzir; Sic, Sicani; SJa, Slavonia–Jadar; SMa, Serbo– Macedonian; SPi, Sitia–Pindos; Sre,
Srednogorie; SS, Sanandaj –Sirjan; Str, Stranja; SBe, Sub–Betic; Tau, Taurus; Tel, Tell; Tis, Tisia; Tor, Talea–Ori;
TrD, Transdanubian; Tro, Troodos; Tus, Tuscan; Tyr, Tyrrhenian; UMr, Umbria– Marches; WCa, West Carpathian;
Zon, Zonguldak –Küre.
Fig. 4. 180–165 Ma. Toarcian; Bathonian. Spreading is now active in the Central Atlantic and the segment of the
Alpine Tethys located south of Iberia (Lig). Eastward in the Carpathians domain, the Alpine Tethys successfully opened
in the Bajocian as it was able to connect with a plate limit through the Dobrogean transform (DoS, DoN) and from
there to the active subduction zone around the Sakarya plate. The Izmir– Ankara ocean is a back-arc of the Neotethys,
whereas opening of the Vardar corresponds to subduction progradation from the Küre domain toward the Maliac
domain. Accelerating roll-back of the Maliac sea floor generated northwestward spreading of the Vardar basin in a
scenario of intra-oceanic subduction. Closure of the Küre basin was entering a collisional stage around the Rhodope
promontory. Around the Alpine Tethys, aborted rifts become rim basins (sub-Betic, Helvetic–Vocontian, Lombardian)
and rifting is still active from Italy to Greece (Ionian zone).
Continued rotation of Africa provoked ridge failure Jurassic times, the northeastern part of the Vardar
in the Vardar and large-scale Late Jurassic ophiolitic arc-trench system collided with the northern
obduction onto the Dinaride –Hellenide passive Meliata passive margin (the Northern Calcareous
margin of the Pelagonian terrane (e.g. Laubscher & Alps: NCA, Fig. 4) (Bernoulli 1981), the Western
Bernoulli 1977; Dercourt et al. 1986). Roll-back of Carpathian domain (Csontos & Vörös 2004) and
the oceanic Maliac-Meliata slab was a centrifugal the Rhodope. In the latter, remnants of the Vardar
phenomenon that controlled successive collision of arc are found in northern Greece and Bulgaria as
the Vardar arc with all passive margins surrounding tectonic klippen (Bonev & Stampfli 2003, 2008).
the Meliata –Maliac basin. Accompanying the Closure of the Balkan rift system between
obduction on the western Dinaride margin in Late Moesia and the Rhodope (Figs 5 & 6) controlled
Fig. 5. 155– 142 Ma. Kimmeridgian; Berriasian. Spreading in the Alpine Tethyan ocean now reached the Carpathian
domain. In the process, Moesia was detached from Europe by only a few hundreds of kilometres, the point of
rotation of Gondwana being located close to Moesia. Northward extension of the Central Atlantic triggered active rifting
between Iberia and Newfoundland, and the northern limit of Iberia was affected by rifting extending eastward into the
Briançonnais domain, through the Pyrenees and Provence. Extension is affecting also the Helvetic-Dauphinois rim basin.
The Küre ocean was totally closed; collision of its arc-trench system with the Rhodope was causing the first phases of the
Balkan orogeny, accompanied by inversion of former rift zones. Roll-back of the Meliata–Maliac slab allowed rapid
westward expansion of the Vardar back-arc basin, its arc trench system colliding with the Pelagonian and Dinaric (SJa,Tis)
landmass. This is accompanied by east–west shortening in the Maliac– Vardar domain due also to the anti-clockwise
rotation of Gondwana with respect to Europe. We regard this event as creating a change in spreading direction in the
western Neotethyan domain, at the origin of the mid-ocean ridge failure, south of the Sanandaj– Sirjan block (SS).
subsequent development of the Balkan orogen, Along the NCA margin, elements of the AA
accompanied by large-scale Early Cretaceous micro-continent were scraped off and incorporated
northward nappe emplacement and metamorphism into the accretionary wedge, to form the different
(Georgiev et al. 2001). This circum-Vardar orogenic internal units of the AA –Carpathian orogen
event commenced in the Middle Jurassic and (Kozur 1991; Plasienka 1996; Faupl & Wagreich
was sealed by Albian to Cenomanian molasse-type 1999; Csontos & Vörös 2004). This event was
sediments in the Balkans (Georgiev et al. 2001), accompanied by Early Cretaceous HP–LT meta-
and mid-to late Cretaceous platforms in the morphism (e.g. Thöni & Jagoutz 1992; Ivan
Hellenides– Dinarides. 2002). Subsequently, the enlarged AA orogenic
Fig. 6. 131– 121 Ma. Hauterivian; Aptian. Accelerating anti-clockwise rotation of Gondwana was responsible for the
obduction of part of the Vardar mid-ocean ridge system onto the Pelagonia, Dinaride (SJa) and Tisia blocks. Collision
of the Vardar arc-trench system with the AA block was taking place at this time, detaching the future NCA domain
from its basement (internal AA). Collision of the Vardar arc-trench system continued also in the Balkan, where parts of
the Rhodope cover and basement were thrust northward . The major changes affecting the Neotethyan domain
brought to an end the opening of the Izmir– Ankara back-arc basin system. The Izmir– Ankara slab started retreating
eastward, allowing the opening of a new supra-subduction spreading centre. In the western Neotethys, ridge failure
generated a new intra-oceanic subduction zone along the former spreading centre; this new oceanic domain will
eventually obduct onto Arabia (e.g. Semail ophiolites). The Iberian plate was totally detached from Laurasia, whereas
spreading stopped in the Ligurian-Piemont part of the Alpine Tethys. In the eastern Penninic segment of the Alpine
Tethys, subduction progradation brought the exotic AA terranes onto the Alpine Tethys sea floor. Orogenic processes
were soon to come to an end in the Balkan (sealed by Albian molasses), whereas the Vardar intra-oceanic system
extended southward and eastward.
wedge began to move over the eastern segment of had commenced during the Carnian in the Küre
the Alpine Tethys (Figs 6 & 7) (Penninic–Vahic domain in Turkey (Stampfli & Kozur 2006) and
ocean), to finally collide with its northern passive continued into the Neogene period in the Eastern
margin (Helvetic domain s.l., Magura rim basin) Carpathians.
thus forming the present Eastern Alps and Car- The Cretaceous AA accretionary wedge was
pathian orogen (Wortel & Spakman 1993). This affected by large-scale collapse in the Late Cretac-
process involved continuous slab roll-back that eous (Figs 8 & 9), accompanied by the development
Fig. 7. 112– 103 Ma. Albian. The absence of magnetic anomalies between the M0 and C34 anomalies reduces the
constraints on the plate tectonics evolution during the intervening 37 Ma. However, the reconstructions are strongly
constrained by the transform movement between the African and Indian plates during that time, leaving very little place
for speculation. The Pyrenean ‘ocean’ opened and closed rapidly due to the rotation of Africa. The Briançonnais
peninsula becomes an upper plate promontory slowly advancing onto the Penninic ocean. A flexural basin developed
along the northern margin of the Izmir– Ankara ocean, with emplacement of ophiolitic mélanges on the Sakarya–East
Pontides domain, preceding subduction reversal. The southern Vardar and Semail intra-oceanic back-arc system
was expanding southward following slab retreat of the Huglu and Neotethys oceans.
of the Gosau basins (e.g. Faupl & Wagreich 1999). the necessary east–west convergence between the
Before that, an obstacle prevented such a collapse, Iberian plate and the Penninic ocean (Figs 7 & 8),
and we propose that spreading lasted in the it is more likely that these gabbros belong to the
Penninic –Vahic ocean until mid-Cretaceous times, latter and were underplated below the Piemont/
the collapse of the accretionary wedge being only Penninic Zermatt–Saas ophiolites.
possible when the mid-oceanic ridge had been
subducted. Recently dated gabbros in the Monte Cretaceous oceans: North Atlantic and
Rosa nappe have given Cenomanian ages (Liati & the Pyrenean domain
Froitzheim 2006). These gabbros are interpreted
by these authors as belonging to the ‘Valais From magnetic anomalies in the Atlantic domain,
ocean’, but in view of their structural position, and at least from the Aptian (M0 magnetic anomaly,
Fig. 8. 95– 84 Ma. Cenomanian; Santonian. The northern limit of the African plate in the Pyrenees became a zone of
convergence/subduction, partly extending into the Biscay–Atlantic domain, and generating large-scale inversion
from the Pyrenees to the western Alps. The Penninic mid-ocean ridge is subducted and the supra-subduction Gosau
rift started to open within the Austroalpine accretionary wedge. East– west shortening in the Alpine-Vardar region
brought Adria behind the Austroalpine prism, whereas the north– eastward subducting Vardar remnant ocean generated
an active margin setting in the Balkan, accompanied by the opening of the Srednogorie (Sre) and western Black Sea (BS)
back-arc basins. The southern Vardar ridge obducted southward on the Anatolide block, and linked eastward with
the Troodos–Semail obduction system. Finally, the Anatolian-Tauric plate was nearly totally covered by an
ophiolitic-type mélange. The accelerated rotation of Africa narrowed down significantly the Neotethyan domain south
of the Semail ocean. The Eratostene seamount could be related to volcanic activity in the Levant during the Cretaceous.
Fig. 7) to the Maastrichtian (Fig. 9) (Stampfli & African plate limit in this region had become con-
Borel 2002) and most likely up to the Thanetian vergent, due to intra-oceanic subduction in the Neo-
(anomaly C25, Fig. 9), it is observed that Iberia tethys and shortening affecting the Vardar region.
rotated with Africa and Apulia– Adria, without Also, the African plate was now separated from
appreciable North– South shortening (Fig. 2); it South America and India and started to rotate
also implies that spreading stopped in the western more on itself. As a result, the Pyrenean-Biscay
branch of the Alpine Tethys. Starting in the Early ocean opened, with a rifting phase starting in the
Cretaceous (Figs 6 & 7) the Pangaean break-up Oxfordian, and the onset of ocean spreading in the
had increasing difficulty in linking up eastward Aptian for the Portuguese–Galician ocean and
with another plate boundary. From divergent, the Pyrenean area, and in the Albian for the Gulf of
Fig. 9. 70–57 Ma. Maastrichtian; Thanetian. In the Alps, the northern sinistral Adriatic plate boundary was
extending westward into the Piemont ocean, defining a temporary Corsica-Briançonnais plate, which started subducting
southward beneath Adria. This was also allowing the AA prism to collapse westward and the supra-subduction
Gosau rift to continue to open. North–south shortening was taking place along the northern boundary of the Iberian
plate before a slight phase of extension in the Palaeocene. East –west shortening was still very active in the Alpine
and Vardar domains, passing into continental subduction. Roll-back of the remnant oceanic slab allows the opening of
the narrow east Black Sea back-arc basin. Counter-clockwise rotation of Africa was responsible for the obduction of the
Semail ocean on the Arabia margin (from Oman to Syria, Hatay), whereas intra-oceanic subduction persisted in front of
the Troodos plate.
Biscay, as clearly shown by peri-Iberian subsidence Gulf of Biscay (Vergés & Garcia-Senz 2001), due
curves (Borel & Stampfli 1999; Stampfli et al. 2002) to the accelerated rotation of the Iberian plate
and a clear mid-Cretaceous thermal event in the together with Africa (Fig. 8).
Pyrenees (e.g. between 112 and 97 Ma; Schärer It is uncertain how wide the Pyrenean ‘ocean’
et al. 1999). was, and whether it was limited to mantle denudation
By the end of the Santonian (84 Ma, Fig. 8), the as indicated by the Palaeozoic lherzolites at Lherz
break-up between North America and Greenland (Fabries et al. 1998; Lagabrielle & Bodinier 2008).
took place, and in the Campanian the Biscay spread- In order to control the geometry of the plate limit,
ing aborted (Olivet 1996). Closing of the Pyrenean a tentative ridge is put on the reconstructions for
domain already began during the opening of the the Pyrenean ‘ocean’, but as was the case in large
parts of the Alpine Tethys (see above), we think that ophiolitic obduction (Pillevuit et al. 1997) and
only mantle denudation took place in this domain. A slab detachment.
minimum extension of 60 –80 km can be calculated This NE directed subduction of the African
from restored cross-sections through the Pyrenees plate in the Cretaceous, brought Africa –Iberia far
(Vergés & Garcia-Senz 2001). However, from our to the east, inducing a relative westward escape of
reconstructions we come to a larger amount (c. the AA wedge into the Piemont oceanic corridor.
200 km), similar to the present northern Red Sea, The southern margin of the Piemont ocean (Southern
where indeed no sea-floor spreading has yet taken Alps–AA domain of the western Alps) was affected
place. Obviously, most of the distal margins have by tectonic movements since the Coniacian –
been subducted/eroded in the Pyrenean domain, as Santonian, as evidenced by the onset of flysch
is the case in the Alps. deposition in the Piemont (Gets and Dranse flysch,
The rotation of Iberia finally placed the Brian- Caron et al. 1989; Ligurian, Argnani et al. 2006
çonnais peninsula, by intra-oceanic subduction of and Lombardian basins, Bernoulli & Winkler
the Piemont/Penninic ocean, beside the distal 1990). We relate such tectonism to large-scale
Helvetic margin. This created a repetition of the sinistral strike-slip movements that affected the
European margin in the western Alps domain boundary between Adria and the Alpine domain.
(Figs 8 & 9) (Ringgenberg et al. 2002). The space These very large-scale lateral movements are
between these two similar margin segments is gener- well known also in the eastern Alps (Trümpy
ally referred to as the Valais ‘ocean’, a domain that 1988, 1992) and finally placed Adria –Tizia south
formerly belonged to the Piemont (Alpine Tethys) of the AA accretionary prism. In this process, the
ocean, trapped by the eastward displacement of the Piemont oceanic lithosphere was progressively
Iberian –Briançonnais block during the opening of detached from the northern margin of Adria and
the North Atlantic. The Middle Jurassic age subducted beneath it. Frontal pieces of the western
(161 Ma) of potential parts of the Valais (Piemont) Adria –AA margin were dragged into the subduc-
sea floor from the Misox zone was recently con- tion zone, as evidenced by the Late Cretaceous –
firmed (Liati et al. 2003). As in the Ligurian and Early Palaeocene HP– LT metamorphism recorded
Piemont domains, where mantle rocks associated in the Sesia domain (Oberhänsli et al. 1985;
to Alpine ophiolites have shown Permian and even Rubatto 1998).
Precambrian ages (Rampone & Piccardo 2000), the
Versoyen magmatic complex in the internal Valais The Pyrenean cycle
domain also has Permian ages, but was clearly
reheated during the Cretaceous (110 –100 Ma). But The ‘Valais trough’, as recognized in the western
it is not yet clear if this complex moved with the Alps today, is actually the remnant of trapped
Briançonnais, or if it belonged to the toe of the Piemont sea floor and of the toe of the Helvetic
Helvetic margin (Fügenschuh et al. 1999). The pres- margin (see above, Stampfli et al. 2002). The
ence of the Cretaceous thermal event there would ‘Valais ocean’ (as defined in Stampfli 1993) was
place them as part of the northern Briançonnais located south of France and we refer to it here as
margin. Other gabbros in the Monte Rosa nappe the ‘Pyrenean ocean’ to avoid confusion between
have been dated as Cenomanian (Liati & Froitzheim Valais ocean and Valais trough. No direct traces
2006), but as discussed above, we would rather see of this ocean have been found so far because its
them as pertaining to the Penninic ocean. suture was located exactly where the Algero–
Provençal ocean re-opened in Oligo –Miocene
Along strike shortening times (Roca 2001). A large part of the southern
margin of the Pyrenean ocean was the Briançonnais
A relatively large remnant of the Vardar ocean sub- peninsula (Figs 5 & 6); its northern margin was the
ducted northeastward under Moesia during the Late Corbières –Provençal domain from the Pyrenees to
Cretaceous, as evidenced by the large Srednogorie the Maures-Estérel massifs. Elements of the Pyre-
volcanic arc of the Balkans and the Late Cretaceous nean margin of the Briançonnais are found in the
opening of the Black Sea (Nikishin et al. 2003), Galibier region of the French Alps (Toury 1985),
representing the third generation of back-arc well known for its Late Jurassic Brèche du Télé-
opening in that region. This subduction zone and graphe. Recent investigation there (Luzieux &
its extension eastward up to Iran, represents a Ferrari 2002) showed that a pull-apart type basin
major slab pull force moving Africa in that direction rapidly deepened under the continental crust defor-
until the closure of the Vardar in the Late Maastrich- mation (CCD) in Late Jurassic times. This area is
tian–Palaeocene. This is clearly confirmed by the regarded as the most external Briançonnais
change of convergence direction and velocity element known so far, unless the Versoyen
decrease of the African plate at that time (Fig. 2), complex is also seen as a part of this margin (see
also related to the synchronous peri-Arabian above). Its conjugate northern Provençal margin
area is characterized by important erosion during in the Biscay ocean (Olivet 1996). We propose
the Oxfordian (rift shoulder uplift) and the develop- here a uniform southward subduction during the
ment of Albian basins deepening southward late Cretaceous Pyrenean phase (Figs 8 & 9),
towards the ocean, followed by the accumulation replaced in the Pyrenees by northward subduction
of thousands of metres of upper Cretaceous clastics of the Iberian indenter during the second Eocene
in a northward migrating fore-deep type basin Pyrenean phase (Fig. 10). This dominating last
(Debrand-Passard & Courbouleix 1984). This event gave to the Pyrenees its final lithospheric
records a southward closure of the Pyrenean structure, as seen on the ECORS transect. In
‘ocean’ on a Provençal transect, whereas a north- between the two phases, and according to the Atlan-
ward subduction is usually proposed on a Pyrenean tic magnetic anomalies, the Palaeocene position of
transect (Vergés & Garcia-Senz 2001), also seen in Iberia seems to retreat slightly (10 –20 km) from
Sardinia (Barca & Costamagna 1997, 2000), but Europe. This has created some large-scale extension
again a southward subduction took place westward superimposed on the Late Cretaceous orogenic
Fig. 10. 48– 40 Ma. Lutetian-Bartonian. The Briançonnais terrane was entering into collision with the AA. The latter
has closed a large part of the Alpine Tethys. Behind the prism, extension generated large rift zones, such as on
the Tisia terrane (Pannonian basin). Similarly, the subduction of the Pindos and Troodos oceanic slabs, triggered
the opening of numerous rifts and associated core-complexes in Turkey, in the Cyclades and in the Balkans. In
the Pyrenees, retro-wedge thrusting is now very active; collision of the Iberian and European domains triggered the
subduction of the remnant Alpine Tethys south of Iberia.
wedge and is responsible for the deposition of the The Alpine cycle
Palaeocene red beds found all along the Pyrenean
chain (Bilotte & Canérot 2006). After several phases of rifting, as described above,
The Pyrenean orogen can be followed from the and a new thermal subsidence stage developing
present Pyrenees eastward to southern France (Prov- during the Cretaceous, the Alpine region entered a
ence), and continues in the Alps in the form of a phase of convergence between the African and
large-scale uplift of the Helvetic margin and local European plate (Fig. 11). The tectonic evolution of
inversion of the Jurassic tilted blocks, well the western Alps started with the formation of an
expressed by the deposits of the Niesen Flysch accretionary prism related to the closure of the
(Ackermann 1986) (mainly Maastrichtian) and Alpine Tethys where different geological objects,
Meilleret Flysch (Middle Eocene) (Homewood corresponding to different stages of accretion, can
1974), sedimented on a structured Mesozoic base- be recognized:
ment. Recent investigations have shown that † the Adriatic back-stop, comprising an aborted
similar turbiditic deposits of Late Cretaceous age Jurassic rifted basin (Lombardian basin);
are also found in ultra-Dauphinois units of the † the oceanic accretionary prism of the Piemont/
French Alps, such as the Pelat units (De Paoli & Penninic ocean (the western Alps portion of the
Thum 2008) and the already known Quermoz unit Alpine Tethys), including crustal elements
(Homewood et al. 1984). These so-called flysches from the former toe of the southern passive
clearly predate the Alpine syn-collisional event in margin (lower AA elements);
the Helvetic domain, characterized by the depo- † accreted material of the Briançonnais terrane
sition of the classical Alpine flysch sequences not derived from the Iberian plate; and
before the earliest Oligocene (Kempf & Pfiffner † accreted material of the Valais trough, represent-
2004). These Cretaceous turbiditic deposits clearly ing the toe of the European passive margin; and
point to an orogenic event affecting the distal † accreted material of the former European conti-
Dauphinois– Helvetic margin and were deposited nental margin and rim basin (Dauphinois –
in the eastern prolongation of the Provençal foreland Helvetic domain).
basin. On a Provence transect, the closure of the In time, one passes from the oceanic accretion-
remnant Pyrenean basin certainly took place at the ary prism to the formation of the orogenic wedge
end of the Cretaceous phase (Figs 8 & 9), whereas (Escher et al. 1996; Pfiffner et al. 1997; Ford et al.
in the western Alps, this closure was not total, as 2006) that we place after the detachment or delami-
witnessed by the continuing deposit of coarse nation of the subducting slab in the Early Oligocene
clastic turbidites in the Valais trough up to the (e.g. Stampfli & Marchant 1995; von Blanckenburg
Late Eocene (Bagnoud et al. 1998). This suggests & Davies 1995). The resulting heat flux allowed
the presence of a subduction/inversion zone some more units to be detached from the European
between the Valais basin and the Niesen – continental slab, triggering large-scale subduction
Quermoz–Pelat foreland basin, both basins being of continental material (Marchant & Stampfli
separated by a relief where basement was outcrop- 1997; Ford et al. 2006) and Oligocene to Pliocene
ping (the so-called Tarine cordillera). overthrusting of the most external units, such as:
Eastward, the uplift of the Bohemian massif (e.g.
Tanner et al. 1998) was certainly related to the same † the external Variscan massifs and their cover;
late Cretaceous inversion event and triggered the † the Subalpine fold and thrust belts;
deposition of the Rheno –Danubian turbiditic † the North Alpine foreland basin (molassic
sequence. Along the European margin it is interest- basin); and
ing to note the similarity of facies between these † the Jura mountains fold belt.
Rheno–Danubian turbiditic deposits and the These late events were accompanied by retro-wedge
Valais trough sequence (the Valais trilogy) from thrusting of the southern Alps on the Po Plain, a flex-
Albian to Late Cretaceous (Stampfli 1993, and ural basin located above the former Jurassic Lom-
references therein). These deep-water clastic facies bardian basin. The retro-wedge becomes wider
were located along the toe of the European margin, eastward, forming the Southern Alps thrust belt
often referred to as the North Penninic basin (Schönlaub & Histon 2000; Schmid et al. 2004).
(Figs 7 & 8). The presence of contourites and This younger S-verging thrust system linked two
strong and changing current directions along the north-dipping subduction systems: the Adriatic –
basin (Hesse 1974) suggests a connection-with Dinaride continental subduction to the east, and
major oceanic domains. So the Valais trough, the subduction of the remnant Alpine Tethys under
together with the Pyrenean–Biscay ocean, must be the Iberian plate in the west (Figs 10 & 12). In the
regarded as connections between the Eastern Adriatic –Dinaride subduction zone, a large part of
Alpine Tethyan realm and the north Atlantic ocean the northern Apulian promontory was subducted
during the Cretaceous. under the AA domain (Schmid et al. 2004; Kissling
Fig. 11. 33– 20 Ma. Rupelian; Burdigalian. The Alps have entered the main orogenic phase. The Briançonnais was
subducted; parts of its cover was detached and accreted to the orogenic prism. Slab roll-back in the Carpathians allowed
the Pannonian basin to enlarge and many other rifts and core-complexes were still active in the Balkans, Greece and
Turkey until the Pindos was totally closed. Then a collision of the Hellenic orogenic prism took place with the Greater
Apulia block (PIM) and triggered the subduction of the East Mediterranean basin (eastern part). Nowadays, this
subduction front has reached the Apulian block (Fig. 1). On the Iberian side, the collapse of the active margin onto the
remnant Alpine Tethys segment has induced large-scale rifting in the upper plate and the detachment of the Corso–
Sardinian block. This rift system link northward with the Bresse–Rhine graben system. A nearly continuous
south-vergent subduction/collision front runs from the southern border of Iberia to Turkey, and the north-vergent
Alpine prism will progressively become less active.
et al. 2006), whereas the northwestern corner still converging African plate, creating maximum
of Adria developed as an indenter (Handy et al. shortening in the western Alps.
2005). As we have seen above, the southward subduc-
Since the Miocene, the Alps and Carpathian tion of the Alpine Tethys ocean is related to the
wedges moved mainly due to remnant roll-back of history of the Meliata –Maliac and Vardar domain,
the European plate. The weakly buoyant lithosphere and was inherited in the western Alps from the
under the rim basins present in that margin allowed pre-existing northward vergence of the AA accre-
further subduction to take place. However, in the tionary wedge. This northward vergence is unique
western Alps, the Adriatic indenter was pushed in the whole Alpine and Tethyan domain, where
against the orogenic prism as it was attached to the most orogens are south vergent.
Fig. 12. North– south palinspatic cross-sections through the western Alpine segment. The Briançonnais block (Bri,
on the reconstruction) is fixed; the other segments move in and out of the picture. Arrows show major uplift and
subsidence events.
The change in vergence of Alpine Tethys sub- Alpine accretionary prism stranded in the eastern
duction is found at the connecting region between border of Corsica started to collapse backward
the Alps and the northern Apennines, south of (eastward), following the build-up of the new
Corsica. It must be emphasized here that the Penni- Apenninic accretionary prism; these Piemont –
nic Alpine accretionary prism is older (Late Cretac- Ligurian units are now found in an upper structural
eous–Eocene) than the Apenninic one (Oligocene – position, whereas they were in a lower one during
Pliocene); actually one started when the other one the Alpine collision. The northern and southern
stopped. During this process, the Apenninic prism Apenninic prisms collapsed into pre-existing
re-mobilized parts of the Piemont– Penninic prism depressions, the Lombardian rift in the north and
as exotic elements (e.g. the Bracco ophiolitic the Ionian sea oceanic corridor (the western most
ridge, Elter et al. 1966; Hoogenduijn Strating part of the East Mediterranean Neotethys ocean)
1991). The Alpine prism collided with the Iberian in the south, with its still active back-arc opening
plate in Corsica (Malavieille et al. 1998) in (Savelli 2000; Argnani & Savelli 2001). In the
Eocene times (Figs 10 –12). The remnant Alpine eastern part of the East Mediterranean basin, sub-
Tethys oceanic domain (Ligurian basin) south of duction had already started in the late Eocene
the Iberian plate started to subduct northward in (Fig. 12); it then prograded westward until its
the Late Eocene (producing HP metamorphism present day position in Greece (Fig. 1).
dated at 25 –21 Ma, Michard et al. 2006) in an
ongoing process of shortening between Europe Conclusions
and Africa and following continent-continent
collision in the Pyrenees. The proposed reconstructions require a tight fit of
On the Miocene reconstruction (Fig. 12), nearly the Iberian– Alboran microplates with respect to
all of this remnant ocean is already subducted. This North America and Europe, and also imply a long
SE-directed roll-back of the Tethyan slab triggered lasting rifting phase in the Atlantic regions now sep-
the opening of the Algero –Provençal back-arc arating these domains. The final position of the
basin starting in the Oligocene (Fig. 12). Thus, the Apulia –Adria plate and its geometry close to the
Oligocene volcanism of Sardinia can be regarded present-day one, was reached through Jurassic
as subduction related (e.g. Monaghan 2001). The rifting events, which finally gave birth to the
narrow Alpine Tethys. The latter is dominated by A RGNANI , A. & S AVELLI , C. 2001. Magmatic signature
stretching of the lithospheric mantle of its borders. of episodic back-arc rifting in the southern Tyrrhenian
On its western side, the Alpine area is largely influ- Sea. In: Z IEGLER , P. A., C AVAZZA , W., R OBERTSON ,
enced by the opening Atlantic ocean, and the north- A. H. F. & C RASQUIN -S OLEAU , S. (eds) PeriTethys
Memoir 6: Peritethyan Rift/wrench Basins and
ward extension of the latter in Early Cretaceous Passive Margins, IGCP 369. Mémoires du Museum
times. On its eastern side, the Alpine domain is Nationale d’Histoire Naturelle, 186, 735– 754.
dominated by the evolution of the supra-subduction A RGNANI , A., F ONTANA , D., S TAFANI , C. & Z UFFA , G. G.
Vardar ocean. The latter collided with its border in 2006. Palaeogeography of the Upper Cretaceous-
latest Jurassic –early Cretaceous times. Its northern Eocene carbonate turbidites of the Northern Apen-
border consisted of the AA micro-continent, not nines from provenance studies. In: M ORATTI , G. &
buoyant enough to resist subduction. Large portions C HALOUAN , A. (eds) Tectonics of the Western
of its crust and cover were underplated and trans- Mediterranean and North Africa. Geological Society,
ported northward onto the eastern Alpine Tethys London, Special Publications, 262, 259– 275.
B AGNOUD , A., W ERNLI , R. & S ARTORI , M. 1998. Décou-
(Penninic– Vahic ocean). verte de foraminifères planctoniques dans la zone de
It is then clear that two specific geodynamic Sion-Courmayeur à Sion. Eclogae geologicae Helve-
scenarios presided at the evolution of the Alpine tiae, 91(3), 421–429.
belt. The western branch was affected by a second B ARCA , S. & C OSTAMAGNA , L. C. 1997. Compressive
phase of rifting, detaching the Iberian plate and its ‘Alpine’ tectonics in Western Sardinia (Italy): geody-
Briançonnais promontory from Europe, finally namic consequences. Comptes Rendus Académie des
duplicating the southern European margin in the Sciences, 325, 791–797.
western Alps. The eastern branch is dominated by B ARCA , S. & C OSTAMAGNA , L. C. 2000. Il Bacino paleo-
an advancing large accretionary prism that included genico del Sulcis-Iglesiente (Sardegna SW): nuovi dati
stratigraphico-strutturali per un modello geodinamico
the AA blocks. The prism was collapsing forward nell’ambito dell’orogenesi pirenaica. Bollettino della
but also westward into the oceanic Piemont part of Società geologica Italiana, 119, 497– 515.
the western branch. This was made possible by B ERNOULLI , D. 1981. Ancient continental margins of the
large-scale east –west shortening, related to the kin- Tethyan Ocean. In: B ALLY , A. W., W ATTS , A. B.,
ematics of the African plate with regards to Europe G ROW , J. A., M ANSPEIZER , W., B ERNOULLI , D.,
(Fig. 2). Thus, the Vardarian oceanic back-stop of S CHREIBER , C. & H UNT , J. M. (eds) Geology of
the AA prism was replaced in time by the continen- passive continental margins; history, structure and
tal larger Adriatic back-stop. The Piemont oceanic sedimentologic record (with special emphasis on the
space was, at the same time, forced to subduct south- Atlantic margin). American Association of Petroleum
Geologists, Education course note series, 19/5, 1–36.
ward in front of the advancing AA prism. At the B ERNOULLI , D. & W INKLER , W. 1990. Heavy mineral
close of the Cretaceous, all the major elements are assemblages from Upper Cretaceous South Alpine
in place, and after the Palaeocene lull, following and Austro-alpine flysch sequences (N-Italy and S-
the change of wander path of Africa, they were Switzerland): source terranes and paleotectonic impli-
thrust over each other to form the Alpine cations. Eclogae geologicae Helvetiae, 83(2),
orogenic wedge. 287– 310.
If this wedge is tectonically relatively cylindri- B ERTOTTI , G., P ICOTTI , V., B ERNOULLI , D. &
cal, its components are not closely related. The C ASTELLARIN , A. 1993. From rifting to drifting: tec-
Alpine orogen is a clear example of juxtaposition tonic evolution of the South-Alpine upper-crust from
the Triassic to the Early Cretaceous. Sedimentary
of far travelled tectonic elements and cannot be Geology, 86, 53–76.
understood from a ‘fixist’ point of view. B ILL , M., B USSY , F., C OSCA , M., M ASSON , H. &
The Swiss National Funds supported C. Hochard PhD H UNZIKER , J. C. 1997. High precision U– Pb and
thesis (grants n. 2000-068015 and 200020-109549). 40Ar/39Ar dating of an Alpine ophiolite (Gets
M. Ford and A. Argnani thoroughly reviewed this manu- nappe, French Alps). Eclogae geologicae Helvetiae,
script, their suggestions and corrections are warmly 90, 43– 54.
acknowledged. B. Murphy strongly encouraged us to B ILOTTE , M. & C ANEROT , J. 2006. Rôles respectifs des
publish this work, we thank him for his patience. tectoniques fini-crétacée et éocène dans la partie orien-
tale de la chaı̂ne des Pyrénées. Le ‘Garumnien’ de
Cucugnan et ses relations avec le Chevauchement
Frontal Nord-Pyrénéen (Corbières méridionales,
References France). Eclogae geologicae Helvetiae, 99(1), 17–27.
B ONEV , N. & S TAMPFLI , G. M. 2003. New structural
A CKERMANN , A. 1986. Le flysch de la nappe du Niesen. and petrologic data on Mesozoic schists in the
Eclogae geologicae Helvetiae, 79(3), 641–684. Rhodope (Bulgaria): geodynamic implications.
A RGNANI , A. 2002. The Northern Apennines and the kin- C. R. Geoscience, 335, 691–699.
ematics of Africa – Europe Convergence. Bolletino B ONEV , N. & S TAMPFLI , G. M. 2008. Petrology, geo-
della Società Geologica Italiana, Volume Speciale, 1, chemistry and geodynamic implications of Jurassic
47– 60. island arc magmatism as revealed by mafic volcanic
rocks in the Mesozoic low-grade sequence, eastern F AUPL , P. & W AGREICH , M. 1999. Late Jurasic to Eocene
Rhodope, Bulgaria. Lithos, 100, 210–233. Palaeogeography and geodynamic evolution of the
B OREL , G. & S TAMPFLI , G. M. 1999. Geodynamic evol- Eastern Alps. Mitteilungen der Österreichischen Geo-
ution of the Briançonnais domain and the individuali- logischen Gesellschaft, 92, 79– 94.
sation of the Iberian plate. 4th workshop Alpine F AVRE , P. 1995. Analyse quantitative du rifting et de la
Geol. Studies, Tübingen, 52, 13–14. relaxation thermique de la partie occidentale de la
C ARON , C., H OMEWOOD , P. & W ILDI , W. 1989. The marge transformante nord-africaine: le Rif externe
original Swiss Flysch: a reappraisal of the type deposits (Maroc). Geodinamica Acta, 8(2), 59– 81.
in the Swiss Prealps. Earth-Science Reviews, 26, 1– 45. F AVRE , P. & S TAMPFLI , G. M. 1992. From rifting to
C HANNELL , J. E. T. 1992. Paleomagnetic data from passive margin: the example of the Red Sea, Central
Umbria (Italy): implications for the rotation of Adria Atlantic and Alpine Tethys. Tectonophysics, 215,
and Mesozoic apparent polar wander paths. Tectono- 69–97.
physics, 216, 365– 378. F AVRE , P., S TAMPFLI , G. & W ILDI , W. 1991. Jurassic
C HANNELL , J. E. T. 1996. Paleomagnetism and Paleogeo- sedimentary record and tectonic evolution of the north-
graphy of Adria. In: M ORRIS , A. & T ARLING , D. H. western corner of Africa. Paleogeography, Paleoeco-
(eds) Paleomagnetism and Tectonics of the Mediterra- lology, Paleoclimatology, 87, 53–73.
nean Region. Geological Society, London, Special F ERRARI , O. M., H OCHARD , C. & S TAMPFLI , G. M.
Publications, 119 –132. 2008. An alternative plate tectonic model for the
C SONTOS , L. & V ÖRÖS , A. 2004. Mesozoic plate tectonic Palaeozoic– Early Mesozoic: palaeotethyan evolution
reconstruction of the Carpathian region. Palaeoclima- of Southeast Asia (Northern Thailand– Burma). Tecto-
tology, Palaeoecology, Palaeogeography, 210, 1– 56. nophysics, 451, 346 –365.
DE J ONGE , M. R., W ORTEL , M. J. R. & S PAKMAN , W. F INETTI , I. R. 2005. CROP Project: deep seismic explora-
1994. Regional scale tectonic evolution and the tion of the central Mediterranean and Italy. Elsevier,
seismic velocity structure of the lithosphere and Amsterdam, Boston.
upper mantle: the Mediterranean region. Journal of F ORD , M., D UCHENE , S., G ASQUET , D. & V ANDER-
Geophysical Research, 99(B6), 12091–12108. HAEGHE , O. 2006. Two-phase orogenic convergence
D E P AOLI , R. & T HUM , L. 2008. Implications géo- in the external and internal SW Alps. Journal of the
dynamiques de l’étude pétrographique des flyschs du Geological Society, London, 163(5), 815–826.
Crétacé supérieur du Piolit, du Pelat et de Baiardo F RISCH , W. 1979. Tectonic progradation and plate tec-
(Alpes occidentales). MSc, University of Lausanne, tonic evolution of the Alps. Tectonophysics, 60,
Switzerland, 105. 121–139.
D EBRAND -P ASSARD , S. & C OURBOULEIX , S. 1984. F ROITZHEIM , N. & M ANATSCHAL , G. 1996. Kinematics
Synthèse géologique du S-E de la France. Mémoires of Jurassic rifting, mantle exhumation, and passive-
du Bureau de Recherches Géologiques et Minières margin formation in the Austroalpine and Penninic
Orléans, 126, 1 –28. nappes (Eastern Switzerland). Geological Society of
D ERCOURT , J., R ICOU , L. E. & V RIELINCK , B. 1993. America Bulletin, 108(9), 1120– 1133.
Atlas Tethys, paleoenvironmental maps. Gauthier- F ÜGENSCHUH , B., L OPRIENO , A., C ERIANI , S. &
Villars, Paris, 307. S CHMID , S. M. 1999. Structural analysis of the Sub-
D ERCOURT , J., Z ONENSHAIN , L. P. ET AL . 1985. Présen- briançonnais and Valais units in the area of Moutiers
tation des 9 cartes paléogéographiques au 1/ (Savoy, Western Alps): paleogeographic and tectonic
20 000 000 s’étendant de l’Atlantique au Pamir pour consequences. International Journal of Earth
la période du Lias à l’Actuel. Bulletin de la Société Sciences, 88, 201– 218.
géologique de France, 8(T I, No 5), 637 –652. G ARFUNKEL , Z. & D ERIN , B. 1984. Permian-early Meso-
D ERCOURT , J., Z ONENSHAIN , L. P. ET AL . 1986. Geologi- zoic tectonism and continental margin formation in
cal evolution of the Tethys from the Atlantic to the Israel and its implications for the history of the
Pamirs since the Lias. Tectonophysics, 123, 241–315. Eastern Mediterranean. In: D IXON , J. E. & R OBERT-
E LTER , G. P. E., S TURANI , C. & W EIDMANN , M. 1966. SON , A. H. F. (eds) The Geological Evolution of the
Sur la prolongation du domaine ligure de l’Apennin Eastern Mediterranean. Geological Society, London,
dans le Montferrat et les Alpes et sur l’origine de la Special Publications, 17, 187–201.
nappe de la Simme s.l. des Préalpes romandes et cha- G EORGIEV , G., D ABOVSKI , C. & S TANISHEVA -
blaisiennes. Archives des Sciences, Genève, 19(3), V ASSILEVA , G. 2001. East Srednogorie-Balkan rift
279– 377. zone. In: Z IEGLER , P. A., C AVAZZA , W., R OBERT-
E SCHER , A., H UNZIKER , J. C., M ARTHALER , M., SON , A. H. F. & C RASQUIN -S OLEAU , S. (eds) Peri-
M ASSON , H., S ARTORI , M. & S TECK , A. 1996. Geo- Tethys memoir 6: Peritethyan rift/wrench basins and
logic framework and structural evolution of the passive margins, IGCP 369. Mémoires du Museum
Western Swiss Alps. In: P FIFFNER , O. A., L EHNER , Nationale d’Histoire Naturelle, 186, 259– 294.
P., H EITZMANN , P., M ÜLLER , S. & S TECK , A. (eds) H ANDY , M., B ABIST , J., W AGNER , R., R OSENBERG , C.
Deep structure of Switzerlannd – Results from the & K ONRAD , M. 2005. Decoupling and its relation
National Research Program 20 (NRP20). Birkhäuser, to strain partitioning in continental lithosphere;
Basel, 205–221. insight from the Periadriatic Fault system (European
F ABRIES , J., L ORAND , J. P. & B ODINIER , J. L. 1998. Pet- Alps). In: G APAIS , D., B RUN , J.-P. & C OBBOLD ,
rogenetic evolution of orogenic lherzolite massifs in P. R. (eds) Deformation mechanisms, rheology and tec-
the central and western Pyrenees. Tectonophysics, tonics; from minerals to the lithosphere. Geological
292(1 –2), 145– 167. Society, London, Special Publications, 243, 249– 276.
H ESSE , R. 1974. Long-distance continuity of turbidites: and Central Alps; geochronological evidence for
possible evidence for an early Cretaceous trench re-rifting of oceanic crust. Contributions to Mineral-
abyssal plain in the East Alps. Geological Society of ogy and Petrology, 149(4), 446– 461.
America Bulletin, 85, 859– 870. L UZIEUX , L. & F ERRARI , O. M. 2002. Etude géologique
H OCHARD , C. 2008. GIS and Geodatbases applications to et minéralogique du secteur du Col du Galibier
global scale plate tectonics modelling. PhD thesis, (Savoie, France): implications géodynamiques. MSc
University of Lausanne, Switzerland. thesis, University of Lausanne, Lausanne.
H OMEWOOD , P. 1974. Le flysch du Meilleret (Préalpes M ALAVIEILLE , J., C HEMENDA , A. & L ARROQUE , C.
romandes) et ses relations avec les unités l’encadrant. 1998. Evolutionary model for the Alpine Corsica:
Eclogae geologicae Helvetiae, 67(2), 349–401. mechanism for ophiolite emplacement and
H OMEWOOD , P., A CKERMANN , T., A NTOINE , P. & exhumation of high-pressure rocks. Terra Nova, 10,
B ARBIER , R. 1984. Sur l’origine de la nappe du 317– 322.
Niesen et la limite entre les zones ultrahelvétiques et M ARCHANT , R. H. & S TAMPFLI , G. M. 1997. Subduction
valaisanne. Comptes Rendus des Séances de l’Acadé- of continental crust in the Western Alps. Tectonophy-
mie des Sciences de Paris, série II, 299(15), sics, 269(3–4), 217 –235.
1055–1059. M ICHARD , A., N EGRO , F., S ADDIQI , O., B OUY-
H OOGENDUIJN S TRATING , E. H. 1991. The evolution of BAOUENE , M. L., C HALOUAN , A., M ONTIGNY , R.
the Piemonte-Ligurian ocean, a structural study of & G OFFÉ , B. 2006. Pressure–temperature–time
ophiolite complexes in Liguria (NW Italy). Geologica constraints on the Maghrebide mountain building:
ultraiectina, 74, 1– 127. evidence from the Rif–Betic transect (Morocco,
I VAN , P. 2002. Relict magmatic minerals and textures in Spain), Algerian correlations, and geodynamic impli-
the HP/LT metamorphosed oceanic rocks of the cations. Comptes Rendus Geosciences, 338(1– 2),
Triassic-Jurassic Meliata Ocean, Inner Western Car- 92–114.
pathians. Slovak Geological Magazine, 8(2), 109– 122. M OIX , P., B ECCALETTO , L., K OZUR , H. W., H OCHARD ,
K EMPF , O. & P FIFFNER , A. 2004. Early Tertiary evolution C., R OSSELET , F. & S TAMPFLI , G. M. 2008. A new
of the North Alpine foreland basin of the Swiss Alps classification of the Turkish terranes and sutures and
and adjoining areas. Basin Research, 16(4), 549– 567. its implication for the paleotectonic history of the
K ISSLING , R., S CHMID , S. M., L IPPITSCH , R., A NSORGE , region. Tectonophysics, 451, 7 –39.
J. & F UEGENSCHUH , B. 2006. Lithosphere structure M ONAGHAN , A. 2001. Coeval extension, sedimentation
and tectonic evolution of the Alpine Arc; new evidence and volcanism along the Cainozoic rift system of
from high-resolution teleseismic tomography. In: G EE , Sardinia. In: Z IEGLER , P. A., C AVAZZA , W.,
D. G. & S TEPHENSON , R. (eds) European lithosphere R OBERTSON , A. H. F. & C RASQUIN -S OLEAU , S.
dynamics. Geological Society, London, Memoirs, 32, (eds) PeriTethys memoir 6: Peritethyan rift/wrench
129–145. basins and passive margins, IGCP 369. Mémoires du
K OZUR , H. 1991. The geological evolution at the western Museum Nationale d’Histoire Naturelle, 186,
end of the Cimmerian ocean in the Western Car- 707– 734.
pathians and Eastern Alps. Zentralblatt für Geologie M UTTONI , G., G ARZANTI , E., A LFONSI , L., C IRILLI , S.
und Paläontologie Teil I, 1(H.1), 99–121. & G ERMANI , D. 2001. Motion of Africa and Adria
L AGABRIELLE , Y. & B ODINIER , J. L. 2008. Submarine since the Permian: paleomagnetic and paleoclimatic
reworking of exhumed subcontinental mantle rocks: constraints from northern Libya. Earth Planetary
field evidence from the Lherz peridotites, French Sciences Letter, 192, 159– 174.
Pyrenees. Terra Nova, 20, 11–21. N IKISHIN , A. M., K OROTAEV , M. V., E RSHOV , A. V. &
L AGABRIELLE , Y., P OLINO , R. ET AL . 1984. Les témoins B RUNET , M. F. 2003. The Black-Sea basin: tectonic
d’une tectonique intraocéanique dans le domaine history and Neogene-Quaternary rapid subsidence
téthysien: analyse des rapports entre les ophiolites et modelling. In: B RUNET , M. F. & C LOETINGH , S.
leurs couvertures métasédimentaires dans la zone pié- (eds) Integrated Peri-Tethyan Basins Studies. Sedi-
montaise des Alpes franco-italiennes. Ofioliti, 9(1), mentary Geology, 156, 149– 168.
67– 88. O BERHÄNSLI , R., H UNZIKER , J. C., M ARTINOTTI , G. &
L AUBSCHER , H. P. & B ERNOULLI , D. 1977. Mediterra- S TEN , W. B. 1985. Geochemistry, geochronology
nean and Tethys. In: N AIRN , A., K ANES , W. & and petrology of Monte Mucrone: an example of
S TEHLI , F. (eds) The Ocean Basins and Margins. eo-alpine eclogitisation of Permian granitoı̈ds in the
Plenum Press, New York, 1–28. Sesia Lanzo zone, western Alps, Italy. Chemical
L IATI , A. & F ROITZHEIM , N. 2006. Assessing the Geology, 52, 165–184.
Valais ocean, Western Alps: U–Pb SHRIMP zircon O LIVET , J.-L. 1996. Lacinématique de la plaque Ibérique.
geochronology of eclogite in the Balma unit, on top Bulletin des Centres de recherches exploration-
of the Monte Rosa nappe. European Journal of production Elf-Aquitaine, 20, 131–195.
Mineralogy, 18, 299 –308. P FIFFNER , O. A., L EHNER , P., H EITZMAN , P. Z.,
L IATI , A., G EBAUER , D., F ROITZHEIM , N. & F ANNING , M UELLER , S. & S TECK , A. 1997. Deep structure of
M. 2003. Origin and geodynamic significance of meta- the Swiss Alps – Results from NRP 20. Birkhäuser
basic rocks from the Antrona ophiolites (western AG., Basel, 380.
Alps): new insights from SHRIMP-dating. EGS- P ILLEVUIT , A., M ARCOUX , J., S TAMPFLI , G. M. &
AGU-EUG joint assembly, Nice, EAE03-A, 12648. B AUD , A. 1997. The Oman exotics: a key to the under-
L IATI , A., F ROITZHEIM , N. & F ANNING , M. 2005. Juras- standing of the Neotethyan geodynamic evolution.
sic ophiolites within the Valais Domain of the Western Geodinamica Acta, 10(5), 209– 238.
P LASIENKA , D. 1996. Mid-Cretaceous (120– 80 Ma) oro- Alpine-Himalayan Orogen. Journal of the Virtual
genic processes in the central Western Carpathians: Explorer, 8, 149– 168.
brief review and interpretation of data. Slovak Geologi- S CHMID , S. M., F UEGENSCHUH , B., K ISSLING , E. &
cal Magazine, 3– 4, 319–324. S CHUSTER , R. 2004. Transects IV, V and VI; the
R AMPONE , E. & P ICCARDO , G. 2000. The Alps and their forelands. In: C AVAZZA , W., R OURE ,
ophiolite-oceanic lithosphere analogue: new insights F., S PAKMAN , W., S TAMPFLI , G. M. & Z IEGLER , P.
from the northern Apennines (Italy). In: D ILEK , Y., (eds) The TRANSMED Atlas: The Mediterranean
M OORES , E. M., E LTHON , D. & N ICOLAS , A. (eds) Region from Crust to Mantle. Springer Verlag,
Ophiolites and oceanic crust: new insights from field Germany, unpaginated.
studies and the ocean drilling project. Geological S CHÖNLAUB , H. P. & H ISTON , K. 2000. The Palaeozoic
Society of America Special Paper, 349, 21–34. evolution of the Southern Alps. Mitteilungen der öster-
R INGGENBERG , Y., T OMASSI , A. & S TAMPFLI , G. M. reichischen geologischen Gesellschaft, 92, 15–34.
2002. The Jurassic sequence of the Niesen nappe in S ENGÖR , A. M. C. 1979. Mid-Mesozoic closure of Permo-
the region of Le Sépey-La Forclaz (Switzerland): Triassic Tethys and its implications. Nature, 279,
witness of the Piemont rifting in the Helvetic paleogeo- 590–593.
graphic domain. Bulletin de Géologie de l’Université S ENGÖR , A. M. C. 1985. The story of Tethys: how many
de Lausanne, 87(4), 353–372. wives did Okeanos have. Episodes, 8(1), 3– 12.
R OBERTSON , A. H. F., D IXON , J. E. ET AL . 1996. Alterna- S ENGÖR , A. M. C., Y ILMAZ , Y. & S UNGURLU , O. 1984.
tive tectonic models for the Late Palaeozoic-Early Tectonics of the Mediterranean Cimmerides: nature
Tertiary development of Tethys in the Eastern Mediter- and evolution of the western termination of Paleo-
ranean region. In: M ORRIS , A. & T ARLING , D. H. Tethys. In: D IXON , J. E. & R OBERTSON , A. H. F.
(eds) Palaeomagnetism and Tectonics of the Eastern (eds) The Geological Evolution of the Eastern Mediter-
Mediterranean Region. Geological Society, London, ranean. Geological Society, London, Special Publi-
Special Publications, 105, 239–263. cations, 17, 77–112.
R OCA , E. 2001. The Northwest Mediterranean Basin S RIVASTAVA , S. P. & T APSCOTT , C. R. 1986. Plate kin-
(Valencia trough, Gulf of Lions and Liguro-Provençal ematics of the North Atlantic. In: V OGT , P. R. &
baisns): structure and geodynamic evolution. In: T UCHOLKE , B. E. (eds) The Western North Atlantic
Z IEGLER , P. A., C AVAZZA , W., R OBERTSON , region. Geological Society of America, M, 379 –404.
A. H. F. & C RASQUIN -S OLEAU , S. (eds) PeriTethys S RIVASTAVA , S. P., R OEST , W. R., K OVACS , L. C.,
memoir 6: Peritethyan rift/wrench basins and O AKAY , G., L ÉVESQUE , S., V ERHOEF , J. &
passive margins, IGCP 369. Mémoires du Museum M ACNAB , R. 1990. Motion of Iberia since the Late
Nationale d’Histoire Naturelle, 186, 671–706. Jurassic: results from detailed aeromagnetic measure-
R UBATTO , D. 1998. Dating of pre-Alpine magmatism, ments in the Newfoundland Basin. Tectonophysics,
Jurassic ophiolites and Alpine subductions in the 184, 229 –260.
Westen Alps. ETHZ, Zürich, 1– 173. S TAMPFLI , G. M. 1993. Le Briançonnais, terrain exotique
S AVELLI , C. 2000. Subduction-related episodes of K’alka- dans les Alpes? Eclogae geologicae Helvetiae, 86(1),
line magmatism (15–0.1 Ma) and geodynamic 1– 45.
implications in the north Tyrrhenian – central Italy S TAMPFLI , G. M. 2000. Tethyan oceans. In: B OZKURT ,
region: a review. Journal of Geodynamics, 30, E., W INCHESTER , J. A. & P IPER , J. D. A. (eds) Tec-
575– 591. tonics and magmatism in Turkey and surrounding
S CHARDT , H. 1889. Sur l’origine des Prealpes Romandes. area. Geological Society, London, Special Publi-
Eclogae Geologicae Helvetiae, 4, 129–142. cations, 173, 1 –23.
S CHARDT , H. 1898. Les régions exotiques du versant Nord S TAMPFLI , G. M. 2005. Plate tectonic of the Apulia-Adria
des Alpes Suisse. Préalpes du Chablais et du Stockhorn microcontinents. In: F INETTI , I. R. (ed.) CROP
et les Klippes. Bulletin de la Société Vaudoise des PROJECT deep seismic exploration of the Mediterra-
Sciences Naturelles, 34(128), 113–219. nean and Italy. Elsevier, The Netherlands, 747–766.
S CHARDT , H. 1900. Encore les régions exotiques. Répli- S TAMPFLI , G. M. & M ARCHANT , R. H. 1995. Plate con-
que aux attaques de M. Emile Haug. Bulletin de la figuration and kinematics in the Alpine region. Accade-
Société Vaudoise des Sciences Naturelles, 36, mia Nazionale delle Scienze, Scritti e Documenti, 14
147– 169. (Atti del congresso ‘Rapporti tra Alpi e Appennino’),
S CHARDT , H. 1907. Les vues modernes sur la tectonique et 147–166.
l’origine de la chaı̂ne des Alpes. Archives de physique S TAMPFLI , G. M. & M ARCHANT , R. H. 1997. Geody-
et des sciences naturelles de Genève, 4(23), 356–385/ namic evolution of the Tethyan margins of the
483– 496. Western Alps. In: P FIFFNER , O. A., L EHNER , P.,
S CHÄRER , U., DE P ARSEVAL , P., P OLVE , M. & DE S AINT H EITZMAN , P. Z., M UELLER , S. & S TECK , A. (eds)
B LANQUAT , M. 1999. Formation of the Trimouns talc- Deep structure of the Swiss Alps – Results from NRP
chlorite deposit (Pyrenees) from persistent hydrother- 20. Birkhaüser, Basel, 223– 239.
mal activity between 112 and 97 Ma. Terra Nova, S TAMPFLI , G. M. & B OREL , G. D. 2002. A plate tectonic
11(1), 30–37. model for the Paleozoic and Mesozoic constrained by
S CHETTINO , A. & S COTESE , C. 2002. Global kinematic dynamic plate boundaries and restored synthetic
constraints to the tectonic history of the Mediterranean oceanic isochrons. Earth and Planetary Science
region and surrounding areas during the Jurassic Letters, 196, 17– 33.
and Cretaceous. In: R OSENBAUM , G. A. & L ISTER , S TAMPFLI , G. M. & B OREL , G. D. 2004. The
G. S. (eds) Reconstruction of the evolution of the TRANSMED transects in space and time: constraints
on the Paleotectonic evolution of the Mediterranean T ANNER , D. C., B EHRMANN , J. H., O NCKEN , O. &
Domain. In: C AVAZZA , W., R OURE , F., S PAKMAN , W EBER , K. 1998. Three-dimensional retro-modelling
W., S TAMPFLI , G. M. & Z IEGLER , P. (eds) The of transpression on a linked fault system; the Upper
TRANSMED Atlas: The Mediterranean Region from Cretaceous deformation on the western border of
Crust to Mantle. Springer Verlag, Germany, 53– 80. the Bohemian Massif, Germany. Geological Society,
S TAMPFLI , G. M. & K OZUR , H. 2006. Europe from the London, Special Publications, 135, 275– 287.
Variscan to the Alpine cycles. In: G EE , D. G. & S TE- T HÖNI , M. & J AGOUTZ , E. 1992. Some new aspects of
PHENSON , R. (eds) European Lithosphere Dynamics. dating eclogites in orogenic belts: Sm– Nd, Rb–Sr
Geological Society, London, Memoirs, 32, 57–82. and Pb–Pb isotopic results from the Austroalpine
S TAMPFLI , G. M., M OSAR , J., M ARCHANT , R., Saualpe and Koralpe type-locality (Carinthia/Styria,
M ARQUER , D., B AUDIN , T. & B OREL , G. 1998. Sub- SE Austria). Geochimica Cosmochimica Acta, 56,
duction and obduction processes in the western Alps. 347– 368.
In: V AUCHEZ , A. & M EISSNER , R. (eds) Continents T OURY , A. 1985. Etude géologique de la Vallée de la
and Their Mantle Roots. Tectonophysics, 296(1 –2), Valloirette entre Valloire, le Col du Galibier et les
159–204. Aiguilles d’Arves (Alpes occidentales – Savoie).
S TAMPFLI , G. M., B OREL , G., C AVAZZA , W., M OSAR , J. Savoie University, Chambéry, 207.
& Z IEGLER , P. A. 2001a. The Paleotectonic Atlas of T RÜMPY , R. 1988. A possible Jurassic-Cretaceous trans-
the Peritethyan Domain. CD ROM; European Geophy- form system in the Alps and the Carpathians. Geologi-
sical Society. cal Society of America Special Paper, 218, 93– 109.
S TAMPFLI , G. M., M OSAR , J., F AVRE , P., P ILLEVUIT , A. T RÜMPY , R. 1992. Ostalpen und Westalpen-Verbindendes
& V ANNAY , J.-C. 2001b. Permo-Mesozoic evolution und Trennendes. Journal of Geology, 135(4),
of the western Tethyan realm: the Neotethys/East- 875– 882.
Mediterranean connection. In: Z IEGLER , P. A., V AI , G. B. 1994. Crustal evolution and basement elements
C AVAZZA , W., R OBERTSON , A. H. F. & C RASQUIN - in the Italian area: paleogeography and characteriz-
S OLEAU , S. (eds) PeriTethys memoir 6: Peritethyan ation. Bollettino di Geofisica Teorica ed Applicata,
rift/wrench basins and passive margins, IGCP 369. 36(141–144), 411–434.
Mémoires du Museum Nationale d’Histoire Naturelle, VAN H INSBERGEN , D. J. J., H AFKENSCHEID , E.,
186, 51–108. S PAKMAN , W., M EULENKAMP , J. E. & W ORTEL , R.
S TAMPFLI , G. M., B OREL , G., M ARCHANT , R. & 2005. Nappe stacking resulting from subduction of
M OSAR , J. 2002. Western Alps geological constraints oceanic and continental lithosphere below Greece.
on western Tethyan reconstructions. Journal Virtual Geology, 33(4), 325–328.
Explorer, 8, 77– 106. V ERGÉS , J. & G ARCIA -S ENZ , J. 2001. Mesozoic evol-
S TAMPFLI , G. M., V AVASSIS , I., D E B ONO , A., R OSSE- ution and Cainozoic inversion of the Pyrenean Rift.
LET , F., M ATTI , B. & B ELLINI , M. 2003. Remnants In: Z IEGLER , P. A., C AVAZZA , W., R OBERTSON ,
of the Palaeotethys oceanic suture-zone in the A. H. F. & C RASQUIN -S OLEAU , S. (eds) PeriTethys
western Tethyan area. In: C ASSINIS , G. & D ECANDIA , memoir 6: Peritethyan rift/wrench basins and
F. A. (eds) Stratigraphic and Structural Evolution on passive margins, IGCP 369. Mémoires du Museum
the Late Carboniferous to Triassic Continental and Nationale d’Histoire Naturelle, 186, 187– 212.
Marine Successions in Tuscany (Italy): Regional VON B LANCKENBURG , F. & D AVIES , J. H. 1995. Slab
Reports and General Correlation. Bolletino della breakoff: a model for syncollisional magmatism and
Società Geologica Italiana, Volume speciale, 2, 1– 24. tectonics in the Alps. Tectonics, 14, 120–131.
S TEINER , C. W., H OBSON , A., F AVRE , P., S TAMPFLI , W ORTEL , M. J. R. & S PAKMAN , W. 1992. Structure and
G. M. & H ERNANDEZ , J. 1998. The Mesozoic sequence dynamics of subducted lithosphere in the Mediterra-
of Fuerteventura (Canary islands): witness of an Early nean region. Proceedings of the koninklijke neder-
to Middle Jurassic sea-floor spreading in the Central landse Akademie van Wetenschappen. Series B:
Atlantic. Geological Society of America Bulletin, Palaeontology, Geology, Physics, Chemistry, Anthro-
110(10), 1304– 1317. pology, 95(3), 325– 347.
S TÖCKLIN , J. 1968. Structural history and tectonics of W ORTEL , M. J. R. & S PAKMAN , W. 1993. The dynamic
Iran: a review. American Association of Petroleum evolution of the Apenninic-Calabrian, Helvetic, and
Geologists Bulletin, 52(7), 1229– 1258. Carpathian arcs: an unifying approach. Terra
S TÖCKLIN , J. 1974. Possible ancient continental margin in Abstracts, 5, 97.
Iran. In: B URK , C. A. & D RAKE , C. L. (eds) The W ORTMANN , U. G., W EISSERT , H., F UNK , H. & H AUCK ,
Geology of Continental Margins. Springer Verlag, J. 2001. Alpine plate kinematics revisited: the Adria
Germany, 873–887. Problem. Tectonics, 20, 134– 147.
The Calabrian Orocline: buckling of a previously more linear orogen
STEPHEN T. JOHNSTON1* & STEFANO MAZZOLI2
1
School of Earth & Ocean Sciences, University of Victoria, PO Box 3055
STN CSC, Victoria, British Columbia, Canada
2
Dipartimento Scienze della Terra, Università di Napoli ‘Federico II’, Largo San
Marcellino 10, 10138 Napoli (NA), Italy
*Corresponding author (e-mail: stj@uvic.ca)
Abstract: Recent structural studies of the Apennines and the Calabrian orocline and a compilation
of structural, stratigraphic, GPS and palaeomagnetic data from the central and western
Mediterranean region show that beginning in the Late Miocene a N– S trending ribbon continent
that had been previously deformed, and which we now recognize as the Apennine– Sicilian
thrust belt, buckled eastward in response to northward movement of Africa relative to stable
Europe. A simple geometric model is consistent with available data and shows how eastward buck-
ling of an originally north– south continental beam explains: (1) opening of the Tyrrhenian Sea
basin from 7 –2 Ma, at which point sea-floor spreading ceases and the basin begins to shrink by
southward subduction beneath Sicily; (2) the coeval development of an east-verging fold-and-
thrust belt along the length of the Apennine–Sicilian belt in response to overthrusting of the
autochthon to the east, followed by extension beginning at 1 Ma as the Apennine portion of the
beam begins to retreat to the SW; and (3) subduction of continental and oceanic lithosphere east
of the buckling beam into a trench that migrates eastward through time due to ‘push back’ by
the buckling upper plate.
Introduction respectively) of the Apennine –Sicilian peninsula

(Kastens et al. 1988) provides no explanation for
The Tyrrhenian Sea (Fig. 1) is amongst the world’s formation of the intervening Apennine –Sicilian
youngest oceanic basins, having opened at rates of fold-and-thrust belt coeval with opening of the
60– 100 km/Ma over the past 7–9 my (Rosenbaum Tyrrhenian Sea. In addition, bend formation in
& Lister 2004). The opening of the Tyrrhenian Sea response to roll-back would require significant
region was coeval with significant crustal shorten- extension parallel to the length of the Apennine –
ing in the Apennine and Sicilian thrust belts that Sicilian peninsula (perpendicular to the thrust
bound the sea to the east and south, respectively belt), for which there is no evidence.
(Fig. 1). Together these mountain chains define a We ask ‘Can the coeval opening of the Tyrrhe-
highly arcuate continental orogen, the bend of nian Sea, shortening in the Apennine –Sicilian
which is commonly referred to as the Calabrian Oro- fold-and-thrust belt, and formation of the Calabrian
cline. There is concensus that the Calabrian orocline orocline, be explained in a single tectonic model?’
developed after 10 Ma, and is a tectonic feature that Toward answering this question we describe the
developed as a result of strain of an originally more major tectonic elements of the southern Italian
linear orogen (Cavazza et al. 2004). However, the region and summarize the neo-tectonic setting of
tectonic setting that gave rise to the bend, its 3D geo- the region. We review the data pertaining to the
metry, and the role the bend played in the opening of nature of the Calabrian Orocline, and end with the
the Tyrrhenian Sea remain matters of debate presentation of, and a discussion regarding, a
(Mantovani et al. 1996; Tapponnier 1977). For model of lithospheric-scale orocline development.
instance, models of the bend as a thin-skinned Our orocline model explains opening and the
feature, limited to the allochthonous Apennine– already initiated closing of the Tyrrhenian Sea, the
Sicilian thrust sheets and not present at deeper coeval crustal shortening within the Apennine –
crustal or lithospheric levels (Eldredge et al. Sicilian fold-and-thrust belt, and the recently
1985), provides no explanation for the coeval initiated extension in the Apennine portion of the
formation of the bend and opening of the Tyrrhenian fold-and-thrust belt. Our model implies that the
Sea. Models of the bend as a product of roll-back of southeastward retreat of the trench bounding
a west- to north-dipping subduction zone along the east side of the Apennine –Sicilian orogen was
the east to south margin (from north to south, driven, at least in part, by the eastward buckling
114 S. T. JOHNSTON & S. MAZZOLI
Fig. 1. A simplified geological map of the Italian peninsula, Sicily and the Tyrrhenian Sea. Inset cross-section
shows schematic relationships between units. Black islands north of Sicily are volcanic islands of the Aeolian magmatic
arc. The thick black line with teeth pointing into the hanging wall shows the Apennine (Italy)–Maghrebide (Sicily)
thrust front (open teeth) that connect through the Calabrian subduction zone (black teeth) and which continues west
through Tunisia (T). Black lines crossing the southernmost Apennines and NE Sicily are the sinistral Sangineto (SL) and
dextral Taormina (TL) lines. The Tyrrhenian Sea is locally floored by oceanic crust (grey) that was generated at the
Vavilov (V) and Marsili (M) north-trending spreading ridges (parallel lines), and is bound to the west by continental
crust of Corsica (C) and Sardinia (S). Sicily is beginning to overthrust the south margin of the Tyrrhenian Sea along a
north-verging thrust fault (line with small open teeth pointing into the hanging wall). Large white arrows show modelled
crustal velocities based on geodetic data (Oldow et al. 2002).
of the Apennine –Sicilian fold-and-thrust belt. In Apulian–Hyblean foreland, all of which have
this model, opening of the Tyrrhenian Sea is a different tectonic histories and which are currently
product, rather than a cause of orocline formation. moving with respect to one another (described
Orocline formation has probably been driven by below). Throughout the following discussion the
the northward drift of Africa relative to Europe, Apennine geology is first introduced, and then the
buckling the intervening Apennine –Sicilian Sicilian equivalents (Fig. 2).
ribbon continent.
The Apennine– Sicilian fold-and-thrust belt
Tectonic elements of southern Italy Three main tectono-stratigraphic units comprise

the fold-and-thrust belt. These are, in ascending
The southern Italian region consists of four main structural order, a Triassic –Palaeogene continental
tectono-geographic regions: (1) the Apennine – ribbon, referred to as the Internal Platform; an
Sicilian fold-and-thrust belt; (2) the Apulian– overlying Jurassic–Cretaceous basinal sequence,
Hyblean foreland; (3) the Tyrrhenian Sea; and (4) known as the Ligurides –Sicilides, that include
the Calabrian subduction zone. Hence, the dismembered ophiolite; and an uppermost exotic
Italian peninsula and its continuation through Alpine continental margin on which sit crystal-
Calabria into Sicily, is divisible into a western to line Variscan thrust sheets, referred to as the
northern Tyrrhenian passive margin; a central Calabrian –Peloritanian nappes. The narrow
Apennine– Sicilian mountain belt, and an eastern (200 –400 km wide) continental Internal Platform
THE CALABRIAN OROCLINE: BUCKLING OF A PREVIOUSLY MORE LINEAR OROGEN 115
Fig. 2. Schematic structural–stratigraphic columns showing Sicilian (at left)–Apennine (at right) correlations around
the Calabrian orocline. Continental platform (white) and basinal (grey) units are arranged in structural stacking order
with the highest structural levels at the top and the lowest at the bottom. All units are bound by foreland-verging
thrust faults (teeth point into the hanging wall). Post-early Tortonian imbrication of basinal strata of the Lagonegro–
Immerese–Sicanian basin was preceded by an earlier inversion event of probable Early Miocene age (indicated by *).
was continuous to the north into the northern basin in the Neogene included removal of the base-
Apennines and to the south and then west ment beneath the basin, while the Ionian basin
through Sicily and into the Maghrebides of north- remained open and intact. The presence of one or
ern Africa (Elter et al. 2003). The structural more straits through which the basins may have
emplacement of the Alpine –Variscan nappes been linked, for example between Apulia and
above the Internal Platform along east-south- Hyblea (Gattacceca & Speranza 2002; Stampfli &
verging thrust faults was facilitated by Borel 2004), cannot be ruled out.
Langhian (14–16 Ma) closure of the intervening Closure of the Lagonegro– Imerese– Sicanian
Liguride-Sicilide (Alpine –Tethys) basin (Stampfli basin was complicated and included pre-Tortonian
& Borel 2002) (Fig. 2). (.12 Ma) imbrication (Mazzoli et al. 2001;
To the east and south, the Internal Platform Shiner et al. 2004). Post-Tortonian (,7 Ma)
passes out into a condensed Triassic and locally NE –S-verging thrust faults resulted in overthrust-
Permian–Eocene basinal sequence of chert and ing of the Lagonegro– Imerese– Sicanian basinal
deep-water shales of the Lagonegro –Imerese– sequences by the Internal Platform and its structu-
Sicanian basin. Although the Ionian Sea has been rally overlying units, imbrication of the basinal
interpreted as a preserved remnant of the Neotethys sequence and the removal of its transitional to
(Stampfli & Borel 2002), two observations argue oceanic lithospheric basement, and structural
against such a correlation: (1) the Lagonegro – emplacement of the Internal Platform above conti-
Imerese –Sicanian basinal strata today lies west nental Apulia –Hyblea foreland (Cello & Mazzoli
and north of, is thrust over, and is separated from 1999). These thrust faults steepen to the west, root
the Ionian, by shelf and platformal strata of the con- into the sub-crustal lithospheric mantle, and offset
tinental Apulian–Hyblean foreland, making it the Moho (Barchi et al. 1998; Butler et al. 2004).
unlikely that the two basins were continuous; and Tomographic imaging shows a discontinuous,
(2) closure of the Lagonegro –Imerese –Sicanian steeply west-dipping, relatively low velocity slab
beneath the Apennines (Lucente et al. 1999; Eldredge et al. (1985), palaeomagnetic declinations
Spakman & Wortel 2004), which likely consists of vary around the bend, implying counter-clockwise
transitional and entrained continental lithosphere rotation of the southern Apennines of up to 568,
(Faccenna et al. 2004). and clockwise rotation of Sicily of up to 1408, but
Overthrusting of the Apulia –Hyblea ended after with a significant scatter. In addition to varying
the Miocene, probably in the Early Pleistocene around the Calabrian bend, vertical axis rotations
(Hippolyte et al. 1994; Patacca & Scandone 2001). increase structurally up-section, such that the struc-
A new tectonic regime, characterized by NE–SW turally highest thrust sheets exhibit the greatest
extension (Cello et al. 1982; Hippolyte et al. rotations (Channell et al. 1980). Subsequent palaeo-
1994; Montone et al. 1999) was established within magnetic studies (Gattacceca & Speranza 2002;
the mountain chain and the adjacent foothills to Rosenbaum & Lister 2004; Speranza et al. 1998,
the east. The result has been the recent and 1999, 2003) have served to confirm the results
ongoing development of extensional and related of the early palaeomagnetic studies. Rotations
transcurrent faults that post-date and dissect the attributable to landslides have, at least locally,
thrust belt (Butler et al. 2004; Cello et al. 1982). significantly complicated the story (Catalano et al.
The regional strike of Apennine –Sicilian 1984; Grasso et al. 1983; Incoronato & Nardi
nappes and thrust faults varies smoothly around 1989). Significant rotation of Pleistocene rocks
the Calabrian arc (Fig. 3). As demonstrated by (Scheepers & Langereis 1994) indicates that bend
Fig. 3. A series of maps (a –d) showing, at left, the current configuration of southern Italy and, at right, maps
depicting the palaeogeography at 10 Ma in which the Calabrian orocline has been restored according to available
palaeomagnetic data. Unit fills are as in Figure 1. (a) Palinspastic restoration (dashed lines) of Oligocene –Miocene
thrust sheets according to their vergence directions (arrows) results in a significant overlap of the restored southern
Apennine and Sicilian thrust sheets – presenting a significant space problem (at left) that is ameliorated by undoing of
the Calabrian orocline (at right). (b) Propagation of thrust sheets toward the foreland requires .225% extension around
the Calabrian orocline (at left – note increase in length of dashed line as it propagates to the south and east), whereas
pre-oroclinal thrusting does not require extension (at right). (c) The regional strike of structures, which varies smoothly
around the orocline (at left) restores to a linear belt of structures (at right). (d) Palaeomagnetic directions vary as a
function of change in strike [see inset at left from (Eldredge et al. 1985)] consistent with oroclinal bending. Declinations
point uniformly north upon restoration of the orocline (at right).
formation is ongoing. Contrary to the palaeo- the sea is the youngest in the Mediterranean region
magnetic data, geodetic data imply fairly uniform (Sartori 2003). Crust underlying the basin varies
northward movement, at rates of .5 mm a21, of between ‘transitional’ (15-km thick) to true oceanic
the Apennine –Sicilian thrust belt (also referred to crust (6-km thick) (Sartori 2003). The north-
as Apulia) relative to a fixed Eurasian block tapering, triangular shape of the basin implies that
around the length of the Calabrian orocline it opened about a pole of rotation to the north, consist-
(Fig. 1) (Jimenez-Munt et al. 2003; Mattei et al. ent with the restricted development of oceanic
2007; Nocquet & Calais 2003; Oldow et al. 2002). crust in the southern Tyrrhenian sea (the region farth-
est from the Euler pole), the north–south trend of
The Apulian – Hyblean Foreland spreading ridges (e.g. the Vavilov and Marsili
ridges) and the east –west strike of transform faults
The foreland to the Apennine– Sicilian fold- (Fig. 1) (Marani & Gamberi 2004). The basin is
and-thrust belt consists of a west- to north-facing characterized by a positive Bouguer gravity
continental margin sequence of Permian to anomaly and a geoid high (Cavazza et al. 2004) con-
Pleistocene age, consisting largely of platformal sistent with its relative youth. Tomographic studies
carbonates (Fig. 2). Unlike the Internal Platform of imply that anomolously hot asthenosphere extends
the Apennine –Sicilian fold-and-thrust belt, the to depths of 300 km (Spakman & Wortel 2004).
foreland carbonate platform was little affected by There is, however, no active spreading within the
Neogene tectonism, and is first overrun by orogenic basin. The westerly Vavilov ridge was active from
flysch deposits in the Late Miocene. Apulia –Hyblea 7 –3.5 Ma; spreading along the younger and more
form part of a coherent continental lithospheric easterly Marsilli ridge started at about 2 Ma and
block referred to as Adria (Channell & Horvath ceased by 1 Ma (Kastens & others 1988).
1976). Adria may be a northern promontory of Eastward-younging of the basin is consistent with
Africa (Channell et al. 1979; McKenzie 1972) or a an eastward increase in heatflow to relatively high
microplate that originated between Africa and values (.120 mWm22) (Mongelli et al. 2004).
Europe (Anderson & Jackson 1987; Dercourt et al. First motion analyses of earthquake data and
1986). GPS data indicate that the Apulia –Hyblea direct stress measurements (Jimenez-Munt et al.
portion of Adria is currently moving with and 2003) imply convergence between and dextral,
appears to be a part of the African continent transpressional overthrusting of the southern
(Nocquet & Calais 2003; Oldow et al. 2002). Tyrrhenian Sea by the north coast of Sicily
Despite now being separated by Calabria, (Anderson & Jackson 1987). Actively developing
continuity of the Apulian–Hyblean continental NW-trending grabens that are rotating clockwise
margin beneath the Calabrian nappes is suggested characterize the north coast of Sicily (Guarnieri
by their shared stratigraphy and tectonic evolution 2004; Somma 2006). The southern margin of the
(Catalano et al. 1976). Palaeomagnetic studies of Tyrrhenian sea is being overthrust along a seismi-
Neogene strata confirm that Apulia has not rotated cally active, north-verging thrust system which
with respect to Hyblea, consistent with these plat- roots beneath Sicily (Giunta et al. 2003). Regional
forms having developed on, and forming part of a compilations of earthquake seismic data indicate
coherent Adria lithospheric block (Besse et al. southeastward subduction of the southern Tyrrhe-
1984; Eldredge et al. 1985; Scheepers 1992). nian Sea lithosphere beneath Sicily and Calabria
Interpretations involving the presence of a strait (Anderson & Jackson 1987; Neri et al. 1996).
floored by Ionian oceanic lithosphere separating
Apulia and Hyblea require that subduction of the Calabrian Subduction Zone
oceanic lithosphere underlying the strait occurred
without resulting in rotation or translation of The marine region SE of Calabria is characterized
Apulia relative to Hyblea (Rosenbaum et al. 2002) by an accretionary prism bound to the SE by a
and are inconsistent with Apulia– Hyblea moving poorly defined trench, SE of which lies Permian or
coherently with and forming a coherent continu- younger oceanic crust of the Ionian Sea (Fig. 2)
ation of the African continent. Post-Tortonian (Stampfli 2000). These observations indicate that
closure of the Lagonegro –Imerese –Sicanian basin the Ionian oceanic crust is being consumed in a
led to overthrusting of the Apulian–Hyblean fore- subduction zone that dips NW beneath Calabria.
land by the Internal Platform along NE- to The volcanic Aeolian Islands, which lie north of
S-verging thrust faults. Sicily and Calabria, are inferred to be an island arc
developed above the dewatering, down-going
The Tyrrhenian Sea Ionian oceanic slab (Fig. 1) (Argnani & Savelli
1999). Tomographic studies indicate the presence
Tortonian rift deposits fringing the sea imply initial of a discontinuous, steeply dipping, high velocity
opening of the basin at 8 –9 Ma and indicate that zone that dips down to depths of 500 km beneath
Calabria. The upper portions of the high velocity originally more linear belt was responsible for the
zone appears continuous with the subducting orocline formation (Fig. 3) (Eldredge et al. 1985).
Ionian slab (Lucente et al. 1999; Spakman & However, the increase in vertical axis rotations
Wortel 2004), implying that what is being imaged structurally up-section led to concensus that the oro-
is, at least in part, the Ionian oceanic lithosphere. cline was thin-skinned, that the vertical axis
Moderate to deep earthquakes occurring within rotations were confined to thin allochthonous
this high velocity zone are consistent with a subduc- thrust sheets, and that the rotations occurred
tion setting (Wortel & Spakman 2000). during thrust sheet emplacement. Timing con-
Subduction beneath Calabria began at 7 Ma, after straints imply that orocline formation commenced
a 7 my period of tectonic quiescence marked by a in the Tortonian, coeval with the initiation of
14–7 Ma magmatic gap that characterizes the opening of the Tyrrhenian Sea (Rosenbaum &
region (Argnani & Savelli 1999; Cavazza et al. Lister 2004; Speranza et al. 1999) and with
2004). Low volumes of arc magmas, characterized thick-skinned crustal shortening along the length
by the eruption of andesite to rhyolite and the intru- of the Apennine –Sicilian fold-and-thrust belt.
sion of granodiorite to monzogranite, occurs
from 7 Ma onwards (Cavazza et al. 2004; Savelli A tectonic model
2002), presumably providing a record of ongoing
subduction throughout this interval. Subduction We assume that the Calabrian bend is an orocline
has consumed a 500 km-wide strip of lithosphere, as that developed in response to lithospheric-scale
indicated by the 500 km length of the tomographi- buckling of a formerly linear portion of a ribbon
cally-imaged subducted slab (Lucente et al. 1999; continent about a vertical axis of rotation. A
Spakman & Wortel 2004). The southern Tyrrhenian model of an eastward-buckling beam (Fig. 4) pre-
Sea is 600 km wide, implying that there is a dicts rapid extension behind (west of ) the beam
one-to-one relationship between the amount of litho- giving rise to a significant basin, representing the
sphere subducted and the width of the Tyrrhenian Tyrrhenian Sea. Basin formation behind the buck-
Sea. Hence eastward retreat of the subduction zone ling beam is coeval with strike-normal compression
was accommodated by the opening of the Tyrrhenian within the beam as it migrates east over, and collides
Sea in its wake. with, autochthonous foreland. A model of buckling
requires counter-clockwise rotation of the Apen-
The calabrian orocline nines and clockwise rotation of the thrust belt in
Sicily, consistent with available GPS and palaeo-
Carey was the first to suggest that the continuity of magnetic data. Our model predicts that the
the geological belts of the Apennine –Sicilian maximum areal extent of the Tyrrhenian Sea was
fold-and-thrust belt around the Calabrian bend achieved at c. 2 Ma (Fig. 4), consistent with the ter-
implies that the bend is an orocline, that the geologi- mination of spreading along the Marselli ridge, the
cal belts were formerly more linear in map-view, youngest spreading ridge in the basin, at 2 Ma.
and that their current geometry is the result of a tec- The subsequent and increasingly rapid reduction in
tonic rotation (Carey 1955). Several other lines of the area of the Tyrrhenian Sea predicted by our
evidence support an oroclinal origin for the bend model explains the recent initiation of southward
(Catalano et al. 1976) (Fig. 3). Oligocene and subduction of the southern Tyrrhenian Sea litho-
Miocene thrust faults around the Calabrian orocline sphere beneath Sicily and Calabria. The Early Plies-
verge E –NE in the Apennines and south in Sicily. tocene switch from crustal shortening to extension
Palinspastic restoration of these thrust sheets along the length of the Apennines is explained as
results in a significant space problem as the restored a result of the predicted southwestward retreat of
Apennine and Sicilian units exhibit considerable the Apennines after reaching a maximum easterly
overlap (Fig. 3). This space problem can also be position at about 2 Ma (Fig. 4). In our model, oro-
thought of as a line length problem: propagation cline formation and the related opening of the Tyr-
of thrust sheets toward the foreland around the rhenian Sea is driven by north –south shortening
length of the Calabrian orocline should have and requires convergence between autochthonous
resulted in enormous (.200%) extension of Africa and Europe since about 9 Ma (Fig. 5).
the thrust sheets along normal faults oriented per- Map-view palinspastic restoration of the conti-
pendicular to the thrust faults (Fig. 3). On the con- nental ribbon formed by the Apennine –Sicilian
trary, extension is restricted to the Apennines, and fold-and-thrust belt into a north–south trending
neotectonic studies imply continued strike-parallel linear belt (Figs 4 & 5) removes the space and line
compression around the Calabrian bend (Di Bucci length problems associated with the fold-and-thrust
& Mazzoli 2002; Di Bucci et al. 2003). belt by restoring the Apennine and Sicilian thrust
Palaeomagnetic declination varies as a function faults into a continuous north–south trending, east-
of structural strike, implying that bending of an verging belt (Figs 4 & 5). Unbuckling accomplished
Fig. 4. A geometric model showing how buckling of an originally linear beam can explain the origin and evolution of a
basin and the coeval development of a spatially-related orocline. The model is organized to represent buckling of an
originally linear Apennine– Sicilian continental ribbon, represented at left by the black and white bar. Throughout the
model, the buckling bar is pinned at the bottom, with a fixed hinge point (black circle) separating a short lower (southerly)
portion of the beam, representing the Sicilian–Calabrian portion of the Calabrian orocline, from a longer upper
(northerly) portion, representing the Apennines. A grey and white unbuckled beam is shown in each panel for reference.
Buckling is constrained to occur by eastward migration of the hinge, is accomplished by southward migration of the
upper (northern) most point of the beam at a constant rate and is assumed to have begun at 9 Ma. The developing buckle is
shown in the three panels at right, at 3 Ma intervals. The grey region behind the eastwardly migrating beam represents the
Tyrrhenian Sea. The relative area of the basin, from 0 –100% (see axis at left) through time is indicated by the solid
grey line (dashed line shows projected future area of basin). The arrow indicates the point at which the basin reaches its
maximum extent (at c. 2 Ma). By 1 Ma, the basin begins to shrink noticeably. The last panel, at right, represents the
current geometry of the Calabrian orocline and Tyrrhenian Sea and requires 30% north–south shortening of the system.
The dashed line shows the maximum easterly extent of the buckling beam, also achieved at c. 2 Ma, requiring significant
post-2 Ma southwestwardly retreat of the northern (Apennine) portion of the buckling beam.
by undoing the palaeomagnetically constrained and 100 km to the south relative to northern Italy
clockwise rotations of Sicily– Calabria and the (Fig. 5).
counter-clockwise rotations of the Apennines Eastward migration of the buckling continental
yields an almost linear peninsula, consistent with beam formed by the Apennine –Sicilian belt
our model. The remaining variability in the restored requires that autochthonous foreland lithosphere
declinations may be attributable to the presence of be removed by subduction or dramatically
additional minor bends which have not been restored thickened. Ongoing subduction of Ionian oceanic
in our model. Small-scale local rotations adjacent lithosphere beneath Calabria accounts for the
to individual faults may also contribute to the varia- foreland lithosphere overrode by the Calabrian–
bility. Land-sliding (Incoronato & Nardi 1989) is Sicilian southern portion of the orocline. The
likely an important factor in contributing to local discontinuous slab imaged beneath the Apennines
palaeomagnetic variability. Unbuckling closes the (Lucente et al. 1999) probably constitutes the
Tyrrhenian Sea and results in juxtaposition of the transitional continental lithosphere that formed the
Apennine –Sicilian thrust belt and the Corsica– basement to the Lagonegro basin and which was
Sardinia peninsula, and restores the Sicilian subsequently subducted beneath the encroaching
portion of the ribbon continent 600 km to the west Apennine peninsula.
Fig. 5. Palaeogeographic evolution of the Apennine–Sicilian belt in response to development of the Calabrian Orocline
from 10 Ma (a) through 6 Ma (b) 3 Ma (c) to present (d). Present-day continental outlines (grey) are shown with the full
extent of continental crust shown in stipple. Oceanic to transitional crust in dark grey. Africa moves north at 1 cm/a
throughout this period. (a) Pre-Calabrian orocline palaeogeography with the Apennine–Maghrebide belt restored to a
linear geometry adjacent to Corsica-Sardinia and separated to the east from the Apulia–Hyblea continental ribbon by the
oceanic to transitional Lagonegro–Imerese–Sicanian (LI) basin. Black lines across the southernmost Apennines and
Sicily are the Sangineto and Taormina lines, respectively. Dashed line with open teeth pointing into the upper plate shows
incipient thrust fault/subduction zone that develops as the Apennine–Maghrebide belt buckles. In this and subsequent
frames, African–European convergence in the Dinarides is indicated by a thrust fault along the eastern margin of the
Adriatic Sea. (b) By 6 Ma, eastward buckling Apennine–Maghrebide belt, accommodated by the counter-clockwise
rotating Apennines and the clockwise rotating Maghrebides, has subducted the Lagonegro–Imerese–Sicanian basin
(black teeth indicating where subduction occurred), with off-scraped basinal sediments having been imbricated beneath
the Internal Platform and thrust over the Apulia–Hyblea continental ribbon, and opened the incipient Tyrrhenian Sea
(TS). (c) Continued buckling results in the Apennine–Maghrebide belt overthrusting oceanic crust of the Ionian sea SE of
Apulia–Hyblea, initating subduction of Ionian (I) oceanic crust (black teeth) with along-strike fold-and-thrust belt
development (open teeth) and further opening the TS. (d) Ongoing buckling of the Apennine–Maghrebide belt with
continued fold-and-thrust belt development (open teeth) and Ionian subduction (black teeth). Sicily has started to rotate
over the southern TS (north-verging thrust fault with small black teeth pointing into hanging wall).
In contrast to the Apennines and the Calabrian Discussion

hinge region of the orocline, the E–NE displacement
of Sicily and the Maghrebides is primarily accom- We explain the opening and already ongoing
plished by strike-slip displacement. Little or no destruction of the Tyrrhenian Sea, compression
subduction beneath the southern side of this portion across the Apennine –Sicilian belt, as well as the
of the belt is required (Fig. 5), consistent with the recent switch to extension along the Apennines, as
lack of a subducted slab imaged beneath this a consequence of formation of the Calabrian
region, the paucity of arc magmatism, and the pre- Orocline. Our model entails the buckling of an
ponderance of Messinian and younger dextral trans- originally linear ribbon continent. Such a model
pressional deformation. An implication of our implies that the orocline is of lithospheric scale,
model is that the Tell thrust front of northern involving the crust and some of the lithospheric
Africa, which bounds the Maghrebides to the south, mantle. The buckling beam is bound to the east
is probably largely a dextral strike-slip boundary and south by the steeply west- and north-dipping
which has been over-ridden by thrust faults with subduction surface, and to the west and north
minor (in comparison) dip-slip displacement. by the Tyrrhenian lithosphere, and is therefore
somewhat wedge-shaped (narrowing at depth) in initiating subduction remains unknown. The

profile. Buckling of a lithospheric-scale beam is Lagonegro–Imerese –Sicanian basin east of the
further implied by the subduction of the lithosphere Apennine –Sicilide beam had already been
lying east of the eastward-migrating buckle; the deformed prior to the initiation of orocline for-
opening and growth of an oceanic basin behind it mation. Subduction may, therefore, have initiated
to the west; and by deep reflection seismic and struc- along pre-existing faults.
tural studies demonstrating that fold-and-thrust belt As with any fold, determining why the hinge
deformation coeval with orocline formation was developed where it did, remains speculative.
thick-skinned, involving thrust faults that rooted Smaller coeval, parasitic bends are present along
steeply to the west into the sub-crustal mantle, and the length of the Apennine –Sicilide beam, includ-
which cut and offset the Moho. Removal of litho- ing the Umbrian arc. It may be that bending initially
sphere east of the eastward buckling beam was resulted in a number of smaller bends continuous
accommodated by the subduction of oceanic litho- with one another, with much of the strain sub-
sphere to the south and transitional lithosphere to sequently being localized within the Calabrian
the east. Eastward migration of the trench was a con- bend. Alternatively, it may be that an external
sequence of the eastward buckling, upper plate, factor dictated the location of the hinge. For
Apennine –Sicilian beam and hence trench retreat example, an oceanic strait separating Apulia from
was at least in part a product of ‘push back’. In Hyblia may have facilitated bending in the Calab-
other words, trench retreat was not a passive rian region, and would explain the subsequent
product of sinking of the lower plate (roll-back), migration of the subduction zone along which the
but was an active product of buckling of a beam Lagonegro–Imerese –Sicanian basin was consumed
within the upper plate about a vertical axis of into the Ionian basin.
rotation (push back). The Calabrian Orocline has previously been
Is it probably not reasonable to consider a conti- interpreted as a thin-skinned feature (Eldredge
nental ribbon as a strong beam relative to adjacent et al. 1985), because palaeomagnetic and structural
oceanic crust. What facilitated buckling of the data indicate that the structurally highest
Apennine –Sicilide beam was the presence of a sub- Apennine –Sicilian thrust sheets are the most
duction zone along the length of its eastern margin. rotated. The larger rotations recorded by the structu-
Hence the buckling Apennine –Sicilide beam need rally highest thrust sheets can, however, be
only have been stronger than the subduction inter- explained as a result of thick-skinned orocline
face bounding the east-margin of the beam. formation (Fig. 6). The line length around the
Whether a buckling continental beam is capable of outside, foreland-portion of an orocline significantly
Fig. 6. (a) Map-view of a linear lithospheric beam. Light-grey strip indicates the foreland-side of the beam, grey
the hinterland side. (b) The same beam buckled into an orocline that has flexed out toward the foreland. The ends of the
beam are kept fixed during orocline formation. Because of the excess line length around the inside of the orocline,
the hinterland-portion of the buckling beam (grey) is thrust out over the foreland (light grey). Note that the hinterland
derived thrust sheet forms the structurally highest element in the orocline, and is characterized throughout by the
largest rotations.
exceeds the distance around its hinterland portion – Geodetic data imply northward motion of the
the wider the buckling beam, the greater the outer to Apennines and correlative Sicilian crust relative to
inner arc line length discrepancy. The result is that a a fixed, stable central Europe, whereas palaeomag-
strip of lithosphere along the inside (hinterland) netic data imply ongoing opposing senses of
portion of a buckling beam will significantly rotation. This apparent dichotomy is, however, con-
exceed the line length of the inside arc of the oro- sistent with a model of orocline formation through
cline. A likely manifestation of this excess line buckling of a linear ribbon continent. In our
length is that lithosphere along the inside of an model, the southern Apennines and Sicily –Calabria
orocline will be thrust out over its more external move north relative to autochthonous stable Europe
foreland regions (Fig. 6), forming the structurally throughout development of the Calabrian Orocline,
highest thrust sheets. These structurally high thrust despite strongly opposing counter-clockwise and
sheets will be everywhere more rotated and clockwise rotations, respectively (Fig. 5). Hence
display a greater arc curvature than the structurally interpreting the Calabrian Orocline as the result of
lower foreland portions of the orocline (Fig. 6). buckling resolves the dichotomy of uniform conver-
Thus the observed up-structural section increase gence of the Apennines and Sicily –Calabria with
in the magnitude of rotations seen in structural stable Europe coeval with opposing senses
and palaeomagnetic data is a predicted consequence of rotation.
of thick-skinned orocline formation. Our expla- Stratigraphy and geological structures of the
nation for the up-section increase in rotation Apennine –Sicilian belt are continuous to the west
requires that this should decrease away from the into the Maghrebian chain of northern African,
hinge region of the orocline. Hence our model and there continuously on through the Tell Moun-
is testable. tains of Algeria, into the Rif Mountains of
Fig. 7. Palaeogeographic evolution of the western Mediterranean from 10 Ma (a) to present (b). The Western
Mediterranean ribbon continent is stippled and includes earlier emplaced nappes of Variscan affinity (wave stipple)
that structurally overly the Internal Platform (spot stipple), and can be traced westward from the Apennines (A) through
the Maghrebides (M) of Sicily, the Tell (T), the Rif (R) and the Betic (B) mountains. (a) With Africa restored .100 km
to the south, the ribbon continent is characterized by a convex to the south orocline whose apex is located in the
Tell region. Dashed lines indicate approximate boundaries of continental lithosphere of the Apulia–Hyblean foreland.
(b) Northward displacement of Africa flattens the Tell orocline, tightens the Betic-Rif orocline, and causes escape
of the ribbon continent to the east, giving rise to the Calabrian orocline.
Morocco and around a bend, the Gibraltar Orocline, continent. Our model predicts that eastward buck-
similar in scale to the Calabrian orocline into the ling of the Apennine –Maghrebide peninsula
Betic mountains of southern Spain (Fig. 7) occurred as a result of the northward encroachment
(Faccenna et al. 2004; Rosenbaum et al. 2002). of Africa on, and subsequent straightening of, the
Hence the Apennine –Sicilian belt is part of a originally convex to the south Tell portion of the
much longer Western Mediterranean ribbon conti- western Mediterranean ribbon continent. Conse-
nent that continues SW into northern Africa quently we are faced with the rather surprising con-
(Fig. 7). Structural continuity along the length of clusion that the Calabrian Orocline owes its
this ribbon continent provides a possible mechanism existence to the unbending of a Tellian orocline.
for producing the rapid eastward translation of the
buckling Apennine –Maghrebian portion of the con- The University of Lausanne and the Geological Museum
tinental ribbon. We postulate that prior to 10 Ma the of Lausanne generously made available office space and
ribbon continent was characterized by a convex to research facilities to STJ. G. Borel and G. Stampfli are
thanked for sharing their expertise on Tethyan geology.
the south geometry with one or more bends in the
This research was funded by an NSERC Discovery
Tell portion of the ribbon continent (Fig. 7). If the Grant, and by a research fellowship from the University
ribbon continent was pinned or fixed in the Gibraltar of Lausanne to STJ. R. Van der Voo, R. Psyklywec and
region to the west, and in the southern Alps to the G. Borel are thanked for their in depth reviews that signifi-
east, then slow northward motion of Africa would cantly improved our manuscript. B. Murphy’s editorial
result in rapid easterly-directed escape of the efforts and encouragement were most appreciated. This
ribbon continent, as proposed by Mantovani et al. is a contribution to IGCP Project 453 Ancient Orogens
(1996) and could explain the brisk rate of opening and Modern Analogues.
of the Tyrrhenian Sea behind the buckling ribbon
continent. A model of easterly escape of the
ribbon continent therefore predicts that the Tell References
portion of the ribbon was formerly characterized
by a far more arcuate geometry that has been pro- A NDERSON , H. & J ACKSON , J. 1987. Active tectonics of
the Adriatic region. Royal Astronomical Society
gressively flattened by the northward motion of Geophysical Journal, 91, 937 –983.
Africa over the past 10 Ma (Fig. 7). In addition, A RGNANI , A. & S AVELLI , C. 1999. Cenozoic volcanism
our model suggests that the Tell thrust front is and tectonics in the southern Tyrrhenian sea: space-
largely a dextral strike-slip boundary. Both these time distribution and geodynamic significance.
predictions, unbending of the Tell, and largely Journal of Geodynamics, 27, 409 –432.
strike-slip motion between the ribbon continent B ARCHI , R., M INELLI , G. & P IALLI , G. 1998. The CROP
and Africa, are testable through palaeomagnetic 03 profile: a synthesis of results on deep structures of
and structural studies. the Northern Apennines. Memorie della Società Geo-
logica Italiana, 52.
B ESSE , J., P OZZI , J.-P., M ASCLE , G. & F EINBERG , H.
1984. Palaeomagnetic study of Sicily: consequences
Conclusions for the deformation of Italian and African margins
over the last 100 million years. Earth and Planetary
Opening of the Tyrrhenian Sea was a direct result of Science Letters, 67, 377– 390.
formation of the Calabrian Orocline. Buckling of B UTLER , R. W. H., M AZZOLI , S. ET AL . 2004. Applying
the Apennine– Maghrebide peninsula was facili- thick-skinned tectonic models to the Apennine thrust
tated by over-riding and subducting lithosphere to belt of Italy – Limitations and implications. In:
the east, and the opening of a basin, now the Tyrrhe- M C C LAY , K. R. (ed.) Thrust Tectonics and Hydro-
carbon Systems. Volume Memoir 82. American
nian Sea, to the west. Hence the Calabrian Orocline
Association of Petroleum Geology, 647– 667.
is a thick-skinned tectonic feature attributable to an C AREY , S. W. 1955. The Orocline concept in
Oroclinal Orogeny that involved buckling of a for- geotectonics – Part 1. Proceedings of the Royal
merly linear ribbon continent. A model of Oroclinal Society of Tasmania, 89, 255–288.
Orogeny for the origin of the Calabrian Orocline C ATALANO , R., C HANNELL , J. E. T., D’A RGENIO , B. &
explains: (1) the rapid opening of the Tyrrhenian N APOLEONE , G. 1976. Mesozoic palaeogeography of
Sea; (2) the cessation of sea-floor spreading within the southern Apennines and Sicily. Memoir Society
the basin at 2 Ma; (3) the subsequent and ongoing Geological Italy, 15, 95–118.
reduction in the size of the basin by southward sub- C ATALANO , R., D ARGENIO , B., G REGOR , C. B., N AIRN ,
duction of Tyrrhenian Sea lithosphere beneath A. E. M. & N ARDI , G. 1984. The Mesozoic Volcanics
of Western Sicily. Geologische Rundschau, 73,
Sicily opening; and (4) the northward displacement 577– 598.
of Sicily– Calabria and the Apennines, despite their C AVAZZA , W., R OURE , F. & Z IEGLER , P. A. 2004. The
respective opposed clockwise and counterclockwise Mediterranean area and the surrounding regions;
rotations. The Apennine –Maghrebide peninsula is active processes, remnants of former Tethyan oceans
part of a larger Western Mediterranean ribbon and related thrust belts. In: C AVAZZA , W., R OURE ,
F., S PAKMAN , W., S TAMPFLI , G. M. & Z IEGLER , Study of the Hyblean Volcanic-Rocks, Sicily. Tectono-
P. A. (eds) The TRANSMED Atlas; the Mediterranean physics, 98, 271 –295.
Region from Crust to Mantle; Geological and Geophy- G UARNIERI , P. 2004. Structural evidence for deformation
sical Framework of the Mediterranean and the by block rotation in the context of transpressive tec-
Surrounding Areas. Berlin, Springer-Verlag, 1– 29. tonics, northwestern Sicily (Italy). Journal of Struc-
C ELLO , G. & M AZZOLI , S. 1999. Apennine tectonics in tural Geology, 26, 207– 219.
southern Italy: a review. Journal of Geodynamics, 27, H IPPOLYTE , J. C., A NGELIER , J. & R OURE , F. 1994. A
191– 211. major geodynamic change revealed by Quaternary
C ELLO , G., G UERRA , I., T ORTORICI , L., T URCO , E. & stress patterns in the southern Apennines (Italy). Tecto-
S CARPA , R. 1982. Geometry of the neotectonic nophysics, 230, 199 –210.
stress-field in southern Italy – geological and seismo- I NCORONATO , A. & N ARDI , G. 1989. Palaeomagnetic evi-
logical evidence. Journal of Structural Geology, 4, dences for a peri-Tyrrhenian orocline. In: B ORIANI ,
385– 393. A., B ONAFEDE , M. & V AI , G. B. (eds) The Litho-
C HANNELL , J. E. T. & H ORVATH , F. 1976. The sphere in Italy. Advances in Earth Science Research
African-Adriatic promontory as a paleogeographical Roma, Accademia Nazionale dei Lincei, 217– 227.
premise for Alpine orogeny and plate movements in J IMENEZ -M UNT , I., S ABADINI , R., G ARDA , A. &
the Carpatho-Balkan region. Tectonophysics, 35, B IANCO , G. 2003. Active deformation in the Mediter-
71–110. ranean from Gibraltar to Anatolia inforred from
C HANNELL , J. E. T., D’A RGENIO , B. & H ORVATH , F. numerical modeling and geodetic and seismological
1979. Adria, the African promontory in Mesozoic Med- data. Journal of Geophysical Research, 108, doi:
iterranean palaeogeography. Earth-Science Reviews, 10: 1029/2001IB001544.
15, 213– 292. K ASTENS , K. ET AL . 1988. ODP Leg 107 in the Tyrrhenian
C HANNELL , J. E. T., C ATALANO , R. & D ARGENIO , B. Sea: insights into passive margin and back-arc basin
1980. Paleomagnetism and deformation of the Meso- evolution. Geological Society of America Bulletin,
zoic Continental-Margin in Sicily. Tectonophysics, 100, 1140–1156.
61, 391–407. L UCENTE , F. P., C HIARABBA , C., C IMINI , G. B. &
D ERCOURT , J., Z ONENSHAIN , L. P. ET AL . 1986. Geologi- G IARDINI , D. 1999. Tomographic constraints on the
cal evolution of the Tethys belt from the Atlantic to the geodynamic evolution of the Italian region. Journal
Pamirs since the Liassic. Tectonophysics, 123, of Geophysical Research-Solid Earth, 104,
241– 315. 20 307–20 327.
D I B UCCI , D. & M AZZOLI , S. 2002. Active tectonics of M ANTOVANI , E., A LBARELLO , D., T AMBURELLI , C. &
the Northern Apennines and Adria geodynamics: new B ABBUCCI , D. 1996. Evolution of the tyrrhenian
data and a discussion. Journal of Geodynamics, 34, basin and surrounding regions as a result of the
687– 707. Africa-Eurasia convergence. Journal of Geodynamics,
D I B UCCI , D., M AZZOLI , S., N ESCI , O., S AVELLI , D., 21, 35–72.
T RAMONTANA , M., D E D ONATIS , M. & B ORRAC- M ARANI , M. P. & G AMBERI , F. 2004. Distribution and
CINI , F. 2003. Active deformation in the frontal part nature of submarine volcanic landforms in the Tyrrhe-
of the Northern Apennines: insights from the lower nian Sea: the arc vs the backarc. Memoir Carta Geolo-
Metauro River basin area (northern Marche, Italy) gica D’Italia, 64, 109–126.
and adjacent Adriatic off shore. Journal of Geody- M ATTEI , M., C IFELLI , F. & D’A GOSTINO , N. 2007. The
namics, 36, 213–238. evolution of the Calabrian Arc: evidence from palaeo-
E LDREDGE , S., B ACHTADSE , V. & V AN DER V OO , R. magnetic and GPS observations. Earth & Planetary
1985. Paleomagnetism and the orocline hypothesis. Science Letters, 263, 259– 274.
Tectonophysics, 119, 153–179. M AZZOLI , S., B ARKHAM , S., C ELLO , G., G AMBINI , R.,
E LTER , P., G RASSO , M., P AROTTO , M. & V EZZANI , L. M ATTIONI , L., S HINER , P. & T ONDI , E. 2001. Recon-
2003. Structural setting of the Apennine-Maghrebian struction of continental margin architecture deformed
thrust belt. Episodes, 26, 205– 211. by the contraction of the Lagonegro Basin, southern
F ACCENNA , C., P IROMALLO , C., C RESPO -B LANC , A., Apennines, Italy. Journal of the Geological Society,
J OLIVET , L. & R OSSETTI , F. 2004. Lateral slab defor- London, 158, 309 –319.
mation and the origin of the western Mediterranean M C K ENZIE , D. 1972. Active tectonics of the Mediterra-
arcs. Tectonics, 23, TC1012. nean region. Royal Astronomical Society Geophysical
G ATTACCECA , J. & S PERANZA , F. 2002. Paleomagnetism Journal, 30, 109–185.
of Jurassic to Miocene sediments from the Apenninic M ONGELLI , F., Z ITO , G., D E L ORENZO , S. & D OGLIONI ,
carbonate platform (southern Apennines, Italy): evi- C. 2004. Geodynamic interpretation of the heat flow in
dence for a 608 counterclockwise Miocene rotation. the Tyrrhenian Sea. Carta Geologica D’Italia, 64,
Earth and Planetary Science Letters, 201, 19–34. 71–82.
G IUNTA , G., N IGRO , F. ET AL . 2003. Il terremoto di M ONTONE , P., A MATO , A. & P ONDRELLI , S. 1999.
Palermo (6 settembre 2002) nel quadro della tettonica Active stress map of Italy. Journal of Geophysical
attiva della Sicilia nord occidentale e del Tirreno Research-Solid Earth, 104(B11), 25 595– 25 610.
meridionale. Studi Geologici Camerti, Nuova Serie, N ERI , G., C ACCAMO , D., C OCINA , O. & M ONTALTO , A.
1/2003, 57–68. 1996. Geodynaimc implications of earthquake data in
G RASSO , M., L ENTINI , F., N AIRN , A. E. M. & V IGLI- the southern Tyrrhenian sea. Tectonophysics, 258,
OTTI , L. 1983. A Geological and Palaeomagnetic 223–249.
N OCQUET , J.-M. & C ALAIS , E. 2003. Crustal velocity AMS evidence for passive clockwise rotations.
field of western Europe from permanent GPS array sol- Journal of Geodynamics, 41, 422 –439.
utions, 1996–2001. Geophysical Journal Inter- S PAKMAN , W. & W ORTEL , R. 2004. A tomographic
national, 154, 72–88. view on western Mediterranean geodynamics. In:
O LDOW , J. S., F ERRANTI , L. ET AL . 2002. Active fragmen- C AVAZZA , W., R OURE , F., S PAKMAN , W., S TAMP-
tation of Adria, the north African promontory, central FLI , G. M. & Z IEGLER , P. A. (eds) The Transmed
Mediterranean orogen. Geology, 30, 779– 782. Atlas: The Mediterranean Region from Crust to
P ATACCA , E. & S CANDONE , P. 2001. Late thrust Mantle. Berlin, Springer, 31–52.
propagation and sedimentary response in the S PERANZA , F., M ATTEI , M., N ASO , G., D I B UCCI , D. &
thrust-belt-foredeep system of the Southern Apennines C ORRADO , S. 1998. Neogene–Quaternary evolution
(Pliocene-Pleistocene). In: V AI , G. B. & M ARTINI , of the central Apennine orogenic system (Italy): a
I. P. (eds) Anatomy of an Orogen: the Apennines and structural and palaeomagnetic approach in the Molise
Adjacent Mediterranean Basins. London, Kluwer region. Tectonophysics, 299, 143 –157.
Academic Publishers, 401–440. S PERANZA , F., M ANISCALCO , R., M ATTEI , M.,
R OSENBAUM , G. & L ISTER , G. S. 2004. Neogene and S TEFANO , D. & B UTLER , R. W. H. 1999. Timing and
Quaternary rollback evolution of the Tyrrhenian Sea, magnitude of rotations in the frontal thrust systems of
the Apennines, and the Sicilian Maghrebides. southwestern Sicily. Tectonics, 18, 1178–1197.
Tectonics, 23. S PERANZA , F., M ANISCALCO , R. & G RASSO , M. 2003.
R OSENBAUM , G., L ISTER , G. S. & D UBOZ , C. 2002. Pattern of orogenic rotations in central-eastern Sicily:
Reconstruction of the tectonic evolution of the implications for the timing of spreading in the
western mediterranean since the Oligocene. Journal Tyrrhenian Sea. Journal of Geological Society,
of the Virtual Explorer, 8, 105–124. London, 160, 183–195.
S ARTORI , R. 2003. The Tyrrhenian back-arc basin and S TAMPFLI , G. M. 2000. Tethyan Oceans. In: B OZKURT ,
subduction of the Ionian lithosphere. Episodes, 26, E., W INCHESTER , J. A. & P IPER , J. D. A. (eds)
217–221. Tectonics and magmatism in Turkey and the surround-
S AVELLI , C. 2002. Time-space distribution of magmatic ing area. Geological Society, London, Special Publi-
activity in the western mediterranean and peripheral cations, 173, 1 –23.
orogens during the past 30 Ma (a stimulus to geody- S TAMPFLI , G. M. & B OREL , G. D. 2002. A plate tectonic
namic considerations). Journal of Geodynamics, 34, model for the Paleozoic and Mesozoic constrained by
99– 126. dynamic plate boundaries and restored synthetic
S CHEEPERS , P. J. J. 1992. No tectonic rotation for the oceanic isochrons. Earth and Planetary Science
Apulia-Gargano foreland in the Pleistocene. Geophysi- Letters, 196, 17– 33.
cal Research Letters, 19, 2275–2278. S TAMPFLI , G. M. & B OREL , G. D. 2004. The
S CHEEPERS , P. J. J. & L ANGEREIS , C. G. 1994. Palaeo- TRANSMED Transects in Space and Time: constraints
magnetic evidence for counter-clockwise rotations in on the paleotectonic evolution of the Mediterranean
the southern Apennines fold-and-thrust belt during domain. In: C AVAZZA , W., R OURE , F. M.,
the Late Pliocene and middle Pleistocene. Tectonophy- S PAKMAN , W., S TAMPFLI , G. M. & Z IEGLER , P. A.
sics, 239, 43–59. (eds) The TRANSMED Atlas: The Mediterranean
S HINER , P., B ECCACINI , A. & M AZZOLI , S. 2004. Region from Crust to Mantle. Berlin, Springer, 53–80.
Thin-skinned versus thick-skinned structural models T APPONNIER , P. 1977. Evolution tectonique du système
for Apulian carbonate reservoirs: constraints from alpin en Méditerranée: poinçonnement et écrasement
the Val d’Agri Fields, S Apennines, Italy. Marine rigide-plastique. Bulletin Geological Society France,
and Petroleum Geology, 21, 805–827, doi: 10.1016/ 7, 437– 460.
jmarpetgeo.2003.11.020. W ORTEL , J. R. & S PAKMAN , W. 2000. Subduction and
S OMMA , R. 2006. The south-western side of the Calabrian slab detachment in the Mediterranean – Carpathian
Arc (Peloritani Mountains): geological, structural and Region. Science, 290, 1910– 1917.
Seismic structure, crustal architecture and tectonic evolution of the
Anatolian –African Plate Boundary and the Cenozoic Orogenic
Belts in the Eastern Mediterranean Region
YILDIRIM DILEK1* & ERIC SANDVOL2
1
Department of Geology, Miami University, Oxford, OH 45056, USA
2
Department of Geology Sciences, University of Missouri-Columbia,
Columbia, MO 65211, USA
*Corresponding author (e-mail: dileky@muohio.edu)
Abstract: The modern Anatolian–African plate boundary is represented by a north-dipping

subduction zone that has been part of a broad domain of regional convergence between Eurasia
and Afro– Arabia since the latest Mesozoic. A series of collisions between Gondwana-derived
ribbon continents and trench-roll-back systems in the Tethyan realm produced nearly East –
West-trending, subparallel mountain belts with high elevation and thick orogenic crust in this
region. Ophiolite emplacement, terrane stacking, high-P and Barrovian metamorphism, and
crustal thickening occurred during the accretion of these microcontinents into the upper plates
of Tethyan subduction roll-back systems during the Late Cretaceous– Early Eocene. Continued
convergence and oceanic lithospheric subduction within the Tethyan realm were punctuated by
slab breakoff events following the microcontinental accretion episodes. Slab breakoff resulted in
asthenospheric upwelling and partial melting, which facilitated post-collisional magmatism
along and across the suture zones. Resumed subduction and slab roll-back-induced upper plate
extension triggered a tectonic collapse of the thermally weakened orogenic crust in Anatolia in
the late Oligocene–Miocene. This extensional phase resulted in exhumation of high-P rocks and
medium- to lower-crustal material leading to the formation of metamorphic core complexes in
the hinterland of the young collision zones. The geochemical character of the attendant magmatism
has progressed from initial shoshonitic and high-K calc-alkaline to calc-alkaline and alkaline
affinities through time, as more asthenosphere-derived melts found their way to the surface with
insignificant degrees of crustal contamination. The occurrence of discrete high-velocity bodies in
the mantle beneath Anatolia, as deduced from lithospheric seismic velocity data, supports our
Tethyan slab breakoff interpretations. Pn velocity and Sn attenuation tomography models indicate
that the uppermost mantle is anomalously hot and thin, consistent with the existence of a shallow
asthenosphere beneath the collapsing Anatolian orogenic belts and widespread volcanism in
this region. The sharp, north-pointing cusp (Isparta Angle) between the Hellenic and Cyprus
trenches along the modern Anatolian–African plate boundary corresponds to a subduction-
transform edge propagator (STEP) fault, which is an artifact of a slab tear within the downgoing
African lithosphere.
Introduction suture zone has been estimated to be c. 16 mm a21

based on global positioning system measurements
The present-day geodynamics of the eastern of present-day central movements in this collision
Mediterranean region is controlled by the relative zone (Reilinger et al. 1997, 2006). These differential
motions of three major plates, Eurasia, Africa and northward motions of Africa (,10 mm a21) and
Arabia, and much of the resulting deformation Arabia (16 mm a21) with respect to Eurasia are
occurs at their boundaries (Fig. 1; Westaway 1994; accommodated along the sinistral Dead Sea fault
Jolivet & Faccenna 2000; McClusky et al. 2000; zone (Fig. 1b). The Anatolian microplate north of
Doglioni et al. 2002; Dilek 2006; Reilinger et al. these convergent plate boundaries is moving SW
2006). The convergence rate between Africa and with respect to Eurasia (Fig. 1a) at c. 30 mm a21
Eurasia is .40 mm a21 across the Hellenic Trench along the North and East Anatolian fault zones
but decreases to ,10 mm a21 across the Cyprus (Reilinger et al. 1997) and is undergoing complex
trench to the east (McClusky et al. 2000; Doglioni internal deformation via mainly strike-slip and
et al. 2002; Wdowinski et al. 2006) as a result of normal faulting. This deformation has resulted in
the subduction of the Eratosthenes seamount the extensional collapse of the young orogenic
beneath Cyprus (Robertson 1998). The Arabia– crust, which developed during a series of collisional
Eurasia convergence across the Bitlis– Zagros events in the region (Dewey et al. 1986; Dilek &
128 Y. DILEK & E. SANDVOL
Fig. 1. (a) GPS velocity vectors for Anatolia and the Aegean Sea plotted in a Eurasia fixed reference frame (modified
from Reilinger et al. 1997). The major mapped faults for western Anatolia are shown as lines and circles. (b) Tectonic
map of the Aegean and eastern Mediterranean region showing the main plate boundaries, major suture zones, fault
systems and tectonic units. Thick, white arrows depict the direction and magnitude (mm a21) of plate convergence; grey
arrows mark the direction of extension (Miocene– Recent). Orange and purple delineate Eurasian and African plate
affinities, respectively. Key to lettering: BF, Burdur fault; CACC, Central Anatolian Crystalline Complex; DKF,
Datça– Kale fault (part of the SW Anatolian Shear Zone); EAFZ, East Anatolian fault zone; EF, Ecemis fault; EKP,
Erzurum–Kars Plateau; IASZ, Izmir– Ankara suture zone; IPS, Intra– Pontide suture zone; ITS, Inner–Tauride suture;
KF, Kefalonia fault; KOTJ, Karliova triple junction; MM, Menderes massif; MS, Marmara Sea; MTR, Maras triple
junction; NAFZ, North Anatolian fault zone; OF, Ovacik fault; PSF, Pampak– Sevan fault; TF, Tutak fault; TGF,
Tuzgölü fault; TIP, Turkish–Iranian plateau (modified from Dilek 2006).
Moores 1990; Yilmaz 1990), giving way to the events prior to the onset of continental collision.
formation of metamorphic core complexes and The 1000 km-long convergent plate boundary
intracontinental basins (Bozkurt & Park 1994; between these two plates comprises two separate
Dilek & Whitney 2000; Jolivet & Faccenna arcs: the Hellenic and the Cyprean (Fig. 1). The
2000; Okay & Satir 2000; Doglioni et al. 2002; intersection of these two subduction zones occurs
Ring & Layer 2003). Extensional deformation in a sharp bend, the Isparta Angle (IA), in the
of the young orogenic belts has been accompanied Tauride block. The Hellenic arc is characterized
by magmatism with varying geochemical finger- by a relatively steep, retreating subduction, where-
prints. The cause-effect relations of the spatial as the Cyprean arc appears to involve a shallow
and temporal interplay between post-collisional subduction with two major seamounts (the Era-
extension and magmatism in the eastern Mediter- tosthenes and Anixamander) impinging on the
ranean region have been a subject of intense scrutiny trench (Kempler & Ben-Avraham 1987; Zitter
and interdisciplinary research over the last twenty et al. 2003).
years. In this paper we examine the Cenozoic evolution
The modern collision zone between the of the African– Anatolian plate boundary utilizing
Anatolian and African plates is an excellent natu- seismic tomography, and use our observations
ral laboratory to study the last stages of subduction, and interpretations from this modern subduction-
subduction roll-back processes, and accretionary collision driven plate boundary to derive
CENOZOIC OROGENIC BELTS IN THE EASTERN MEDITERRANEAN REGION 129
Fig. 1. (Continued)
conclusions for the geodynamic evolution and Pn tomography, however, offers an important
mantle dynamics of the young (Cenozoic) orogenic snapshot of the lithospheric mantle seismic velocity
belts in Anatolia. Our tomographic models of the structure, even in regions with sparse station cover-
upper mantle beneath Anatolia and the Aegean age. Similarly, seismic phase attenuation can also
Sea support the existence of two discrete high- indicate the state of the lithospheric mantle. Speci-
velocity regions. These may be indicative of two fically, the regional seismic phase Sn is very sensi-
separate slab breakoff events that have occurred tive to lithospheric mantle temperature anomalies
during the subduction of the Tethyan oceanic litho- (e.g. Molnar & Oliver 1969; Rodgers et al. 1995;
sphere. We use our own geological observations and Sandvol et al. 2001). Sn attenuation and Pn velocity
data, and the extant literature to re-interpret the are, therefore, two complementary and independent
collision-driven tectonic evolution of the western, measures of the state of the uppermost mantle.
central and eastern Anatolian orogenic belts and to Using these two models we find that the majority
provide an overview of their crustal architecture. of the Anatolian plate is underlain by highly attenu-
ating and seismically slow material (Figs 2 & 3).
Seismic structure of the collision zones in The two images from these independent datasets
confirm that the Anatolian lithosphere is hot and
Anatolia and its environs most probably relatively thin.
Lithospheric seismic velocity structure Pn velocity tomography with an anisotropy
component shows two scales of low Pn velocity
Seismic images of the mantle provide important anomalies (Fig. 2). First, a broader scale
information on the current state of the lithosphere (c. 500 km) low (,7.8 km/s) Pn velocity anomaly
and asthenosphere in the eastern Mediterranean underlies northwestern Iran, eastern Anatolia, the
region. In particular, the location of possible pre- Caucasus, most of the Anatolian plate, and the
served slabs beneath the Anatolian plate could northern Aegean Sea. These broad-scale low Pn
prove useful in understanding the evolution of the velocity anomalies occupy regions within the
African –Anatolian plate boundary and the geody- Eurasian side of the Eurasia– Arabia collision zone.
namics of the Cenozoic orogenic belts in the In central Iran, fast Pn velocities appear to extend
region. In many models, however, the uppermost beyond the Zagros suture line, while in northwestern
mantle picture is not well resolved in regions Iran and eastern Turkey the high Pn velocities
lacking fairly dense station coverage. This is cer- are limited to the region immediately south of the
tainly true for much of western and central Anatolia. Bitlis –Zagros suture zone and the Tauride block.
Fig. 2. Pn travel time tomography using all available phase data in the Middle East (Al-Lazki et al. 2004). The green
line is the approximate boundary of a zone of Sn blockage (Fig. 3). Note the anomalously low Pn velocities along
southern Turkey especially in the vicinity of the Isparta Angle (IA).
The Sn attenuation tomography model reveals and Sn blockage. Throughout the entire Anatolian
regions of blocked and attenuated Sn (Fig. 3; Gök plateau and northern Arabia, young basaltic volcan-
et al. 2000, 2003). In the southern Aegean Sea, a ism correlates well with regions of high Sn attenu-
volcanic arc is present north of Crete and parallel ation (Fig. 3).
to the Hellenic arc and Sn is attenuated. In the north- In the easternmost portion of the model, ineffi-
ern Aegean Sea, there is a high attenuating Sn zone cient Sn is observed in the Greater Caucasus.
with low upper mantle velocities (Al-Lazki et al. Throughout most of the Lesser Caucasus, Sn is
2003) and Sn is partially attenuated or inefficient highly attenuated. Sn is observed as efficient in the
in northern Greece. Conversely, there is efficient western section of the Pontides and inefficient in
Sn throughout most of Greece. In the Black Sea the eastern Pontides. In western Anatolia, there are
there is a very abrupt transition from efficient to some inefficient Sn regions, unlike the regions of
blocked Sn regions. Efficient Sn is observed for Sn blockage observed in eastern Anatolia.
paths within the eastern Arabian plate and the The results of Sn attenuation tomography are
Zagros fold-and-thrust belt. An efficient zone is consistent with those of Pn tomography, both indi-
also seen at the northern part of the fold-and-thrust cating that the uppermost mantle beneath much of
belt in the northwestern corner of the Iranian Anatolia is anomalously hot and thin. In western
plateau. However, this region might be distorted Anatolia, we observe complicated patterns of Sn
because of smearing (Fig. 2). The Pn velocity tom- attenuation, except for the Aegean Sea volcanic arc.
ography of Al-Lazki et al. (2003) shows higher Pn A region of clear and distinctly higher Sn atte-
velocities for a smaller portion of the same region. nuation is observed north of the Hellenic trench that
Clearly, Sn is blocked throughout the Anatolian is an area of active arc-backarc extension above the
plateau, whereas it propagates efficiently within northward subducting African plate. Numerous
the eastern part of the Bitlis –Zagros suture zone. studies have found low velocities in the upper
Sn also propagates efficiently through the Black mantle that are consistent with the Sn efficiency
Sea and the Mediterranean Sea. In general there is results (e.g. Spakman et al. 1993; Alessandrini
a strong correlation between Quaternary volcanism et al. 1997; Piromallo & Morelli 2003). The
Fig. 3. Tomographic map of Sn attenuation within the Anatolian plate and surrounding regions. Green diamonds
indicate Holocene volcanoes. Note good correlation between the Pn tomographic image (Fig. 2) and Sn attenuation,
both indicating that the lithosphere is thin and hot throughout most of the Anatolian plate (Gök et al. 2003).
partial attenuation of Sn in the northern part of the extends to depths that exceed the intermediate
Aegean Sea is related to active extensional defor- depth seismicity in the region, a result that concurs
mation in the backarc setting. The continental with the previous images of deep subduction in this
shortening in NW Greece and Albania does not region (e.g. Spakman et al. 1988, 1993; Wortel &
allow the rotation of the western margin of the Spakman 1992). In contrast, the image of the
region and leads to east –west shortening and Cyprean slab velocity anomaly is much weaker in
north–south extension as the southern Aegean most tomographic models (e.g. Widiyantoro et al.
margin moves towards the Hellenic trench. The effi- 2004; Fig. 5). This largely aseismic structure is
cient Sn zone in part of Greece might be related to located near the Vrancea seismic zone and may be
the subducting slab. related to an earlier episode in the multi-stage
closing history of the Tethyan seaways (Dilek
Upper mantle and transition zone et al. 1999a, 2008). This structure is, however, less
obvious in the Kárason & Van der Hilst (2001)
There have been many tomographic P-wave, model, and further investigation is necessary.
S-wave and bulk sound velocity models constructed We can use the upper mantle velocity structure to
for the upper mantle of the eastern Mediterranean infer the number of slab breakoff events that have
region (Spakman et al. 1993; Bijwaard & occurred in the Neogene. Beneath the Eastern
Spakman 2000; Kárason & Van Der Hilst et al. Mediterranean, Aegean Sea and Anatolian plate,
2001; Wortel & Spakman 1992, 2000; Widiyantoro the seismic velocity structure of the mantle transition
et al. 2004). Features common to many of these zone is characterized by a number of discrete high-
models support their reliability. For example, along velocity bodies. Current global and regional mod-
the Hellenic trench, the subducted slab of largely els for the Eastern Mediterranean clearly show the
oceanic lithosphere of the African plate stands out continuation of subducting ocean lithosphere into
as a relatively steeply dipping structure with a the lower mantle along the Hellenic trench and its
strong high-velocity P-wave (Fig. 4). This anomaly flattening out at the 660 km discontinuity. Many
Fig. 4. Seismic tomography cross-section of the mantle to a depth of 800 km in the Aegean region mantle taken
perpendicular to the Hellenic trench. Colours display the percentage deviation of seismic wave speed. Negative
(positive) anomalies likely represent predominantly higher (lower) temperatures than average. Dashed lateral lines in
the section depict the mantle discontinuities at 410 km and 660 km depths. See text for discussion.
(not all) of the newer tomographic models, however, to the North Anatolian Fault Zone (Fig. 5). It still
show that there are two high-velocity bodies in the remains unclear as to what role the subducting
mantle transition zone beneath the Aegean Sea Cyprean lithosphere had in the very recent uplift
(e.g. Kárason 2002; Widiyantoro et al. 2004; Fig. 4). of the high topography in the Cyprean back arc
Pn and Sn tomographic models provide no evi- (i.e. the Tauride block).
dence for an attached slab along any of the suture
zones within the Anatolian plate. Nor do they The Isparta Angle
provide evidence of a very shallow/flat slab under-
lying southern Anatolia, as suggested by some The Isparta Angle (IA) occurs at the intersection
researchers (i.e. Doglioni et al. 2002; Agostini between the Cyprean (east) and Hellenic (west)
et al. 2007). Many teleseismic velocity models arcs in the western end of the Tauride block
show a relatively shallow high-velocity anomaly (Fig. 1b). Palaeomagnetic investigations suggest
underlying much of central and, in some cases, that there has been very little rotation of the IA in
northern Anatolia. This is in contrast to the upper- the last 10 my (Tatar et al. 2002). Seismogenic
most mantle images that suggest that the Anatolian deformation along the Hellenic arc indicates a tran-
lithosphere is hot and possibly thin. This might sition from compressional to normal faulting near
suggest that the flat or shallow slab of the Cyprean the Isparta Angle. On the other hand, the few avail-
subduction zone may not be in direct contact with able focal mechanisms along the westernmost
the Anatolian lithosphere, and may extend almost Cyprean arc are more consistent with thrust
Fig. 5. Seismic tomography cross-section of the mantle to a depth of 800 km in the Central Anatolian region mantle
taken perpendicular to the Cyprus trench. Colours display the percentage deviation of seismic wave speed. Negative
(positive) anomalies likely represent predominantly higher (lower) temperatures than average. Dashed lateral lines in
the section depict the mantle discontinuities at 410 km and 660 km depths. See text for discussion.
faulting. The GPS velocity vectors of McClusky Lesser Antilles trench, and the northern end of the
et al. (2000) suggest that the lithosphere within South Sandwich trench. In all these cases, STEP
the IA is moving independently of the rest of the faults propagate in a direction opposite to the sub-
Anatolian plate and that it may be attached to the duction direction, and asthenospheric upwelling
African plate or to a piece of it (Barka & Reilinger occurs behind and beneath their propagating tips.
1997; McClusky et al. 2000). This upwelling induces decompressional melting
The sharp cusp between the Hellenic and Cyprus of shallow asthenosphere, leading to linearly distrib-
trenches (Fig. 1b) and the significant differences uted alkaline magmatism younging in the direction
in the convergence velocities of the African litho- of tear propagation.
sphere at these trenches (c. 40 vs ,10 mm a21 at The North–South-trending potassic and ultra-
the Hellenic and Cyprus trenches, respectively) are potassic volcanic fields stretching from the Kirka
likely to have resulted in a lithospheric tear in the and Afyon-Suhut region in the north to the Isparta –
downgoing African plate that allowed the astheno- Gölcük area in the south (Fig. 7) shows an age pro-
spheric mantle to rise beneath SW Anatolia gression from 21– 17 to 4.6–4.0 Ma (Yagmurlu
(Fig. 6; Doglioni et al. 2002; Agostini et al. 2007; et al. 1997; Alici et al. 1998; Savasçin & Oyman
Dilek & Altunkaynak 2008). This scenario is analo- 1998; Francalanci et al. 2000; Kumral et al. 2006;
gous to lithospheric tearing at subduction-transform Çoban & Flower 2006; Dilek & Altunkaynak
edge propagator (STEP) faults described by Govers 2007). This distribution of the potassic and ultra-
& Wortel (2005) from the Ionian and Calabrian arcs, potassic volcanic rocks in SW Anatolia is consistent
the New Hebrides trench, the southern edge of the with a progressive migration of their melt source
zone (Le Pichon & Angelier 1979; Jolivet 2001;

Faccenna et al. 2003; van Hinsbergen et al. 2005).
The slab retreat rate of the subducting African litho-
sphere has been larger than the absolute velocity
of the Eurasian upper plate, causing net c. North–
South extension in the Aegean region since the
early Miocene (Jolivet et al. 1994; Jolivet & Fac-
cenna 2000; Faccenna et al. 2003; Ring & Layer
2003; Dilek & Altunkaynak 2008). Since then, the
thrust front associated with this subduction zone
and its slab retreat has migrated from the Hellenic
trench (south of Crete) to the south of the Mediter-
ranean Ridge (Jolivet & Faccenna 2000; Le
Pichon et al. 2003). Backarc extension in the
Aegean region appears to have started at c. 25 Ma,
long before the onset of the Arabian collision-
driven southwestward displacement of the Anato-
lian microplate in the late Miocene (Barka &
Reilinger 1997; Jolivet & Faccenna 2000).
The continental crust making up the upper plate
of the Hellenic subduction zone south of the North
Fig. 6. A cartoon illustrating the development of a slab
Anatolian fault is composed of the Sakarya and
tear or subduction transform edge propagator (STEP)
fault beneath the Tauride block and the positions of the Anatolide –Tauride continental blocks (Fig. 7).
subducting African lithosphere beneath the Cyprean and These two microcontinents are separated by the
Hellenic arcs (after Barka & Reilinger 1997). Evidence Izmir–Ankara suture zone, which is marked by
for a southward increase in the age of volcanism is Tethyan ophiolites and associated tectonic units.
consistent with model (Yagmurlu et al. 1997; Savasçin & The basement of the Anatolide block is composed
Oyman 1998; Dilek & Altunkaynak 2008). Also, the mainly of the Menderes massif, which is intruded
presence of ignimbrites is consistent with asthenospheric and overlain by Cenozoic granitoid plutons and
upwelling leading to widespread crustal melting. extrusive rocks; Lower Miocene and younger
sedimentary rocks of a series of extensional basins
towards the south and supports a STEP model for overlie the high-grade metamorphic rocks of the
their origin (Fig. 6; Dilek & Altunkaynak 2008). Menderes massif (Bozkurt 2003; Purvis &
Asthenospheric low velocities detected through Pn Robertson 2004; Oner & Dilek 2007).
tomographic imaging in this region (Al-Lazki
et al. 2004) also support the existence of shallow Sakarya continent
asthenosphere beneath the IA at present.
The Sakarya continent consists of a Palaeozoic crys-
talline basement with its Permo –Carboniferous
Regional geology of western Anatolia sedimentary cover and Permo –Triassic ophiolitic
and rift or accretionary-type mélange units (Kara-
The Aegean province is situated in the upper plate of kaya complex) that collectively form a composite
a north-dipping subduction zone at the Hellenic continental block (Tekeli 1981; Okay et al. 1996).
trench (Fig. 1b) and is considered to have evolved The Sakarya continental rocks and the ophiolitic
as a backarc environment above this subduction units of the Izmir–Ankara suture zone (IASZ) are
Fig. 7. Geological map of western Anatolia and the eastern Aegean region, showing the distribution of suture
zones, ophiolite complexes, major Cenozoic igneous provinces discussed in this paper, and the salient fault systems.
Menderes and Kazdag (KDM) massifs represent metamorphic core complexes with exhumed lower continental crust.
Izmir–Ankara suture zone (IASZ) marks the collision front between the Sakarya continental block to the north and the
Anatolide –Tauride block to the south. The Intra–Pontide suture zone (IPSZ) marks the collision front between the
Sakarya continent to the south and the Rhodope–Pontide block to the north. The Eocene and Oligo– Miocene granitoids
(shown in red) represent the post-collisional (post-Eocene) and/or extensional magmatism in the region. Key to
lettering of these granitoids: AG, Alasehir; BG, Baklan; CGD, Çataldag; EP, Egrigöz; GBG, Göynükbelen; GYG,
Gürgenyayla; IGD, Ilica; KG, Kozak; KOP, Koyunoba; OGD, Orhaneli; TG, Turgutlu; TGD, Topuk; SG, Salihli. Much
of western Anatolia is covered by Cenozoic volcanic rocks intercalated with terrestrial deposits. Key to lettering of
major fault systems: AF, Acigöl fault; BFZ, Burdur fault zone; DF, Datça fault; IASZ, Izmir– Ankara suture zone; IPSZ,
Intra– Pontide suture zone; KF, Kale fault; NAFZ, North Anatolian fault zone; SWASZ, SW Anatolian shear zone.
Fig. 7. Continued.
intruded by a series of east –west trending Eocene Imbricate stacking of the Menderes massif beneath
and Oligo –Miocene granitoid plutons (Fig. 7; the Lycian nappes and ophiolitic thrust sheets
Bingöl et al. 1982, 1994; Delaloye & Bingöl 2000; appears to have migrated southwards throughout
Altunkaynak 2007). The Kazdag massif within the Palaeocene – Middle Eocene (Özer et al. 2001;
the western part of the Sakarya continent (KDM in Candan et al. 2005). Unroofing and exhumation of
Fig. 7) represents a metamorphic core complex, the Menderes massif may have started as early as
which is inferred to have been exhumed from a the Oligocene (25–21 Ma) based on the cooling
depth of c. 14 km along a north-dipping mylonitic ages of syn-extensional granitoid intrusions that
shear zone starting at c. 24 Ma (Okay & Satir 2000). crosscut the metamorphic rocks (Bozkurt & Satir
2000; Catlos & Çemen 2005; Ring & Collins
2005; Thomson & Ring 2006). This timing may
Izmir – Ankara Suture Zone (IASZ) and signal the onset of post-collisional tectonic exten-
Tethyan Ophiolites sion in the Aegean region.
The Izmir–Ankara Suture Zone south of the
Sakarya continent includes dismembered Tethyan Tauride Block in Western Anatolia
ophiolites, high-pressure low-temperature (HP–LT)
The Tauride Block south of the Menderes core
blueschist-bearing rocks, and flysch deposits
complex consists of Precambrian and Cambro–
mainly occurring in south-directed thrust sheets
Ordovician to lower Cretaceous carbonate rocks
(Fig. 7; Önen & Hall 1993; Okay et al. 1998;
intercalated with volcano-sedimentary and epiclas-
Sherlock et al. 1999). Late-stage diabasic dykes
tic rocks (Ricou et al. 1975; Demirtasli et al.
crosscutting the ophiolitic units in the Kütahya
1984; Özgül 1984; Gürsu et al. 2004). These rocks
area are dated at c. 92–90 Ma (40Ar/39Ar hornble-
are tectonically overlain by Tethyan ophiolites
nde ages; Önen 2003), indicating a minimum
(i.e. Lycian, Beysehir, Alihoca and Aladag ophio-
late Cretaceous igneous age for the ophiolites. The
lites) along south-directed thrust sheets (Dilek
blueschist rocks along the suture zone in the
et al. 1999a; Collins & Robertson 2003; Çelik &
Tavsanli area have yielded 40Ar/39Ar cooling
Chiaradia 2008; Elitok & Drüppel 2008). Under-
ages (phengite crystallization during exhumation)
thrusting of the Tauride carbonate platform
of 79.7 + 1.6–82.8 + 1.7 Ma (Sherlock et al.
beneath the Tethyan oceanic crust and its partial
1999), suggesting a latest Cretaceous age for the
subduction at a north-dipping subduction zone in
HP–LT metamorphism in the region. The Lycian
the Inner-Tauride Ocean resulted in HP –LT meta-
nappes and ophiolites that occur farther south in
morphism (Dilek & Whitney 1997; Okay et al.
the Tauride block (Fig. 7; Collins & Robertson
1998). Continued convergence caused crustal imbri-
2003; Ring & Layer 2003; Çelik & Chiaradia
cation and thickening within the platform and
2008) represent the tectonic outliers of the Cretac-
resulted in the development of several major over-
eous oceanic crust derived from the IASZ. The
thrusts throughout the Tauride block (Demirtasli
Lycian nappes are inferred to have once covered
et al. 1984; Dilek et al. 1999b).
the entire Anatolide belt in western Anatolia, and
to have been later removed as a result of tectonic
uplift and erosion associated with exhumation of Tectonic evolution of the western
the Menderes core complex during the late Ceno-
zoic (Ring & Layer 2003; Thomson & Ring 2006; Anatolian orogenic belt
Dilek & Altunkaynak 2007). Emplacement of the Cretaceous ophiolites onto the
Anatolide –Tauride block along the Izmir–Ankara
Menderes massif suture zone and partial subduction of the continental
edge that led to its HP–LT metamorphism
The Menderes massif represents a core complex (Sherlock et al. 1999; Okay 2002; Ring & Layer
composed of high-grade metamorphic rocks of 2003; Ring et al. 2003) in the late Cretaceous
Pan-African affinity that are intruded by syn- mark the initial stages of collision tectonics in
kinematic Miocene granitoid plutons (Hetzel & western Anatolia (Fig. 8). The terminal closure
Reischmann 1996; Bozkurt & Satir 2000; Bozkurt of the Northern Tethyan seaway resulted in the
2004; Gessner et al. 2004). Rimmelé et al. (2003) collision of the Sakarya continent with the
estimated the P–T conditions of the metamorphic Anatolide –Tauride block during the early Eocene
peak for the Menderes massif rocks at .10 kbar that, in turn, caused regional deformation, meta-
and .440 8C. The main episode of metamorphism morphism and crustal thickening. This Barrovian-
is inferred to have resulted from the burial regime type, collision-driven regional metamorphism was
associated with the emplacement of the Lycian responsible for the development of high-grade
nappes and ophiolitic thrust sheets (Yilmaz 2002). rocks in the Kazdag and Menderes metamorphic
Fig. 8. Late Mesozoic–Cenozoic geodynamic evolution of the western Anatolian orogenic belt as a result of collisional
and extensional processes in the upper plate of north-dipping subduction zone(s) within the Tethyan realm. See text
for discussion.
massifs. Resistance of the buoyant Anatolide– a natural response to the gravitational settling of
Tauride continental crust to subduction and conse- subducted lithosphere in continental collision
quent arrest of the north-dipping subduction zone zones, as a result of a decrease in the subduction
resulted in isostatic uplift of its partially subducted rate caused by the positive buoyancy of partially
passive margin, exhumation of high-P rocks, and subducted continental lithosphere (Davies & von
rapid denudation of upper crustal rocks, leading to Blackenburg 1995; von Blackenburg & Davies
widespread flysch formation during the early to 1995; Wortel & Spakman 2000; Gerya et al.
middle Eocene (Ring et al. 2003, and references 2004). Our seismic tomography model (Fig. 4)
therein). shows the existence of a second high-velocity
With continued continental collision the leading (cold) slab near the 660 km discontinuity in the
edge of the subducted Tethyan slab possibly broke lower mantle north of the Hellenic slab, which we
off from the rest of the continental lithosphere, interpret as a detached Tethyan slab dipping
resulting in the development of an asthenospheric beneath the western Anatolian orogenic belt. This
window (Fig. 8). Slab detachment and breakoff is implies a punctuated evolution of Tethyan
subduction zones in the eastern Mediterranean, its extensional collapse. The Kazdag core complex
rather than the single subduction zone hypothesized in NW Anatolia (Figs 7 & 8) began its initial exhu-
for much of the Mesozoic and the entire Cenozoic mation in the latest Oligocene–Early Miocene
in some recent geodynamic models (i.e. van (Okay & Satir 2000) and the Menderes core
Hinsbergen et al. 2005; Jolivet & Brun 2008). complex in central western Anatolia (Fig. 7) under-
As the downgoing oceanic plate breaks off, the went its exhumation in the earliest Miocene
asthenosphere rises rapidly and is juxtaposed (Isik et al. 2004; Thomson & Ring 2006; Bozkurt
against the thickened mantle lithosphere in the col- 2007). Some of the collision-generated thrust
lision zone (Fig. 8, c. 54 Ma). Conductive heating faults may have been reactivated during this time
of this overriding lithosphere results in melting of as crustal-scale low-angle detachment faults
the metasomatized and hydrated layers, producing (e.g. Simav detachment fault, SW Anatolian shear
potassic, calc-alkaline magmas. Crustal melting zone), facilitating regional extension (Thomson &
at shallow depths, induced by asthenospheric Ring 2006; Çemen et al. 2006).
upwelling, causes granitic/rhyolitic magmatism. Starting in the middle Miocene, both lithospheric
Hence, the middle to late Eocene emplacement and asthenospheric mantle melts were involved in
of widespread granitoid plutons in northwestern the evolution of bimodal volcanic rocks in western
Anatolia, mainly through the IASZ and the Anatolia with the lithospheric input diminishing
Sakarya continent (Orhaneli, Topuk, Göynükbelen, in time (Akay & Erdogan 2004; Aldanmaz et al.
etc.), is thought to be a direct result of the heat flux 2000, 2006). The timing of this magmatism
through this window and the associated thermal per- coincides with widespread lower crustal exhumation
turbation that caused melting of the metasomatized and tectonic extension across the Aegean region
continental lithospheric mantle (Dilek & Altunkay- (c. 14 Ma in Fig. 8). This extensional phase and
nak 2007). Major and trace element compositions the attendant mildly alkaline volcanism are attri-
of I-type granitoid rocks in NW Anatolia, dated buted to thermal relaxation associated with possible
at c. 54 –35 Ma (40Ar/39Ar hornblende and mica delamination of the subcontinental lithospheric
separates, and SHRIMP zircon dating; Dilek & mantle beneath the northwestern Anatolian oro-
Altunkaynak, unpublished data) indicate their genic belt (Altunkaynak & Dilek 2006; Dilek &
origin by melting of a subduction-modified mantle Altunkaynak 2007, and references therein).
source that had been enriched in mobile incompati- Lithospheric delamination may have been triggered
ble elements (Altunkaynak 2004, 2007; Köprübasi by peeling of the base of the subcontinental
& Aldanmaz 2004), a signature that is consistent lithosphere as a result of slab roll-back at the
with post-collisional magmatism driven by slab Hellenic trench.
breakoff in other orogenic belts (Schliestedt et al. During the advanced stages of extensional tec-
1987; Hansmann & Oberli 1991; Davies & von tonism in the late Miocene–Quaternary, the deve-
Blackenburg 1995; Nemcok et al. 1998). lopment of regional graben systems (i.e. Gediz,
Geological evidence in support of the inferred Büyük Menderes, Fig. 7) further attenuated the con-
slab breakoff magmatism in western Anatolia tinental lithosphere beneath the region. This exten-
includes: (1) the linear distribution of the plutons sional phase was accompanied by upwelling of the
in a narrow belt straddling the IASZ where asthenospheric mantle and consequent decompres-
ophiolitic and high-P blueschist rocks are exposed sional melting (Fig. 8). Lithospheric-scale exten-
(Figs 7 & 8) – a pattern that suggests a focused sional fault systems acted as natural conduits for
heat source likely derived from an asthenospheric the transport of uncontaminated alkaline magmas
window; and (2) the attempted subduction of to the surface (Richardson-Bunbury 1996; Seyitoglu
Anatolide– Tauride continental crust to a depth of et al. 1997; Alici et al. 2002). Asthenospheric
80 km (Okay et al. 1998), evidenced by the flow in the region following the late Miocene
latest Cretaceous blueschist rocks of the Tavsanli (,10 Ma) may also have been driven in part by
zone, that is likely to have clogged the subduction the extrusion tectonics caused by the Arabian col-
zone and caused the detachment of the sinking lision in the east, as suggested by the parallelism
Tethyan oceanic lithosphere. of the SW-oriented shear wave splitting fast polariz-
Continued collision of the Sakarya and ation direction in the mantle with the motion of
Anatolide– Tauride continental blocks led to the the Anatolian plate (Sandvol et al. 2003b; Russo
development of thick orogenic crust, orogen-wide et al. 2001). This SW-directed lower mantle flow
burial metamorphism, and anatectic melting of the beneath Anatolia could have played a significant
lower crust (c. 25 Ma, Fig. 8). This episode of post- role in triggering intra-plate deformation via exten-
collisional magmatism coincided with bimodal sion and strike-slip faulting parallel to the flow
volcanism and a widespread ignimbrite flare-up in direction and horizontal mantle thermal anomalies,
western Anatolia (Ercan et al. 1985; Yilmaz et al. which would have facilitated melting and associated
2001), and caused thermal weakening of the crust basaltic volcanism. This lateral asthenospheric
in the western Anatolian orogenic belt that led to flow may also have resulted in the interaction of
different compositional end-members, contributing units and hypabyssal rocks (Dilek & Whitney
to the mantle heterogeneity beneath western Anato- 2000, and references therein). These sedimentary
lia. Lateral displacement of the asthenosphere as units underwent regional metamorphism as a result
a result of the extrusion of collision-entrapped of burial beneath the Tethyan ophiolites derived
ductile mantle beneath Asia and SE Asia has been from the Izmir– Ankara–Erzincan suture zone to
similarly identified as the cause of post-collisional the north (Seymen 1984; Göncüoglu et al. 1991;
high-K volcanism in Tibet and Indo– China during Whitney & Dilek 1998). Subsequent collisional
the late Cenozoic (Liu et al. 2004; Williams et al. events in the latest Mesozoic and early Cenozoic
2004; Mo et al. 2006). resulted in further crustal imbrication and thicken-
The apparent SW propagation of Cenozoic tec- ing of crystalline rocks in the CACC (Dilek et al.
tonic extension and magmatism through time is 1999b). Three main massifs (Kirsehir, Akdag and
likely to have been the combined result of slab roll- Nigde; Fig. 9) have been delineated in the CACC
back associated with the subduction of the Southern based on different P–T –t paths of their protoliths.
Tethys ocean floor at the Hellenic trench and the The Kirsehir massif in the NW part of the CACC
thermally induced collapse of the western Anatolian consists of marble and calcsilicate rocks inter-
orogenic belt (Fig. 8). The thermal input and melt layered with metapelitic/psammitic schist, amphi-
sources were likely provided first by slab breakoff- bolite (+ garnet), and quartzite (Whitney et al.
generated asthenospheric flow, then by litho- 2001). The metamorphic grade appears to increase
spheric delamination-related asthenospheric flow from 450 8C (garnet zone) in the SE to c. 750 8C
and finally by tectonic extension-driven upward (sillimanite zone) in the NW within the massif.
asthenospheric flow and collision-induced (Arabia- The pressure corresponding to peak temperatures
Eurasia collision) lateral (westward) mantle flow for the garnet zone rocks is estimated to be 2.5–
(Innocenti et al. 2005; Dilek & Altunkaynak 4 kbar based on the structural position of these
2007). Since the late Miocene, subduction zone rocks (Whitney et al. 2001). However, pressures
magmatism related to the retreating Hellenic of c. 6 kbar are estimated from the first appearance
trench has been responsible for the progressive of garnet in amphibolite-facies rocks in the NW
southward migration of the South Aegean Arc part of the massif.
(Fig. 8; Pe-Piper & Piper 2006). The exhumation The Akdag massif represents the largest coher-
of high-P rocks in the Cyclades was likely driven ent block of metamorphic rocks in the NE part of
by upper plate extension and channel flow associ- the CACC (Fig. 9), metamorphic grade of which
ated with this subduction (Jolivet et al. 2003; Ring ranges from chlorite zone to sillimanite –K-feldspar
& Layer 2003). zone. The highest grade rocks occur in an elongate
NE –SW-trending belt along the central axis of the
Akdag massif (Whitney et al. 2001) containing abun-
Regional geology of central Anatolia dant garnet–muscovite–quartz gneiss (Whitney
et al. 2001). Various thermometry and barometry
Much of central Anatolia is occupied by the Central applications using different equilibrium phase
Anatolian Crystalline Complex (CACC), which assemblages in the Akdag metamorphic rocks
consists of the Kirsehir, Akdag and Nigde meta- have revealed peak pressures ranging from 5+1–
morphic massifs, dismembered Tethyan ophiolites, 8+1 kbar and peak temperatures ranging from
and felsic to mafic plutons ranging in age from the 550–600 to 660–675 8C (Whitney et al. 2001).
late Cretaceous to the Miocene (Fig. 9; Güleç The Nigde massif in the southern part of the
1994; Boztug 2000; Düzgören-Aydin 2000; Kadio- CACC (Fig. 9) is exposed in a structural dome
glu et al. 2003, 2006; Ilbeyli 2004; Ilbeyli et al. that has been interpreted as a Cordilleran-type meta-
2004). Several curvilinear sedimentary basins morphic core complex (Whitney & Dilek 1997). A
(Tuzgölü, Ulukisla and Sivas), which initially gently (c. 308) S-dipping detachment fault bounding
evolved as peripheral foreland and/or forearc the Nigde massif along its southern edge juxtaposes
basins in the late Cretaceous, delimit the CACC in multiply deformed marble, quartzite and schist in
the west, the east and the south. The south –central the footwall from clastic sedimentary rocks of the
part of the CACC includes the Cappadocian volcanic Ulukisla basin in the hanging wall. The central
province, containing Upper Miocene to Quaternary part of the Nigde massif consists predominantly of
volcanic-volcaniclastic rocks and polygenetic volca- upper amphibolite facies metasedimentary rocks
nic centers (Toprak et al. 1994; Dilek et al. 1999b). and the peraluminous Uçkapili granite. These meta-
pelitic rocks record an earlier episode of regional,
Metamorphic massifs medium-pressure/high temperature metamorphism
(5– 6 kbar and T . 700 8C), and a younger epi-
Protoliths of the high-grade rocks in the meta- sode of low-pressure/high temperature metamorph-
morphic massifs of the CACC were Palaeozoic to ism (730 –770 8C) (Whitney & Dilek 1998; Fayon
Mesozoic pelitic sediments with local mafic lava et al. 2001). The earlier event was most likely
Fig. 9. Continued.
related to burial and heating after emplacement of the northern CACC undergoing erosional unroofing
the Tethyan ophiolites southward onto the CACC nearly 20 Ma before the tectonic exhumation of the
during the late Cretaceous and during the sub- southern CACC starting in the late Miocene.
sequent collision between the CACC and Tauride
blocks, whereas the younger event was associated
with shallow intrusion of the Uçkapili granite, Inner-Tauride suture zone (ITSZ) and
which crosscuts the metamorphic and structural Tethyan ophiolites
fabrics in the Nigde massif.
The P–T –t paths of the Akdag and Kirsehir Discontinuous exposures of the Tethyan ophiolites
massifs indicate that the northern part of the CACC and mélanges define two major suture zones sur-
where they are now exposed was deformed, meta- rounding the CACC in the north and the south.
morphosed and unroofed during tectonic events The Izmir–Ankara –Erzincan suture zone to the
associated with the collision of the CACC with a north and the Inner-Tauride suture (ITSZ) zone
south-facing intraoceanic arc in the northern branch to the south (Fig. 1b) mark the obliteration sites of
of the Tethys, followed by a collision with the the Tethyan seaways, which had evolved between
Pontide arc to the north in the early to mid-Eocene. the Gondwana-derived continental fragments
The sedimentary rocks of the Akdag and Kirsehir (Sengör & Yilmaz 1981; Robertson & Dixon
massifs were buried to moderate depths (c. 20 km) 1984; Dilek & Moores 1990). Subduction of the
beneath the Tethyan ophiolite nappes derived Tethyan oceanic lithosphere that evolved in these
from the Izmir –Ankara–Erzincan suture zone and seaways resulted in the development of incipient
were thickened via folding and thrust faulting. The arc-forearc complexes that subsequently formed
existence of kyanite in the Akdag massif rocks the ophiolites, in the mantle heterogeneity beneath
suggests that the eastern part of the CACC either the continental fragments, and eventually, in
experienced greater degrees of crustal thickening continental collisions in the Eocene (Dilek &
or was exhumed from deeper crustal levels than Flower 2003).
the rest of the massifs in the CACC. The northern Ophiolite complexes within the Izmir– Ankara–
part of the CACC was slowly exhumed via erosion Erzincan suture zone include serpentinized upper
by the Eocene (Fayon et al. 2001). mantle peridotites and gabbros that are crosscut by
The SW part of the CACC experienced relatively dolerite and plagiogranite dikes and overlain by
high-temperature metamorphism associated with pillow lavas (Tankut et al. 1998; Dilek & Thy
extensive Andean-type arc magmatism at relatively 2006). Both dolerite and plagiogranite dykes show
shallow crustal levels that are represented by the negative Ta–Nb patterns typical of arc-related
CACC plutons (see below). The southern part of petrogenesis and zircon ages of c. 179 + 15 Ma,
the CACC, where the Nigde massif is now indicating that the ophiolitic basement in the
exposed, experienced Barrovian metamorphism at Izmir –Ankara–Erzincan suture zone is at least
mid-crustal pressures (c. 5 –6 kbar) and at high early Jurassic in age or older (Dilek & Thy 2006).
temperatures (.700 8C), possibly associated with The Inner-Tauride ophiolites to the south (i.e.
orogenic crustal thickening (Whitney & Dilek Alihoca, Aladag, Mersin) consist mainly of tecto-
1998). The Nigde core complex was exhumed to nized harzburgites, mafic-ultramafic cumulates and
,2 km depth mainly by tectonic unroofing along gabbros, and commonly lack sheeted dykes and
low-angle detachment faults. Apatite fission track extrusive rocks (Parlak et al. 1996, 2002; Dilek
ages from the Nigde rocks range from c. 9 –12 Ma et al. 1999a). They include thin (c. 200 m) thrust
and indicate slow to moderate cooling via exhuma- sheets of metamorphic sole rocks beneath them,
tion at rates of 30 –8 8C/Ma (Fayon et al. 2001). and both the ophiolitic units and the sole rocks are
Thus, different unroofing mechanisms appear to intruded by mafic dyke swarms composed of basal-
have affected the CACC since the early Oligocene: tic to andesitic rocks with island arc tholeiite (IAT)
Fig. 9. Geological map of the central Anatolian region, showing the distribution of Tethyan ophiolites, suture zones,
major tectonostratigraphic units within the Tauride block and the Central Anatolian Crystalline Complex (CACC)
(including metamorphic massifs and major plutons), and major faults systems. Key to lettering of major granitoid
plutons in the CACC: AG, Agaçören granitoid; BAP, Bayindir pluton; BDP, Baranadag pluton; BHP, Behrek Dag
pluton; CFP, Cefalik pluton; CP, Çelebi pluton; TP, Terlemez pluton; UKG, Uçkapili granite; YB, Yozgat batholith.
Key to lettering of major fault systems: BSZ, Bitlis suture zone; DSF, Dead Sea Fault; EAF, East Anatolian Fault; EF,
Ecemis Fault; ITSZ, Inner-Tauride Suture Zone; TGF, Tuzgölü Fault. Major ophiolites: ADO, Aladag ophiolite; AHO,
Alihoca ophiolite; BO, Beysehir ophiolite; KDO, Kizildag ophiolite; MO, Mersin ophiolite. Other symbols: ALM,
Alanya massif; MTJ, Maras Triple Junction; NM, Nigde massif; (data are from Dilek & Moores 1990; Dilek et al.
1999a, b; Kadioglu et al. 2003, 2006). Notice that the ITSZ, marked by a dark-blue, dashed thick line, is truncated and
offset by the sinistral Ecemis fault (EF).
affinities. 40Ar/39Ar hornblende ages of 92– 90 and are composed mainly of subalkaline quartz
and 90 –91 Ma from the metamorphic sole and monzonite and monzonite. Both the Granite and
dyke rocks, respectively, indicate Cenomanian– Monzonite Supersuite rocks show enrichment in
Turonian ages for the Inner-Tauride ophiolites LILE and depletion in HFSE relative to ocean
(Dilek et al. 1999a; Parlak & Delaloye 1999; ridge granites (ORG). They display isotopic and
Çelik et al. 2006). These age brackets suggest that trace element signatures suggesting a crustal com-
the Tethyan ophiolites rooted in the Inner-Tauride ponent that reflect subduction-influenced source
suture zone are possibly younger than those enrichment and crustal contamination resulting
derived from the Izmir–Ankara –Erzincan suture from assimilation fractional crystallization (AFC)
zone farther north. processes during magma transport through the
In addition to the Cenomanian –Turonian ophio- CACC crust (Bayhan 1987; Aydin et al. 1998;
lites, the ITSZ is also marked by discontinuous Güleç & Kadioglu 1998; Kadioglu et al. 2003,
exposures of blueschist-bearing mafic-ultramafic 2006). There is a progression from high-K calc-
and carbonate rocks along the northern edge of the alkaline and high-K shoshonitic compositions
Tauride block. The occurrence of sodic amphibole- in the Granite Supersuite to typical shoshonitic
containing metasedimentary and metavolcanic compositions in the Monzonite Supersuite rocks.
rocks in the Bolkar Mountains region (Blumenthal The Syenite Supersuite represents the youngest
1956; van der Kaaden 1966; Gianelli et al. 1972; phase of plutonism in the late Cretaceous
Dilek & Whitney 1997) extends into the Tavsanli (c. 69 Ma) and generally occurs in the inner part
zone in NW Anatolia and into the Pinarbasi of the CACC. Rocks of this supersuite are composed
zone in the Eastern Taurides in East –Central of silica saturated (quartz syenite and syenite) and
Anatolia. These HP/LT rock assemblages showing silica under-saturated, nepheline and pseudoleucite
counterclockwise P–T –t trajectories of their meta- bearing alkaline rocks, which show more enrich-
morphic evolution indicate increasing P/T ratio ment in LILE and a slight enrichment in HFSE in
with cooling that was associated with continuous comparison to the other two supersuites (Boztug
subduction within the Inner-Tauride Ocean et al. 1997; Kadioglu et al. 2006). Isotopic and
(Dilek & Whitney 1997, 2000). Late Cretaceous – trace element signatures of the Syenite Supersuite
Palaeocene calc-alkaline plutonic rocks (Dilek, plutons suggest that their magmas were more
unpublished data) intruding into the Ulukisla basin enriched in within-plate mantle components com-
strata north of the Bolkar Mountains point out that pared to the Granite and Monzonite Supersuite
this subduction activity had continued into the plutons (Kadioglu et al. 2006). 40Ar/39Ar age data
early Cenozoic. from these Granite, Monzonite and Syenite Super-
suite plutons yield ages of 77.7 + 0.3, 70 + 1.0,
Syn-collisional CACC plutons and 69.8 + 0.3 Ma, respectively (Kadioglu et al.
2006). Thus, the subduction zone influence on
The late Cretaceous plutonic rocks in the CACC melt evolution beneath the CACC appears to
were emplaced after obduction of the Cretaceous have decreased rapidly (within c. 7 –8 Ma) from
Tethyan ophiolites, which were derived from the the earlier calc-alkaline granitic magmatism to the
Izmir–Ankara– Erzincan suture zone to the north, later alkaline, syenitic magmatism during the
and before its collision with the Pontide and latest Cretaceous.
Tauride continental blocks during the middle
Eocene. Therefore, the magmatic evolution of Tauride block
these plutons preceded the terminal continental col-
lisions in the region (Akiman et al. 1993; Erler & The Tauride block south of the CACC consists of
Bayhan 1995; Kadioglu et al. 2006). Palaeozoic to upper Cretaceous carbonate, silici-
The CACC plutons can be grouped into three clastic and volcanic rocks (Özgül 1976; Demirtasli
supersuites based on their field occurrences and dis- et al. 1984) and represents a ribbon continent
tinct differences in their mineral and chemical com- rifted off from the northwestern edge of Gondwana
positions (Kadioglu et al. 2006). Plutonic rocks of (Robertson & Dixon 1984; Garfunkel 1998). The
the Granite Supersuite commonly occur in a curvi- Palaeozoic –Jurassic tectonostratigraphic units in
linear belt along the western edge of the CACC the Tauride block are tightly folded and imbricated
(east of the Salt Lake-Tuzgölü, Fig. 9) and consist along major thrust faults that developed first during
of calc-alkaline rocks ranging in composition the obduction of the Inner-Tauride ophiolites from
from tonalite, granodiorite and biotite granite to the north in the late Cretaceous, and subsequently
amphibole biotite granite and biotite-alkali feldspar during the collision of the Tauride block with the
granite. Plutons of the Monzonite Supersuite (i.e. CACC in the latest Palaeocene – Eocene. The buoy-
Terlemez, Cefalik, Baranadag plutons) occur ancy of the Tauride continental crust in the lower
immediately east of the Granite Supersuite plutons plate eventually arrested the subduction process
and caused the isostatic rebound of the partially sub- of the CACC around 90 –85 Ma was facilitated by
ducted platform edge, leading to block-fault uplift- a subduction zone dipping northward beneath the
ing of the Taurides during the latest Cenozoic Pontides and away from the CACC (Tankut et al.
(Dilek & Whitney 1997, 2000). The entire Tauride 1998; Dilek & Whitney 2000; Floyd et al. 2000).
block experienced gradual uplift in the footwall of This subduction zone could not have had any
a north-dipping frontal normal fault system along effect on the mantle dynamics and heterogeneity
its northern edge starting in the Miocene, and devel- beneath the CACC or on the evolution of the Late
oped as a southward-tilted, asymmetric mega-fault Cretaceous magmatism on and across the CACC
block with a rugged, alpine topography (Dilek (as suggested in Boztug 2000; Ilbeyli et al. 2004;
et al. 1999b). Apatite fission track ages of Köksal et al. 2004; Ilbeyli 2005). Instead, the sub-
23.6 + 1.2 Ma from the 55 Ma-old Horoz granite duction zone that was involved in the evolution
that is intrusive into the Bolkar carbonates are con- of the CACC magmatism was located to the
sistent with this uplift history (Dilek et al. 1999b). SW, dipping northeastward beneath the CACC
(in present coordinate system) and consuming the
Cappadocian volcanic province oceanic lithosphere of the Inner-Tauride Ocean
(Fig. 10a; Erdogan et al. 1996; Dilek et al. 1999a;
The Cappadocian volcanic province defines a Kadioglu et al. 2003). The inferred chemical
c. 300 km-long volcanic belt extending NE– SW modification of the mantle wedge beneath the
across the CACC (Fig. 9). The earliest volcanism CACC was associated with this ITSZ.
in the province started in the mid-Miocene There are several independent lines of evidence
(c. 13.5 Ma) and continued into the Quaternary for the existence of this Inner-Tauride basin
(Pasquaré et al. 1988; Ercan et al. 1994). The between the CACC and the Tauride block and the
initial volcanic products include 13.5 –8.5 Ma ande- northward dipping subduction zone within this
sitic lavas, tuff and ignimbrites in the form of effu- Tethyan seaway. The Cenomanian age supra-
sive centres and endogeneous domes. This volcanic subduction zone ophiolites in the Tauride belt
phase was followed by widespread eruption of were derived from the arc-forearc setting in the
ignimbrites, volcanic ash, lapilli and agglomerates Inner-Tauride Ocean and were emplaced southward
between 8.5 and 2.7 Ma (Pasquaré et al. 1988; onto the continental edge of the Tauride block
Ercan et al. 1994). The most recent phase of volcan- during its collision with an intra-oceanic arc-trench
ism produced central volcanoes oriented parallel system (Dilek et al. 1999a; Parlak et al. 1996).
to the NE– SW axis of the province, consisting of Partial subduction of the Tauride edge beneath the
basaltic andesite, andesite, dacite, rhyodacite and ophiolite nappes resulted in HP/LT metamorphism
basaltic lavas. Geochemically, these rocks collec- of the platform carbonates and in the formation of
tively have an A-type granitic melt origin showing blueschists (Fig. 10b), which are currently exposed
post-collisional, within-plate affinities (Innocenti discontinuously along the northern periphery of
et al. 1975; Ercan et al. 1994; Toprak et al. 1994). the Tauride belt (Okay 1984; Okay et al. 1996,
The Cappadocian volcanic province broadly cor- 1998; Dilek & Whitney 1997; Önen 2003). The
responds to a structurally controlled topographic calculated P–T conditions of metamorphism and
depression filled with Upper Miocene to Pliocene 40
Ar/39Ar ages of phengites and glaucophanes
fluvial and lacustrine deposits (Dilek et al. 1999b). from blueschist rocks in northwestern Anatolia
The volcanic edifices appear to have been built at suggest that this HP/LT metamorphism occurred
the intersections of major strike-slip fault systems around 88 Ma (Harris et al. 1994; Okay et al. 1996).
(i.e. Mt. Erciyes volcano) and/or in local graben The older peraluminous granitoid plutons (c.
structures associated with tectonic subsidence 100–85 Ma) along the western edge of the CACC
and a c. NNW–SSE-oriented regional extension represent a magmatic arc complex that developed
during the late Cenozoic (Toprak et al. 1994; above a NE-dipping subduction zone within the
Dilek et al. 1999b). These relations indicate that Inner-Tauride Ocean (Fig. 10b; Görür et al. 1984).
faulting and volcanism in the Cappadocian volcanic During the closure of this basin as a result of the
province were spatially and temporally associated, subduction of the Tethyan oceanic lithosphere
and that magma transport and extrusion were in beneath the western edge of the CACC, the melts
part facilitated by crustal-scale fault systems derived from the metasomatized upper mantle
(Dilek et al. 1999b). were injected into the continental crust, causing its
partial melting. This led to interaction of mantle-
Tectonic evolution of the central and crustal-derived magmas that involved AFC,
Anatolian orogenic belt mixing and mingling which collectively produced
the calc-alkaline granitoid plutons (Fig. 10b, c;
The southward emplacement of the Tethyan supra- Erdogan et al. 1996; Kadioglu et al. 2003; Ilbeyli
subduction zone ophiolites along the northern edge et al. 2004).
Fig. 10. Tectonic model for the evolution of the CACC and the central Anatolian orogenic belts in the late Mesozoic
(modified after Kadioglu et al. 2006). See text for discussion. Ellipses beneath SMM depict melting in the asthenosphere.
Stippled (in white) pattern characterizes subduction-metasomatized mantle. Key to lettering: AFC, Assimilation
Following the demise of the Inner-Tauride within the CACC prior to the Eocene (Gautier
oceanic lithosphere at the NE-dipping subduction et al. 2002).
zone and the emplacement of the incipient arc- Strike-slip faulting also played a major role in the
forearc ophiolites onto the northern edge of the late Cenozoic tectonic evolution of the CACC. A
Tauride block, subduction was arrested by the series of NE –SW and NW –SE trending strike-slip
underplating of the buoyant Tauride continental fault systems crosscut both the crystalline basement
crust. The leading edge of the subducted Tethyan rocks and the Palaeogene–Quaternary sedimentary
slab likely broke off from the rest of the Tauride strata and volcanic rock units (Fig. 9). The sinistral
continental lithosphere, resulting in the develop- Ecemis fault (EF) truncates and offsets the Tauride
ment of an asthenospheric window (Fig. 10c). The block and the ITSZ by as much as 100 km, and
juxtaposition of this asthenospheric heat source forms mainly a transtensional fault system with
against the overlying continental lithosphere may several releasing bends and pull-apart structures.
have caused melting of the metasomatized mantle It appears to have accommodated top-to-the
layers, producing the high-K shoshonitic, adakite- south extension and crustal exhumation along the
like magmas of the monzonitic plutons and then south-dipping detachment fault of the Nigde core
the more-enriched alkaline magmas of the syenitic complex. It may have, therefore, facilitated the
plutons (Fig. 10d). This process is similar to that vertical displacement and unroofing of high-grade
inferred for slab breakoff-related collisional mag- metamorphic rocks in the eastern part of the
matism in other orogenic belts (Davies & von Nigde massif during the Oligo –Miocene (Dilek &
Blackenburg 1995, and references therein) and in Whitney 1998). The NW–SE trending Tuzgölü
the early Cenozoic of western Anatolia, as discussed fault (TGF) bounding the Cappadocian volcanic pro-
earlier. vince on the west represents a dextral transpressional
As the asthenosphere moved upwards through fault system (Saroglu et al. 1992). The latest Cretac-
the window in the slab, partial melting occurred at eous granitoid plutons and their high-grade meta-
the boundary between the lithosphere and astheno- morphic host rocks in the western part of the
sphere producing basaltic melts with transitional CACC have been uplifted on the eastern shoulder
characteristics between those of calc-alkaline and of the TGF. The Ecemis and Tuzgölü faults together
alkaline basalts. The mantle sources for these form a triangular-shaped crustal flake within the
primary basaltic melts may have been metasoma- CACC that has been moving southwards while
tized garnet peridotites and/or spinel lherzolites undergoing rifting and subsidence, which has
and phlogopite-bearing lherzolites of an upper given way into the development of the Cappadocian
mantle wedge origin (Conceiçao & Green 2004). volcanic province throughout the latest Cenozoic.
Such slab breakoff-related magmatism with
similar calc-alkaline to alkaline products has been
documented from other orogenic belts, such as Regional geology of eastern Anatolia
the Neogene –Quaternary Carpathian –Pannonian
region (Nemcok et al. 1998; Seghedi et al. 2004), Much of eastern Turkey is occupied by the East Ana-
the Late Palaeogene Periadriatic –Sava –Vardar tolian High Plateau, which is bounded in the north
magmatic zone in the Dinaride–Hellenide moun- by the Eastern Pontide arc and in the south by the
tain belt (Pamic et al. 2002), the Neogene Maghre- Bitlis –Pütürge massif (Fig. 11). The mean surface
bian orogenic belt in northern Africa (Maury et al. elevation of the plateau is about 2– 2.5 km above
2000) and the late Oligocene –early Miocene sea level with scattered Plio– Quaternary volcanic
central Aegean Sea region (Pe-Piper & Piper 1994). cones over 5 km high (e.g. Mt. Ararat or Mt. Agri).
Thermal perturbation of the continental litho- The basement geology of the plateau is composed
sphere and the alkaline magmatism weakened the of ophiolites and ophiolitic mélange, flysch and
orogenic crust, leading to tectonic extension in and molasse deposits, and the eastward extension of
across the CACC (Fig. 10d). The results of recent the Tauride platform carbonates. The Tethyan
studies of metamorphic massifs and core complexes ophiolites and ophiolitic mélanges, flysch deposits
in central Anatolia suggest that significant crustal and volcanic arc units collectively constitute the
extension and unroofing might have occurred East Anatolian Subduction– Accretion Complex.
Fig. 10. (Continued) fractional crystallization; CACC, Central Anatolian Crystalline Complex; CACC LM, Central
Anatolian Crystalline Complex lithospheric mantle; IAESZ, Izmir–Ankara–Erzincan suture zone; ITO, Inner Tauride
Ocean; SMM, Subduction modified mantle; SSZ, Suprasubduction zone; TGB, Tuzgölü basin; TLM, Tauride
lithospheric mantle; TO, Tauride ophiolites (including the Aladag, Alihoca and Mersin ophiolites); UKB, Ulukisla basin.
GSS, MSS and SSS magmatism represent the Granite, Monzonite and Syenite Supersuites, respectively, of the
syncollisional CACC plutons. SSZ ophiolites of the Inner-Tauride Ocean (ITO) include the Beysehir, Alihoca, Aladag,
and Mersin ophiolites.
Fig. 11. Simplified geological map of Eastern Anatolia and the Arabian foreland. Munzur Platform constitutes the
eastern extension of the platform carbonates and basement rocks of the Tauride block. Bitlis–Pütürge massif is a rifted
off fragment of the Arabian plate, analogous to the Tauride block. The East Anatolian Plateau is covered by Miocene–
Quaternary volcanic rocks; its basement is composed of Tethyan ophiolites and ophiolitic mélanges, flysch and
molasses deposits, and platform carbonates of the Tauride block. Key to lettering: EAFZ, East Anatolian fault zone;
NAFZ, North Anatolian fault zone.
Eastern Pontide arc Sanandaj–Sirjan continental block to the SE in

Iran represents the eastern continuation of the
The Pontide block north of the plateau represents Bitlis–Pütürge massif in the peri-Arabian region.
a south-facing early Cretaceous – late Eocene The Sanandaj– Sirjan continental block to the SE
volcano-plutonic arc that developed over a subduc- in Iran represents the eastern continuation of the
tion zone dipping northward beneath the Eurasian Bitlis–Pütürge massif in the peri-Arabian region.
continental margin (Yilmaz et al. 1997). The col- The Pütürge massif is composed of pre-Triassic
lision of the Pontide arc with the Tauride microcon- gneisses and micaschists, and granitoids (Michard
tinent in the late Eocene terminated the subduction et al. 1984; Aktas & Robertson 1990) and is inter-
zone magmatism in the Pontide terrane and pro- preted as a pre-Triassic continental sliver of Afro–
duced extensive flysch deposits with intense Arabian origin, similar to the Bitlis massif to the
folding in the collision zone (Dewey et al. 1986). east. The Pütürge massif and the overlying volcanic
and ophiolitic rocks are structurally underlain in
Bitlis– Pütürge massif the south by an upper Cretaceous –early Tertiary
mélange, which is underthrust to the south by the
The Bitlis–Pütürge massif is a nearly East –West- foreland sedimentary sequences of the Arabian
trending microcontinent that is surrounded on all plate (Fig. 12).
sides by ophiolitic rocks, mélanges, and volcanic The Bitlis massif is a composite tectonic unit,
and volcaniclastic rocks of an arc origin. The which is composed of a metamorphic basement
CENOZOIC OROGENIC BELTS IN THE EASTERN MEDITERRANEAN REGION
Fig. 12. Simplified tectonic map of Eastern Anatolia showing the basement geology, which consists of metamorphic massifs, ophiolites, ophiolitic mélanges, magmatic arcs, flysch
deposits, Tauride platform carbonates, main fault zones, and the distribution of the major Plio–Quaternary volcanic eruptive centres and the Miocene– Pliocene volcanic rocks in
the East Anatolian High Plateau. EAFZ and NAFZ mark the East and North Anatolian fault zones, respectively. Symbols for the ophiolites: GO, Guleman; IPO, Ispendere;
147
KHO, Kömürhan; OOM, Ovacik ophiolitic mélange.
(Precambrian?), overlying metamorphosed Palaeo- and gabbros of an island-arc. These arc rocks
zoic to Triassic carbonate rocks (Göncüoglu & intrude the oceanic rocks and mafic extrusives of
Turhan 1984; Helvaci & Griffin 1984), and the Ispendere –Kömürhan ophiolites of Yazgan
Palaeozoic to late Mesozoic granitoids (Fig. 12). (1984) and the Guleman nappe of Michard et al.
Oberhänsli et al. (2008) recently reported a region- (1984). The Ispendere ophiolite is unconformably
ally distributed LT/HP metamorphic overprint in overlain by a flysch deposit, which contains fossils
the thermal evolution of the Bitlis massif and of Upper Campanian–Lower Maastrichtian age
suggested that the massif is composed of a stack (Yazgan 1984). The Kömürhan ophiolite structu-
of nappes formed during the closure of the Southern rally beneath the Ispendere ophiolite contains
Tethys between Arabia and the Anatolide –Tauride highly deformed, amphibolite-grade mafic, ultrama-
continental block. The whole massif displays a fic and intermediate rocks. The metamorphic rocks
doubly plunging, multiply folded anticlinorium are intruded by diorites and granodiorites of the
with overturned limbs both to the north and the Baskil arc complex (Elazig –Palu nappe). The Gule-
south (Dilek & Moores 1990). The relatively young- man ophiolite constitutes the eastern extension of
est thrust faults are south-vergent and synthetic to the Ispendere –Kömürhan ophiolite and is in tec-
the Bitlis suture that represents the collision zone tonic contact with the Bitlis metamorphic massif.
between the Arabian and Eurasian plates. The contact relationships indicate, in general, that
the oceanic rocks of the Ispendere –Kömürhan
ophiolites and the Guleman nappe form the ocea-
East Anatolian subduction-accretion nic basement on which the volcanic arc rocks of
complex the Elazig-Palu nappe (including the Baskil arc)
were deposited (Dilek & Moores 1990).
The Jurassic (?) –Cretaceous ophiolites underlying In the eastern part of the East Anatolian High
the molasse deposits and the Tertiary volcanic Plateau the Bitlis massif is tectonically overlain to
cover in the western part of the East Anatolian the north and west by the Gevas and Guleman
High Plateau represent the remnants of Mesozoic ophiolites, respectively. The Gevas ophiolite,
Tethys and are commonly directed southwards exposed in an east –west trending narrow belt
onto the margins of the Eastern Tauride platform immediately south of Lake Van (Fig. 12), consists
and the Pütürge massif (Dilek & Moores 1990). of serpentinized ultramafic rocks, cumulate and iso-
The Maastrichtian Ovacik ophiolitic mélange over- tropic gabbros, microgabbros and plagiogranites
lies tectonically the Upper Triassic –Cretaceous overlain by extrusive rocks and pelagic sediments
Munzur platform (Fig. 12) that is interpreted to be (Dilek 1979). These mafic –ultramafic rocks tecto-
the northeastward extension of the calcareous axis nically rest on the Bitlis massif along a south-
of the Tauride block (Özgül & Tursucu 1984). The vergent thrust fault. This thrust fault and the
Ovacik mélange consists of blocks of serpentinites, Gevas ophiolite are further deformed and thrust
metamorphic rocks and pelagic limestones in a over by the Bitlis massif along north-vergent faults
fine-grained, phyllitic matrix and is unconformably that are depositionally overlain by Palaeocene –
overlain by Maastrichtian clastic rocks (Özgül & Eocene flysch deposits (Dilek 1979).
Tursucu 1984). Both the Ovacik mélange and Mafic –ultramafic rocks and extrusives of similar
Munzur carbonates are thrust to the south over character crop out farther west, south of the East
the Keban metamorphic rocks that consist of Anatolian Fault and west of the Bitlis massif,
Permian to Cretaceous metamorphosed platform where the Guleman ophiolite overlies tectonically
carbonates (Fig. 12; Michard et al. 1984). The the metamorphosed carbonates of the Bitlis massif
Keban metamorphics and Munzur carbonates (Fig. 11). The Guleman ophiolite includes serpenti-
display north-vergent, upright folds that have been nized peridotites, banded gabbro, microgabbro,
subsequently folded by south-vergent folds (Dilek metamorphosed basalt, tuff and agglomerate and
& Moores 1990). radiolarian mudstone (Göncüoglu & Turhan 1984;
The Keban metamorphic rocks overlie tectoni- Aktas & Robertson 1990). The Upper Jurassic –
cally the Elazig –Palu nappe to the south along Lower Cretaceous Guleman ophiolite is separated
south-vergent thrust faults (Fig. 12). The Elazig – from the underlying Bitlis massif by an intensely
Palu nappe consists of calc-alkaline intrusive mylonitized zone and overlain by an unmetamor-
and extrusive rocks, and overlying Campanian– phosed Upper Maastrichtian flysch deposit that
Maastrichtian volcaniclastic and flysch deposits contains blocks of both the Guleman ophiolite and
(Michard et al. 1984; Yazgan 1984; Aktas & the Bitlis massif. This depositional relationship con-
Robertson 1990). The Elazig –Palu nappe includes strains the ophiolite emplacement age as pre-late
Yazgan’s (1984) Baskil magmatic rocks that Maastrichtian. The rocks of the Gevas– Guleman
consist of Coniacian-Santonian granodiorites, tona- ophiolite belt continue farther west and are intruded
lites, quartz monzonites, monzodiorites, diorites and overlain by the calc-alkaline intrusives and
volcaniclastic rocks of the Elazig –Palu nappe collision zone took up much of the convergence
(Ozkaya 1978; Michard et al. 1984; Aktas & between Arabia and Eurasia. East–West-oriented
Robertson 1990). thrust faults and folds in the Upper Miocene lavas
The Bitlis massif is underlain to the south by a and pyroclastic rocks of the Solhan volcanic rocks
late Cretaceous – early Tertiary tectonic mélange, in the Mus basin area (Fig. 11; Yilmaz et al. 1987)
which directly overlies the foreland deposits of indicate that crustal shortening continued after the
Arabian plate (Fig. 12). The mélange is composed Arabia-Eurasia collision and affected the post-
of ophiolitic material tectonically interleaved with collisional volcanic rocks in the plateau. However,
hemipelagic and clastic rocks (Aktas & Robertson progressive thickening of the crust has been
1990). In places, however, the Bitlis massif is under- accompanied by major strike-slip faulting on the
lain by thrust sheets of ophiolitic rocks consisting of dextral North Anatolian and the sinistral East Ana-
serpentinized peridotites, gabbro, diabase and basal- tolian faults that have accommodated the westward
tic andesites that structurally overlie the tectonic escape of the Anatolian plate at a rate of 0.5 cm a21
mélange. Both the mélange and the ophiolitic (Jackson & McKenzie 1984). NE- and NW-striking
thrust sheets constitute the Killan Imbricate Unit conjugate strike-slip faults (sinistral and dextral,
of Aktas and Robertson (1990) that comprises, respectively) with a significant compressional com-
with the Guleman ophiolite, their Maden complex ponent also occur within the plateau (i.e. Tutak
(Fig. 12). The Maden complex and the overlying fault; Fig. 11), and some of these faults are seismi-
Elazig –Palu nappe directly rest on the Arabian plat- cally active (Tan & Taymaz 2006).
form units along south-vergent thrust faults wher- The widespread volcanism in the Turkish–
ever the Bitlis massif is absent (Dilek & Moores Iranian Plateau started between 8 and 6 Ma,
1990). South of the Bitlis Suture Zone, the lower 4 –5 Ma years after the initial collision of Arabia
Palaeozoic to Miocene shelf sequences of the with Eurasia in Serravallian-Tortonian times (Inno-
Arabian foreland display a south-vergent fold-and- centi et al. 1982; Yilmaz et al. 1987; Pearce et al.
thrust belt architecture. 1990). In general, volcanism in the southern seg-
East-southeast of the Bitlis massif, south-vergent ment of the plateau is characterized by the construc-
thrust sheets composed of a complete ophiolite tion of stratovolcanoes with significant peaks (i.e.
sequence (the Cilo ophiolite) and arc-related calc- Nemrut, Suphan, Tendürek, Ararat; Fig. 11), where-
alkaline rocks rest tectonically on the Mesozoic as in the north it forms an extensive (5000 km2) and
Arabian platform (Yilmaz 1985). The Cilo ophiolite relatively flat volcanic field (Erzurum-Kars plateau;
contains, from bottom to top, peridotites, cumulate Fig. 11) with an average elevation of c. 1.5 km
and isotropic gabbros, diorite, quartz diorite, above sea level. This volcanic field consists
sheeted dykes, pillow and massive lava flows, and mainly of lava flows intercalated with subordinate
ribbon cherts and shales (Yilmaz 1985). The sedi- ignimbrite units and sedimentary layers giving
mentary rocks associated with the ophiolite give ages from 6.9 + 0.9–1.3 + 0.3 Ma (Innocenti
Jurassic to Upper Cretaceous fossil ages; the entire et al. 1982; Keskin et al. 1998). Pleistocene scoriac-
ophiolitic sequence is depositionally overlain by eous spatter cones locally overlie this lava-
volcanic-pyroclastic rocks and is intruded by ignimbrite sequence. The initial eruptive phase of
granitic–granodioritic intrusions (Fig. 11; Yilmaz post-collisional volcanism in the plateau is charac-
1985). The lower thrust sheet underlying the Cilo terized by basic and intermediate alkaline rocks
ophiolite consists mainly of calc-alkaline lavas, and was followed by widespread eruptions of ande-
pyroclastics, and blocks of radiolarian chert and sitic to dacitic calc-alkaline magma during the Plio-
recrystallized limestone. These relations suggest cene; the last volcanic phase involved the eruption
that the Cilo ophiolite may constitute the oceanic of alkaline and transitional lavas throughout the
basement of an ensimatic arc complex represented Plio-Pleistocene and Quaternary (Yilmaz et al.
by the calc-alkaline intrusions and volcanic- 1987). Most of the major stratovolcanoes in the
pyroclastic rocks. region were built during this last phase of volcan-
ism, which continued until historical times. Alkaline
basaltic lavas of the late volcanic phase appear
Deformation and volcanism in the East to predominate mainly in the northern part of the
Anatolian High Plateau plateau, whereas the calc-alkaline rocks of the
second major volcanic phase occur most extensively
Emplacement of the Tethyan ophiolites in the late in the south.
Cretaceous and the subsequent continental col- The East Anatolian High Plateau has under-
lisions in the Eocene and mid-Miocene played a gone significant uplift since the Arabian collision
major role in the construction of the East Anatolian in the Mid-Miocene (c. 13 Ma). The Lower to
High Plateau. North –South shortening and crustal Middle Miocene fossiliferous marine marl and
imbrication via thrust faulting and folding in the reefal carbonates around Lake Van indicate that
Fig. 13. Maastrichtian –Cenozoic geodynamic evolution of the East Anatolian High Plateau and the subduction-
accretion complex through subduction and collisional processes in the upper plate of north-dipping subduction zone(s)
within the Tethyan realm (data and interpretations are derived in part from Dilek & Moores 1990; Keskin 2003; Sengör
et al. 2003; and Dilek et al. 2010). Guleman ophiolite (GO) together with the Ispendere–Kömürhan ophiolites and
Baskil arc units represent backarc and arc oceanic crust tectonically emplaced onto the northern margin of the Bitlis–
Pütürge metamorphic massif by the early Eocene. See text for discussion. Key to lettering: ALM, Arabian lithospheric
mantle; BPM, Bitlis– Pütürge massif; BSZ, Bitlis suture zone; EKP, Erzurum–Kars plateau; GO, Guleman ophiolite;
MS, Mus suture; PLM, Pontide lithospheric mantle; SMAM, Subduction metasomatized asthenospheric mantle; TIP,
Turkish–Iranian high plateau.
the area remained under the sea until the Serraval- and thickened within the closing basin (Fig. 13c).
lian (Gelati 1975). Upper Miocene lacustrine and North– South contraction across the Northern
fluvial sedimentary rocks unconformably overlie Tethyan realm and vertical thickening of the East
these marine sedimentary rocks, suggesting the Anatolian subduction-accretion complex created a
emergence of land and the onset of terrestrial ‘tectonic bumper’ between the converging Eurasia
conditions by the late Miocene (Altinli 1966). A and Bitlis–Pütürge–Arabia plates that had
late Miocene– early Pliocene erosional surface trun- reached the average thickness of continental crust
cating the fluvial rocks was covered with Pliocene by the late Oligocene–early Miocene (c. 24 Ma;
andesitic– dacitic lavas and was subsequently Sengör et al. 2003). Further steepening and south-
deeply dissected by streams and rivers due to rapid ward retreat of the subducting Tethyan lithosphere
block-uplift of the western and central segments of might have triggered lithospheric delamination
the plateau (Altinli 1966; Innocenti et al. 1976). beneath the southern margin of the Eastern
These stream valleys were then filled with Pleisto- Pontide arc and the northern part of the East Anato-
cene lava flows fed by the alkaline volcanic phase lian high plateau, resulting in remobilization and
(Erinç 1953). partial melting of the subduction-metasomatized
asthenospheric mantle (Fig. 13d). This event pro-
duced the initial stages of calc-alkaline magmatism
Tectonic Evolution of the East Anatolian in the Erzurum-Kars Plateau by the middle Miocene
High Plateau (Keskin et al. 2006).
Arrival of the Bitlis–Pütürge– Arabia composite
The late Mesozoic geodynamic evolution of eastern continental plate at the trench and the continent-
Anatolia was controlled by subduction zone tec- trench collision by c. 13 Ma slowed down and tem-
tonics in two separate Tethyan seaways (Fig. 13a). porarily arrested the northward subduction beneath
The Northern Tethys seafloor was being consumed the East Anatolian subduction-accretionary com-
at a subduction zone dipping northward beneath plex. However, the continued sinking of the ocea-
the Eastern Pontide arc, and the Black Sea was nic lithosphere in this subduction zone must have
opening up as a back-arc basin behind this arc caused the detachment of the subducting slab,
around 75–70 Ma (Yilmaz et al. 1997). Subduction leading into slab breakoff and development of
of the Southern Tethys seafloor beneath the Tauride an asthenospheric window (Fig. 13d). Rising hot
microcontinent farther south developed a magmatic asthenosphere beneath the subduction-accretion
arc on the Bitlis–Pütürge microcontinent and the complex resulted in widespread partial melting
arc-backarc oceanic crust presently represented by both in the upwelling and convecting asthenosphere
the Ispendere –Kömürhan and Guleman ophiolites and in the overlying crust (Fig. 13d; Sengör et al.
tectonically overlying the Bitlis –Pütürge massifs. 2003; Keskin 2003) that produced bimodal volcan-
A similar tectonic scenario has been suggested for ism throughout the uplifted plateau. Extensive
the Cretaceous arc-ophiolite duo in the Malatya – strike-slip and extensional normal faulting in the
Maras region farther west in the Tauride block Turkish–Iranian high plateau (TIP) facilitated the
(Parlak 2006). rise and eruption of asthenosphere-derived alkaline
The collision of the Arabian plate with the olivine basalts at the surface without much conti-
Bitlis–Pütürge magmatic arc occurred in the early nental contamination in the late Miocene –Pliocene
Eocene (Yilmaz 1993) and produced the mélange (Fig. 13e).
and flysch deposits along the Bitlis suture zone Widespread volcanism across the entire East
(Fig. 13b). This continental collision led to slab Anatolian high plateau (.250 km wide) throughout
breakoff and development of an asthenospheric the late Cenozoic and until historic times indicates a
window, which in turn facilitated partial melting significant heat source beneath it, resulting in exten-
of the subduction-metasomatized lithospheric sive melting. The findings of the recent Eastern
mantle beneath the Bitlis– Pütürge massifs, pro- Turkey Seismic Experiment (ETSE) and tomo-
ducing the shoshonitic magmatism in the Maden graphic models have shown the lack of mantle litho-
Complex (Fig. 13b; Elmas & Yilmaz 2003). sphere, an average continental crustal thickness
Continued subduction of the Northern Tethyan (c. 40–45 km), lack of earthquakes deeper than
seafloor beneath Eurasia farther north and slab stee- c. 30 km, and very low Pn velocity zones indicating
pening and roll-back produced southward migrating the presence of partially molten material beneath the
magmatism in the Eastern Pontide arc during the region (Sandvol et al. 2003a; Al-Lazki et al. 2003;
Eocene –Oligocene, while the subduction-accretion Gök et al. 2003; Zor et al. 2003; Angus et al.
complex widened toward the south (Sengör et al. 2006). These observations collectively suggest that
2003). As the Tethyan lithosphere continued its the East Anatolian high plateau is likely supported
subduction beneath the Pontide arc, the widening by hot asthenospheric mantle, not by overthickened
East Anatolia accretionary complex was shortened crust (Dewey et al. 1986) or subducted Arabian
continental lithosphere (Rotstein & Kafka 1982) as The collision-driven tectonic evolution of the
previously inferred. Anatolian –African plate boundary and the young
orogenic belts in the eastern Mediterranean region
Conclusions is typical of the geodynamic development of the
Alpine –Himalayan orogenic system. Successive
The modern Anatolian– African plate boundary is collisions of Gondwana-derived microcontinents
characterized by subduction zone tectonics and is with trench-roll-back cycles in the Tethyan realms
in the initial stages of collision-driven orogenic of the Alpine –Himalayan system caused basin
buildup. The Anatolian microplate itself is made collapse, ophiolite emplacement and continental
of young orogenic belts (Eocene and younger) accretion, producing subparallel mountain belts.
that evolved during a series of collisions between Subduction of the Tethyan mantle lithosphere was
Gondwana-derived ribbon continents and trench- nearly continuous throughout these accretionary
roll-back systems within the Tethyan realm. The processes, only temporarily punctuated by slab
collision of the Eratosthenes seamount with the breakoff events.
Cyprus trench since the late Miocene is a smaller-
scale example of this accretionary process and has Part of this study was undertaken during the Eastern
affected the slab geometry and kinematics of the Turkey Seismic Experiment (ETSE), which was supported
subducting African lithosphere. by the NSF under Grant No EAR-9804780; additional
support for this experiment was provided by the Bogaziçi
Pn velocity and Sn attenuation tomography
University Research Fund under Grant No 99T206. We
results show that the uppermost mantle beneath gratefully acknowledge these grants. The geological field-
much of the young orogenic belts in Anatolia is work in Turkey has been funded over the years by the NSF
anomalously hot and thin. This is consistent with Tectonics Program, Miami University Committee on
the surface geology, which is dominantly controlled Faculty Research and Hampton International Initiatives
by strike-slip and extensional tectonics and wide- Funds, Scientific & Technical Research Council of
spread volcanism in western, central and eastern Turkey (TUBITAK), and Istanbul Technical University
Turkey. In all these areas, the extension was well Research Funds. Discussions with S. Altunkaynak,
under way by the late Oligocene –Miocene, follow- M. Barazangi, E. Bozkurt, C. Genç, R. Gök, C. Helvaci,
Y. K. Kadioglu, A. Polat, Y. Savasçin, F. Yagmurlu,
ing the main episodes of continental collisions.
Y. Yilmaz, and E. Zor on various aspects of the
Pinning of subduction hinge zones by the accreted Mesozoic– Cenozoic geology, geodynamics and seismic
ribbon continents arrested slab roll-back processes, structure of Anatolia have been most helpful to us in
causing terrane stacking and crustal thickening, developing the ideas and interpretations presented in this
and resulted in slab breakoff because of continued paper. Thorough and constructive reviews by D. Nance,
convergence of the lithospheric mantle. Slab B. Murphy and P. T. Robinson helped us improve
breakoff-induced asthenospheric upwelling pro- the paper.
vided the necessary heat and melt to produce the
first phases of post-collisional magmatism in these
young orogenic belts. Renewed subduction and References
slab roll-back in the Tethyan realm triggered A GOSTINI , S., D OGLIONI , C., I NNOCENTI , F., M ANETTI ,
lithospheric-scale extension in the upper plate, and P., T ONARINI , S. & S AVASÇIN , M. Y. 2007. The tran-
the thermally weakened orogenic crust started sition from subduction-related to intraplate Neogene
collapsing. These processes resulted in rapid exhu- magmatism in the Western Anatolia and Aegean
mation of recently formed high-pressure meta- area. In: B ECCALUVA , L., B IANCHINI , G. &
morphic rocks and in the formation of metamorphic W ILSON , M. (eds) Cenozoic Volcanism in the Mediter-
ranean Area. Geological Society of America Special
core complexes.
Paper, 418, 1– 16.
The Cenozoic geodynamic evolution of the A KAY , E. & E RDOGAN , B. 2004. Evolution of Neogene
western, central and eastern Anatolian orogenic calc-alkaline to alkaline volcanism in the Aliaga-
belts indicates that the asthenospheric mantle Foça region (Western Anatolia, Turkey). Journal of
beneath collision zones responds swiftly to crustal Asian Earth Science, 24, 367–387.
tectonics on time scales of just a few million A KIMAN , O., E RLER , A., G ÖNCÜOĞLU , M. C., G ULEÇ ,
years. Slab breakoff, lithospheric delamination and N., G EVEN , A., T URELI , T. K. & K ADIOGLU , Y. K.
slab tearing were common processes that resulted 1993. Geochemical characteristics of granitoids along
directly from collision-induced events, and caused the western margin of the Central Anatolian Crystal-
line Complex and their tectonic implications. Geologi-
convective remobilization of the asthenosphere
cal Journal, 28, 371– 382.
leading to magmatism. Asthenospheric upwelling A KTAS , G. & R OBERTSON , A. H. F. 1990. Tectonic
and partial melting played a major role in a evolution of the Tethys suture zone in SE Turkey:
geochemical progression of post-collisional mag- evidence from the petrology and geochemistry of
matism from initial shoshonitic, calc-alkaline to Late Cretaceous and Middle Eocene extrusives.
late-stage alkaline affinities through time. In: M ALPAS , J., M OORES , E. M., P ANAYIOTOU , A.
& X ENOPHONTOS , C. (eds) Ophiolites, Oceanic A YDIN , N. S., G ONCUOGLU , M. C. & E RLER , A. 1998.
Crustal Analogues, Proceedings of the Symposium Latest Cretaceous magmatism in the Central Anatolian
‘Troodos 1987’. The Geological Survey Department, Crystalline Complex: review of field, petrographic and
Nicosia, Cyprus, 311–328. geochemical features. Turkish Journal of Earth
A LDANMAZ , E., P EARCE , J. A., T HIRWALL , M. F. & Sciences, 7, 259– 268.
M ITCHELL , J. 2000. Petrogenetic evolution of late B ARKA , A. & R EILINGER , R. 1997. Active tectonics of the
Cenozoic, post-collision volcanism in western Anato- Eastern Mediterranean region: deduced from GPS,
lia, Turkey. Journal of Volcanology & Geothermal neotectonic and seismicity data. Annali Di Geofisica,
Research, 102, 67– 95. XL, 587–610.
A LDANMAZ , E., K ÖPRÜBASI , N., G ÜRER , Ö. F., K AY- B AYHAN , H. 1987. Cefalıkdag ve Baranadag plütonlarının
MAKÇI , N. & G OURGAUD , A. 2006. Geochemical con- (Kaman) petrografik ve kimyasal-mineralojik özelik-
straints on the Cenozoic, OIB-type alkaline volcanic leri. Jeoloji Mühendisligi, 30–31, 11–16 (in Turkish).
rocks of NW Turkey: implications for mantle sources B IJWAARD , H. & S PAKMAN , W. 2000. Non-linear global
and melting processes. Lithos, 86, 50–76. P-wave tomography by iterated linearized inversion.
A LESSANDRINI , B., B ERANZOLI , L., D RAKATOS , G., Geophysical Journal International, 141, 71– 82.
F ALCONE , C., K ARANTONIS , G., M ELE , F. M. & B INGÖL , E., D ELALOYE , M. & A TAMAN , G. 1982. Grani-
S TAVRAKAKIS , G. 1997. Back arc basins and tic intrusions in Western Anatolia: A contribution of
P-wave crustal velocity in the Ionian and Aegean the geodynamic study of this area. Eclogae Geologica
regions. Geophysical Research Letters, 24, 527–530. Helvetica, 75, 437– 446.
A LICI , P., T EMEL , A., G OURGAUD , A., K IEFFER , G. & B INGÖL , E., D ELALOYE , M. & G ENÇ , S. 1994. Magma-
G ÜNDOGDU , M. N. 1998. Petrology and geochemistry tism of northwestern Anatolia. International Volcano-
of potassic rocks in the Gölcük area (Isparta, SW logical Congress, IAVCEI 1994, Excursion Guide
Turkey): genesis of enriched alkaline magmas. (A3), 56.
Journal of Volcanology and Geothermal Research, B LUMENTHAL , M. M. 1956. Yüksek Bolkardaginin kuzey
85, 423–446. kenar bölgelerinin ve bati uzantilarinin jeolojisi
A LICI , P., T EMEL , A. & G OURGAUD , A. 2002. Pb– Nd– (Güney Anadolu Toroslari). Mineral Research and
Sr isotope and trace element geochemistry of Quatern- Exploration Institute of Turkey (MTA). Special Publi-
ary extension-related alkaline volcanism: a case study cation Series D, 7, 153.
of Kula region (western Anatolia, Turkey). Journal of B OZKURT , E. 2003. Origin of NE-trending basins in
Volcanology & Geothermal Research, 115, 487– 510. western Turkey. Geodinamica Acta, 14, 61– 81.
A L -L AZKI , A. I., S EBER , D., S ANDVOL , E., T ÜRKELLI , B OZKURT , E. 2004. Granitoid rocks of the southern
N., M OHAMAD , R. & B ARAZANGI , M. 2003. Tomo- Menderes Massif (Southwest Turkey): field evidence
graphic Pn velocity and anisotropy structure beneath for Tertiary magmatism in an extensional shear
the Anatolian plateau (eastern Turkey) and the sur- zone. International Journal of Earth Sciences, 93,
rounding regions. Geophysical Research Letters, 52–71.
30(24), 8043, doi: 10.1029/2003GL017391. B OZKURT , E. 2007. Extensional v. contractional origin for
A L -L AZKI , A. I., S ANDVOL , E., S EBER , D., B ARAZANGI , the southern Menderes shear zone, SW Turkey: tec-
M., T ÜRKELLI , N. & M OHAMMAD , R. 2004. Pn tomo- tonic and metamorphic implications. Geological
graphic imaging of mantle lid velocity and anisotropy Magazine, 144, 191–210.
at the junction of the Arabian, Eurasian and African B OZKURT , E. & P ARK , R. G. 1994. Southern Menderes
plates. Geophysical Journal International, 158, Massif: an incipient metamorphic core complex in
1024–1040. western Anatolia, Turkey. Journal of the Geological
A LTINLI , E. 1966. Dogu ve güneydogu Anadolu’nun jeo- Society, London, 151, 213– 216.
lojisi (in Turkish). Bulletin of the Mineral Research & B OZKURT , E. & S ATIR , M. 2000. The southern Menderes
Exploration Institute of Turkey, 66, 35–74. Massif (western Turkey): geochronology and exhuma-
A LTUNKAYNAK , Ş. 2004. Post collisional multistage tion history. Geological Journal, 35, 285– 296.
magmatism in northwest Anatolia (Turkey): Geochem- B OZTUG , D. 2000. S-I-A-type intrusive associations:
ical and isotopic study of Orhaneli magmatic associ- geodynamic significance of synchronism between
ations. International Geological Congress, Florence, metamorphism and magmatism in Central Anatolia,
Italy, August 20–28, 2004. Abstracts 1298. Turkey. In: B OZKURT , E., W INCHESTER , J. A. &
A LTUNKAYNAK , Ş. 2007. Collision-driven slab breakoff P IPER , J. D. A. (eds) Tectonics and Magmatism in
magmatism in northwestern Anatolia, Turkey. Turkey and the Surrounding Area. Geological
Journal of Geology, 115, 63– 82. Society, London, Special Publications, 173, 441–458.
A LTUNKAYNAK , Ş. & D ILEK , Y. 2006. Timing and nature B OZTUG , D., D EBON , F., I NAN , S., T UTKUN , Z. S., A VCI ,
of postcollisional volcanism in western Anatolia and N. & K ESGIN , Ö. 1997. Comparative geochemistry of
geodynamic implications. In: D ILEK , Y. & P AVLIDES , four plutons from the Cretaceous– Palaeogene central
S. (eds) Postcollisional Tectonics and Magmatism in Anatolian alkaline province (Divrigi region,
the Mediterranean Region and Asia. Geological Sivas, Turkey). Turkish Journal of Earth Sciences,
Society of America Special Paper, 409, 321– 351. 6, 96–115.
A NGUS , D. A., W ILSON , D. C., S ANDVOL , E. & N I , J. F. C ANDAN , O., Ç ETINKAPLAN , M., O BERHÄNSLI , R.,
2006. Lithospheric structure of the Arabian and R IMMELÉ , G. & A KAL , C. 2005. Alpine high-P/
Eurasian collision zone in eastern Turkey from low-T metamorphism of the Afyon Zone and impli-
S-wave receiver functions. Geophysical Journal Inter- cations for the metamorphic evolution of Western
national, 166, 1335–1346. Anatolia, Turkey. Lithos, 84, 102–124.
C ATLOS , E. J. & Ç EMEN , I. 2005. Monazite ages and the Geological Society of America Special Paper, 409,
evolution of the Menderes Massif, western Turkey. 1– 13.
International Journal of Earth Sciences, 94, 204–217. D ILEK , Y. & A LTUNKAYNAK , S. 2007. Cenozoic crustal
Ç ELIK , Ö. F. & C HIARADIA , M. 2008. Geochemical and evolution and mantle dynamics of post-collisional
petrological aspects of dike intrusions in the Lycian magmatism in western Anatolia. International
ophiolites (SW Turkey): a case study for the dike Geology Review, 49, 431–453.
emplacement along the Tauride Belt Ophiolites. Inter- D ILEK , Y. & A LTUNKAYNAK , S. 2008. Geochemical
national Journal of Earth Sciences, 97, 1151– 1164, and temporal evolution of Cenozoic magmatism in
doi: 10.1007/s00531-007-0204-0. western Turkey: mantle response to collision, slab
Ç ELIK , Ö. F., D ELALAOYE , M. & F ERAUD , G. 2006. breakoff, and lithospheric tearing in an orogenic belt.
Precise 40Ar– 39Ar ages from the metamorphic sole In: VAN H INSBERGEN , D. J. J., E DWARDS , M. A. &
rocks of the Tauride Belt Ophiolites, southern G OVERS , R. (eds) Collision and Collapse at the
Turkey: implications for the rapid cooling history. Africa-Arabia-Eurasia Subduction Zone. Geological
Geological Magazine, 143, 213–227. Society, London, Special Publications, 311,
Ç EMEN , I., C ATLOS , E. J., G ÖGÜS , O. & Ö ZERDEM , C. 213–233, doi: 10.1144/SP311.8.
2006. Postcollisional extensional tectonics and exhu- D ILEK , Y. & F LOWER , M. F. J. 2003. Arc-trench roll
mation of the Menderes massif in the Western Anatolia back and forearc accretion: 2. Model template
extended terrane. In: D ILEK , Y. & P AVLIDES , S. (eds) for Albania, Cyprus, and Oman. In: D ILEK , Y. &
Postcollisional tectonics and magmatism in the Medi- R OBINSON , P. T. (eds) Ophiolites in Earth History.
terranean region and Asia. Geological Society of Geological Society of London Special Publications,
America Special Paper, 409, 353– 379. 218, 43– 68.
Ç OBAN , H. & F LOWER , M. F. J. 2006. Mineral phase com- D ILEK , Y. & M OORES , E. M. 1990. Regional Tectonics of
positions in silica-undersaturated ‘leucite’ lamproites the Eastern Mediterranean ophiolites. In: M ALPAS , J.,
from the Bucak area, Isparta, SW Turkey. Lithos, 89, M OORES , E. M., P ANAYIOTOU , A. & X ENOPHON-
275– 299. TOS , C. (eds) Ophiolites, Oceanic Crustal Analogues,
C OLLINS , A. & R OBERTSON , A. H. F. 2003. Kinematic Proceedings of the Symposium ‘Troodos 1987’. The
evidence for Late Mesozoic–Miocene emplacement Geological Survey Department, Nicosia, Cyprus,
of the Lycian Allochthon over the Western Anatolide 295–309.
Belt, SW Turkey. Geological Journal, 38, 295– 310. D ILEK , Y. & T HY , P. 2006. Age and petrogenesis of
C ONCEIÇAO , R. V. & G REEN , D. H. 2004. Derivation of plagiogranite intrusions in the Ankara mélange,
potassic (shoshonitic) magmas by decompression central Turkey. Island Arc, 15, 44– 57.
melting of phlogopite þ pargasite lherzolite. Lithos, D ILEK , Y. & W HITNEY , D. L. 1997. Counterclockwise
72, 209–229, doi: 10.1016/j.lithos.2003.09.003. P-T-t trajectory from the metamorphic sole of a
D AVIS , J. H. & V ON B LANCKENBURG , F. 1995. Slab Neo-Tethyan ophiolite (Turkey). Tectonophysics,
breakoff: a model of lithosphere detachment and its 280, 295 –310.
test in the magmatism and deformation of collisional D ILEK , Y. & W HITNEY , D. L. 1998. Syn-metamorphic to
orogens. Earth and Planetary Science Letters, 129, neotectonic evolution of the Ecemis strike-slip fault
85–102. zone (Turkey). EOS Transactions, American Geo-
D ELALOYE , M. & B INGÖL , E. 2000. Granitoids from physical Union, 80, F915.
western and northwestern Anatolia: geochemistry and D ILEK , Y. & W HITNEY , D. L. 2000. Cenozoic crustal
modeling of geodynamic evolution. International evolution in central Anatolia: extension, magmatism
Geology Review, 42, 241– 268. and landscape development. In: P ANAYIDES , I.,
D EMIRTASLI , E., T URHAN , N., B ILGIN , A. Z. & S ELIM , X ENOPHONTOS , C. & M ALPAS , J. (eds) Proceedings
M. 1984. Geology of the Bolkar Mountains. In: of the Third International Conference on the Geology
T EKELI , O. & G ÖNCÜOGLU , M. C. (eds) Geology of of the Eastern Mediterranean. Geological Survey
the Taurus Belt, Proceedings of the International Sym- Department, September 1998, Nicosia, Cyprus,
posium, Geology of the Taurus Belt, 1984. Ankara, 183–192.
Mineral Research & Exploration Institute of Turkey D ILEK , Y., T HY , P., H ACKER , B. & G RUNDVIG , S. 1999a.
(MTA), Ankara, 125– 141. Structure and petrology of Tauride ophiolites and
D EWEY , J. F., H EMPTON , M. R., K IDD , W. S. F., mafic dike intrusions (Turkey): implications for the
S AROGLU , F. & S ENGÖR , A. M. C. 1986. Shortening Neo-Tethyan ocean. Bulletin of the Geological
of continental lithosphere: the neotectonics of Society of America, 111, 1192–1216.
Eastern Anatolia – a young collision zone. In: D ILEK , Y., W HITNEY , D. L. & T EKELI , O. 1999b. Links
C OWARD , M. P. & R IES , A. C. (eds) Collision Zone between tectonic processes and landscape morphology
Tectonics. Geological Society, London, Special Publi- in an alpine collision zone, south-central Turkey.
cations, 19, 3 –36. Annals of Geomorphology (Z. Geomorph. N.F.), 118,
D ILEK , Y. 1979. Geology of Gevas (Van Province, Turkey) 147–164.
and surrounding areas. Diploma thesis, Institute of D ILEK , Y., F URNES , H. & S HALLO , M. 2008. Geochem-
Applied Geology, Faculty of Science, Istanbul Univer- istry of the Jurassic Mirdita ophiolite (Albania) and
sity, Turkey, 1– 63. the MORB to SSZ evolution of a marginal basin
D ILEK , Y. 2006. Collision tectonics of the Mediterranean oceanic crust. Lithos, 100, 174– 209, doi: 10.1016/
region: causes and consequences. In: D ILEK , Y. & j.lithos.2007.06.026.
P AVLIDES , S. (eds) Postcollisional Tectonics and D ILEK , Y., I MAMVERDIYEV , N. & A LTUNKAYNAK , S.
Magmatism in the Mediterranean Region and Asia. 2010. Geochemistry and tectonics of Cenozoic
volcanism in the Lesser Caucasus (Azerbaijan) and F RANCALANCI , L., I NNOCENTI , F., M ANETTI , P. &
the Peri-Arabian region: Collision-induced mantle S AVASÇIN , M. Y. 2000. Neogene alkaline volcanism
dynamics and its magmatic fingerprint. International of the Afyon-Isparta area, Turkey: petrogenesis and
Geology Review, in press. geodynamic implications. Mineralogy & Petrology,
D OGLIONI , C., A GOSTINI , S., C RESPI , M., I NNOCENTI , 70, 285 –312.
F., M ANETTI , P., R IGUZZI , F. & S AVASÇIN , Y. G ARFUNKEL , Z. 1998. Constraints on the origin and
2002. On the extension in western Anatolia and the history of the Eastern Mediterranean basin. Tectono-
Aegean Sea. In: R OSENBAUM , G. & L ISTER , G. S. physics, 298, 5 –35.
(eds) Reconstruction of the Evolution of the G AUTIER , P., B OZKURT , E., H ALLOT , E. & D IRIK , K.
Alpine-Himalayan Orogen. Journal of the Virtual 2002. Dating the exhumation of a metamorphic
Explorer, 8, 169– 183. dome: geological evidence for pre-Eocene unroofing
D ÜZGÖREN - AYDIN , N. S. 2000. Post-collisional granitoid of the Nigde massif (Central Anatolia, Turkey).
magmatism: case study from the Yozgat Batholith, Geological Magazine, 139, 559– 576.
central Anatolia, Turkey. Proceedings of the Third G ELATI , R. 1975. Miocene marine sequence from Lake
International Conference on the Geology of the Van, eastern Turkey. Rivista italiana di paleontologia
Eastern Mediterranean. Geological Survey Depart- e stratigrafia, 81, 477–490.
ment, September 1998, Nicosia, Cyprus, 171– 181. G ERYA , T. V., Y UEN , D. A. & M ARESCH , W. V. 2004.
E LITOK , Ö & D RUPPEL , K. 2008. Geochemistry and tec- Thermomechanical modeling of slab detachment.
tonic significance of metamorphic sole rocks beneath Earth and Planetary Science Letters, 226, 101–116.
the Beysehir-Hoyran ophiolite (SW Turkey). Lithos, G ESSNER , K., C OLLINS , A. S., R ING , U. & G ÜNGÖR , T.
100, 322– 353. 2004. Structural and thermal history of poly-orogenic
E LMAS , A. & Y ILMAZ , Y. 2003. Development of an basement: U–Pb geochronology of granitoid rocks in
oblique subduction zone-Tectonic evolution of the the southern Menderes Massif, western Turkey.
Tethys suture zone in southeast Turkey. International Journal of the Geological Society, London, 161,
Geology Review, 45, 827– 840. doi: 10.2747/ 93–101.
0020-6814.45.9.827. G IANELLI , G., P ASERINI , P. & S GUAZZONI , G. 1972.
E RCAN , T., S ATIR , M. ET AL . 1985. Bati Anadolu Seno- Some observations on mafic and ultramafic complexes
zoyik volkanitlerine ait yeni kimyasal, izotopik ve north of the Bolkardag (Taurus, Turkey). Bolletino
radyometrik verilerin yorumu. Bulletin of the Geologi- della Societad Geologica Italiana, 91, 439– 488.
cal Society of Turkey, 28, 121–136 (in Turkish). G ÖK , R., T ÜRKELLI , N., S ANDVOL , E., S EBER , D. &
E RCAN , T., T URKECAN , A. & K ARABIYIKOGLU , M. B ARAZANGI , M. 2000. Regional wave propagation
1994. Neogene and Quaternary Volcanics of Cappado- in Turkey and surrounding regions. Geophysical
cia. International Volcanological Congress, Ankara, Research Letters, 27(3), 429 –432.
Turkey, 17– 22 September 1994, Post Congress Excur- G ÖK , R., S ANDVOL , E., T ÜRKELLI , N., S EBER , D. &
sion Guidebook, 1– 28. B ARAZANGI , M. 2003. Sn attenuation in the Anatolian
E RDOGAN , B., A KAY , E. & U GUR , M. S. 1996. Geology and Iranian plateau and surrounding regions. Geo-
of the Yozgat region and evolution of the collisional physical Research Letters, 30(24), doi: 10.1029/
Çankiri basin. International Geology Review, 38, 2003GL018020.
788–806. G ÖNCÜOGLU , M. C. & T URHAN , N. 1984. Geology of
E RINÇ , S. 1953. Dogu Anadolu Cografyasi. Istanbul the Bitlis Metamorphic Belt. In: T EKELI , O. &
Universitesi Yayini, 572 (in Turkish). G ÖNCÜOGLU , M. C. (eds) Geology of the Taurus
E RLER , A. & B AYHAN , H. 1995. Orta Anadolu granitoi- Belt. Proceedings of the International Symposium,
dlerinin genel degerlendirilmesi ve sorunları. Yerbilim- Ankara, Turkey, 1983, 237 –244.
leri, 17, 49–67 (in Turkish). G ÖNCÜOGLU , M. C., T OPRAK , V., K USCU , I., E RLER , A.
F ACCENNA , C., J OLIVET , L., P IROMELLI , C. & & O LGUN , E. 1991. Orta Anadolu Masifinin batı bölü-
M ORALLO , A. 2003. Subduction and the depth of con- münün jeolojisi,Bölüm 1: Güney Kesim. TPAO(Turkish
vection in the Mediterranean mantle. Journal of Petroleum Corporation) Report No. 2909, 134.
Geophysical Research, 108, 2099, doi: 1029/ G ÖRÜR , N., O KTAY , F., S EYMEN , I. & S ENGÖR , A. M. C.
2001JB001690. 1984. Palaeotectonic evolution of the Tuzgölü basin
F AYON , A. K., W HITNEY , D. L., T EYSSIER , C., G ARVER , complex, central Turkey: sedimentary record of a Neo-
J. I. & D ILEK , Y. 2001. Effects of plate convergence Tethyan closure. In: D IXON , J. E. & R OBERTSON ,
obliquity on timing and mechanisms of exhumation A. H. F. (eds) Geological Evolution of the Eastern
of a mid-crustal terrain, the Central Anatolian Crystal- Mediterranean. Geological Society, London, Special
line Complex. Earth and Planetary Science Letters, Publications, 17, 467–482.
192, 191– 205. G OVERS , R. & W ORTEL , M. J. R. 2005. Lithosphere tearing
F LOYD , P. A., G ÖNCÜOGLU , M. C., W INCHESTER , J. A. at STEP faults: response to edges of subduction zones.
& Y ALINIZ , K. 2000. Geochemical character and Earth and Planetary Science Letters, 236, 505–523.
tectonic environment of Neotethyan ophiolitic G ÜLEÇ , N. 1994, Rb–Sr isotope data from the Agaçören
fragments and metabasites in the Central Anatolian granitoid (East of Tuz Gölü): geochronological and
Crystalline Complex, Turkey. In: B OZKURT , E., genetical implications. Turkish Journal of Earth
W INCHESTER , J. A. & P IPER , J. D. A. (eds) Tectonics Sciences, 3, 39–43.
and Magmatism in Turkey and the Surrounding Area. G ÜLEÇ , N. & K ADIOGLU , Y. K. 1998. Relative involve-
Geological Society, London, Special Publications, ment of mantle and crustal components in the Agaçö-
173, 183– 202. ren granitoid (central Anatolia-Turkey): estimates
from trace element and Sr-isotope data. Chemie der complex: implications for the initiation of extensional
Erde – Geochemistry, 58, 23– 37. tectonics in western Turkey. Journal of Asian Earth
G ÜRSU , S., G ÖNCÜOGLU , M. C. & B AYHAN , H. 2004. Sciences, 23, 555– 566.
Geology and geochemistry of the Pre-early Cambrian J ACKSON , J. & M C K ENZIE , D. 1984. Active tectonics of
rocks in the Sandikli area: implications for the the Alpine-Himalayan belt between western Turkey
Pan-African evolution of NW Gondwanaland. Gond- and Pakistan. Geophysical Journal International, 77,
wana Research, 7, 923–935. 185–264.
H ANSMANN , W. & O BERLI , F. 1991. Zircon inheritance J OLIVET , L. 2001. A comparison of geodetic and finite
in an igneous rock suite from the southern Adamello strain pattern in the Aegean, geodynamic implica-
batholith (Italian Alps). Contributions to Mineralogy tions. Earth and Planetary Science Letters, 187,
and Petrology, 107, 501– 518. 95–104.
H ARRIS , N. B. W., K ELLEY , S. & O KAY , A. I. 1994. Post- J OLIVET , L. & B RUN , J.-P. 2008. Cenozoic geodynamic
collisional magmatism and tectonics in northwest evolution of the Aegean. International Journal of
Anatolia. Contributions to Mineralogy and Petrology, Earth Sciences, doi: 10.1007/s00531-008-0366-4.
117, 241– 252. J OLIVET , L. & F ACCENNA , C. 2000. Mediterranean exten-
H ELVACI , C. & G RIFFIN , W. L. 1984. Rb– Sr geochro- sion and the Africa-Eurasia collision. Tectonics, 19,
nology of the Bitlis massif, Avnik (Bingöl) area, 1095– 1106.
SE Turkey. In: D IXON , J. E. & R OBERTSON , J OLIVET , L., B RUN , J. P., G AUTIER , S., L ELLEMAND , S.
A. H. F. (eds) Geological Evolution of the Eastern & P ATRIAT , M. 1994. 3-D kinematics of extension
Mediterranean. Geological Society, London, Special in the Aegean from the early Miocene to the
Publications, 17, 403–413. present: insight from the ductile crust. Bulletin de la
H ETZEL , R. & R EISCHMANN , T. 1996. Intrusion age of Societe Geologique de France, 165, 195– 209.
Pan-African augen gneisses in the southern Menderes J OLIVET , L., F ACCENNA , C., G OFFÉ , B., B UROV , E. &
Massif and the age of cooling after Alpine ductile A GARD , P. 2003. Subduction tectonics and exhuma-
extensional deformation. Geological Magazine, 133, tion of high-pressure metamorphic rocks in the Medi-
565– 572. terranean orogen. American Journal of Science, 303,
I LBEYLI , N. 2004. Petrographic and geochemical charac- 353–409.
teristics of the Hamit Alkaline Intrusion in the K ADIOGLU , Y. K., D ILEK , Y. & F OLAND , K. A. 2006.
Central Anatolian Crystalline Complex, Turkey. Slab breakoff and syncollisional origin of the Late
Turkish Journal of Earth Sciences, 13, 269– 286. Cretaceous magmatism in the Central Anatolian Crys-
I LBEYLI , N. 2005. Mineralogical-geochemical constraints talline Complex, Turkey. In: D ILEK , Y. & P AVLIDES ,
on intrusives in central Anatolia, Turkey: tectono- S. (eds) Postcollisional tectonics and magmatism in the
magmatic evolution and characteristics of mantle Mediterranean region and Asia. Geological Society of
source. Geological Magazine, 142, 187– 207. America Special Paper, 409, 381–415, doi: 10.1130/
I LBEYLI , N., P EARCE , J. A., T HIRWALL , M. F. & 2006/2409(19).
M ITCHELL , J. G. 2004. Petrogenesis of collision- K ADIOGLU , Y. K., D ILEK , Y., G ULEÇ , N. & F OLAND ,
related plutonics in Central Anatolia, Turkey. Lithos, K. A. 2003. Tectonomagmatic evolution of bimodal
72, 163–182. plutons in the Central Anatolian Crystalline Complex,
I NNOCENTI , F., M AZZUOULI , G., P ASQUARE , F., R ADI- Turkey. Journal of Geology, 111, 671–690.
CATI DI B ROZOLA , F. & & V ILLARI , L. 1975. The K ÁRASON , H. 2002. Constraints on mantle convection
Neogene calcalkaline volcanism of Central Anatolia: from seismic tomography and flow modeling. PhD
geochronological data on Kayseri-Nigde area. Geo- thesis, Massachusetts Institute of Technology,
logical Magazine, 112, 349– 360. Cambridge, USA.
I NNOCENTI , F., M AZZUOULI , G., P ASQUARE , F., R ADI- K ÁRASON , H. & V AN D ER H ILST , R. D. 2001. Tomo-
CATI DI B ROZOLA , F. & V ILLARI , L. 1976. Evolution graphic imaging of the lowermost mantle with differ-
of volcanism in the area between the Arabian, Anato- ential times of refracted and defracted core phases
lian, and Iranian plates (Lake Van, Eastern Turkey). (PKP, Pdiff). Journal of Geophysical Research, 106,
Journal of Volcanological and Geothermal Research, 6569– 6587.
1, 103– 112. K EMPLER , D. & B EN -A VRAHAM , Z. 1987. The tectonic
I NNOCENTI , F., M AZZUOULI , G., P ASQUARE , F., R ADI- evolution of the Cyprean Arc. Annales Tectonicae, 1,
CATI DI B ROZOLA , F. & V ILLARI , L. 1982. Tertiary 58–71.
and Quaternary volcanism of the Erzurum-Kars area K ESKIN , M. 2003. Magma generation by slab steepening
(Eastern Turkey): geochronological data and geody- and breakoff beneath a subduction-accretion
namic evolution. Journal of Volcanological and complex: an alternative model for collision-related
Geothermal Research, 13, 223–240. volcanism in Eastern Anatolia, Turkey. Geophysical
I NNOCENTI , F., A GOSTINI , S., D I V INCENZO , G., Research Letters, 30, 8046, doi: 10.1029/2003GL018
D OGLIONI , C., M ANETTI , P., S AVASÇIN , M. Y. & 019.
T ONARINI , S. 2005. Neogene and Quaternary K ESKIN , M., P EARCE , J. A. & M ITCHELL , J. G. 1998.
volcanism in Western Anatolia: magma sources Volcano-stratigraphy and geochemistry of collision-
and geodynamic evolution. Marine Geology, 221, related volcanism on the Erzurum-Kars Plateau,
397– 421. North Eastern Turkey. Journal of Volcanology and
I SIK , V., T EKELI , O. & S EYITOGLU , G. 2004. The Geothermal Research, 85, 355 –404.
40
Ar/39Ar age of extensional ductile deformation and K ESKIN , M., P EARCE , J. A., K EMPTON , P. D. & G REEN-
granitoid intrusion in the northern Menderes core WOOD , P. 2006. Magma-crust interactions and magma
plumbing in a postcollisional setting: geochemical N EMCOK , M., P OSPISIL , L., L EXA , J. & D ONELICK , R. A.
evidence from the Erzurum-Kars volcanic plateau, 1998. Tertiary subduction and slab break-off model of
eastern Turkey. In: D ILEK , Y. & P AVLIDES , S. (eds) the Carpathian-Pannonian region. Tectonophysics,
Postcollisional tectonics and magmatism in the Medi- 295, 307–340.
terranean region and Asia. Geological Society of O BERHANSLI , R., B OUSQUET , R. & C ANDAN , O.
America Special Paper, 409, 475 –505. 2008. Bitlis Massif – East Anatolia Plateau connec-
K ÖKSAL , S., R OMER , R. L., G ÖNCÜOGLU , M. C. & tion? IGC-33, Session T-18, Oslo-Norway, August
T OKSOY -K ÖKSAL , F. 2004. Timing of post- 2008.
collisional H-type to A-type granitic magmatism: O KAY , A. I. 1984. Distribution and characteristics of the
U–Pb titanite ages from the Alpine central Anatolian northwest Turkish blueschists. In: D IXON , J. E. &
granitoids (Turkey). International Journal of Earth R OBERTSON , A. H. F. (eds) The Geological Evolution
Sciences (Geologische Rundschau), 93, 974– 989. of the Eastern Mediterranean Region. Geological
K ÖPRÜBASI , N. & A LDANMAZ , E. 2004. Geochemical Society, London, Special Publications, 17, 455– 466.
constraints on the petrogenesis of Cenozoic I-type O KAY , A. I. 2002. Stratigraphic and metamorphic inver-
granitoids in Northwest Anatolia, Turkey: evidence sions in the central Menderes massif. A new structural
for magma generation by lithospheric delamination model. International Journal of Earth Sciences, 91,
in a post-collisional setting. International Geology 173– 178.
Review, 46, 705– 729. O KAY , A. I. & S ATIR , M. 2000. Coeval plutonism and
K UMRAL , M., Ç OBAN , H., G EDIKOGLU , A. & K ILINÇ , A. metamorphism in a latest Oligocene metamorphic
2006. Petrology and geochemistry of augite trachytes core complex in northwest Turkey. Geological Maga-
and porphyritic trachytes from the Gölcük volcanic zine, 137, 495–516.
region, SW Turkey: a case study. Journal of Asian O KAY , A. I., S ATIR , M., M ALUSKI , H., S IYAKO , M.,
Earth Sciences, 27, 707 –716. M ONIE , P., M ETZGER , R. & A KYÜZ , S. 1996.
L E P ICHON , X. & A NGELIER , J. 1979. The Hellenic arc Paleo- and Neo-Tethyan events in northwest Turkey:
and trench system: a key to the evolution of the geological and geochronological constraints. In: Y IN ,
Eastern Mediterranean area. Tectonophysics, 60, A. & H ARRISON , M. T. (eds) Tectonics of Asia.
1–42. Cambridge University Press, 420–441.
L E P ICHON , X., L ALLEMANT , S. J., C HAMOT -R OOKE , O KAY , A. I., H ARRIS , N. B. W. & K ELLEY , S. P. 1998.
N., L EMEUR , D. & P ASCAL , G. 2003. The Mediterra- Exhumation of blueschists along a Tethyan suture in
nean Ridge backstop and the Hellenic nappes. Marine northwest Turkey. Tectonophysics, 285, 275– 299.
Geology, 186, 111–125. Ö NEN , P. 2003. Neotethyan ophiolitic rocks of the Anato-
L IU , M., C UI , X. & L IU , F. 2004. Cenozoic rifting and vol- lides of NW Turkey and comparison with Tauride
canism in eastern China: a mantle dynamic link to the ophiolites. Journal of the Geological Society,
Indo-Asian collision? Tectonophysics, 393, 29– 42, London, 160, 947–962.
doi: 10.1016/j.tecto.2004.07.029. O NER , Z. & D ILEK , Y. 2007. Depositional and tec-
M AURY , R. C., F OURCADE , S. ET AL . 2000. Post- tonic evolution of the Late Cenozoic Alasehir sup-
collisional Neogene magmatism of the Mediterranean radetachment basin, western Anatolia (Turkey).
Maghreb margin: a consequence of slab breakoff. Geological Society of America Abstracts with
Earth and Planetary Sciences, 331, 159–173. Programs, 39(6), 228.
M CCLUSKY , S., B ALASSANIAN , S. ET AL . 2000. Global Ö NEN , A. P. & H ALL , R. 1993. Ophiolites and related
Positioning System constraints on plate kinematics metamorphic rocks from the Kütahya region, north-
and dynamics in the eastern Mediterranean and Cauca- west Turkey. Geological Journal, 28, 399– 412.
sus. Journal of Geophysical Research, 105, Ö ZER , S., S ÖZBILIR , H., Ö ZKAR , I., T OKER , V. & S ARI ,
5695–5719. B. 2001. Stratigraphy of Upper Cretaceous-Palaeogene
M ICHARD , A., W HITECHURCH , H., R ICOU , L.-E., M ON- sequences in the southern and eastern Menderes Massif
TIGNY , R. & Y AZGAN , E. 1984. Tauric subduction (Western Turkey). International Journal of Earth
(Malatya-Elazig provinces) and its bearing on tec- Sciences, 89, 852–866.
tonics of the Tethyan realm in Turkey. In: D IXON , Ö ZGÜL , N. 1976. Some geological aspects of the Taurus
J. E. & R OBERTSON , A. H. F. (eds) The Geological orogenic belt (Turkey). Bulletin of the Geological
Evolution of the Eastern Mediterranean Region. Geo- Society of Turkey, 19, 65–78.
logical Society, London, Special Publications, 17, Ö ZGÜL , N. 1984. Stratigraphy and tectonic evolution of
361–374. the Central Taurides. In: T EKELI , O. & G ÖNCÜOĞLU ,
M O , X., Z HAO , Z. ET AL . 2006. Petrology and geochemis- M. C. (eds) Geology of the Taurus Belt, Proceedings
try of postcollisional volcanic rocks from the Tibetan of the International Symposium on the Geology of the
plateau: implications for lithospheric heterogeneity Taurus Belt, 1983. Ankara, Turkey, Mineral Research
and collision-induced asthenospheric mantle flow. In: and Exploration Institute of Turkey, Ankara, 77–90.
D ILEK , Y. & P AVLIDES , S. (eds) Postcollisional tec- Ö ZGÜL , N. & T URSUCU , A. 1984. Stratigraphy of the Mes-
tonics and magmatism in the Mediterranean region ozoic carbonate sequence of the Munzur Mountains
and Asia. Geological Society of America Special (Eastern Taurides). In: T EKELI , O. & G ÖNCÜOĞLU ,
Paper, 409, 507–530. M. C. (eds) Geology of the Taurus Belt. Proceedings
M OLNAR , P. & O LIVER , J. 1969. Lateral variations of of the International Symposium, Geology of the
attenuation in the upper mantle and discontinuities in Taurus Belt, 1984. Ankara, Mineral Research &
the lithosphere. Journal of Geophysical Research, 74, Exploration Institute of Turkey (MTA), Ankara,
2648–2682. 173– 180.
Ö ZKAYA , I. 1978. Stratigraphy of Ergani-Maden area, SE Holocene alkali basalt province and the adjacent
Turkey. Bulletin of the Geological Society of Turkey, normal-faulting graben. Geological Magazine, 133,
21, 129–139. 275–283.
P AMIC , J., B ALEN , D. & H ERAK , M. 2002. Origin and R ICOU , L. E., A RGYRIADIS , I. & M ARCOUX , J. 1975.
geodynamic evolution of Late Paleogene magmatic L’axe calcaire du Taurus, un alignement de fenetres
associations along the Periadriatic-Sava-Vardar mag- arabo-africaines sous des nappes radiolaritiques,
matic belt. Geodinamica Acta, 15, 209–231. ophiolitiques et métamorphiques. Bulletin de la
P ARLAK , O. 2006. Geodynamic significance of granitoid Société de Geologie de France, 16, 107–111.
magmatism in the southeast Anatolian orogen: R IMMELÉ , G., O BERHANSLI , R., G OFFE , B., J OLIVET , L.,
geochemical and geochronological evidence from C ANDAN , O. & Ç ETINKAPLAN , M. 2003. First evi-
Göksun-Afsin (Kahramanmaras, Turkey) region. dence of high-pressure metamorphism in the ‘Cover
International Journal of Earth Sciences, 95, 609–627. Series’ of the Southern Menderes Massif. Tectonic
P ARLAK , O. & D ELALOYE , M. 1999. Precise 40Ar-39Ar and metamorphic implications for the evolution of
ages from the metamorphic sole of the Mersin ophiolite SW Turkey. Lithos, 71, 19–46.
(Southern Turkey). Tectonophysics, 301, 145– 158. R ING , U. & C OLLINS , A. S. 2005. U–Pb SIMS dating
P ARLAK , O., D ELALOYE , M. & B INGÖL , E. 1996. of synkinematic granites: timing of core-complex for-
Mineral chemistry of ultramafic and mafic cumulates mation in the northern Anatolide belt of western
as an indicator of the arc-related origin of the Mersin Turkey. Journal of the Geological Society, London,
ophiolite (southern Turkey). Geologische Rundschau, 162, 289 –298.
85, 647–661. R ING , U. & L AYER , P. W. 2003. High-pressure meta-
P ARLAK , O., H ÖCK , V. & D ELALOYE , M. 2002. The supra- morphism in the Aegean, eastern Mediterranean:
subduction zone Pozanti-Karsanti ophiolite, southern Underplating and exhumation from the Late Cretac-
Turkey: evidence for high-pressure crystal fraction- eous until the Miocene to Recent above the retreating
ation of ultramafic cumulates. Lithos, 65, 205–224. Hellenic subduction zone. Tectonics, 22(3), 1022,
P ASQUARÉ , G., P OLI , S., V EZZOLI , L. & Z ANCHI , A. doi: 10.1029/2001TC001350.
1988. Continental arc volcanism and tectonic setting R ING , U., J OHNSON , C., H ETZEL , R. & G ESSNER , K.
in Central Anatolia, Turkey. Tectonophysics, 146, 2003. Tectonic denudation of a Late Cretaceous-
217– 230. Tertiary collisional belt: regionally symmetric
P EARCE , J. A. & B ENDER , J. F. ET AL . 1990. Genesis of cooling patterns and their relation to extensional
collision volcanism in eastern Anatolia, Turkey. faults in the Anatolide belt of western Turkey. Geologi-
Journal of Volcanology and Geothermal Research, cal Magazine, 140, 421–441.
44, 189–229. R OBERTSON , A. H. F. 1998. Tectonic significance of
P E -P IPER , G. & P IPER , D. J. W. 1994. Miocene magnesian the Eratosthenes Seamount: a continental fragment in
andesites and dacites, Evia, Greece: adakites associ- the process of collision with a subduction zone in the
ated with subducted slab detachment and extension. eastern Mediterranean (Ocean Drilling Program Leg
Lithos, 31, 125– 140. 160). Tectonophysics, 298, 63– 82.
P E -P IPER , G. & P IPER , D. J. W. 2006. Unique features of R OBERTSON , A. H. F. & D IXON , J. E. 1984. Introduction:
the Cenozoic igneous rocks of Greece. In: D ILEK , Y. & aspects of the geological evolution of the eastern Med-
P AVLIDES , S. (eds) Postcollisional Tectonics and iterranean: In: D IXON , J. E. & R OBERTSON , A. H. F.
Magmatism in the Mediterranean Region and Asia. (eds) The Geological Evolution of the Eastern Mediter-
Geological Society of America Special Paper, 409, ranean. Geological Society, London, Special Publi-
259– 282. cations, 17, 1– 74.
P IROMALLO , C. & M ORELLI , A. 2003. P wave tomogra- R OGERS , A. J., N I , J. F. & H EARN , T. M. 1995. Propoga-
phy of the mantle under the Alpine-Mediterranean tion characteristics of short-period Sn And Lg in The
area. Journal of Geophysical Research, 108, 2065, Middle East. Bulletin of The Seismological Society of
doi: 10.1029/2002JB001757. America, 87, 396 –413.
P URVIS , M. & R OBERTSON , A. H. F. 2004. A pulsed R OTSTEIN , Y. & K AFKA , A. L. 1982. Seismotectonics
extension model for the Neogene-Recent E-W- of the southern boundary of Anatolia, eastern Medi-
trending Alasehir Graben and the NE-SW-trending terranean region; subduction, collision, and arc
Selendi and Gördes Basins, western Turkey. Tectono- jumping. Journal of Geophysical Research, 87,
physics, 391, 171– 201. 7694– 7706.
R EILINGER , R. E., M C C LUSKY , S. C. & O RAL , M. B. R USSO , R. M., D ILEK , Y. & F LOWER , M. F. J. 2001.
1997. Global positioning system measurements of Collision-driven mantle flow and crustal response in
present-day crustal movements in the Arabia-Africa- the eastern Mediterranean region during the late Ceno-
Eurasia plate collision zone. Journal of Geophysical zoic. 4th International Turkish Geology Symposium,
Research, 102, 9983–9999. Work in Progress on the Geology of Turkey and Its Sur-
R EILINGER , R. E., M C C LUSKY , S. C. ET AL . 2006. roundings. Adana, Turkey, 24–28 September 2001, 98.
GPS constraints on continental deformation in the S ANDVOL , E., A L -D AMEGH , K. ET AL . 2001. Tomographic
Africa-Arabia-Eurasia continental collision zone and imaging of Lg and Sn Propagation in the Middle East.
implications for the dynamics of plate interactions. Pure and Applied Geophysics, 158, 1121– 1163.
Journal of Geophysical Research, 111, V05411, doi: S ANDVOL , E., T ÜRKELLI , N. & B ARAZANGI , M. 2003a.
10.1029/2005JB004051. The Eastern Turkey Seismic Experiment: the study of a
R ICHARDSON -B UNBURY , J. M. 1996. The Kula young continent-continent collision. Geophysical
volcanic field, western Turkey: the development of a Research Letters, 30(24), doi: 10.1029/2003GL018912.
S ANDVOL , E., T ÜRKELLI , N. ET AL . 2003b. Shear wave T ATAR , M., H ATZFELD , D., M ARTINOD , J., W ALPERS-
splitting in a young continent-continent collision: an DORF , A., G HAFORI -A SHTIANY , M. & C HÉRY , J.
example from eastern Turkey. Geophysical Research 2002. The present day deformation of the central
Letters, 30(24), doi: 10.1029/2003GL017390. Zagros from GPS measurements. Geophysical Research
S AROGLU , F., E MRE , Ö & K USÇU , I. 1992. Active fault Letters, 29, 1927, doi: 10.1029/2002GL015427.
map of Turkey (3 sheets, 1:100,000 scale). General T EKELI , O. 1981. Subduction complex of pre-Jurassic age,
Directorate of Mineral Research and Exploration, northern Anatolia, Turkey. Geology, 9, 68–72.
MTA, Ankara, Turkey. T HOMSON , S. N. & R ING , U. 2006. Thermochronologic
S AVASÇIN , M. Y. & O YMAN , T. 1998. Tectono-magmatic evaluation of postcollision extension in the Anatolide
evolution of alkaline volcanics at the Kirka-Afyon- orogen, western Turkey. Tectonics, 25, TC3005,
Isparta structural trend, SW Turkey. Turkish Journal doi: 10.1029/2005TC001833.
of Earth Sciences, 7, 201–214. T OPRAK , V., K ELLER , J. & S CHUMACHER , R. 1994.
S CHLIESTEDT , M., A LTHERR , R. & M ATTHEWS , A. Volcano-tectonic features of the Cappadocian Volca-
1987. Evolution of the Cycladic Crystalline nic Province. International Volcanological Congress,
Complex: petrology, isotope geochemistry and geo- Ankara, Turkey, 17–22 September 1994, Post Con-
chronology. In: H ELGESON , H. C. (ed.) Chemical gress Excursion Guidebook, 58.
Transport in Metasomatic Processes. Reidel, V AN D ER K AADEN , G. 1966. The significance and distri-
Dordrecht, 389–428. bution of glaucophane rocks in Turkey. Bulletin of the
S EGHEDI , I., D OWNES , H. ET AL . 2004. Neogene- Mineral Research and Exploration Institute of Turkey,
Quaternary magmatism and geodynamics in the 67, 37– 67.
Carpathian-Pannonian region: a synthesis. Lithos, 72, V AN H INSBERGEN , D. J. J., H AFKENSCHEID , E.,
117–146, doi: 10.1016/j.lithos.2003.08.006. S PAKMAN , W., M EULENKAMP , J. E. & W ORTEL , R.
S ENGÖR , A. M. C. & Y ILMAZ , Y. 1981. Tethyan evol- 2005. Nappe stacking resulting from subduction of
ution of Turkey, a plate tectonic approach. Tectono- oceanic and continental lithosphere below Greece.
physics, 75, 181– 241. Geology, 33, 325–328.
S ENGÖR , A. M. C., Ö ZEREN , S., G ENÇ , T. & Z OR , E. V ON B LANCKENBURG , F. & D AVIES , J. H. 1995. Slab
2003. East Anatolian high plateau as a mantle- breakoff: a model for syncollisional magmatism and
supported, north-south shortened domal structure. tectonics in the Alps. Tectonics, 14, 120–131.
Geophysical Research Letters, 30, 8045, doi: W DOWINSKI , S., B E -A VRAHAM , Z., A RVIDSSON , R. &
10.1029/2003GL017858. E KSTRÖM , G. 2006. Seismotectonics of the Cyprian
S EYITOGLU , G., A NDERSON , D., N OWELL , G. & S COTT , Arc. Geophysical Journal International, 164,
B. 1997. The evolution from Miocene potassic to Qua- 176– 181.
ternary sodic magmatism in western Turkey: impli- W ESTAWAY , B. 1994. Present-day kinematics of the
cations for enrichment processes in the lithospheric Middle East and eastern Mediterranean. Journal of
mantle. Journal of Volcanology and Geothermal Geophysical Research, 99, 12071–12090.
Research, 76, 127– 147. W HITNEY , D. L. & D ILEK , Y. 1997. Core complex devel-
S EYMEN , I. 1984. Kırsehir metamorfitlerinin jeolojik opment in central Anatolia, Turkey. Geology, 25,
evrimi: Ketin Simpozyumu. Türkiye Jeoloji Kurumu 1023– 1026.
Yayını, 133–148 (in Turkish). W HITNEY , D. L. & D ILEK , Y. 1998. Metamorphism
S HERLOCK , S., K ELLEY , S. P., I NGER , S., H ARRIS , N. & during crustal thickening and extension in central Ana-
O KAY , A. I. 1999. 40Ar-39Ar and Rb–Sr geochronol- tolia: the Nigde metamorphic core complex. Journal of
ogy of high-pressure metamorphism and exhumation Petrology, 39, 1385– 1403.
history of the Tavsanli Zone, NW Turkey. Contri- W HITNEY , D. L.,T EYSSIER , C., D ILEK , Y. &F AYON , A. K.
butions to Mineralogy and Petrology, 137, 46–58. 2001. Metamorphism of the central Anatolian crystal-
S PAKMAN , W., W ORTEL , M. J. R. & V LAAR , N. J. 1988. line complex, Turkey: influence of orogen-normal
The Hellenic subduction zone: a tomographic image collision vs. wrench-dominated tectonics on P-T-t.
and its geodynamic implications. Geophysical paths. Journal of Metamorphic Geology, 19, 411–432.
Research Letters, 15, 60–63. W IDIYANTORO , S., V AN D ER H ILST , R. D. & W ENZEL ,
S PAKMAN , W., V AN D ER L EE , S. & V AN D ER H ILST , R. F. 2004. Deformation of the Aegean slab in the mantle
1993. Travel-time tomography of the European- transition zone. International Journal of Tomography
Mediterranean mantle down to 1400 km. Physics of & Statistics, D04, 1 –14.
the Earth and Planetary Interiors, 79, 3–74. W ILLIAMS , H. M., T URNER , S. P., P EARCE , J. A.,
T AN , O. & T AYMAZ , T. 2006. Active tectonics of the K ELLEY , S. P. & H ARRIS , N. B. W. 2004. Nature of
Caucasus: Earthquake source mechanisms and the source regions for post-collisional, potassic mag-
rupture histories obtained from inversion of teleseismic matism in southern and northern Tibet from geochem-
body waveforms. In: D ILEK , Y. & P AVLIDES , S. (eds) ical variations and inverse trace element modelling.
Postcollisional Tectonics and Magmatism in the Med- Journal of Petrology, 45, 555–607.
iterranean Region and Asia. Geological Society of W ORTEL , M. J. R. & S PAKMAN , W. 1992. Structure and
America Special Paper, 409, 531 –578. dynamics of subducted lithosphere in the Mediterra-
T ANKUT , A., D ILEK , Y. & Ö NEN , P. 1998. Petrology nean region. Proceedings of the Koninklijke Neder-
and geochemistry of the Neo-Tethyan volcanism landse Akademie van Wetenschappen, 95, 325–347.
as revealed in the Ankara Melange, Turkey. Journal W ORTEL , M. J. R. & S PAKMAN , W. 2000. Subduction and
of Volcanological & Geothermal Research, 85, slab detachment in the Mediterranean-Carpathian
265–284. region. Science, 290, 1910– 1917.
Y AGMURLU , F., S AVASÇIN , M. Y. & E RGUN , M. 1997. Quaternary. Geological Society of America Abstracts
Relation of alkaline volcanism and active tectonism with Programs, 34(6), 179.
within the evolution of Isparta Angle, SW Turkey. Y ILMAZ , Y., S AROGLU , F. & G ÜNER , Y. 1987. Initiation
Journal of Geology, 105, 717– 728. of the neomagmatism in East Anatolia. Tectonophy-
Y AZGAN , E. 1984. Geodynamic evolution of the Eastern sics, 134, 177– 199.
Taurus region. In: T EKELI , O. & G ÖNCÜOĞLU , Y ILMAZ , Y., T ÜYSÜZ , O., Y IGITBAS , E., G ENÇ , S. C. &
M. C. (eds) Geology of the Taurus Belt. Proceedings S ENGÖR , A. M. C. 1997. Geology and tectonic evol-
of the International Symposium, Geology of the ution of the Pontides. In: R OBINSON , A. G. (ed.)
Taurus Belt. 1984, Ankara, Mineral Research & Regional and Petroleum Geology of the Black Sea
Exploration Institute of Turkey (MTA), Ankara, and Surrounding Region. AAPG Memoir, 68, 183– 226.
199– 208. Y ILMAZ , Y., G ENÇ , S. C., K ARACIK , Z. & A LTUNKAY-
Y ILMAZ , Y. 1985. Geology of the Cilo ophiolite: an NAK , S. 2001. Two contrasting magmatic associations
ancient ensimatic island arc fragment on the Arabian of NW Anatolia and their tectonic significance.
platform, SE Turkey. Ofioliti, 10, 457– 484. Journal of Geodynamics, 31, 243– 271.
Y ILMAZ , Y. 1990. Comparison of young volcanic associ- Z ITTER , T. A. C., W OODSIDE , J. M. & M ASCLE , J. 2003.
ations of western and eastern Anatolia under compres- The Anaximander Mountains: a clue to the tectonics of
sional regime; a review. Journal of Volcanology and Southwest Anatolia. Geological Journal, 38, 375– 394,
Geothermal Research, 44, 69–87. doi: 10.1002/gj.961.
Y ILMAZ , Y. 1993. New evidence and model on the evol- Z OR , E., S ANDVOL , E., G ÜRBÜZ , C., T ÜRKELLI , N.,
ution of the Southeast Anatolia orogen. Geological S EBER , D. & B ARAZANGI , M. 2003. The crustal struc-
Society of America Bulletin, 105, 251–271. ture of the East Anatolian plateau (Turkey) from recei-
Y ILMAZ , Y. 2002. Tectonic evolution of western Anato- ver functions. Geophysical Research Letters, 30(24),
lian extensional province during the Neogene and TUR7.1–TUR7.4, doi: 10.1029/2003GL018192.
The evolution of the Uralian orogen
VICTOR N. PUCHKOV
K. Marx st. 16/2, Institute of Geology, Ufimian Scientific Centre,
Russian Academy of Sciences, 450 000 Ufa, Russia
(e-mail: puchkv@anrb.ru)
Abstract: The Uralian orogen is located along the western flank of a huge (.4000 km long)
intracontinental Uralo-Mongolian mobile belt. The orogen developed mainly between the Late
Devonian and the Late Permian, with a brief resumption of orogenic activity in the Lower
Jurassic and Pliocene–Quaternary time. Although its evolution is commonly related to the
Variscides of Western Europe, its very distinctive features argue against a simple geodynamic
connection. To a first order, the evolution of Uralian orogen shows similarities with the
‘Wilson cycle’, beginning with epi-continental rifting (Late Cambrian– Lower Ordovician)
followed by passive margin (since Middle Ordovician) development, onset of subduction and
arc-related magmatism (Late Ordovician) followed by arc– continent collision (Late Devonian
in the south and Early Carboniferous in the north) and continent– continent collision (beginning
in the mid-Carboniferous). In detail, however, the Uralides preserve a number of rare features.
Oceanic (Ordovician to Lower Devonian) and island– arc (Ordovician to Lower Carboniferous)
complexes are particularly well preserved as is the foreland belt in the Southern Urals, which
exhibits very limited shortening of deformed Mesoproterozoic to Permian sediments. Geophysi-
cal studies indicate the presence of ‘cold’, isostatically equilibrated root. Other characteristic
features include a Silurian platinum-rich belt of subduction-related layered plutons, a simul-
taneous development of orogenic and rift-related magmatism, a succession of collisions that
are both diachronous and oblique, and a single dominant stage of transpressive deformation
after the Early Carboniferous. The end result is a pronounced bi-vergent structure. The Uralides
are also characterized by Meso-Cenozoic post-orogenic stage and plume-related tectonics in
Ordovician, Devonian and especially Triassic time. The evolution of the Uralides is consistent
with the development and destruction of a Palaeouralian ocean to form part of a giant Uralo-
Mongolian orogen, which involved an interaction of cratonic Baltica and Siberia with a
young and rheologically weak Kazakhstanian continent. The Uralides are characterized by
protracted and recurrent orogenesis, interrupted in the Triassic by tectonothermal activity
associated with the Uralo-Siberian superplume.
Introduction Early Palaeozoic, these continents were separated

by the Palaeouralian and Central Asian oceans.
The last general review of the structural and tectonic The continents and oceans were partly inherited
evolution of the Uralian orogen was published in from the Proterozoic (Vernikovsky et al. 2004;
English more than 10 years ago (Puchkov 1997). Puchkov 2005). The Central Asian (or Palaeo-
Since then, the stimulus created by the international Asian) ocean existed before the Late Neoproterozoic
EUROPROBE Uralides Project has resulted in con- as the portion of the Palaeo-Pacific Ocean, which
siderable advances in the understanding of the surrounded a considerable part of the Siberian conti-
geology of the Urals. However, most publications nent. A complete separation of this continent from
in English are concerned with the evolution of the Rodinia by 750 Ma resulted in the birth of the
Southern and Middle Urals (Brown et al. 2006a, Palaeo-Asian ocean (Li et al. 2008).
2008, and references therein). This publication pro- The E –NE margin of Baltica was modified by
vides an overview of the tectonic evolution of the the Ediacarian deformational and accretionary
Uralian orogen as a whole, incorporating a wealth events of the Timanian orogeny (Puchkov 1997;
of recently published data, and compares this evol- Gee et al. 2006). A close resemblance of Timanides
ution with that of the European Variscides and and Cadomides pointed out by the author (Puchkov
Mesozoic-Cenozoic orogens. 1997) has led to an idea of their immediate lateral
The Uralian orogen sensu stricto partly connection, supporting a notion of a supercontinent
coincides geographically with the young, neotec- (Pannotia, Dalziel 1992; or Panterra, Puchkov
tonic (Pliocene– Quaternary) Urals mountains, and 2000), welded by Ediacarian orogenies. Several var-
occurs between three former, Palaeozoic continents: iants of such a connection have been suggested,
Baltica, Kazakhstania and Siberia (Fig. 1). In the depending on what side of Baltica was thought to
162 V. N. PUCHKOV
Fig. 1. Position and linkages of the Urals in the structure of the Central Eurasia.
THE EVOLUTION OF THE URALIAN OROGEN 163
be attached to Gondwana (e.g. Linneman et al.

1998; Puchkov 2000; Cocks & Torsvik 2006).
Unfortunately the APWP of Baltica is still very
poorly constrained by existing palaeomagnetic
data and additional work is required to distinguish
between these hypotheses.
The Timanian/Cadomian orogeny was followed
by Late Cambrian –Early Ordovician rifting and
passive margin development (Puchkov 2005;
Pease et al. 2008). These events led to the develop-
ment of the Palaeouralian ocean, which contained
some younger microcontinents (the continent that
rifted-away from Baltica is not identified yet). The
Mid- to Late Palaeozoic history of the Uralian
orogen can be described in terms of formation of
an island-arc crust (Late Ordovician –Devonian)
and followed by Late Devonian –Early Carbonifer-
ous accretion of the island-arcs and other microcon-
tinents with the continental margin of Baltica, which
by that time had merged with Laurentia to form
Laurussia (Brown et al. 1997, 2006a–c). Kazakh- Fig. 2. A tentative reconstruction of the Central Eurasia
stania formed in the Ordovician–Silurian as a for the Late Devonian (after YUGGEO2002, simplified).
result of subduction-driven accretion of crust
around small Neoproterozoic microcontinents
(Puchkov 1996a). Its collision with the terranes pre- into the Southern Tyan-Shan have been thwarted
viously accreted to the margins of Laurussia and by the higher degree of shortening and lack of
Siberia started in the Late Bashkirian, after the inter- Devonian island arc development in the latter fold-
vening oceanic crust was completely subducted. belt. However, palaeocontinental reconstructions
This continent–continent collision continued into and modern structural connections (Figs 1 & 2)
Permian time, and resulted in the formation of suggest a genetic linkage. The evolution of Southern
huge foldbelt, variously known as Uralo-Mongolian Tyan-Shan is related to the subduction of the
(Muratov 1979) or Uralo-Okhotskian (Khain 2001) Turkestanian ocean (Burtman 2006), that in the
belt, a fundamental event in the assembly of Ordovician to Carboniferous time was a direct
Pangaea. The narrow western flank of the orogen continuation of the Palaeouralian ocean (Puchkov
is known as the Uralides, and the more extensive 2000; Didenko et al. 2001; YUGGEO 2002; Biske
central and eastern flank of the belt is called the 2006). The Middle to Late Palaeozoic evolution of
Altaides (Sengör et al. 1993). However the position- the Uralian and Southern Tyan-Shan continental
ing of the Kazakhstanian continent as a separate margins has a striking resemblance (Puchkov
Mid-Palaeozoic structure (Figs 1 & 2) suggests 1996a). The end of orogenic activity in the Southern
that the Kazakhstanides should be excluded from Tyan-Shan is accompanied by Upper Permian and
the Altaides. Lower Triassic continental molasse followed by
An attempt to formalize the term ‘Uralides’ was the development of a stable platform in the whole
made by Puchkov (2003), who noted that the north- Tyan-Shan, which lasted until strong intra-plate
ern continuation of the western zones of the Urals is deformation of the Pliocene– Quaternary age
represented by the Pay-Khoy-Novozemelian propagated from the Alpine-Himalayan orogenic
foldbelt, which was the result of the younger system, resulting in renewed mountain building
(Early Jurassic) collision between Laurussia and (Burtman 2006; Biske 2006; Trifonov 2008).
Siberia. More southern parts of the Urals were
also deformed by this collision, though the intensity Comparison between the Uralides,
and structural expression are not as prominent. Variscides and Appalachians
To the south, Late Palaeozoic collisional struc-
tures of the Urals continue in the south-southwest- The development of the Ural orogenic belt has tra-
wards vergent fold-and-thrust belt of the Southern ditionally been interpreted to be one of the Late
Tyan-Shan. Like the Urals, this Late Palaeozoic Palaeozoic Variscide (or Hercynide) orogenic belts
belt is bordered to the NE and north by the Early (Shatsky 1965; International Committee 1982).
Caledonides of Kazakhstania, where the Northern However, results of geological studies of the last
Tyan-Shan occurs (Fig. 1). Attempts to trace Late few decades, enhanced by EUROPROBE, have
Palaeozoic structural features from the Uralides shown such fundamental differences with the
164 V. N. PUCHKOV
Variscides that this interpretation is no longer con- followed by Triassic terrigenous and evaporitic
sidered valid (Brown et al. 2006a). Instead the deposits in graben structures. The Permian–Triassic
Urals is treated as a separate orogenic belt, the events reflect post-orogenic extension (Schwab
main part of the Uralides. 1984; Ziegler 1999). The Alleghanian orogeny in
In western Europe, tectonic events leading to the the southern Appalachians and Ouachita continued,
development of the Variscan (also known as Hercy- albeit in a weakened manner, into the Early Permian
nian) orogenic belts (sensu stricto) occurred (Engelder 2007). In the Triassic, the Variscide and
between the Famennian and Late Carboniferous Appalachian orogenic belts were characterized by
(Fig. 3). Famennian flysch deposits, thought to rep- horsts and grabens that strongly influenced the for-
resent the onset of orogenesis, continued until the mation of platform oil and gas deposits, especially
end of the Lower Carboniferous, and were followed in the North Sea (Lützner et al. 1979; Beutler
by molasse deposits which continued until the end 1979; Schwab 1984; Ziegler 1999; Khain 2001;
of the Carboniferous. Devono-Carboniferous tecto- Franke 2000).
nothermal activity accompanied a pre-collisional The Uralides have a much more protracted
closure of the Rheic (Saxo-Thuringian) ocean, history of orogenesis than the Variscides, with tec-
collisional deformation, granitoid intrusions and tonic events continuing into the Jurassic. In the
metamorphism (including HP–LT type; Franke southern Uralides, syn-orogenic flysch deposition
2000; Ricken et al. 2000). Permian rocks are charac- began in the Late Frasnian (Brown et al. 2006b).
terized by rift-related magmatism and the formation Further to the north, collision started in the Visean
of ensialic basins within a wrench regime, and are (Puchkov 2002a). This strongly diachronous
Fig. 3. To the left: correlation of orogenic and post-orogenic (intraplate) events in the Variscides and Uralides. To
the right: a comparison of idealized sections across the foreland flysch and molasse basins: Preuralian (Southern Urals)
after Puchkov (2000) and Central European Variscides, after Ricken et al. (2000). PUF, Preuralian foredeep; WUZ,
West Uralian zone; ZS, Zilair synclinorium, Stages; art, Asselian to Artinskian stages; v-s, Vizean and Serpukhovian
(¼ c. Lower Namurian); t, Tournaisian; fm, Famennian; fr-fm, Frasnian and Famennian; SVMB, Subvariscan molasse
basin; RHTB, Renohercynian turbidite basin; WC/D, Westphalian stage (¼ c. the Moscovian stage of the Urals) the
upper part; WA/B, Westphalian stage, the lower part; NA/B, Namurian stage (¼ c. Lower Bashkirian and
Serpukhovian stages of the Urals); NA, Namurian stage, the lower part (mostly Serpukhovian stage); a, b, g, the units of
the Upper Visean substage. Thick lines, the lower boundary of molasse.
arc-continent collision was accompanied by HP –

LT metamorphism (Brown et al. 2006b; Puchkov
2008, and references therein). In Bashkirian time,
continent–continent (Laurussia –Kazakhstania) col-
lision commenced and continued intermittently
until the end of the Permian. The early stage of this
collision was accompanied by flysch deposition,
periodically interrupted by relatively deep-water
sediments of a starved type. In the Kungurian
(Uppermost Priuralian), evaporite deposition
occurred in the southern to northern parts of the
Urals, but in its polar area the Kungurian deposits
filling the Preuralian foredeep are dominated by
flysch and coal-bearing sediments. Late Permian
strata are dominated by alluvial to lacustrine
molasse deposits and are weakly deformed in the
frontal part of the foreland fold-and-thrust belt
(Puchkov 2000). At the beginning of the Triassic,
tectonic movements connected with Large Igneous
Province (LIP) formation led to the accumulation
of molasse-like sediments. Voluminous mafic
magmatism began in the earliest Triassic, approxi-
mately at 250 Ma according to Rb –Sr, Sm–Nd
(Anderichev et al. 2005) and Ar –Ar (Reichov
et al. 2007) age data, suggesting a correlation of
these events with the contemporaneous events in
Siberia and the hypothesis of a single Uralo-
Siberian superplume.
The latest collisional events in the Uralides took
place in the Early Jurassic, with increasing intensity
from the South to the North, and represents the term-
inal collision between Baltica (Laurussia) and
Siberia loose blocks united into Pangaea. The
effects of this collision are well exposed in the
Pai-Khoy Range, Novaya Zemlya islands and
Taymyr.
Tectonic zones of the Urals (Fig. 4)

The Uralides are divided into several north–south
striking structural zones, giving the Urals a
general appearance of an approximately linear fold-
belt, in contrast to the more strongly oroclinal and
more mosaic chains of the European Variscides,
Alps or Kazakhstanides (Franke 2000; Khain
2001; Agard & Lemoine, 2005).
The Urals is divided into the following structural
zones, which are from west to east (Fig. 4; Puchkov
1997, 2000):
(1) A – the Preuralian foredeep, which inherited
the western part of a bigger and long-living
orogenic basin. It is filled mostly by Permian
preflysch (deep-water condensed sediments),
Fig. 4. Tectonic zones of the Urals (explanations in the
flysch and molasse. text). Abbreviations: PBB, Platinum-bearing Belt;
(2) B – the West Uralian megazone, predomi- MGA, Main Granitic Axis; MUF, Main Uralian Fault;
nantly consisting of Palaeozoic shelf and EMF, East Magnitogorsk Fault; SMF, Serov–Mauk
deep-water passive margin sediments. This Fault; KRF, Kartaly (‘Troitsk’) Fault. URSEIS and
zone was affected by intense fold-and-thrust ESRU– SB, lines of seismic profiles described in the text.
166 V. N. PUCHKOV
deformation, and includes klippe containing 2000). As ophiolites with MORB signatures are
easterly-derived ophiolites and arc volcanics poorly preserved in most orogens, the abundance
(Puchkov 2002a). of ophiolites in the eastern zones of the Uralian
(3) C – the Central Uralian megazone, where the orogen is a rather anomalous feature, compared to
Precambrian (predominantly Meso- and Neo- many other orogens.
proterozoic) crystalline basement of the Urals All megazones are either exposed or are near the
is exhumed. This basement is traced by geo- Earth’s surface only in the Southern Urals. To the
physical data under the A, B and C megazones. north, the easternmost zones are covered by
Basically, these megazones were formed as a the Mesozoic and Cenozoic strata of the West
result of deformation of the continental Siberian basin, and in the Northern and Polar areas
margin of Baltica, although some allochthons, only the Preuralian foredeep, West Uralian,
and partly the Ural-Tau antiform (UTA) were Central Uralian and western part of the Tagil-
derived from more eastern oceanic structures. Magnitogorskian megazone are exposed.
(4) D – the Tagilo–Magnitogorskian megazone,
bordered to the west by serpentinitic Structural development of the Urals (Fig. 5)
mélange within the Main Uralian Fault zone In general, the Urals comprise the following
(MUF) and to the east by the East Magnito- first-order structural stages: (1) Archaean –
gorsk Fault (EMF) and Serov-Mauk Fault Palaeoproterozoic development of cratonic base-
zones (SMF). This megazone predominantly ment; (2) Meso-Neoproterozoic rift and basin devel-
consists of Ordovician–Lower Carboniferous opment, followed by orogenesis that culminated
complexes of oceanic crust and ensimatic with the formation of Timanide orogen along
island arc, including the Platinum-bearing the periphery of Baltica; (3) Palaeozoic–Lower
Belt of layered basic-ultramafic massifs Jurassic development of the Uralides; (4) Middle
(PBB), overlain by platformal carbonate and Jurassic–Palaeogene –Miocene (platform); and
rift-related volcanic rocks. (5) Pliocene– Quaternary (neo-orogenic) activity
(5) E – the East Uralian zone, bordered to the which is a far-field effect of Alpine-Himalayan
west by the East Magnitogorskian mélange orogenesis. In this paper we focus on the
zone (EMF) and to the east by the Kartaly
(Troitsk) Fault (KRF) (Fig. 4). This zone
comprises Proterozoic gneisses and schists
overlain by weakly metamorphosed Ordovi-
cian to Devonian sedimentary clastic strata
and by tectonically emplaced sheets of
Palaeozoic (Ordovician–Lower Carbonifer-
ous) oceanic and island arc complexes. The
East Uralian Zone is intruded by voluminous
Late Palaeozoic granite bodies which define
the Main Granitic Axis (MGA) of the Urals
(Puchkov et al. 1986).
(6) F – the Transuralian zone, the easternmost
zone of the Urals has probably an accretionary
nature. It contains pre-Carboniferous com-
plexes which preserve a variety of tectonic
settings, including Proterozoic blocks of
gneisses, crystalline schists and weakly meta-
morphosed sediments, Ordovician rift (coarse
terrigenous and volcanic) and oceanic
(ophiolite) deposits, Silurian island-arc com-
plexes and Devonian deep-water deposits
overlain unconformably by the Lower Carbon-
iferous suprasubductional volcanogenic strata,
which form a post-accretionary overstep
complex.
Zones D –F, together with MUF, are traditionally
interpreted to comprise of vestiges of the palaeo-
ceanic component of the Urals, relics of the
Palaeouralian ocean (Peyve et al. 1977; Puchkov Fig. 5. Structural stages of the Urals.
Palaeozoic– Lower Jurassic stage as an example of a truncated by the N –NE-trending magnetic

full Wilson cycle leading to formation of the anomalies which correspond to the MUF and
Uralide orogen. palaeo-oceanic structures of the Uralides. Second,
the lower age limit of the ophiolites attributed to
the Palaeouralian ocean, as determined by recent
The Uralides (Fig. 6) studies of conodonts, is Arenig–Llandeilian (corre-
lated with two unnamed stages between Tremado-
The Uralides consist of the following stages of
cian and Darriwilian (Gradstein et al. 2004) and is
development: (1) rifting of Baltica continental
clearly younger than the Ediacarian–Tremadocian
crust, composed of a cratonic crystalline basement
age that would be expected if there was uninter-
and the Neoproterozoic Timanide foldbelt;
rupted ocean development (Puchkov 2005;
(2) formation of an oceanic basin and micro-
Borozdina et al. 2004; Borozdina 2006; Smirnov
continents; (3) subduction of the oceanic crust and
et al. 2006). These ages are very different from
consequent arc generation; (4) arc-continent
the relics of the Palaeoasian ocean in Altaides and
collision, followed by (5) continent–continent
Kazakhstanides, where Cambrian ophiolites occur.
(Laurussia –Kazakhstania and then Laurussia–
Third, the presence of the Ordovician rift complexes
Siberia) collisions.
along the margins of the former Baltica continent
from one side, and the microcontinent(s) incorpor-
Rifting stage of the Uralides: a precursor of ated into the East Uralian and Transuralian zones
the Palaeouralian ocean support the former interpretation (Fig. 8).
A detailed description of the Uralian Early
On a global scale, rifting and development of the Palaeozoic rift facies in the western slope of the
Palaeouralian ocean episode was preceded by a Urals is given in Puchkov (2002a, and references
series of pene-contemporaneous collisions and oro- therein). Typically, the rift facies consists of Upper-
genies (Cadomian, Timanian, Brasilian, Panafrican) most Cambrian –Tremadocian to Middle Ordovi-
associated with the assembly of the supercontinent cian coarse terrigenous sediments (conglomerates,
Pannotia (or Panterra) in the Ediacarian time. The sandstones, siltstones of very variable thickness,
Palaeouralian ocean was formed as a result of a combined with interlayered subalkaline flows and
breakup of this supercontinent in the Late tuffs). They overlie unconformably the crystalline
Cambrian –Ordovician time (Puchkov 2000). Two basement and are overlain either by shelf or deep-
other possible scenarios for the origin of the ocean water facies, reflecting the development of eastern
have been suggested. According to Zonenshain passive continental margin of Baltica. Although the
et al. (1990), a system of rifts formed in the Early rift facies of the eastern zones of the Urals (Kliuzhina
Ordovician at the eastern margin of East European 1985; Snachev et al. 2006) resemble them lithologi-
continent. Rifting gradually changed to oceanic cally, their age is restricted to the Middle Ordovician.
spreading and the generation of a series of micro- Alkaline carbonatite-bearing complexes (mostly
continental fragments (Uvat-Khantymansian, miaskites) in the western part of the Middle Urals
Uraltau, Mugodzharian) which formed adjacent to (Levin et al. 1997), originally thought to be a
the boundary between the new-formed Palaeoura- manifestation of this rifting event (Samygin et al.
lian ocean and the older, Asiatian ocean. Alterna- 1998; Puchkov 2000) have been re-interpreted to
tively, some researchers (Scarrow et al. 2001; post-date rifting, on the basis of Rb–Sr and U –Pb
Samygin & Ruzhentsev 2003) maintain that isotopic data which indicate a latest Ordovician to
Palaeouralian ocean was inherited from the Protero- Silurian age (Puchkov 2006a; Nedosekova et al.
zoic time, implying no distinction between the 2006, and references therein). These intrusions are
development of the eastern and western flanks of oblique to the Uralian structural grain, and may be
the Uralo-Mongolian orogenic belt in the Early analogous to the Early Cretaceous Monteregian
Palaeozoic. However, there are several strong argu- alkaline intrusions in eastern Canada or the more
ments against the latter interpretation. First, the con- or less contemporaneous hot spot tracks of Eastern
trasting orientation of the structural grain of the Brazil (Bell 2001; Cobbold ft al. 2001). Probable
Timanides and Uralides (especially in the North) Late Ordovician–Early Silurian plume-related
support the idea of a continental breakup preceding complexes also occur in more northern parts of
the formation of the Uralian ophiolites. This the Urals (e.g. monzogabbro-syenite-porphyry as
interpretation is supported by the pattern of strong indicated by the 447 + 8 Ma U –Pb (SHRIMP)
NW-trending magnetic anomalies of the Timan- age of the Verkhneserebryansky complex (Petrov
Pechora province, positive anomalies reflecting 2006) or REE-rich phases of subalkaline grani-
vast fields of rift and island-arc volcanics, and nega- toids in the North of the Urals, dated as 420–
tive anomalies associated with granites and meta- 460 Ma by Rb–Sr and U –Pb methods (Udoratina &
morphic rocks (Fig. 7). These anomalies are Larionov 2005).
168
V. N. PUCHKOV
Fig. 6. A tectonostratigraphic chart of the Uralides in the Southern Urals. All the formations are tentatively restored to their initial, autochthonous positions. Arrows show a
provenance of terrigenous material. (After Maslov et al. 2008, strongly modified.)
Fig. 7. Magnetic anomalies of eastern Baltica (after Jorgensen et al. 1995, with data processed by CONOCO Inc.,
USA). White dotted line, Timanian deformation front; white solid line, Uralide deformation front; white dot-and-dash
line, the Main Uralian Fault.
170 V. N. PUCHKOV
The origin of the Palaeouralian

ocean ophiolites
The aforementioned continental rifting within
Baltica led to oceanic spreading and formation of
ophiolites. The unusual abundance of ophiolites is
a special feature of the Uralian orogen. Several
studies report on the petrology, geochemistry, struc-
ture and metallogeny of the Uralian ophiolites (e.g.
Savelieva 1987; Savelieva & Nesbitt 1996; Saveliev
1997; Melcher et al. 1999; Spadea et al. 2003;
Savelieva et al. 2006a, b). The general consensus
is that the ‘ideal’ section of the Uralian ophiolites
consists of (from top to bottom):
(1) Tholeiitic basalts (mostly pillow lavas) with
layers of pelagic sediments (typically cherts
containing relics of half-dissolved radiolar-
ians). The age of these basalts, constrained
by many occurrences of conodonts, is never
older than Arenigian–Llandeilian (see
earlier comment). In the Tagil zone, ophiolitic
complexes are overlain by Upper Ordovician
island-arc formations, whereas in the Magni-
togorskian zone, the condensed oceanic sedi-
ments overlie Ordovician– Llandoverian
basalts, and persist until the onset of island
arc magmatism in the Early Devonian;
(2) Dyke-in-dyke sheeted complexes, which are
common in the Urals, in contrast to some
other orogens (e.g. the Alps);
(3) Alpine-type gabbro;
(4) Banded dunite-wehrlite-clinopyroxenite com-
plexes, interpreted to reflect a fossil MOHO
boundary; and
(5) Peridotite complexes, represented by lherzo-
lites, harzburgites and dunites in different pro-
portions and combinations.
However, in detail not all ophiolitic complexes
display this simple sequence. For example, Ishki-
nino, Ivanovka and Dergamysh Ni– Co-rich pyrite
deposits in the MUF zone are attributed to ocean-
floor black smokers and overlie and partly penetrate
peridotitic host-rock, devoid of the several ‘stan-
dard’ members of the ophiolite section. Similar
features occur in modern Atlantic thermal fields
(e.g. Logachev and Rainbow fields), although their
geodynamic setting may not support the direct
analogy. The deposits rather belong to the relics of
Magnitogorsk forearc (Jonas 2004; Melekesceva
2007).
Fig. 8. The position of the Early Palaeozoic rift/plume

related complexes. I–VI, tectonic zones of the Urals Fig. 8. (Continued) 2, Bardym; 3, Lemva; 4, Baydarata;
(correspond to A, B, C, D, E, F in Fig. 4, with the same 5, Samar; 6, Sargaza; Uvelka; 8, Mayachnaya.
symbols); 1 –8, localities where the Lower Palaeozoic Dashed-line ellipse, location of possible plume-related
graben formations are developed; 1, Sakmarian; alkaline complexes Vishnevogorsk and other).
Savelieva et al. (2006a) classified the Uralian age and the Voykar sheeted dykes, described
ophiolites into three groups according to their by Remizov (2004) as island arc complex,
inferred geodynamic setting: were dated at 426–444 Ma (Didenko et al.
2001). Khain et al. (2004) dated zircons from
(1) Complete sections of ophiolites (e.g. Kempir-
a plagiogranite dyke in the parallel dyke
say massif) or their fragments in the south of
complex of the Voykar-Synya massif at
Magnitogorsk, East Uralian or Denisovka
490 + 7 Ma. The above data indicate that the
zones), include restite peridotite and overlying
generation of these ophiolites is probably
succession of plutonic gabbro, parallel diabase
more complicated than the current geody-
dyke complexes and tholeiitic lavas formed in
namic models purported to explain them.
a MOR setting (Savelieva & Nesbitt 1996).
However, a supra-subduction geochemical This complicated scenario is highlighted by
component has been documented in ophiolites recent age data (Gurskaya & Smelova 2003;
in the southern part of the Kempirsay massif Savelieva et al. 2006a; Tessalina et al. 2005; Bata-
(Melcher et al. 1999); nova et al. 2007; Krasnobaev et al. 2008), which
(2) Massifs of a lherzolite type representing fairly yield Neoproterozoic– Ediacarian (536 –885 Ma)
low depleted lithospheric mantle (e.g. Kraka, and some older dates for many of the ultramafic
Nurali) have a simple evolutionary history and alpine-type gabbro complexes that were pre-
consisting of enriched peridotite and dunite viously regarded as Palaeozoic. These data include
associated with less abundant amphibole Re –Os and Sm –Nd mineral isochrons and U –Pb
gabbro (Savelieva 1987). These character- analyses of zircons. In general, however, most of
istics are thought to reflect a low degree of published age data support also an Ordovician–
partial melting of a mantle diapir, followed Lower Devonian age of alteration processes for
by rapid uplift, a scenario typical of rifting most of the complexes (summarized by Puchkov
that immediately precedes oceanic spreading. 2000 and well illustrated by Krasnobaev et al.
Alternatively, Spadea et al. (2003) propose a 2008). Two contrasting explanations have been pro-
more dynamic history for these massifs, invol- posed. According to Tessalina et al. (2005), the
ving re-fertilization of a depleted mantle by ultramafic complexes are Neoproterozoic ophiolites
basaltic magma, by analogy with Lanzo and represent relics of the oceanic crust developed
massif of Alps, which however is also during Timanide orogenesis. Alternatively,
thought to be indicative of pre-spreading Puchkov (2006b) suggests that the Neoproterozoic
rifting (Müntener et al. 2005); and dates in the Palaeozoic ophiolites reflect a relict sig-
nature preserved in the mantle part of the younger
(3) According to the general geodynamic recon-
ophiolites (peridotites and partly ex-eclogitic
structions of Saveliev (1997), the huge Polar
mantle gabbro), notwithstanding the overprinting
Urals massifs such as Voykar-Synya, Ray-Iz
during subsequent processes of ophiolite formation.
and Syum-Keu, are integrated into a system
The possibility of preservation of ancient mantle
of allochthons, composed of complexes of
zircons and their contamination of younger MOR
two island arcs – Tagil-Schuchya (O3 –S1)
and island arc volcanics has recently been under-
and Voykar (S2 –D3). The restites in these
lined (Sharkov et al. 2004; Bortnikov et al 2005;
massifs are strongly depleted and preserve evi-
Puchkov et al. 2006).
dence of interaction with basaltic magma. The
Despite the above complexities, the age of
sections consist of multi-phase intrusions of
ophiolite basalts determined by conodonts, is
gabbro and diabase, interpreted by Savelieva
never older in the Urals than Arenigian–Llandeilian
et al. (2006a) to reflect the development of a
(see again the earlier comment).
marginal basin when an island arc rifted apart
in the Late Silurian-Early Devonian. Therefore The passive margin of the continent
this spreading was related to the development
of a second island arc. This interpretation, Simultaneously with oceanic development, the con-
however, may be an over-simplification as tinental margin of Baltica started to develop by
the presence of two island arcs suggests the rifting in the Ordovician (Fig. 9). The identity of
existence of older (Ordovician– Lower Silur- the conjugate margin to this rift is not known. By
ian?) ophiolites corresponding to the older of the end of the Silurian, Baltica had collided with
the two arcs. Indeed, Ar –Ar data from the Laurentia to form Laurussia (Ziegler 1999). The
banded complex of Voykar-Synya and development of the margin is described in detail
Khadata massifs (primary amphibole, fresh by Puchkov (2002a), and only a general summary
plagioclase and clinopyroxene of gabbro) is given here. Typically, the succession starts with
yield 420 –490 Ma ages. In addition the uppermost Cambrian– Lower Ordovician coarse
Khadata spreading dikes yield a c. 423 Ma terrigenous deposits, in some cases accompanied
172 V. N. PUCHKOV
Fig. 9. Major structural elements and complexes of Laurussia/Baltica passive margin involved into the Urals (from
Puchkov 2002, with minor changes).
by minor volcanic rocks (see above). The margin is limestones, dolomites) and terrigenous sediments
classified as a non-volcanic type (in a classification with west-derived (Smirnov 1957) oligomictic,
of Geoffroy 2005). quartz sandstones. Regressions are marked by
Two facies were established early in the passive barrier reefs at the outer margin of the shelf zone,
margin development – an inner (shelf ) and an outer while the transgressions favour a formation of deep-
(continental slope, grading to continental rise) water, starved basins with condensed facies of marls
(Fig. 6). Generally, shelf sediments are represented and oil shales (called ‘domanik’ in Russia),
by shallow-water carbonates (limestones, dolomitic surrounded by reefs and bioherms.
The outer, continental slope and rise (bathyal) by andesites, dacites, basalts and very abundant
sections consist of thick westerly-derived quartz tuffs. In the Wenlockian, the volcanic rocks are lat-
sandstones, and thin, condensed units consisting of erally equivalent to reefal limestones (All-Russian
shales, cherts and minor limestones (Puchkov Committee 1993). These biohermal deposits per-
1979, 2000). Fauna are mostly pelagic: radiolarians, sisted until the Pridolian as an unstable, narrow car-
conodonts, graptolites and rare goniatites. The bonate shelf on the perimeter of the island arc. The
paucity of limestones indicate a transition to above Silurian complexes are substituted laterally
abyssal conditions. by a volcanic association (S1ll3 –S2ld1), represented
The upper strata of the outer facies consist of by basalts, andesites and tuffs with rare layers of
polymictic, flysch deposits signifying a sharp cherty siltstones. After the Late Ludlovian, the
change of provenance that is connected with the marine conditions partly changed to continental
start of orogenesis (Puchkov 1979; Willner et al. conditions: the Pridolian association is represented
2002, 2004). This change of provenance is diachro- by predominant coarse-grained red-coloured poly-
nous: it is earlier in the east and south of the western mictic terrigenous-volcanogenic deposits with frag-
slope of the Urals; in the Southern Urals it occurs in ments of the older rocks; such as volcanics, with
the uppermost Frasnian, but in the Polar Urals it sublakalic and alkalic basalts being predominant.
starts in the Early Visean. The island arc succession is terminated by a very
Puchkov (1979) drew attention to the similarity specific association (S2pr– D1lh), preserved in the
of these deposits with analogous geodynamic set- axial part of the Tagil synform, which resembles
tings in other orogens. For example, eastern Lauren- the underlying Pridolian association, and includes
tia was bordered by deep-water sediments in the shoshonitic mafic to intermediate volcanic rocks
Ordovician that are preserved in the allochthons (Narkissova 2005) and minor flysch-like volcani-
formed during the generation of the Appalachian clastic deposits.
orogen. In some cases (e.g. Ouachita), the deep- The volcanism of the Tagil arc evolved from a
water facies persisted, like in the Polar Urals, uniformly tholeiitic affinity to a differentiated
through most of the Palaeozoic (from the lowermost calc-alkaline and then to subalkalic shoshonitic
Ordovician until the Carboniferous). affinity, suggesting deeper levels of partial
melting in mantle with time, a trend that is opposite
to the typical trend in rift and superplume zones
Subduction of oceanic crust (Dobretsov et al. 2001). Geochemically, the vol-
canics are typical of ensimatic island arcs
The Urals is characterized by an exceptionally good
(Narkissova 2005; Borozdina 2006). Basalts
preservation of subduction complexes, which
retrieved from the superdeep SG-4 borehole
permits reconstruction of the development of at
exhibit a distinct Ta–Nb minimum. In general, the
least three subduction zones in place and time:
volcanics are depleted in Nb,Ta, Zr, Ti, Y and
Tagil (Late Ordovician–Early Devonian), Magnito- enriched in K, Rb, Ba, Pb relatively to N –MORB.
gorsk (Early –Late Devonian) and Valerianovka
The geochemical trends of contemporaneous volca-
(Tournaisian –Early Bashkirian) (Fig. 10).
nics suggest an eastward (in modern co-ordinates)
dipping subduction zone (Narkissova 2005).
The Late Ordovician – Early Devonian Tagil The Tagil arc is also known for the presence of
arc (Fig 10, to the left) gabbro-ultramafic massifs composing a gigantic
(c. 1000 km) linear, platinum-bearing belt (PBB
The oldest (Tagil) arc complexes are developed on Fig. 4). The concentric-zonal massifs consist of
in the Middle, Northern, Cis-Polar and Polar parts dunites, clinopyroxenites, gabbro and plagiogra-
of the Tagilo– Magnitogorskian zone. The best sec- nites, and mafic rocks comprise up to 80% of the
tions, well constrained by conodonts, are preserved belt. Disseminated platinum is hosted by dunites,
in the southern part of the Tagil synclinorium and industrial deposits are represented mostly by
(synform). According to recent stratigraphic and modern (or reworked Meso-Cenozoic) placers.
petrochemical studies (Narkissova 2005; Borozdina The geodynamic significance of the belt is contro-
2006), the oldest ensimatic island-arc succession is versial: models vary from a rift (Efimov 1993) to a
predominately represented by the basaltic (O3), supra-subduction zone setting (Ivanov et al. 2006).
basalt-plagiorhyolitic (O3) and basalt–andesite – The age of the belt, determined by several
plagiodacite (S1) volcanic associations. The latter methods as 420–430 Ma, and the similarity of
two have calc-alkaline affinities (Narkissova petrogenetic-indicating trace and rare earth
2005). The overlying Silurian association (S1ll2 – elements with island-arc tholeiites (Ivanov et al.
S2w1) is represented by flysch consisting of 2006) supports the supra-subduction zone model.
interbedded black cherty siltstones, tuffites and Such belts are rare in modern arc environments,
tuffaceous sandstones (arc slope deposits), overlain but possible analogues occur in the northern part
174 V. N. PUCHKOV
Fig. 10. The distribution of magmatism of three main stages of subduction: Ordovician– Early Devonian (Tagil): Early
Devonian –Famennian (Magnitogorsk); Tournaisian–Early Bashkirian (Valerianovka and contemporaneous to it).
of the circum-Pacific ring, and are known as includes also olistostromes developed locally with
Alaskan type (Burns 1985). and within serpentinitic mélanges of the Main
Ordovician–Early Devonian island-arc volcan- Uralian fault. The bioherms and olistostromes are
ism was followed by the development of a relatively local indicators of buckling of the oceanic crust at
stable carbonate shelf which caps the western part of the onset of subduction and are related to an early,
the island arc complexes. After that time the Tagil non-volcanic stage of subduction.
arc was dismembered and by the Emsian, it was a Volcanic complexes of the Magnitogorsk arc in
terrane that accreted to the Magnitogorsk island the Southern Urals are well studied (Brown et al.
arc, an event that coincided with the formation of 2001, 2006b; Kosarev et al. 2005, 2006, and
the Magnitogorsk arc itself (see below). The references therein). The volcanic succession,
location of the subduction zone changed and the represented by a characteristic interlayering of tho-
region became characterized by presence of two leiitic basalts, bimodal basalt-rhyolite and regularly
sub-zones: the western, Petropavlovsk and the differentiated basalt-andesite-dacite-rhyolite series,
eastern, Turyinsk sub-zones. comprises the following units (the local names
The Petropavlovsk sub-zone contains Lower to are partly shown in Figure 6, and numbered
Middle Devonian shallow-water limestones and correspondingly):
bauxite, followed in the Late Devonian by deep-
(1) A bimodal rhyolite-basalt series that overlies a
water cherty shales and polymictic terrigenous sedi-
tholeiite-boninite unit (Emsian);
ments. In the Turyinsk sub-zone, sedimentary strata
(2) Basalt-andesite-dacite-rhyolitic series (Upper
(shallow-water limestones, shales and cherty shales)
Emsian);
are interlayered with andesites, basalts, tuffs and
(3) Andesite-basalt series (uppermost Emsian–
volcanogenic sandstones (All-Russian Committee
Lower Eifelian);
1993). Yazeva & Bochkarev (1993) point out that
(4) Bimodal rhyolite-basalt series (Upper Eifelian);
these thick (up to 4–5 km) Devonian volcanic
(5) Basalt-andesite-dacite-rhyolite series (Givetian–
layers occur with comagmatic intrusions in volcano-
Lower Frasnian);
plutonic complexes. Geochemical parameters (in
(6) Basalt-andesite formation (Upper Frasnian);
particular, Rb and Sr contents) indicate that the
(7) Local shoshonite-absarokite formation
thickness of the crust was c. 30 km (Yazeva &
(Famennian). In addition, subduction-related
Bochkarev 1993), which implies that the new arc
370– 350 Ma granitoid intrusions of calc-
was ensialic, in contrast with the Ordovician–
alkaline affinity are developed in the northern
Lower Devonian ensimatic Tagil island arc.
part of Magnitogorsk synclinorium (Bea et al.
2002); and
Devonian Magnitigorsk arc (Fig. 10, centre) (8) These intrusions are unconformably overlain
by Lower Carboniferous volcanics dominated
The development of the Magnitogorsk arc in the
by tholeiitic basalt in the west and by
Southern Urals was more or less synchronous with
more widely developed subalkaline bimodal
the dismemberment of the Tagil arc. The location
basalt-rhyolite in the east. They are accom-
of the Turinsk zone of Tagil arc along the extension
panied by a chain of coeval (335–315 Ma,
of the Magnitogorsk arc suggests that the Magnito-
Bea et al. 2002) bimodal gabbro-granitoid
gorsk subduction zone was inherited from the Tagil
intrusions (Magnitogorsk-type plutons). Both
zone. The Middle–Upper Devonian calc-alkaline
volcanic and intrusive bimodal series, accord-
complexes can be traced northward to the Polar
ing to their field relationships, mineralogy and
Urals.
geochemistry, suggest an extensional or
However, in the Southern Urals, island arc
passive within-plate non-arc origin and are
development was preceded by a long period of
probably produced by undepleted lherzolites
quiescence, expressed by the deposition of deep-
(Bochkarev & Yazeva 2000; Fershtater et al.
water oceanic cherts and carbonaceous cherty
2006). This magmatism may be related to a
shales accompanied by basalts in the Ordovician
slab break-off of the Magnitogorsk subduction
and Llandoverian. Most of the Silurian is rep-
zone which gave way to a melt from the less
resented by 300 m of distal, condensed cherty
depleted, deeper mantle under it (Kosarev
shales. They are considered to represent the
et al. 2006).
sedimentary cover of the ophiolites. Non-volcanic
sections of the Lower Devonian (Lochkovian – Notwithstanding differences in composition,
lowermost Emsian) are represented by either magmatic and tectonic affinities, all the Devonian
deep-water terrigenous chert, argillaceous cherty volcanic series share geochemical traits typical of
sediments of Masovo, Turatka, Ishkinino and other a supra-subduction zone origin, such as negative
formations or bioherm limestones (Artiushkova & Nb, Ta, Zr, Hf, Y anomalies, and elevated concen-
Maslov 2003; Fig. 6). This stratigraphic level trations of large ion lithophile (LIL) elements
176 V. N. PUCHKOV
(K, Rb, Ba, Cs) and LREE. They show no signs (6) formation of the suture zone of the Main
of contamination by continental crust and are Uralian Fault, which divides the accretionary
interpreted as ensimatic arc complexes. prism on the continent from the remnant of
There are many parallels with a development of the island arc.
the Tagil arc, including the alkaline trend towards
the upper member of the succession. Trace In summary, the Southern Urals Magnitogorsk
element abundances for contemporaneous volcanic arc was formed in the Early Devonian upon
rocks are consistent with an eastward-dipping Ordovician– Early Devonian oceanic crust
subduction zone. To the west and east of the main (Puchkov 2000; Snachev et al. 2006) and developed
volcanic body of the arc, mostly in mélanges of until it collided with the passive continental margin
the MUF and EMF zones and associated alloch- of Laurussia in the Late Devonian. Due to the
thons, condensed cherty-terrigenous series contem- oblique orientation of the subduction zone relative
poraneous to the arc are developed, corresponding to the continental margin, collision at this time did
to the forearc and backarc basins (Fig. 6). not take place along the whole length of the
It looks like the ophiolite basement of the passive margin of Laurussia continent. To the
arc is mostly Ordovician in age, except the north of the Ufimian promontory, the margin of
Mugodzhary section, where the Emsian basalts Laurussia shows no evidence of the Late Devonian
and cherts of Mugodzhary and Kurkuduk for- arc –continent collision.
mations overlie a large-scale Aktogay sheeted Two remarkable Late Devonian events in the
dyke complex, gabbro and serpentinites, composing Southern Urals can be regarded (along with direct
a Lower Devonian ophiolite (Fig. 6). The ophiolite structural data) as important indicators of the tran-
is tentatively interpreted as a result of a backarc sition from subduction to collision. The first is the
spreading. deposition of the Zilair greywacke flysch formation
of the eastern provenance (uppermost Frasnian –
The collision of the Magnitogorsk arc with Famennian) which overlies Frasnian deep-water
the passive margin of Laurussia and Famennian shallow-water deposits of the
continental margin of Laurussia. The second is a
The collision of the Magnitogorsk arc with the culmination of a HP–LT metamorphism at
passive margin of Laurussia (former Baltica) has 372–378 Ma that provides additional evidence for
been described in several recent publications the end of subduction and the onset of collision.
(Brown & Puchkov 2004; Brown et al. 2006b, and As for the metamorphism, its age range in the
references therein) and is briefly summarized here subduction zone should be broadly contempora-
(Figs 11 & 12). Since the Early Devonian, an neous with the subduction, and its oldest products
island arc formed within the Uralian palaeocean should be older than the supra-subductional volcan-
above an east-dipping subduction zone. Collision ism. However, it is not clear if the products of this
of the arc with Laurussia occurred in the Late early metamorphism were preserved and then
Devonian and was accompanied by the following exhumed or if they were completely entrained by
events: the slab. The Magnitogorsk arc appears to have
these products preserved and then exhumed. In the
(1) scraping-up of the deep-water sediments of serpentinitic mélange of MUF, along the margin
the continental passive margin by the rigid of the Magnitogorsk arc, garnet pyroxenites (meta-
wedge (backstop) of the arc and the formation morphic basites) occur. The best documented occur-
of an accretionary prism; rence where P–T parameters of their origin are
(2) jamming of the subduction zone followed by a determined as 1.5–2 GPa, 800– 1200 8C is in the
jump in the location of the subduction zone; Mindyak peridotite massif of MUF (Pushkarev
(3) slab break-off and opening of a slab window 2001). Its age was determined by two methods:
permitting the deeper, more fertile and hotter Sm– Nd isochron is 406–399 Ma (Gaggero et al.
mantle to produce subalkaline, non- 1997); whereas U –Pb analysis of zircons yield an
subduction volcanics; age of 410 + 5 Ma, which is interpreted as a meta-
(4) uplift of the buoyant continental part of the morphic age (Saveliev et al. 2001). A Pb –Pb analy-
slab, exhumation and erosion of UHP(?) and sis of zircon cores yield an age of 467 Ma, and is
HP–LT metamorphic complexes and their interpreted as a protolith age (Gaggero et al.
erosion; 1997). The age of garnet pyroxenite from the
(5) formation of the accretionary cordillera of Bayguskarovo occurrence is 416 + 6,1 U –Pb
Uraltau antiform (comparable to accretionary SHRIMP (Tretyakov et al. 2008). A series of
avolcanic arc of Indonesia) and two flysch Ar –Ar age determinations of phengite, whose
basins flanking both sides of it: forearc and interpretation depends on mineral dimensions and
foredeep basins (Fig. 13); and temperature conditions of equilibration, has been
Fig. 11. Time/process evolutionary diagram for intra-oceanic subduction and arc-continent collision in the Southern
Urals (after Brown et al. 2006b, with added information given in bold).
178 V. N. PUCHKOV
The younger ages of the metamorphic rocks

cluster around 375–380 Ma and probably are con-
sistent with the general exhumation of the HP –LT
metamorphic rocks of the Southern Urals. The
rocks are divided into two units, established by
Zakharov & Puchkov (1994) and by many
later researchers.
The age of the start of general cooling and exhu-
mation (reviewed by Brown et al. 2006b) for the
eclogite facies metamorphism of the lower unit of
the Maksutovo Complex is thought to be Frasnian
in age, with a mean value of 378 + 6 Ma according
to many isotopic determinations (Matte et al. 1993;
Lennykh et al. 1995; Beane & Connelly 2000;
Hetzel & Romer 2000; Glodny et al. 1999, 2002).
The upper unit was metamorphosed together with
the lower unit, suggesting juxtaposition during
exhumation at a higher crustal level by
360 + 8 Ma (Rb– Sr and Ar– Ar methods; Beane
& Connelly 2000; Hetzel & Romer 2000).
In the Polar Urals, isotope dating of HP –LT
metamorphism of the Marun-Keu complex of eclo-
gites and related metamorphic rocks was reviewed
recently by Petrov et al. (2005). According to
Shatsky et al. (2000), the Sm–Nd isotopic analyses
of garnet, clinopyroxene and whole rock gave
366 + 8.5 Ma for the hornblende eclogite and
339 + 16 Ma for the kyanite eclogite. Rb–Sr
whole-rock dating of the eclogites (Glodny et al.
1999) gave 358 + 3 Ma. According to the data of
Glodny et al. (2003, 2004), the concordant U –Pb
age data for the metamorphic zircon domains are
Fig. 12. A model for development of the Magnitogorsk between 353 and 362 Ma, coincident with the age
arc and subduction zone (Kosarev et al. 2006, slightly of metamorphism as inferred from Rb–Sr internal
modified). Dotted lenses, supposed zones of melting of mineral isochrons (an average value of
initial magmas of different petrogenetic types: T, 355.5 þ 1.4 Ma).
tholeiitic; BON, boninitic; TMg, tholeiitic magnesial; The eclogite –glaucophane Nerka-Yu and Parus-
CA, calc-alkaline; ASh, absarokite-shoshonite; SA, Shor complexes in the southernmost Polar Urals
subalkaline. Stages of the Devonian: em, Emsian; ef, yielded 351 + 3.6 and 352 + 3.6 Ma (40Ar – 39Ar
Eifelian; gv, Givetian; f, Frasnian; fm, Famennian. ages, Ivanov et al. 2000). Sm–Nd dating of glauco-
phane schists of the Salatim belt (Northern Urals)
gave 370 + 35 Ma. Taken together, these dates
done for a succession of samples across the contact characterize an Upper Famennian-Middle Tournai-
between eclogite and garnet glaucophane schist sian age of HP–LT metamorphism and the begin-
from Maksiutovo complex. The age range of phen- ning of its exhumation. These data are supported
gites is from 400 Ma at c. 500 8C to c. 379 Ma at the by the Lower Visean age of the oldest known
final closure temperature of the system (c. 370 8C). Palaeozoic easterly-derived polymictic sandstones
The Ar –Ar age of glaucophanes from the same and conglomerates on the continental margin of
sample is 411–389 Ma (Lepesin et al. 2006). The Laurussia, to the west of the Main Uralian Fault,
peak of Ar –Ar ages obtained from detrital phengites in the Polar and Northern Urals (Puchkov 2002a,
of the Zilair series clusters around 400 Ma (Willner and references therein).
et al. 2004). U –Pb SHRIMP dating of zircons Developing the oblique collision model of
from Maksiutovo eclogites yielded 388 + 4 Ma Puchkov (1996b), Ivanov (2001) calculated an
(Leech & Willingshofer 2004). These older average rate of subduction, which led to a gradual
(Lower and Middle Devonian) dates of meta- northward-shifting collision, of 2.75 –2.80 cm.
morphism are consistent with the cooling action of However according to recent data, the diachroneity
the subducting slab, causing the closure of isotopic of events at the end of the Devonian and beginning
systems. of the Carboniferous show no gradual south– north
Fig. 13. Reconstructed geological section across the Magnitogorsk arc, in the Famennian time.
pattern, implying that behaviour of the subduction the lack of data may be because the eastern limb
zone is more complex. of this arc is concealed in the Northern to Polar
The collision of the Magnitogorsk arc with Urals under the Mesozoic –Cenozoic cover of the
Laurussia may have occurred in two discrete West Siberian plate.
stages (Fig. 14). First, collision in the Southern The above-described Early Carboniferous
Urals occurred by the Famennian, and a triangular- (Tournaisian –Visean) stage of subduction was
shaped gap was left between the arc and the conti- followed in the Middle Urals by a Serpukhovian
nent, similar to the modern Bengal Bay, Northwest stage, as indicated by the Verkhisetsk chain of
Australian Bay or the South China Sea. Second, in granitoids (320 Ma) (Fig. 10), related by Fershtater
the Early Carboniferous, the northern half of the et al. (2006) to another east-dipping subduction
arc was bent to the west and docked to the conti- zone.
nental margin. At this stage subduction in the
south virtually ceased, but in the north, increasing Early Carboniferous-Bashkirian Valerianovka
velocity of subduction resulted in increasing inten- subduction zone(s) (Fig. 10, to the right)
sity of the HP –LT metamorphism in the same
direction (from glaucophane schists of the By the middle of the Lower Carboniferous, the
Salatim belt and Cis-Polar Urals to eclogites of suture zone was established along the whole
the Polar Urals). In the Middle Urals, these length of the MUF (Puchkov 2000, 2002a). The
events were immediately followed by intrusion of above-mentioned chain of 335–330 Ma (mid-
Turgoyak-Syrostan group of granitoids (335– Visean) massifs (Turgoyak –Syrostan group of
330 Ma), that was described by Fershtater et al. granites) intrude the suture zone (Fershtater et al.
(2006) as ‘granitoids connected with tensional 2006) and therefore post-date the Magnitogorsk
structures’. This group can be correlated by the subduction, providing an upper age limit for MUF.
age with the Magnitogorsk within-plate gabbro- This conclusion is supported by the age of the
granite series (see below). Ufaley intrusion (concordant U –Pb, 316 + 1 Ma,
Unfortunately there is no support for this model Early Bashkirian), which seals the Main Uralian
from the data on the Early Carboniferous volcanism Fault in the Northern part of the Ufimian promon-
in the northern part of the Magnitogorsk arc. But tory (Hetzel & Romer 1999).
180 V. N. PUCHKOV
Fig. 14. A model for a two-stage Upper Devonian– Lower Carboniferous arc-continent collision in the Urals. 1,
continental crust; 2a, transitional crust; 2b, oceanic crust; 3, Tagil island arc; 4 and 5, Magnitogorsk island arc; 4;
ensialic (epi-tagilian); 5; ensimatic (Magnitogorsk arc sensu stricto); 6, subduction zone; 7, continent –ocean boundary;
8, suture zone of the Main Uralian fault.
With the demise of the Magnitogorskian arc, suprasubductional tonalite–granodiorite massifs by

subduction did not terminate in the Urals as a the end of the Famennian or the beginning of the
whole. Ensialic subduction (either island arc or Tournaisian (c. 360 Ma), when the southern part
Andean-type or maybe two subduction zones of of the Magnitogorsk subduction zone ceased to
different type) of uncertain polarity began in the operate (Bea et al. 2002; Fershtater et al. 2006).
latest Devonian and lasted until the Mid-Bashkirian Simultaneously, immediately to the east, a wide
in the eastern Urals. The Main Granitic Axis of the NNE-trending band of calc-alkaline and partly
Urals (Fig. 4) developed first as a chain of within-plate volcanic rocks and associated plutonic
complexes ranging up to mid-Bashkirian in age shelf zones and intensely eroded uplifted crustal
were formed, suggesting a close affinity with the blocks. As uplift continued, the basins contracted
massifs of the Main Granitic Axis (All-Russian and inverted. By the Kasimovian time, the territory
Committee 1993; Tevelev et al. 2005; Fershtater east of MUF was dominated by erosion and subaer-
et al. 2006). The Lower Carboniferous (320 Ma) ial deposits. To the west of MUF, a deep-water
tonalite-granodiorite massifs occur in the Middle foredeep trough was filled by easterly-derived
Urals, situated to the east of the Serov-Mauk flysch, prograding to the west (Puchkov 2000). A
suture zone (i.e. Verkhisetsk massif and others westerly prograding foreland fold-and-thrust belt
located to the east of the former Magnitogorsk was developed along the eastern margin of the fore-
volcano-plutonic arc). deep, which deformed and uplifted flysch of its
According to Fershtater et al. (2006), increases in eastern limb. The diachroneity of these processes
K2O and REE abundances in the granodiorites to the is documented by detailed studies of resedimented
east indicate that subduction had an eastern polarity. conodonts within these strata (Gorozhanina &
However, Kosarev & Puchkov (1999) point out that Pazukhin 2007). In the Gzhelian– Sakmarian, the
K2O concentrations in the Lower Carboniferous vol- thrusting and crustal thickening created a hot
canic rocks in the eastern Urals increase westward, crustal root in the Southern Urals, which resulted
suggesting a western polarity for the subduction in generation of 305–290 Ma syn-collisional gran-
zone. Of the same opinion are Tevelev et al. (2005) ites of the Main Granitic axis (Fig. 2), followed by
for the Uralian Lower Carboniferous volcanics, but 10 –15 km of erosion (Fershtater et al. 2006).
they propose that the volcanism occurred in a Crustal thickness may reach 65 km (the modern
wrench regime and that the easternmost Valerya- crust thickness of the East Uralian zone is up to
novka volcanic band belonged to Kazakhstanides 50 km; see below), similar to modern orogens.
and developed over a separate subduction zone with Tuff layers in deep-water sections of Gzhelian to
eastern polarity. Brown et al. (2006a) also suggest Lower Kungurian preflysch and distal flysch of Pre-
two oppositely dipping subduction zones, whereas uralian foredeep may represent volcanic equivalents
Matte (2006) suggests a westerly dip for several sub- of this magmatic activity (Davydov et al. 2002).
duction zones. Such a difference in opinion is Syn-collisional granite magmatism migrated
explained by the complicated nature of the process, northward, and is thought to be a manifestation of
involvement of wrench tectonics, and the rather the oblique, diachronous character of collision
poor exposure of the complexes. (305 –290 Ma for the southern Uralian granites,
Calc-alkaline volcanic complexes in the Urals 265 Ma for the Kisegach massif, 250–255 Ma for
abruptly stopped forming by the mid-Bashkirian, Murzinka and Adui massifs of the Middle Urals,
signifying the end of a wide-scale subduction of Fershtater et al. 2006). The idea of the transpressive
an oceanic crust and transition to a continent- character of orogenic deformation is supported by
continent-type collision. structural studies (e.g. Pliusnin 1966; Znamensky
2007). A change from thrust-dominated tectonics
to sinistral transpression occurred in the Southern
Continent –continent collision and Urals (Znamensky 2007), explaining the K-rich
formation of the orogen concentric-zoned post-tectonic c. 283 Ma rift-
related granite-monzonite massifs at the northern
The main Late Bashkirian to Permian end of Magnitogorsk synclinorium (Ferstater et al.
stage of collision 2006) and the occurrence of c. 301–310 Ma lam-
proite dykes in the Southern Urals (Pribavkin
The collision between Laurussia and Kazakhstania et al. 2006).
that resulted in mountain building in the southern Post-collisional granite magmatism in each
and middle Urals has been described recently by region was followed by uplift and erosion, and the
Brown et al. (2006a), and its main events are summar- diachroneity of the magmatism is exemplified by
ized here. The external (palaeogeographic and mag- the presence of the Late Permian marine sediments
matic) expressions of the orogeny, including the with Tethyan fauna in the Southern Urals that is
northern-to-polar and eastern (epi-Kazakhstanian) coeval with granite magmatism in the Middle
regions are emphasized in this synthesis. Urals (e.g. Chuvashov et al. 1984).
By the Mid-Bashkirian, subduction had ceased
and collisional processes between Laurussia and An interlude: LIP formation and
Kazakhstania began first as a formation of linear localized rifting
uplifts and basins, documented in the Southern
Urals (Puchkov 2000). In the Late Bashkirian and At the Permian–Triassic boundary, the waning
Moscovian, widespread marine flysch were depos- effects of orogenesis were overprinted by the for-
ited in troughs separated by more slowly subsiding mation of the vast Uralo-Siberian LIP, extending
182 V. N. PUCHKOV
from Taymyr in the north to the Kuznetsk and latter two (Fig. 1). According to palaeomagnetic
Turgay basins in the south and from the Tunguska data (Kazansky et al. 2004), Siberia rotated 308
basin in the east to the Urals in the west. Volcanism clockwise between the Triassic and the Late
started locally with alkaline basalts and minor rhyo- Cretaceous.
lites. Ar –Ar data (Ivanov et al. 2005) suggest that
the bulk of the volcanism initiated in Siberia at the Peneplain formation
Permian–Triassic boundary but probably continued
for 22–26 Ma, with several short surges. Recent Rapid uplift and erosion of the Uralide orogen
Ar– Ar dates for plagioclase from basalts in the resulted in the formation of a Cretaceous-
Polar Urals (249.5 + 0.7 Ma) and in the east of Palaeogene peneplain (Papulov 1974; Sigov 1969;
the Southern Urals (243.3 + 0.6 Ma) (Reichov Amon 2001), and by the Late Jurassic or Early Cre-
et al. 2007) support the simultaneous beginning of taceous, there was no topographic barrier dividing
the LIP formation over a vast region followed by a Europe and Siberia. Buried river-bed deposits
more protracted period of reduced magmatism. along the eastern slope of the Southern and Middle
In contrast with eastern Siberia, the Early Trias- Urals have a north-eastern direction, as revealed
sic history of the Urals is dominated by considerable by a shallow prospecting drilling. For the Late Cre-
uplift, erosion and formation of thick coarse-grained taceous and Eocene, the existence of short-lived
alluvial to proluvial sediments that resemble oro- straits connecting the European and Siberian
genic molasse but are attributed to the effects of marine basins is hypothesized.
the Uralo –Siberian distributed rifting and LIP Along the eastern slope of the Southern and
magmatism. Examples include the huge Triassic Middle Urals, thin marine sediments occur only
Koltogorsk–Urengoy graben of Western Siberia, during maximal transgressions (in the Late Cretac-
the newly-identified Severososvinsky graben in the eous and Middle Eocene); more generally, fluviatile
subsurface of the Cis-Polar Urals (Ivanov et al. deposits occur. To the north, Upper Jurassic and
2004), and the eastern parts of the LIP (Kurenkov younger marine sediments occur adjacent to the
et al. 2002; Ryabov & Grib 2005). eastern foothills of the modern Ural mountains.
The difference between the southern and northern
The Cimmerian orogeny parts of the modern Urals (as expressed by better
exposure of the eastern zones in the south), was
A short pulse of orogeny occurred at the end of the probably inherited from this time.
early Jurassic, and its effects differ along the strike
of the Uralides. The Triassic deposits of the Neo-orogeny
Southern Urals are affected by this orogeny only
in the Trans-Uralian zone (Chelyabinsk and other In the Pliocene –Quaternary time, a modern chain of
graben-like depressions), where Upper Triassic moderately high mountains was uplifted, forming a
and older deposits are deformed by thrusting natural drainage divide between Europe and Asia.
(Rasulov 1982), followed by uplift and peneplana- These mountains are formed in an intra-plate
tion during the Middle and Upper Jurassic, and setting, having no precursory suture zone. Conver-
deposition of Upper Cretaceous marine deposits. gence is indicated by studies which show that
In the Northern Urals, three ‘grabens’ (Mostovskoi, maximum stresses are oriented perpendicular to,
Volchansky, Bogoslovsk-Veselovsky) (Tuzhikova or at a high angle to, the strike of the belt and
1973) containing Upper Triassic coal-bearing by the identification of a zone with anomalously
sediments were complexly deformed. In the low heat flow (Golovanova 2006). According to
Polar Urals, the Triassic deposits of the foredeep Mikhailov et al. (2002), mountain building is
and the Chernyshov and Chernov range are all accompanied by westward-directed thrusting.
deformed, and are unconformably overlain by The timing of mountain building is controver-
Middle Jurassic strata. However, in the nearby sial. Until recently, there was a consensus that oro-
Severososvinsky graben to the east, Triassic and genesis began in the Late Oligocene and continued
Jurassic deposits are not deformed (Ivanov et al. into the Quaternary inclusive (Trifonov 1999;
2004), attesting to the localized nature of Cimmer- Rozhdestvensky & Zinyakhina 1997) and that
ian orogenic events. ancient peneplains formed in the Triassic– Jurassic
The Pay-Khoy and Novozemelsky ranges were were preserved (Sigov 1969; Borisevich 1992).
formed in the Cimmerian (Korago et al. 1989; On the contrary, Puchkov (2002b) pointed out
Yudin 1994). Cimmerian orogenesis is attributed that models proposing Late Oligocene uplift of the
to a large-scale intra-Pangaean strike-slip faulting, Urals are inconsistent with: (1) the occurrence of
accompanied by block rotation, possibly reflecting Miocene oligomictic quartz sands and sandstones
lateral escape of Kazakhstania between Laurussia (Yakhimovich & Andrianova 1959; Kozlov 1976),
and Siberia and an immediate collision of the which indicate stable non-orogenic conditions of
weathering, erosion and deposition in both foredeep Sobornov & Bushuev 1992), combined with drilling,
terrane and the Urals itself. The first appearance of helped to solve some structural problems in the
polymictic sediments, indicating more rapid uplift Uralian foreland. In 1993, the commencement of
and erosion, is Late Pliocene in age (Verbitskaya the EUROPROBE Programme ‘Uralides’, involved
1964); (2) deep Miocene river incision is best acquisition and interpretation of seismic data along
explained by a drop of the Caspian sea level, Messi- two regional profiles (the Southern and Middle
nian crisis, as well documented in the Mediterr- Urals) and re-interpretation of some existing
anean (Milanovsky 1963), rather than by crustal shorter profiles. A combined geological and multi-
uplift; (3) no well-documented Miocene –Early component geophysical URSEIS-95 project in the
Pleistocene terraces occur in the river valleys of Southern Urals (Berzin et al. 1996; Carbonell et al.
the Urals; (4) no cave deposits older than 1996; Echtler et al. 1996; Knapp et al. 1996;
Middle-Upper Neopleistocene are found; and (5) Suleimanov 2006), included a c. 500 km-long
the velocities of the modern uplift of the Urals seismic reflection line across most of the orogen at
surface are 5 –7 mm a21, which is an order of mag- a latitude of Kraka and Gebyk massifs. The
nitude faster than the time needed to build the Urals ESRU-SB profile, ultimately c. 440 km long,
mountains since Oligocene. crossed the Middle Urals where the ‘superdeep’
On the other hand, fission-track (Seward et al. SG-4 borehole (currently c. 5.5 km deep) is located
1997; Glasmacher et al. 2002) and unpublished (Kashubin et al. 2006, and references therein).
U/Th –He data show that the relief of the axial
part of the Southern Urals was not completely The URSEIS profile (Fig. 15)
stabilized by the Late Cretaceous. Puchkov &
Danukalova (2004) demonstrated that the altitudes The interpretation given here is based on combined
of the base of shallow marine Upper Cretaceous (vibroseis and explosion) seismic section along the
deposits increase progressively in the direction of geotraverse profile, after Suleimanov (2006), Spets-
the mountain ridge, disappearing at elevations of Geofizika. The coherency-filtered, depth-migrated
500 m. Therefore no Triassic– Jurassic planation vibroseis data by Tryggvason et al. (2001) were
surfaces can be preserved in the modern surface. also used as an alternative source of information
The depth of erosion since the Cretaceous is for the upper and middle crust. Along with generally
between 1000 and 2000 m (depending on the accepted conclusions (Brown et al. 2008, and refer-
thermal gradient), and these numbers are several ences therein), the following interpretations contain
times greater than the previous estimates. some latest original inferences of the author.
From 500 km (in the Preuralian foredeep), to the
Main Uralian Fault at c. 275 km, the survey charac-
The deep structure of the Urals terizes the structure of the foreland fold-and-thrust
The main milestones of the study belt (Fig. 15). From 500–c. 420 km, subhorizontal,
moderately coherent reflectivity in the upper 5 km
Fifteen regional deep seismic survey (DSS) profiles, corresponds to weakly deformed Palaeozoic fore-
made between 1961 and 1993 permit the definition land basin (foredeep) and platform margin shelf
of the Moho surface beneath the Urals and demon- rocks of Ordovician –Lower Permian age (Brown
strated its layered seismic structure and the anoma- et al. 2006b). Below this, to approximately 20 km
lous character of its crust. In particular, these depth, strongly coherent, subhorizontal reflectivity
surveys suggest the presence of a crustal ‘root’ is interpreted to represent undeformed Meso- and
under the Tagil –Magnitogorsk zone and a Neoproterozoic strata of the SSE prolongation of
complex compositional transition zone in the the Kama-Belsk aulacogen. The base of the reflec-
lower crust, with Vp velocities between 7.2 and tivity here is thought to represent the unconformity
7.8 km/s (Druzhinin et al. 1976). Puchkov & Sve- between undeformed and low-metamorphic
tlakova (1993) placed these results into a plate tec- Mesoproterozoic strata and the non-reflective
tonic context for the first time, by interpreting a Archaean-Palaeoproterozoic crystalline basement
DSS profile in the Middle Urals as an indicator of (Dianconescu et al. 1998; Echtler et al. 1996). The
the bi-vergent character of the Uralian orogen. sedimentary prism, almost 20 km thick, has a
Reflection profiles made between 1964 and the convex lens-like shape, consistent with the
early 1990s in the Magnitogorsk and Tagil zones interpretation that the prism represents an inverted
(e.g. Menshikov et al. 1983; Sokolov 1992) revealed aulacogen. Beneath the crystalline basement, at
the inclined reflectors that define synclinoria and c. 430 km the Moho is cut by Makarovo normal
anticlinoria, and along-strike variations in the mor- (?) fault, with up to 5 km of amplitude – probably
phology of the Main Uralian fault. In the 1980s, related to the rift nature of the aulacogen (Fig. 15).
state oil company surveys along the western slope To the east, the upper and middle crust has weak,
of the Urals (e.g. Skripiy & Yunusov 1989; gently east-dipping reflectivity that, between
184
V. N. PUCHKOV
Fig. 15. (a) Uninterpreted combined (vibro- and explosion) seismic section along the URSEIS geotraverse profile (the seismic data after Suleimanov 2006, SpetsGeofizika).
(b) Geological interpretation, overlain on the profile. See Figure 4 for location.
420 km and the MUF is concave downward. This between the East Uralian and Trans-Uralian zones
reflectivity is associated with the Precambrian is thought to be a regional fault called Kartaly or
rocks in the Bashkirian Anticlinorium which, in its (wrongly) Troitsk fault located immediately to the
eastern part, was deformed during the Timanide east of Dzhabyk massif and traced in the SSW and
orogeny (Puchkov 2000). The base of the reflectiv- NNE directions where it is interpreted as a wrench
ity is usually interpreted to be the basal detachment fault of a considerable amplitude.
contact between Mesoproterozoic strata and the In the western and eastern parts of the profile, the
Archaean– Palaeoproterozoic crystalline basement. Moho is imaged in the URSEIS combined (vibro-
However, according to structural studies, large and dynamite) reflection data to a depth of
anticlines of the central and eastern part of the c. 50 km but cannot be traced in the deeper,
Bashkirian anticlinoria have detached blocks of central portion of the profile. The Moho has been
the crystalline basement beneath them, close to the determined from wide-angle data to occur at a
surface. Taratash anticline in the north exposes maximum depth of 55 km (Carbonell et al. 1998).
such Precambrian core in the surface. Although there is some bias between the wide-angle
The lower crust beneath the foreland and CDP data, a cloudy reflection under this depth
fold-and-thrust belt is weakly- to non-reflective, at c. 250 km, can be tentatively interpreted as a
though the Moho boundary can be traced by wedge of the lower crust protruding into the
explosion seismic data further to the east, towards mantle (compare with a much better imaged
the MUF. Close to the MUF, the deep-seated wedge of the lower crust in the ESRU-SB profile,
Uraltau antiform is clearly imaged and the MUF see below) (Fig. 16).
fault is traced as a gently concave structure,
mainly by a loss of reflections from the Precambrian The ESRU-SB profile (Fig. 16)
rocks and assuming that the base of the island-arc
complex in the hanging wall of MUF is transparent. The latest interpretation of crustal structure of the
The Zilair synform and Uraltau antiform immedi- Middle Urals at latitude 56 –628 based on seismic
ately to the west of the MUF form a dynamic reflection data obtained by yearly installments
couple, with the antiform making a tectonic wedge since 1993 (ESRU-SB profile) was given recently
downthrusted to the west under the antiform. by Kashubin et al. (2006), Rybalka et al. (2006)
From the MUF to c. 180 km, the Magnitogorsk and Brown et al. (2008). From –100 km in the Pre-
arc is almost non-reflective in the upper crust, ruralian foredeep to c. –25 km in the east, the upper
though the east-vergent Kizil thrust is clearly crust has a flat-lying reflectivity, interpreted to rep-
imaged in the coherency-filtered vibroseis data resent an undeformed foredeep orogenic basin and
(Tryggvason et al. 2001) and its interpretation is platformal deposits (up to c. –65 km). In contrast,
supported by deep drilling and recent short seismic to the east, the steeply east-dipping shallow reflec-
reflection profiles. The middle and lower crust tivity of the ‘thin-skinned’ fold-and-thrust foreland
is relatively transparent. The contact between occurs (Fig. 16). Both undeformed and steeply-
the Magnitogorsk arc and the East Uralian Zone dipping reflections are underlain by a gently east-
at c. 180 km (the East Magnitogorsk Fault and dipping zone of reflections at a depth of c. 5–8 km
suture zone) is imaged by a sharp change from that is interpreted as a low-angle unconformity
almost transparent crust in the west to coherent, surface between the platform cover of Ediacarian
highly reflective, middle crust to the east. In the and Palaeozoic age and the older Neoproterozoic
East Uralian Zone, from c. 180 –100 km, the upper (Upper Riphean), that is transformed in the east
crust is nearly transparent down to about 8 km, into a basal detachment of the fold-and-thrust belt
corresponding to the Gebyk granite. Below this, a (Brown et al. 2006c). At c. –25 km this reflectivity
series of short east-dipping and subhorizontal is abruptly truncated by a series of steep east-
reflectors are descending into the middle crust. dipping concave reflectors corresponding probably
The lower crust is almost transparent or semi- to listric-like faults, traced into the middle crust to
transparent, except in the east, where a zone a depth of 25–30 km. This type of reflectivity
with strong west-dipping reflectivity extends persists from a distance mark of 25 km to the
downward and westward from the Trans-Uralian MUF zone. The zone is characterized by several
Zone (a continuation of Kartaly reflections; see pronounced closely-spaced reflectors dipping to
below). the east at angles of 458 –608 between 0 and
The crust of the Trans-Uralian Zone is imaged as 10 km, imaging an imbrication zone of the
west-dipping, strongly coherent reflectivity called strongly deformed margin of Laurussia continent.
the Kartaly Reflection Sequence (KRS, Fig. 15), From –20 –10 km, the steeply east-dipping reflec-
which merges with the Moho in a system of tors represent ‘thick-skinned’ deformation in the
thrusts; in this region the Moho appears as a near- Precambrian-cored Kvarkush anticlinorium and
horizontal detachment fault. The boundary Early Palaeozoic hanging wall.
186
V. N. PUCHKOV
Fig. 16. Seismic cross-sections (a) Uninterpreted, and (b) Interpreted line drawings of the coherency filtered, depth-migrated ESRU-SB data (Kashubin et al. 2006 Rybalka
et al. 2006). See Figure 4 for location. Abbreviations in the Figure A: MUFZ, Main Uralian Fault zone; SMZ, Serov-Mauk Fault zone; MAMC, Murzinka-Aduy
metamorphic complex.
Below the undeformed foreland basin and the The next zone to the east, Murzinka-Adui zone
basal detachment of the fold-and-thrust belt, the (103 –120 km), is represented at the surface by Neo-
middle crust exhibits wave-like concave to convex proterozoic metamorphic rocks and Permian gran-
reflectivity down to approximately 25 km depth. ites, and together with the Salda zone belongs to
From 25 km to c. 42 km depth, the lower crust is the East Uralian megazone. The character of its
characterized by more coherent and strong, sub- reflectivity in the upper crust is incoherent and
horizontal reflectivity. The middle crustal reflectivity patchy, and does not permit recognition of its
probably images the Neoproterozoic and Mesoproter- detailed structure. Further east, from 120–180 km
ozoic sedimentary rocks, and the lower crustal within the Trans-Uralian zone, the upper crustal
reflectivity images Archaean–Palaeoproterozoic structure is more difficult to interpret owing to
crystalline basement rocks of the Laurussian both poor surface exposure and the almost complete
margin, though there is a striking difference absence of coherent reflectors in the upper 10-km of
between the reflectance character of the crystalline the profile.
basement here and in the URSEIS. The character From c. 180 km to the end of the profile
of crustal reflectivity in the deep part of the (260 km), the platformal Cretaceous and Cenozoic
section suggests that it is unaffected by the strata of the West Siberian Basin appear to be
Uralide deformation which also makes a difference characterized by a zone of good subhorizontal
between the profiles (Brown et al. 2006c; Kashubin reflectivity which thickens to the east up to 1.5 km
et al. 2006). On the other hand, the reflectivity at 260 km. The details of the profile imaged
pattern of the middle crust suggests that at – 100 by Rybalka et al. (2006) show a relief of the
to c. –20 km, the 15 or more kilometre-thick Palaeozoic –Cretaceous unconformity surface,
Meso- and Neoproterozoic strata form a large which is strongly uneven probably due to pre-
synform that is underthrust by an antiformal tectonic Cretaceous grabens, flexures and river-bed
wedge composed of the rocks of the same age. The incisions. Under the platformal cover, the structure
structure is characteristic of a Timanian foreland of the Trans-Uralian zone reveals a series of west-
deformation, but also resembles the wedge-like dipping reflectors, merging with the Moho at a
relationships between the Uraltau antiform and depth of c. 40 km, in a manner similar to that seen
Zilair synform imaged by URSEIS, though their in the KRS of the URSEIS profile, although not
relationships had been formed by the Uralide as bright.
orogeny. In general, the Moho is very well defined along
The Moho surface is traced here as a gently east- the whole profile as a sharp boundary between
dipping boundary between highly reflective lower highly reflective lower crust and an almost transpar-
crust and almost transparent mantle at a depth of ent mantle. The crust thickens from c. 42 –43 km in
42– 45 km. From about c. 10 –50 km, the upper both the west and east to nearly 60 km beneath the
crust of the Tagil arc is imaged as an open Central Uralian zone. The above-mentioned wedge
synform, thrust to the west over the Meso- of the lower crust protruding into the mantle gives
Neoproterozoic rocks of Kvarkush aniclinorium. here the impression that the eastern limb of the
The Tagil synform is asymmetric, and its eastern orogen is thrust under the former Laurussian lower
limb is limited by a serpentinite mélange of the crust and Moho.
Serov-Mauk fault zone, separating the synform
from the Salda metamorphic complex. The Discussion
mélange zone is transparent and its western bound-
ary can be traced along abrupt truncations of Tagil The development of the Uralides in the Palaeozoic
reflectors, suggesting a westerly dip of the zone at preserves many of the characteristics of a full
an angle of 608. The zone can be traced tentatively Wilson cycle. However, if the concept of such a
into the lower crust along weak and diffuse reflec- cycle is restricted to a classical ‘accordion-type’
tors, changing the steep western dip of the zone to development, wherein the continent that rifted
a gentler 308 close to Moho surface. away returns back to collide (as it was in the case
The Salda metamorphic complex of probable of Rheic ocean), then the development of the Ura-
island-arc nature (Rybalka et al. 2006), situated lides does not conform with the idea. According to
between 55 and 103 km, is characterized by a palaeomagnetic data and geodynamic reconstruc-
series of west-dipping reflections, which can be tions (e.g. Puchkov 2000; Kurenkov et al. 2002;
traced under the Tagil synform and Central Svyazhina et al. 2003; Levashova et al. 2003),
Uralian zone together with the Serov-Mauk fault. between Ordovician and Early Permian time, the
This portion of the crust, correlated with the Salda microcontinents of the East Uralian zone and Kok-
zone, is wedge-like, protruding down into the chetav block of Kazakhstania were transported at
mantle under the western slope of the Urals to a least 2000 km from north to south parallel to the
depth of 60 km. margin of Baltica. Kazakhstania as a whole
188 V. N. PUCHKOV
intervened between Siberia and Baltica in the Devo- deformation of the whole crust with the Moho as a
nian, and Siberia rotated clockwise at 908 – prob- detachment, in the young, eastern limb of the Ura-
ably because of the arrival of Kazakhstania. lides (see above), absorbed a considerable part of
Kazakhstania was squeezed between accreted strain. The preservation of ophiolites may depend
margins of Siberia and Laurussia, forming a pro- on strain. In zones of higher strain, ophiolites are
nounced horseshoe-like orocline (Fig. 1). It looks squeezed from sutures as allochthonous sheets
more like a rock’n’roll dance than an accordion-like aided by the formation of rheologically weak ser-
motion. On the other hand, only a minority of pentinites which may also have acted as a lubricant.
orogens belongs to the regular ‘accordion’ type, This interpretation is supported by experiments that
and therefore we prefer to follow to a more liberal demonstrate the rheological weakness of serpenti-
understanding of Wilson cycle. nite (e.g. Escartin et al. 1997; Hilairet et al. 2007)
The rift processes at the beginning of the cycle and by structural studies in the foreland
are demonstrated by a profound difference fold-and-thrust belt of the Southern and Middle
between the strikes of Timanide structures and the Urals, where shortening deduced from balanced
Main Uralian Fault to the north of the Poliud geological sections is anomalously low (14–17%,
Range (Fig. 7). The accompanying sedimentary Brown et al. 1997). Shortening increases to the
deposits are characterized by irregularly distributed north, and in the Mikhailovsk and Serebryansk pro-
coarse-grained polymictic to arkosic sediments files it is c. 30% (calculated after Brown et al.
unconformably overlying the basement and 2006c). In the Cis-Polar and Polar Urals, however,
accompanied by subalkaline basaltic volcanism. shortening can be still much greater, judging by
Similar formations on the sialic blocks of the East the upper section of figure 4 in Puchkov (1997);
Uralian and Transuralian zones are c. 15– 20 Ma see also Yudin (1994). This may be explained by
younger than those in the Western zones, indicating the wedging-out of Kazakstania to the north,
that the East Uralian microcontinent was not pre- where two rigid cratons, Laurussia and Siberia,
viously detached from the same place where it come into contact.
finally docked. Modern analogies to the arc-continent collision
The suggested Early Palaeozoic plume magma- in the Urals include the Indonesian, Taiwan,
tism described above is considerably younger than Tyrrhenian and Greater Antilles arcs. The evolution
the rifting event, and is only indirectly connected, is similar to that proposed for the Taconian arc in
possibly in an analogous manner to the relationship the Appalachian orogen, and the mid-Devonian
between the transverse chain of the Cretaceous arc collision with Baltica along the margins of the
Monteregian alkaline with carbonatite intrusions Rheic Ocean.
of eastern North America and the Mesozoic rifts Another striking feature of the Uralides, is a long
along the margin of the Atlantic Ocean. duration and recurrence of orogenic events – from
Widespread development of ophiolite com- the Devonian until the Early Jurassic, which is com-
plexes, almost unprecedented among the Palaeozoic parable to the duration of Appalachian orogenic
or earlier foldbelts, is one of the most important activity. The oblique, transpressive character and
characteristics of the Uralides. Ophiolites appear the diachroneity of the continent–continent col-
to represent anomalous portions of the oceanic lision in the Urals is similar to that of the Alleghe-
tract (Aden and Red sea-type, supra-subduction nian orogeny (Engelder 2007). But the absence of
zone basalts, Lanzo-type mantle blocks) whereas Caledonian collisions, from the Ordovician to the
most typical MORB appears to have been almost Middle Devonian, is another striking feature of the
eliminated by subduction. Island-arc complexes Uralides.
are also widely spread in the Uralides, and sub- The characteristics of two regional profiles trans-
sequent uplift and erosion offers a rare possibility ecting the orogen, confirm the bi-vergent symmetry
of studying the deeper structure of an island arc, of it. Alternatively, uniformly vergent orogens may
than is available for study in modern island arcs. be part of a bigger, bi-vergent one, that has been dis-
The internal structure of the arcs exhibits a moderate membered and dispersed by subsequent tectonic
strain (Brown et al. 2001). According to Alvarez- events (e.g. Greenland and Scandinavian Caledo-
Marron (2002) the Uralides ‘may be seen as a nides) or its lacking limb overlain by a younger
factory for “making” new continental crust in con- orogen (e.g. a Variscan basement of Alps).
trast to the Variscides which is a factory for “recy- In the western Uralides, where the foreland is
cling” existing continental crust’. cratonic (i.e. ancient and characterized by a thick
One of the possible explanations for the excep- and rigid lithosphere), the crystalline basement
tionally good preservation of oceanic complexes and Moho surface are not affected, or only partly
in the Uralides is the low rigidity of the Kazakhsta- affected, by the main Uralide orogenesis. In the
nian plate, which became continental crust only in eastern Uralides, however, the Moho is interpreted
the Silurian (Puchkov 1996a). The deep-seated as a detachment, and the crust is deformed to great
depths, an interpretation which supports the idea of Earth’s crust formation must be also highly appreciated.
a comparative weakness of the Palaeozoic crust and The last, but not the least of my expression of gratitude
lithosphere as a whole. is to the reviewers, V. Ramos and R. Ernst and to
In the west, the regional and local profiles show B. Murphy, the editor, who helped greatly at the last
stages of the work.
an abrupt transition from ‘thin-skinned’ – to
thick-skinned tectonics along a sharp ramp due to
an abrupt change in the plasticity of the rocks. In References
the east we do not see any ‘thin-skinned’ tectonics
A GARD , P. & L EMOINE , M. 2005. Faces of the Alps:
at all. This may reflect the incompleteness of the Structure and Geodynamic Evolution. Commission of
eastern part of the profiles (URSEIS had been the geological map of the world, Paris.
stopped at a state border with Kazakhstan and A LEKSEIEV , D. V. 2008. Kinematics of the Variscan Ural
ESRU – in swamps of Siberia) (Fig. 4). On the and Tian-Shan junction area. Geologie de la France,
other hand, Late Palaeozoic orogenic processes 2, 75.
reached much farther to the east than the initial ALL -R USSIAN S TRATIGRAPHIC C OMMITTEE . 1993.
boundary between Uralides and Kazakhstanides Stratigraphic Schemes of the Urals: Precambrian
(Fig. 1). In the Central Kazakhstanian Caledonides, and Paleozoic. Ekaterinburg, 151 scheme (in Russian).
orogenic processes are documented by Permian A LVAREZ -M ARRON , J. 2002. Tectonic processes during
collisional orogenesis from comparison of the
deformation of different styles, voluminous Late Southern Urals with the Central Variscides. In:
Palaeozoic syn-orogenic granite and deposition of B ROWN , D., J UHLIN , C. & P UCHKOV , V. (eds) Moun-
the contemporaneous molasse in huge intermontane tain Building in the Uralides: Pangea to Present.
basins such as Chu-Sarysu and Teniz (YUGGEO American Geophysical Union, Geophysical Mono-
2002). Where the Upper Palaeozoic epi-Caledonian graph, 132, 83– 100.
sedimentary cover is preserved, as in the Greater A MON , E. O. 2001. The marine equatorial basin of the
Karatau, the east-vergent ‘thin-skinned’ tectonics Uralian region in the Mid- and Late Cretaceous
was developed during the Upper Carboniferous epoch. The Geology and Geophysics, 42, 471–483
and Permian (Alekseiev 2008). (in Russian).
A NDREICHEV , V. L., R ONKIN , YU . K., L EPIKHINA , O. P.
& L ITVINENKO , A. F. 2005. Rb–Sr and Sm–Nd
Concluding remarks Isotopic-geochronometric Systems in Basalts of the
Polar Cis-Urals. Syktyvkar, Geoprint (in Russian).
Extremely good preservation of oceanic and A RTIUSHKOVA , O. V. & M ASLOV , V. A. 2003. Lower
island-arc complexes and low degree of shortening Devonian (pre-Upper Emsian) deposits of the Magni-
in the foreland belt are unprecedented in the Palaeo- togorsk megazone. Geological Sbornik, IG USC
RAS, 3, 80–87 (in Russian).
zoic orogens and give the Uralides a real individua-
B ATANOVA , V. G., B RUEGMANN , G., B ELOUSOV , I. A.,
lity. Many other features of the Urals are also rare, S AVELIEVA , G. N. & S OBOLEV , A. V. 2007. The pro-
such as its well-preserved bi-lateral structure, cesses of melting and migration of melts in mantle
island-arc related platinum-bearing belt, demonstra- based on a study of highly siderophile elements and
tive arc-continent collision, diachroneity of col- their isotopes. In: G ALIMOV , E. M. (ed.) XVIII Vino-
lisions, a combination of orogenic and rift-related gradov Symposium on Geochemistry of Isotopes. V.,
magmatism in a single stage of transpressive defor- GEOKHI, 40– 41 (in Russian).
mation of the lithosphere, and the preservation of a B EA , F., F ERSHTATER , G. & M ONTERO , P. 2002. Grani-
heavy, relatively ‘cold’, isostatically equilibrated toids of the Urals: implications for the evolution of the
root. On a plate scale, however, the history of the orogen. In: B ROWN , D., J UHLIN , C. & P UCHKOV , V.
(eds) Mountain Building in the Uralides: Pangea to
Uralides follows the main stages of a Wilson Present. American Geophysical Union, Geophysical
cycle, modified somewhat by episodes of plume- Monograph, 132, 211–232.
related tectonics and magmatism in the Early B EANE , R. J. & C ONNELLY , J. H. 2000. 40Ar– 39Ar,
Palaeozoic and Triassic. U– Pb and Sm–Nd constraints on the timing of meta-
morphic events in the Maksyutov Complex, Southern
Fieldwork campaigns, workshops and informal discus- Ural Mountains. Journal of the Geological Society,
sions in more than a decade of co-operation with prominent London, 157, 811–822.
European and American geologists during the activity of B ELL , K. 2001. Carbonatites: relationships to mantle-
the EUROPROBE Commission in the Urals have consider- plume activity. In: E RNST , R. E. & B UCHAN , K. L.
ably stimulated the work of the author. The author (eds) Mantle Plumes: Their Identification Through
acknowledges also the importance of his participation in Time. Geological Society of America Special Paper,
the IGCP-453 ‘Uniformitarianism revisited: a comparison 352, 267–290.
between modern and ancient orogens’, which ultimately B ERZIN , R., O NCKEN , O., K NAPP , J. H., P EREZ -E STAUN ,
inspired him to write this paper. The financial support of A., H ISMATULIN , T., Y UNUSOV , N. & L IPILIN , A.
the Department of Earth’s Sciences of the Russian 1996. Orogenic evolution of the Ural Mountains:
Academy of Sciences, Program No. 10 The Central results from an integrated seismic experiment.
Asian mobile belt: geodynamics and stages of the Science, 274, 220–221.
190 V. N. PUCHKOV
B EUTLER , G. 1979. Verbreitung und Charakter der altkim- B URTMAN , V. S. 2006. Tien Shan and High Asia.
merischen Hauptdiskordanz in Mitteleuropa. Zeit- Tectonics and Geodynamics in Paleozoic. Moscow,
schrift für Geologische Wissenschaften, 5, 617–632. GEOS (in Russian).
B ISKE , YU . S. 2006. The Foldbelts of the Northern C ARBONELL , R., P EREZ -E STAÚN , A. ET AL . 1996. A
Eurasia. Tyanshan Foldbelt. St. Petersburg University crustal root beneath the Urals: Wide-angle seismic
(in Russian). evidence. Science, 274, 222–224.
B OCHKAREV , V. V. & Y AZEVA , R. G. 2000. Subalkaline C ARBONELL , R., L ECERF , D., I TZIN , M., G ALLART , J. &
Magmatism of the Urals. Ekaterinburg (in Russian). B ROWN , D. 1998. Mapping the Moho beneath the
B ORISEVICH , D. V. 1992. Neotectonics of the Urals. Geo- Southern Urals. Geophysical Research Letters, 25,
tectonics, 26, 41–47 (in Russian). 4229– 4233.
B OROZDINA , G. N. 2006. The history of geological devel- C HUVASHOV , B. I., I VANOVA , R. M. & K OLCHINA , A. N.
opment of the Ragil megazone of the Middle and 1984. The Upper Paleozoic of the Eastern Slope of
Southern part of the Middle Urals during the Lower the Urals. Uralian Scientific Centre, Sverdlovsk (in
Paleozoic. Candidate of Science thesis, Ekaterinburg, Russian).
IGG Uralian Branch of Russian Academy (in Russian). C OBBOLD , P. R., M EISLING , K. E. & M OUNT , V. S. 2001.
B OROZDINA , G. N., I VANOV , K. S., N ASSEDKINA , V. A. Reactivation of obliquely rifted margin, Campos and
& S NIGIREVA , M. P. 2004. On the age and volume of Santos basins, southeastern Brazil. AAPG Bulletin,
the Shemur formation of the Tagil megazone. 85, 1925–1944.
Ezhegodnik-2003, Institute of Geology and Geophy- C OCKS , L. R. M. & T ORSVIK , T. H. 2006. European
sics, Sverdlovsk, 10– 13 (in Russian). geography in a global context from the Vendian to
B ORTNIKOV , P. S., S AVELIEVA , G. N., M ATUKOV , D. I., the end of Paleozoic. In: G EE , D. G. & S TEPHENSON ,
S ERGEEV , S. A., B AREZHNAYA , N. G., L EPEKHINA , R. A. (eds) European Lithosphere Dynamics. Geologi-
E. N. & A NTONOV , A. V. 2005. The age of a zircon cal Society, London, Memoirs, 32, 83–95.
from plagiogranites and gabbro after SHRIMP data: D ALZIEL , I. 1992. On the organization of American plates
pleistocene intrusion in the rift valley of the Mid- in the Neoproterozoic and the breakout of Laurentia.
Atlantic Ridge, 58300 600 –58320 400 N. Doklady of GSA Today, 11, 237–241.
Russian Academy of Science, Geology, 404, 94– 99 D AVYDOV , V. I., C HERNYKH , V. V., C HUVASHOV , B. I.,
(in Russian). N ORTHRUP , K. J. & S NYDER , V. S. 2002. Volcanic
B ROWN , D. & P UCHKOV , V. 2004. Arc-continent Col- tuff layers in the Upper Paleozoic of the Southern
lision in the Uralides. An IGCP-453 ‘Uniformitarian- Urals and prospects of creation of exactly calibrated
ism revisited: a comparison between modern and time scale of the Carboniferous. In: C HUVASHOV ,
ancient orogens’ conference and fieldtrip. Design Poli- B. I. (ed.) Stratigraphy and Paleogeography of the
graph Service, Ufa. Carboniferous of Eurasia. Ekaterinburg, Institute of
B ROWN , D., A LVAREZ -M ARRON , J., P EREZ -E STAUN , A., Geology and Geochemistry, 112 –123 (in Russian).
G OROZHANINA , Y., B ARYSHEV , V. & P UCHKOV , V. D IANCONESCU , C. C., K NAPP , J. H., B ROWN , L. D.,
1997. Geometric and kinematic evolution of the S TEER , D. N. & S TILLER , M. 1998. Precambrian
foreland thrust and fold belt in the southern Urals. Moho offset and tectonic stability of the East European
Tectonics, 16, 551– 562. platform from the URSEIS deep seismic profile.
B ROWN , D., A LVAREZ -M ARRON , J., P EREZ -E STAUN , Geology, 26, 211 –214.
A., P UCHKOV , V., A YARZA , P. & G OROZHANINA , D IDENKO , A. H., K URENKOV , S. A. ET AL . 2001. Tectonic
Y. 2001. Structure and evolution of the Magnitogorsk History of the Polar Urals. M. Nauka (in Russian).
forearc basin: identifying upper crustal processes D OBRETSOV , N. L., K IRDYASHKIN , A. G. & K IRDYASH-
during arc-continent collision in the Southern Urals. KIN , A. 2001. The Deep Geodynamics. Novosibirsk,
Tectonics, 20, 364– 375. GEO (in Russian).
B ROWN , D., P UCHKOV , V. N., A LVAREZ -M ARRON , J. & D RUZHININ , V. S., R YBALKA , V. M. & S OBOLEV , I. D.
P EREZ -E STAUN , A. 2006a. Tectonic processes in the 1976. The Connection of the Tectonics and Magmatism
Southern and Middle Urals: an overview. In: G EE , with the Deep Structure after DSS Data. Sverdlovsk
D. G. & S TEPHENSON , R. A. (eds) European Litho- (in Russian).
sphere Dynamics. Geological Society, London, E CHTLER , H. P., S TILLER , M. ET AL . 1996. Preserved
Memoirs, 32, 407–419. collisional crustal architecture of the Southern Urals
B ROWN , D., S PADEA , P. ET AL . 2006b. Arc-continent – Vibroseis CMP-profiling. Science, 274, 224–226.
collision in the Southern Urals. Earth-Science E FIMOV , A. A., E FIMOVA , L. P. & M AEGOV , V. I. 1993.
Reviews, 79, 261– 287. Tectonics of the Platinum-bearing belt of the Urals:
B ROWN , D., J UHLIN , C., T RYGGVASON , A., F RIBERG , relations of complexes and mechanism of structure for-
M., P UCHKOV , V., R YBALKA , A. & P ETROV , G. mation. Geotectonics, 3, 4– 46 (in Russian).
2006c. Structural architecture of the southern and E NGELDER , T. 2007. The Alleghanian orogeny of
middle Urals foreland from reflection seismic profiles. the Appalachian Mountains. Geologie de la France,
Tectonics, 25, 1– 12. 2, 94.
B ROWN , D., J UHLIN , C. ET AL . 2008. Mountain building E SCARTIN , J., H IRTH , G. & E VANS , B. 1997. Nondilatant
processes during continent-continent collision in the brittle deformation of serpentinites: implications for
Uralides. Earth-Science Reviews, 89, 177– 195. Mohr-Coulomb theory and the strength of faults.
B URNS , L. E. 1985. The Border Ranges ultramafic and Journal of Geophysical Research, 102, B2, 2897.
mafic complex, south-central Alaska: cumulate frac- F ERSHTATER , G. B., B EA , F. & M ONTERO , P. 2006.
tionates of island-arc volcanics. Canadian Journal of Granitoids. In: M OROZOV , A. F. (ed.) Structure
Earth Sciences, 22, 1020–1038. and Dynamics of Lithosphere of the Eastern
Europe, Issue 2. The Results of the EUROPROBE massif (Polar Urals). Geology of Ore deposits, 45,
Research. Moscow, GEOKART, GEOS, 449– 461 (in 353– 371 (in Russian).
Russian). H ETZEL , R. & R OMER , R. L. 1999. U –Pb dating of the
F RANKE , W. 2000. The Mid-European segment of the Verkniy Ufaley intrusion, middle Urals, Russia: a
Variscides: tectonostratigraphic units, terrane bound- minimum age for subduction and amphibolite facies
aries and plate tectonic evolution. In: F RANKE , W., overprintof the East European continental margin.
H AAK , V., O NCKEN , O. & T ANNER , D. (eds) Geological Magazine, 136, 593– 597.
Orogenic Processes: Quantification and Modelling in H ETZEL , R. & R OMER , R. L. 2000. A moderate exhuma-
the Variscan Belt. Geological Society, London, tion rate for the high-pressure Maksyutov Complex,
Special Publications, 179, 35–61. Southern Urals, Russia. Geological Journal, 35,
G AGGERO , L., S PADEA , P. & C ORTESOGNO , L. 1997. 327– 344.
Geochemical investigation of the igneous rocks from H ILAIRET , N., R EYNARD , B., W ANG , YANBIN , D ANIEL ,
the Nurali ophiolite melange zone, Southern Urals. I., M ERKEL , S., N ISHIYAMA , N. & P ETITGIRARD , S.
Tectonophysics, 276, 139–161. 2007. High-pressure creep of serpentine, interseismic
G EE , D. G., B OGOLEPOVA , O. K. & L ORENZ , H. 2006. deformation, and initiation of subduction. Science,
The Timanide, Caledonide and Uralide orogen in the 318(5858), 1910– 1913.
Eurasian high Arctic, and relationships to the palaeo- INTERNATIONAL COMMITTEE OF TECTONIC M APS .
continents Laurentia, Baltica and Siberia European 1982. Tectonics of Europe and adjacent areas. Varis-
Lithosphere Dynamics. In: G EE , D. G. & S TEPHENSON , cides, epi-Paleozoic platforms and Alpides. Explana-
R. A. (eds) European Lithosphere Dynamics. Geologi- tory note to the International tecnonic map of Europe
cal Society, London, Memoirs, 32, 507– 520. and adjacent areas, scale 1:2 5000 000. M. Nauka
G EOFFROY , L. 2005. Volcanic passive margins. Comptes Publishing House.
Rendu Geoscience, 337, 1395–1408. I VANOV , A. V., R ASSKAZOV , S. V., F EOKTISTOV , G. D.,
G LASMACHER , U. A., W AGNER , G. A. & P UCHKOV , H UAIYU , H. E. & B OVEN , A. 2005. 40Ar– 39Ar dating
V. N. 2002. Thermotectonic evolution of the western of Usol’skii sill in the south-eastern Siberian Traps
fold-and-thrust belt, southern Urals, Russia, as Large Igneous Province: evidence for long-lived
revealed by apatite fission-track data. Tectonophysics, magmatism. Terra Nova, 17, 203 –208.
354, 25–48. I VANOV , K. S. 2001. The evaluation of paleovelocities of
G LODNY , J., A USTRHEIM , H., M ONTERO , P. & R USIN , A. subduction and collision in the origin of the Urals.
1999. The Marun-Keu Metamorphic Complex, Polar Transactions (Doklady) of the Russian Academy of
Urals, Russia Protolith ages, Eclogite facies fluid- Sciences/Earth Science Section, 377, 231– 234 (in
rock interaction, and Exhumation History. EUG– 10 Russian).
Abstracts, Cambridge Publications, 80. I VANOV , K. S., K ARSTEN , L. A. & M ALUSKI , G. 2000.
G LODNY , J., B INGEN , B., A USTRHEIM , H., M OLINA , J. F. The first data on the age of subductional (eclogite-
& R USIN , A. 2002. Precise eclogitization ages deduced glaucophane) metamorphism in the Polar Urals. In:
from Rb/Sr mineral systematics: the Maksyutov K OROTEEV , V., I VANOV , K. & B OCHKAREV , V.
Complex, Southern Urals, Russia. Geochimica et Cos- (eds) Paleozones of Subduction: Nectonics, Magma-
mochimica Acta, 66, 1221– 1235. tism, Metamorphism, Sedimentology. Ekaterinburg,
G LODNY , J., A USTERHEM , H., M OLINA , J. F., R USIN , A. Urals Branch of Russian Academy of Sciences,
& S EWARD , D. 2003. Rb –Sr record of fluid-rock inter- 121– 128 (in Russian).
action in eclogites: the Marun-Keu complex, Polar I VANOV , K. S., K OROTEEV , V. A. ET AL . 2004. The struc-
Urals, Russia. Geochimica et Cosmochimica Acta, ture of a conjugation zone between Cis-Urals and West
67, 4353– 4371. Siberian oil-and-gas-bearing basin. Lithosphere, 2,
G LODNY , J., P EASE , V. L., M ONTERO , P., A USTERHEIM , 108– 124 (in Russian).
H. & R USIN , A. I. 2004. Protolith ages of eclogites, I VANOV , K. S., S HMELEV , V. R., R ONKIN , YU . L.,
Marun-Keu Complex, Polar Urals, Russia: impli- S AVELIEVA , G. N. & P UCHKOV , V. N. 2006. Zonal
cations for the pre- and early Uralian evolution of the gabbro-ultramafic complexes. In: M OROZOV , A. F.
northeastern European continental margin. In: G EE , (ed.) Structure and Dynamics of Lithosphere of the
D. G. & P EASE , V. L. (eds) The Neoproterozoic Tima- Eastern Europe, Issue 2. The Results of the EUROP-
nide Orogen of Eastern Baltica. Geological Society, ROBE Research. Moscow, GEOKART, GEOS,
London, Memoirs, 30, 87–105. 437– 445 (in Russian).
G OLOVANOVA , I. V. 2006. The Thermal Field of the J ONAS , P. 2004. Tectonostratigraphy of oceanic crustal
Southern Urals. M. Nauka (in Russian). terrains hosting serpentinite-associated massive sul-
G OROZHANINA , E. N. & P AZUKHIN , V. N. 2007. The phide deposits in the Main Uralian Fault zone (South
stages of activization of geodynamic processes in Urals). Freiberger Forschungshefte, C498, Tech-
the Late Devonian-Mid-Carboniferous time at the nische Universitat Bergakademie Freiberg.
western limb of the Zilair synclinorium and their K ASHUBIN , S., J UHLIN , C. ET AL . 2006. Crustal structure
dating by conodonts. Geologichesky Sbornik, 6, Ufa, of the Middle Urals based on reflection seismic data.
Institute of Geology, 55– 64 (in Russian). In: G EE , D. & S TEPHENSON , R. (eds) European Litho-
G RADSTEIN , F. M., O GG , J. G., S MITH , A. G., B LEEKER , sphere Dynamics. Geological Society, London,
W. & L OURENS , L. J. 2004. A new Geologic time scale Memoirs, 32, 427–442.
with special reference to Precambrian and Neogene. K AZANSKY , A. YU ., M ETELKIN , D. V., B RAGIN , V. YU .,
Episodes, 27, 83–100. M IKHALTSOV , N. V. & K UNGURTSEV , L. V. 2004.
G URSKAYA , L. I. & S MELOVA , L. V. 2003. Platinum- Paleomagnetic data on the Mesozoic complexes of a
metal ore formation and structure of Syum-Keu frame of Siberian platform as a reflection of
192 V. N. PUCHKOV
interpolated transcurrent deformations of the Central south Urals. Earth and Planetary Science Letters,
Asiatic belt. In: S KLYAROV , E. M. (ed.) The Geody- 226, 85– 99.
namic Evolution of Lithosphere of the Central Asiatic L ENNYKH , V. I., V ALISER , P. M., B EANE , R., L EECH , M.
Belt. From Ocean to Continent. 2, Irkutsk, Institute & E RNST , W. G. 1995. Petrotectonic evolution of the
of the Earth Crust, 151–155. Makysutov complex, southern Ural Mountains,
K HAIN , V. E. 2001. Tectonics of Continents and Oceans Russia: implications for ultrahigh-pressure meta-
(2000). Nauchnyi Mir, Moscow (in Russian). morphism. International Geology Review, 37,
K HAIN , E. V., F EDOTOVA , A. A., S ALNIKOVA , E. B., 584–600.
K OTOV , A. B. & Y AKOVLEVA , S. Z. 2004. New L EPEZIN , G. G., T RAVIN , A. V., Y UDIN , D. S.,
U– Pb data on the age of ophiolites of the Polar Urals V OLKOVA , N. V. & K ORSAKOV , A. V. 2006. The
and development of the margins of the Paleoasiatian age and thermal history of the Maksiutovo meta-
ocean in the Late Precambrian and Early Paleozoic. morphic complex (after Ar–Ar data). Petrology, 14,
In: K OROTEEV , V. A. (ed.) Geology and Metallogeny 1– 18.
of the Ultramafic-mafic and Granitoid Associations of L EVASHOVA , N. M., D EGTYAREV , K. E., B AZHENOV ,
Foldbelts. Ekaterinburg, Institute of Geology and Geo- M. L., C OLLINS , A. Q. & V AN DER V OO , R. 2003.
chemistry, 183–186 (in Russian). Middle Paleozoic paleomagnetism of east Kazakhstan:
K LIUZHINA , M. L. 1985. Paleogeography of the Urals in post-Middle Devonian rotations in a large-scale oro-
the Ordovician. M. Nauka (in Russian). cline in the central Ural-Mongol belt. Tectonophysics,
K NAPP , J. H., S TEER , D. N. ET AL . 1996. A 377, 249 –268.
lithosphere-scale image of the Southern Urals from L EVIN , V. YA ., R ONENSON , B. M. & S AMKOV , V. S.
explosion-source seismic reflection profiling in 1997. Alkaline-carbonatite Complexes of the Urals.
URSEIS ‘95. Science, 274, 226– 228. Ekaterinburg (in Russian).
K ORAGO , E. A., K OVALEVA , G. N. & T RUFANOV , G. N. L I , Z. X., B OGDANOVA , S. V. ET AL . 2008. Assembly,
1989. Formations, tectonics and geological history of configuration, and break-up history of Rodinia: a syn-
the Novozemelian Cimmerides. Geotectonics, 6, thesis. Precambrian Research, 160, 179–210.
40–61 (in Russian). L INNEMANN , U., G EHMLICH , V., T ICHOMIROWA , V. &
K OSAREV , A. M. & P UCHKOV , V. N. 1999. The distri- B USCHMANN , B. 1998. Introduction to the Pre-
bution of K, Ti and Zr in the Silurian-Carboniferous Symposium Excursion (part I): the Peri-Gondwanan
volcanic formations of the Southern Urals in connec- basement of the Saxothuringian Composite Terrane.
tion with the position of the Paleozoic subduction Excursion Guide to Saxony, Thüringia, Bohemia.
zone. Ezhegodnik-1977. Ufa, Institute of Geology, Abstracts. Schriften des Staatlichen Museums für
Ufimian Science Centre. 186–191 (in Russian). Mineralogie und Geologie zu Dresden, 7– 13.
K OSAREV , A. M., P UCHKOV , V. N. & S ERAVKIN , I. B. L ÜTZNER , H., F ALK , F. & E LLENBERG , J. 1979. Uber-
2005. Petrological-geochemical features of the Early sicht über die Variszische Molasseentwicklung in Mit-
Devonian-Eifelian island arc volcanites of the Magni- teleuropa und am Ural. Zeitschrift für Geologische
togorsk zone in a geodynamical context. Lithosphere, Wissenschaften, 9, 1157–1167.
4, 24– 40 (in Russian). M ASLOV , V. F., A RTIUSHKOVA , O. V., Y AKUPOV , R. R.
K OSAREV , A. M., P UCHKOV , V. N. & S ERAVKIN , I. B. & M AVRINSKAYA , T. M. 2008. The problems of the
2006. Petrologo-geochemical character of the Middle Lower–Middle Paleozoic stratigraphy of the Southern
Devonian– Early Carboniferous island-arc and colli- Urals. Geologichesky Sbornik, 7, 193– 204 (in
sional volcanic rocks of the Magnitogorsk zone in the Russian).
geodunamical context. Lithosphere, 1, 3– 21 (in M ATTE , P. 2006. The Southern Urals: deep subduction,
Russian). soft collision and weak erosion. In: G EE , D. &
K OZLOV , V. I. 1976. The coal-bearing deposits of the S TEPHENSON , R. (eds) European Lithosphere
Paleogene and Neogene of the Tirlyan syncline. In: Dynamics. Geological Society, London, Memoirs, 32,
Y AKHIMOVICH , V. L. (ed.) The Problems of 421–426.
Stratigraphy and Correlation of Pliocene and M ATTE , P., M ALUSKI , H., C ABY , R., N ICOLAS , A.,
Pleistocene Deposits of the Northern and Southern K EPEZHINSKAS , P. & S OBOLEV , S. 1993. Geody-
Parts of the Cis-Urals. Ufa, Bashkirian Branch of namic model and 39Ar/40Ar dating for the generation
the Academy of Sciences, USSR, 210– 227 (in and emplacement of the high-pressure (HP) meta-
Russian). morphic rocks in SW Urals. Academie de Sciences
K RASNOBAEV , A. A., R USIN , A. I., R USIN , I. A. & Comptes Rendus, 317, 1667– 1674.
B USHARINA , S. V. 2008. Zirconology of lherzolite- M ELCHER , F., G RUM , W., T HALHAMMER , T. V. &
garnet pyroxenite-dunite complex of the Uzyan T HALHAMMER , O. A. R. 1999. The giant chromite
Kraka (S.Urals). In: K ORTOTEEV , V. A. (ed.) The deposits at Kempirsai, Urals: constraints from trace
Structural-material Complexes and Geodynamic Pro- element (PGE, REE) and isotope data. Mineralium
blems of the Precambrian of Phanerozoic Orogens. Deposita, 34, 250– 272.
Ekaterinburg, Uralian Branch of Russian Academy of M ELEKESTSEVA , I. YU . 2007. Heterogenous Co-bearing
Sciences, 58–61 (in Russian). Massive Sulfide Deposits Associated with
K URENKOV , S. A., D IDENKO , A. N. & S IMONOV , V. A. Ultramafites of Paleo-island Arc Structures.
2002. The Geodynamics of Paleospreading. M. Nauka (in Russian).
M. GEOS. (in Russian). M ENSHIKOV , YU . P., K UZNETSOVA , N. V., S HEBU-
L EECH , M. L. & W ILLINGSHOFER , E. 2004. Thermal KHOVA , S. V. & N IKISHEVA , G. N. 1983. The tec-
modeling of the UHP Maksyutov Complex in the tonics of the northern half of the Magnitogorsk
depression after geophysical data. In: K HALEVIN , N. Ore Formation, Magmatism, Sedimentation and Meta-
(ed.) The Faults of the Earth’s Crust of the Urals and morphism. M. GEOS, 123– 125.
the Methods of Their Study. Sverdlovsk, Uralian Scien- P UCHKOV , V. N. 1979. Bathyal Complexes of the Passive
tific Centre, 65–78 (in Russian). Margins of Geosynclines. M. Nauka, Moscow (in
M IKHAILOV , V. O., T EVELEV , A. V., B ERZIN , R. G., Russian).
K ISELEVA , E. A., S MOLYANINOVA , E. I., S ULEIMA- P UCHKOV , V. N. 1996a. The Paleozoic geology of Asiatic
NOV , A. K. & T IMOSHKINA , E. P. 2002. Constraints Russia and adjacent territories. In: M OULLADE , M. &
on the Neogene-Quaternary Geodynamics of the N AIRN , A. (ed.) The Phanerozoic Geology of the
Southern Urals: comparative study of Neotectonic World I. The Paleozoic, B. Elsevier, 1–107.
data and results of Strength and Strain Modelling P UCHKOV , V. N. 1996b. The origin of the Ural-
Along the URSEIS Profile. In: B ROWN , D., J UHLIN , Novozemelian foldbelt as a result of an uneven,
C. & P UCHKOV , V. (eds) Mountain Building in the oblique collision of continents. Geotectonics, 5,
Uralides: Pangea to the Present. AGU Geophysical 66–75 (in Russian).
Monograph Series, 132, 273– 286. P UCHKOV , V. N. 1997. Structure and geodynamics of the
M ILANOVSKI , E. E. 1963. On the paleogeography of the Uralian orogen. In: B URG , J.-P. & F ORD , M. (eds)
Kaspian basin in the Middle and beginning of the Orogeny Through Time. Geological Society, London,
Upper Pliocene. Bulletin Moscow Society of Nature Special Publications, 121, 201–236.
Studies, Section of Geology, 38, 23– 26 (in Russian). P UCHKOV , V. N. 2000. Paleogeodynamics of the Middle
M ÜNTENER , O., P ICCARDO , G. B., P OLINO , R. & and Southern Urals. Ufa, Dauria (in Russian).
Z ANETTI , A. 2005. Revisiting the Lanzo peridotite P UCHKOV , V. N. 2002a. Paleozoic evolution of the East
(NW Italy): “astenospherization” of ancient mantle European continental margin involved into the Urals.
lithosphere. Ofioliti, 30, 111 –124. In: B ROWN , D., J UHLIN , C. & P UCHKOV , V. (eds)
M URATOV , M. V. 1979. The Uralo-Mongolian foldbelt. Mountain Building in the Uralides: Pangea to the
In: M URATOV , M. V. (ed.) Tectonics of the Uralo- Present. AGU Geophysical Monograph Series, 132,
Mongolian Foldbelt. M. Nauka, 5– 11 (in Russian). 9– 32.
N ARKISSOVA , V. V. 2005. Petrochemistry of the P UCHKOV , V. N. 2002b. Neotectonics of the Urals. In:
Late Ordovician-Early Devonian basaltoids of P UCHKOV , V. N. & D ANUKALOVA , G. A. (eds)
the Southern part of the Middle Urals (after data on Upper Pliocene and Pleistocene pf the Southern
the Uralian superdeep borehole and near-hole area). Urals Regions INQUA-SEQS’02, Ufa, 70–72.
Candidate Dissertation thesis. Moscow University (in P UCHKOV , V. N. 2003. Uralides and Timanides: their
Russian). structural relationship and position in the geological
N EDOSEKOVA , I. L., P RIBAVKIN , S. V., S EROV , P. A. & history of the Uralo-Mongolian foldbelt. Russian
R ONKIN , YU . L. 2006. The isotopic compositions and Geology and Geophysics, 44, 28–39.
age of carbonatites of the Ilmeno-Vishnevogorsk alka- P UCHKOV , V. N. 2005. Evolution of lithosphere: from the
line complex. In: G ALIMOV , E. M. (ed.) Isotopic Pechora ocean to Timanian orogen, from the Paleo-
Dating of Processes of Ore Formation, Magmatism, uralian ocean to Uralian orogen. In: L EONOV , YU . G.
Sedimentation and Metamorphism. M. GEOS, 40–54 (ed.) Problems of Tectonics of the Central Asia. M.,
(in Russian). GEOS, 309– 342 (in Russian).
P APULOV , G. N. 1974. The Cretaceous Deposits of the P UCHKOV , V. N. 2006a. The outline of the Uralian miner-
Urals. M. Nauka (in Russian). ageny. In: P YSTIN , M. A. (ed.) The Problems of
P EASE , V., D ALY , J. S. ET AL . 2008. Baltica in the Cryo- Geology and Mineralogy. Syktyvkar, Geoprint,
genian, 850–630 Ma. Precambrian Research, 160, 195– 222 (in Russian).
46– 65. P UCHKOV , V. N. 2006b. On the age of the Uralian ophio-
P ETROV , G. A. 2006. The Geology and Minerageny of the lites. In: K OROTEEV , V. A. (ed.) Ophiolites: Geology,
Main Uralian Fault Zone in the Middle Urals. Ekater- Petrology, Metallogeny and Geodynamics. Ekaterin-
inburg, Uralian Mining University (in Russian). burg, IG USC RAS, 121– 129 (in Russian).
P ETROV , G. A., R ONKIN , YU . L., L EPIKHINA , O. P. & P UCHKOV , V. N. 2008. The diachroneity of arc-continent
P OPOVA , O. YU . 2005. High-pressure metamorphism collision in the Uralides: subduction jump, oblique
in the Urals – two stages? Ezhegodnik-2004. IGG subduction or . . . what? In: National Cheng Kung
Uralian Branch of RAS, Ekaterinburg. 97– 102 (in University Tainan. Arc– continent collision. Taiwan,
Russian). 47–51.
P EYVE , A. V., I VANOV , S. N., N ECHEUKHIN , V. M., P UCHKOV , V. N., R APOPORT , M. B., F ERSHTATER , G. B.
P ERFILYEV , A. S. & P UCHKOV , V. N. 1977. Tectonics & A NANYEVA , E. M. 1986. Tectonic control of the
of the Urals. The explanatory notes for the Paleozoic granitoid magmatism in the eastern slope
1:1 000 000-scale tectonic map. M. Nauka, Moscow of the Urals. In: I VANOV , S. N. (ed.) Studies of
(in Russian). Geology and Metallogeny of the Urals. UNTs AN
P LIUSNIN , K. P. 1966. Wrench faults of the eastern slope SSSR, Sverdlovsk, 85–95 (in Russian).
of the Southern Urals. Geotectonics, 4, 57–68 (in P UCHKOV , V. N. & S VETLAKOVA , A. N. 1993. The struc-
Russian). ture of the Urals in the transect of the Troitsk DSS
P RIBAVKIN , S. V., R ONKIN , YU . L., T RAVIN , A. V. & profile. Transactions (Doklady) of the Russian
P ONOMARCHUK , V. A. 2006. The new data on the Academy of Sciences/Earth Science Section, 333,
age of lamproite magmatism of the Urals. In: 348– 351 (in Russian).
G ALIMOV , E. M. (ed.) III Russian Conference on P UCHKOV , V. N. & D ANUKALOVA , G. A. 2004. New data
Isotope Geochronology: Isotopic Dating of Processes on the character of tectonic deformations of the
194 V. N. PUCHKOV
Cretaceous-Paleogene peneplain in the Southern Urals. ophiolites. Journal of Geological Society, London,
In: P UCHKOV , V. N. (ed.) Geological Sbornik 4. Insti- 153, 525 –537.
tute of Geology, Ufimian Geoscience Center, Russian S AVELIEVA , G. N., P UCHKOV , V. N. & S PADEA , P.
Academy of Science, Ufa, 183– 184 (in Russian). 2006a. The ophiolites of the Urals. Structure and
P UCHKOV , V. N., R OSEN , O. M., Z HURAVLEV , D. Z. & dynamics of lithosphere of the Eastern Europe. In:
B IBIKOVA , E. V. 2006. Contamination of Silurian M OROZOV , A. F. (ed.) Structure and Dynamics of
volcanites of the Tagil synform by Precambrian Lithosphere of the Eastern Europe, Issue 2. The
zircons. Transactions (Doklady) of the Russian Results of the EUROPROBE Research. Moscow,
Academy of Sciences/Earth Science Section, 411, GEOKART, GEOS, 421–436 (in Russian).
1 –4 (in Russian). S AVELIEVA , G. N., S HISHKIN , M. A., L ARIONOV , A. N.,
P USHKAREV , E. V. 2001. Explosive breccias with inclu- S USLOV , P. V. & B EREZHNAYA , N. G. 2006b.
siona of mafic and ultramafic high-pressure rocks in Tectono-magmatic events of the Late Vendian in the
the Mindyak lherzolite massif – composition and mantle complexes of ophiolites of the Polar Urals:
petrogenetic consequences. In: P UCHKOV , V. N. data of U–Pb dating of zircons from chromites. In:
(ed.) Geology and Prospects of the Raw Materials of K OROTEEV , V. A. (ed.) Ophiolites: Geology, Petrol-
Bashkortostan and Adjacent Territories. Ufa, IG ogy, Metallogeny and Geodynamics. Ekaterinburg,
UNC RAS, 1, 155– 168 (in Russian). IG USC RAS, 160–164 (in Russian).
R ASULOV , A. T. 1982. The Tectonics of the Early S AMYGIN , S. G., K UZNETSOV , N. B., P AVLENKO , T. I. &
Mesozoic Depressions of the Eastern Slope of the D EGTYAREV , K. E. 1998. The structure of the
Urals. Sverdlovsk, Institute of Geology and Geochem- Kyshtym-Miass region of the Southern Urals and a
istry, Uralian Scientific Centre, USSR (in Russian). problem of conjugation of Magnitogorsk and Tagil
R EICHOV , M. K., P RINGLE , M. S. ET AL . 2007. New high- complexes. In: L EONOV , YU . G (ed.) The Urals funda-
precision 40Ar– 39Ar ages and geochemical data from mental Problems of Geodynamics and Stratigraphy.
the greater Siberian large igneous province: The M. Nauka, 73– 92 (in Russian).
Biggest gets Bigger. Abstract of AGU meeting, S AMYGIN , S. G. & R UZHENTSEV , V. S. 2003. The
San Francisco, 425. Uralian paleocean: a model of inheritad development.
R EMIZOV , D. N. 2004. The Island-arc System of the Polar Transactions (Doklady) of the Russian Academy of
Urals (Petrology and Evolution of Deep-seated Zones). Sciences/Earth Science Section, 392, 226–229 (in
Ekaterinburg, Uralian Branch of Russian Academy (in Russian).
Russian). S CARROW , J. H., P EASE , V., F LEUTELOT , C. & D USHIN ,
R ICKEN , W., S CHRADER , S., O NCKEN , O. & P LESCH , A. V. 2001. The Late Neoproterozoic Enganepe ophiolite,
2000. Turbidite basin and mass dynamics related to Polar Urals: an extension of the Cadomian arc? Pre-
orogenic wedge growth; the Rheno-Hercynian case. cambrian Research, 110, 255 –275.
In: F RANKE , W., H AAK , V., O NCKEN , O. & S CHWAB , M. 1984. The Harz Mountains. In: B ANKVITZ ,
T ANNER , D. (eds) Orogenic Processes: Quantification P. (ed.) Sedimentary and tectonic structures in
and Modelling in the Variscan Belt. Geological the Saxoturingian and Rhenohercynian Zones.
Society, London, Special Publications, 179, 257–280. Central Institute for Physics of the Earth, Potsdam,
R OZHDESTVENSKY , A. P. & Z INYAKHINA , I. K. 1997. 34–77.
The Development of the Relief of the Southern Urals S ENGÖR , A. M. C., N ATAL ’ IN , B. A. & B URTMAN , V. S.
in the Cenozoic, the Neogene Period. Ufa, IG USC 1993. Evolution of the Altaid tectonic collage and
RAS (in Russian). Palaeozoic crustal growth in Eurasia. Nature, 34
R YABOV , V. V. & G RIB , D. YE . 2005. Multiphase dykes – (6435), 299 – 307.
an example of dissipated spreading in the North of the S EWARD , D., P EREZ -E STAUN , A. & P UCHKOV , V. 1997.
Siberian platform. Geology and Geophysics, 46, Preliminary fission-track results from the southern
471– 485 (in Russian). Urals – Sterlitamak to Magnitogorsk. Tectonophysics,
R YBALKA , A. V., P ETROV , G. A., K ASHUBIN , S. N. & 276, 281 –290.
J UHLIN , C. 2006. The Middle Uralian ESRU transsect. S HARKOV , E. V., B ORTNIKOV , N. S. & B OGATIKOV ,
In: M OROZOV , A. F. (ed.) Structure and Dynamics of O. A. 2004. The Mesozoic zircon from gabbro-norites
Lithosphere of the Eastern Europe, Issue 2. The of the axial zone of the Mid-Atlantic Ridge, 68N (the
Results of the EUROPROBE Research. Moscow, Basin of Markov). Transactions (Doklady) of the
GEOKART, GEOS, 390– 401 (in Russian). Russian Academy of Sciences/Earth Science Section,
S AVELIEV , A. A. 1997. Ultramafic-gabbro formations 396, 675 –679 (in Russian).
in the structure of ophiolites of the Voikar-Synya S HATSKY , N. S. 1965. On the tectonics of Soviet Union
massif (Polar Urals). Geotectonics, 1, 48– 58 (On 10 years after demise of A. D. Archangelsky. In:
(in Russian). SHATSKY , V. S. (ed.) Selected publications. M.
S AVELIEV , A. A., B IBIKOVA , E. V. & S AVELIEVA , G. N. Nauka Publishing House, IV, 76– 84 (in Russian).
2001. Garnet pyroxenites of the Mindyak massif, the S HATSKY , V. S., S IMONOV , V. A. & Y AGOUTZ , E. 2000.
position and age of formation. Bulletin MOIP, New data on the age of eclogites of the Polar Urals.
Section of Geology, 1, 22– 29. Doklady of Russian Academy, 371, 519– 523 (in
S AVELIEVA , G. N. 1987. Gabbro-Ultramafic Complexes Russian).
of Ophiolites of the Urals and their Analogues in the S IGOV , A. P. 1969. The Mesozoic and Cenozoic Metallo-
Modern crust. M. Nauka (in Russian). geny of the Urals. M. Nedra. (in Russian).
S AVELIEVA , G. N. & N ESBITT , R. Q. 1996. A Synthesis of S KRIPIY , A. A. & Y UNUSOV , H. K. 1989. Tension and
the stratigraphic and tectonic setting of the Uralian compression structures in the Articulation zone of the
Southern Urals and the East European Platform. Geo- T RYGGVASON , A., B ROWN , D. & P EREZ -E STAUN , A.
tectonics, 23, 515– 522 (in Russian). 2001. Crustal architecture of the southern Uralides
S MIRNOV , G. A. 1957. Materials for Paleogeography of from true amplitude processing of the URSEIS vibro-
the Urals. The Visean Stage. Sverdlovsk, Mining and seis profile. Tectonics, 20, 1040– 1052.
Geology Institute (in Russian). T UZHIKOVA , V. I. 1973. The History of the Lower Meso-
S MIRNOV , V. N., B OROZDINA , G. N., D ESYATNI- zoic Coal Accumulation in the Urals. M. Nauka (in
CHENKO , L. I., I VANOV , K. S., M EDVEDEVA , Russian).
T. YU . & F ADEICHEVA , I. F. 2006. On the time of U DORATINA , O. V. & L ARIONOV , A. N. 2005. The age of
opening of the Uralian paleocean (biostratigraphic granitoids of Taikeu massif (Polar Urals). In:
and geochemical data). Russian Geology and Geophy- Y USHKIN , N. P. (ed.) The Structure, Geodynamics
sics, 47, 755–761 (in Russian). and Mineragenic Processes in Lithosphere. Syktyvkar,
S NACHEV , A. V., P UCHKOV , V. N., S AVELIEV , D. B. & Geoprint, 346–349 (in Russian).
S NACHEV , V. N. 2006. Geology of the Aramil- V ERBITSKAYA , N. P. 1964. The Regional Specific Fea-
Sukhtelinsk Zone of the Urals. IG USC RAS, Ufa (in tures of Neotectonics of the Urals. Leningrad (in
Russian). Russian).
S OBORNOV , K. O. & B USHUEV , F. S. 1992. The kin- V ERNIKOVSKY , V. A., V ERNIKOVSKAYA , A. A., P EASE ,
ematics of the conjugation zone of the Northern V. L. & G EE , D. G. 2004. Neoproterozoic Orogeny
Urals and the Upper Pechora Basin. Geotectonics, 1, along the margins of Siberia. In: G EE , D. G. &
39– 51 (in Russian). P EASE , V. (eds) The Neoproterozoic Timanide
S OKOLOV , V. B. 1992. The structure of the Earth’s crust Orogen of Eastern Baltica. Geological Society,
of the Urals. Geotectonics, 5, 3– 19 (in Russian). London, Memoirs, 30, 233– 248.
S PADEA , P., Z ANETTI , A. & V ANUCCI , R. 2003. Mineral W ILLNER , A. P., E RMOLAEVA , T. ET AL . 2002. Surface
chemistry of ultramafic massifs of the Southern Ura- Signals of an Arc-Continent Collision: The Detritus
lides orogenic belt (Russia) and the petrogenesis of of the Upper Devonian Zilair Formation in the
the Lower Palaeozoic ophiolites of the Uralian Southern Urals, Russia. In: B ROWN , D., J UHLIN , C.
Ocean. In: D ILEK , Y. & R OBINSON , P. T. (eds) Ophio- & P UCHKOV , V. (eds) Mountain Building in the Ura-
lites in Earth History. Geological Society, London, lides: Pangea to the Present. AGU Geophysical Mono-
Special Publications, 218, 567– 596. graph Series, 132, 183– 209.
S ULEIMANOV , A. K. 2006. The CDP works along W ILLNER , A. P., W ARTHO , J.-A, K RAMM , U. &
the URSEIS profile. In: M OROZOV , A. F. (ed.) Struc- P UCHKOV , V. N. 2004. Laser 40Ar – 39Ar ages of
ture and Dynamics of Lithosphere of the Eastern single detrital white mica grains related to the exhuma-
Europe, Issue 2. The Results of the EUROPROBE tion of Neoproterozoic and Late Devonian high
Research. Moscow, GEOKART, GEOS, 363– 373 pressure rocks in the Southern Urals (Russia). Geologi-
(in Russian). cal Magazine, 141, 161– 172.
T ESSALINA , B. S., B OURDON , G., P USHKAREV , E., Z AKHAROV , O. A. & P UCHKOV , V. N. 1994. On the Tec-
C APMAS , F., B IRCK , J.-L. & G ANNOUN , M. 2005. tonic Nature of the Maksutovo Complex of the Ural-
MinUrals – IPGP contribution. Complex Proterozoic Tau zone. Ufimian Science Centre, Ufa (in Russian).
to Paleozoic history of the upper mantle recorded in Z IEGLER , P. A. 1999. Evolution of the Arctic-North
the Urals lherzolite massifs by Re–Os and Sm–Nd Atlantic and the Western Tethys. AAPG Memoir, 43,
systematics. In: CERCAMS-6 Workshop Mineral 164– 196.
Deposits of the Urals. London, The Natural History Z NAMENSKY , S. E. 2007. The stages of Late Paleozoic
Museum, CD. deformations of the Magnitogorsk megazone
T EVELEV , A. L. V., D EGTYAREV , K. E. ET AL . (Southern Urals). In: Y USHKIN , N. P. & S AZONOV ,
2005. Geodynamic situations of formation of the V. N. (eds) Geodynamics of Formation of Mobile
Carboniferous volcanic complexes of the Southern Belts of the Earth. Ekaterinburg, Institute of Geology
Urals. In: N IKISHIN , A. (ed.) The Studies on the and Geochemistry, Uralian Branch of Russian
Regional Tectonics of the Urals, Kazahstan and Academy, 108 –111 (in Russian).
Tyan-Shan, 1. Southern Urals. Moscow, Nauka, Z ONENSHAIN , L. P., K UZMIN , M. I. & N ATAPOV , L. M.
213–247. 1990. Geology of the USSR, a plate tectonic synthesis.
T RETYAKOV , A. A., R YAZANTSEV , A. V., K UZNETSOV , American Geophysical Union, Geodynamic Series, 21,
N. B. & B ELOVA , A. A. 2008. The structural position 1– 242.
and geochronological dating of granate ultramafics Y AKHIMOVICH , V. L. & A NDRIANOVA , O. S. 1959. The
in the Southern Urals. In: K ARYAKIN , YU . V. (ed.) Southern Uralian Brown Coal Basin. Ufa,
The General and Regional Problems of Tectonics Mining-Geological Institute of the Bashkirian Branch
and Geodynamics, 2. Moscow, GEOS, 343– 349 (in of the Academy of Sciences, USSR.
Russian). Y AZEVA , R. G. & B OCHKAREV , V. V. 1993. Post-
T RIFONOV , V. G. 1999. The neotectonics of the Europe. collisional magmatism of the Northern Urals. Geotec-
Nauchny Mir (in Russian). tonics, 4, 56–65 (in Russian).
T RIFONOV , V. G. 2008. The age and mechanisms of the Y UDIN , V. V. 1994. Orogeny in the Northern Urals and
newest mountain building. In: K ARYAKIN , YU . V. Pai-Khoi. Nauka, Ekaterinburg (in Russian).
(ed.) The General and Regional Problems of Tectonics YUGGEO. 2002. Atlas of Paleogeographic, Structural,
and Geodynamics. 2. Moscow, GEOS, 349– 353 Palinspastic and Geoecological Maps of the Central
(in Russian). Eurasia. Almaty (in Russian).
Timing of dextral strike-slip processes and basement exhumation
in the Elbe Zone (Saxo-Thuringian Zone): the final pulse of
the Variscan Orogeny in the Bohemian Massif constrained
by LA-SF-ICP-MS U –Pb zircon data
M. HOFMANN1*, U. LINNEMANN1, A. GERDES2, B. ULLRICH3 & M. SCHAUER4
1
Staatliche Naturhistorische Sammlungen Dresden, Museum für Mineralogie und Geologie,
Königsbrücker Landstraße 159, D-01109 Dresden, Germany
2
Institut für Geowissenschaften, Johann Wolfgang Goethe-Universität, FE Mineralogie,
Altenhöferallee 1, 60438 Frankfurt am Main, Germany
3
TU Dresden, Institut für Geotechnik, Lehrstuhl für Angewandte Geologie, Neuffer-Bau,
Helmholtzstraße 10, D-01069 Dresden, Germany
4
Am Hexenberg 8, 09224 Grüna, Germany
*Corresponding author (e-mail: mandy.hofmann@senckenberg.de)
Abstract: The final pulse of the Variscan Orogeny in the northern Bohemian Massif (Saxo-
Thuringian Zone) is related to the closure of the Rheic Ocean, which resulted in subduction-
related D1-deformation followed by dextral strike-slip activity (D2-deformation, the Elbe Zone).
Taken together, these deformation events reflect the amalgamation of Pangaea in central Europe.
Lateral extrusion of high-grade metamorphosed rocks from an allochthonous domain (Saxonian
Granulitgebirge) and the top –NW-directed transport of these domains (Erzgebirge nappe
complex, Saxonian Granulitgebirge) are responsible for these dextral strike-slip movements. Geo-
chronological data presented herein, together with published data, allow the timing of the final
pulse of the Variscan Orogeny and related plutonic, volcano-sedimentary and tectonic processes.
Marine sedimentation lasted at least until the Tournaisian (357 Ma). Onset of Variscan strike-slip
along the Elbe Zone is assumed to be coeval with the beginning of the top–NW-directed lateral
extrusion of the Saxonian Granulitgebirge at 342 Ma (D2-deformation). The sigmoidal shape of
the Meissen Massif indicates that strike-slip activity was coexistent with intrusion of the pluton
at c. 334 Ma into the schist belt of the Elbe Zone. In contrast, the intrusion of the Markersbach
Granite provides a minimum age of c. 327 Ma for the termination of D2 strike-slip activity, because
this undeformed pluton cross-cuts all strike-slip related tectonic structures. Geochronological data
of an ash bed from the Permo-Carboniferous Döhlen Basin show clearly that post-orogenic sedi-
mentation of Variscan molasse in that area was already active at 305 Ma. This pull-apart basin
is a local example of regional Permo-Carboniferous extension within Pangaea. The uplift and denu-
dation of the Variscan basement in the Saxo-Thuringian Zone occurred between c. 327–305 Ma.
Introduction suture, see Fig. 1) (Scotese & Barret 1990).

Despite decades of study, a number of important
The Bohemian Massif is the type area of the first-order questions remain unanswered, including
Variscan Orogeny. The name is derived from an the precise age of Variscan orogenic processes and
area called by the Romans ‘Curia Variscorum’, their relationship to the closure of the Rheic
which is situated adjacent to the city of Hof in the Ocean. Palaeocontinental reconstructions suggest
Frankonian Forest (Germany) (Linnemann 2003a, that convergence of the Rheic Ocean started after
and references therein). In the Anglo –American lit- the formation of the supercontinent Laurussia at
erature, the Variscan Orogeny is also known as the c. 420– 400 Ma (Romer et al. 2003; Sánchez
‘Hercynian Orogeny’. Variscan orogenic events Martı́nez et al. 2007). Northward drift of Gondwana
are geodynamically linked to Alleghenian processes started around the same time and subduction of
in the Appalachians (Linnemann et al. 2007b). Rheic oceanic lithosphere terminated at around
Alleghenian –Variscan collisional orogenic activity c. 400– 370 Ma (Romer et al. 2003).
reflects the closure of the Rheic Ocean and the amal- In the Bohemian Massif, continental collision
gamation of the supercontinent Pangaea (Rheic involved subduction of continental crust, which
198 M. HOFMANN ET AL.
Fig. 1. Cadomian and Variscan Massifs in southwestern and central Europe with Variscan oceanic suture of
the Rheic Ocean (after Robardet 2002; Linnemann et al. 2007a, 2008a). IM, Iberian Massif; AM, Armorican
Massif; FMC, French Massif Central; RM, Rhenish Massif; BRM, Brabant Massif; BM, Bohemian Massif; SPZ,
South Portuguese Zone; OMZ, Ossa-Morena Zone; CIZ, Central Iberian Zone; GTOM, Galicia-Trás os Montes
Zone; WALZ, West Asturian Leonese Zone; CZ, Cantabrian Zone; PL, Pulo de Lobo oceanic units; IC, Iberian
Chains; BCSZ, Badajoz-Cordoba Shear Zone; P-Pyrénées; MM, Maures Massif; SXZ, Saxo-Thuringian Zone;
TBU, Teplá-Barrandian Unit; MZ, Moldanubian Zone; S, Sudetes; M, Moravo-Silesian Zone; Li, Lizard Ophiolite;
SL, Śle˛ża ophiolite. Black-oceanic rocks of the Pulo de Lobo suture and ophiolitic units of allochthonous complexes
of Galicia.
started at c. 370 –360 Ma (Kroner et al. 2007), as herein a number of new U –Pb age determinations
the northern margin of Gondwana underthrusted of igneous rocks and one ash bed from a major
the southern margin of Laurussia. Complex sub- late Variscan shear zone, known as the Elbe Zone.
duction processes and stacking of Gondwanan This belt of sheared rocks occurs within the
continental crust occurred between c. 360 and Saxo-Thuringian Zone in the northern part of the
340 Ma. The final stage of the Variscan Orogeny Bohemian Massif (Linnemann 1994, 2003a;
is characterized by orogen-wide transpressional Kroner et al. 2007). As field relationships indicate
tectonics, regional HT/LP metamorphism, late that some plutonic bodies are coeval with shearing,
orogenic granite intrusions, and the formation of whereas other bodies post-date shearing, our data
fold-and-thrust belts in the external parts of the help constrain the age of D2 strike-slip deformation
orogen (Kroner et al. 2007). The related final extru- and therefore constrain the age of Late Variscan
sion of deeper parts of subducted units exhumed tectonic activity in central Europe.
rocks of peri-Gondwanan affinity. Their juxtaposi-
tion with low-grade assemblages in the upper crust Geological setting
is today an important feature of the Variscan
orogen (Kroner et al. 2007) that is due to transport The Saxo-Thuringian Zone and the
directions of allochthonous nappe piles that in part history of the Rheic Ocean
had a large influence on the samples presented in
this study. A detailed description of the geology of the
In order to better characterize the timing of the Bohemian Massif, and the Saxo-Thuringian Zone
final pulse of the Variscan Orogeny, we present in particular, may be found in Franke (2000),
TIMING OF DEXTRAL STRIKE-SLIP PROCESSES AND BASEMENT EXHUMATION 199
Franke & Żelazniewics (2002), Linnemann (2003a, and magmatic quiescence and shelf sedimentation.
b), Kroner et al. (2007), Linnemann et al. (2008b) Linnemann et al. (2007a) interpret this to be a
and Nance et al. (2008). The Elbe Zone occurs in passive margin along the northern Gondwanan
the Saxo-Thuringian Zone, which is located in the margin that is flanked to the north by the Rheic
Bohemian Massif (Kossmat 1927) (Fig. 1). The Ocean. On the opposing flank of the Rheic Ocean
oldest rocks of the Saxo-Thuringian Zone are c. Avalonia collided with Baltica during the closure
570 –565 Ma sedimentary rocks (Fig. 2), which pre- of the Tornquist Sea at c. 450 Ma, which was fol-
dominantly consist of turbiditic greywackes and lowed, at c. 420 Ma, by the collision of Baltica þ
shales, and are inferred to have been deposited in Avalonia with Laurentia, which closed the Iapetus
back-arc and retro-arc basins along the northern Ocean and resulted in the formation of Laurussia
margin of Gondwana during the Cadomian (Fig. 3) (Scotese & Barret 1990; Torsvik &
Orogeny (Linnemann et al. 2007a). Rehnström 2002).
These sedimentary basins were deformed by In the Upper Devonian, closure of the Rheic
arc-continent collision and were intruded by volu- Ocean in central Europe was accompanied by sub-
minous granitoids at c. 540 Ma, which are coeval duction of the Gondwanan margin beneath Lau-
with a switch from an arc to a transform margin russia (Kroner et al. 2007) (Fig. 3). In the Saxo-
setting, analogous with the modern Eastern Pacific Thuringian Zone, this collision was accompanied
(Linnemann et al. 2007a). The granitoid intrusions by deposition of clastic strata until the Lower Car-
were due to slab break-off of the subducted boniferous (Linnemann 2003b; Kroner et al. 2007).
oceanic plate during the change of the marginal The Mid-German Crystalline Zone is interpreted
setting (Linnemann et al. 2007a). The tectonic as the suture of the Rheic Ocean (Fig. 1). Martı́nez
setting may have been similar to that of the present- et al. (2007) suggest the closure of the Rheic
day Basin and Range Province adjacent to Baja Ocean occurred in a northward-dipping supra-
California and the San Andreas Fault (Nance et al. subduction zone setting because of the general
2002). Cadomian orogenic processes provided the absence of large Silurian-Devonian volcanic arcs
structural heterogeneities which played an impor- on both margins of the ocean.
tant role for the opening of the Rheic Ocean In the Bohemian Massif and also in other parts of
during the Late Cambrian (Linnemann et al. Central and Western Europe, rocks of the Variscan
2007a, 2008c). orogen include voluminous Late Devonian –Early
Lower to Middle Cambrian strata in the Carboniferous high-pressure metamorphic units
Saxo-Thuringian Zone (Fig. 2) (Elicki 1997) are (allochthonous domains) tectonically juxtaposed
characterized by carbonates with archaeocyatha, with low-grade Neoproterozoic (Cadomian) and
siliciclastic strata and red beds. Lower Ordovician Palaeozoic successions (autochthonous domains).
deposits in the Saxothuringian Zone are for The allochthonous domains are surrounded by
example quartzites (Frauenbach and Phycodes wrench-and-thrust zones and are interpreted to
groups) that are up to 3000 m thick (e.g. SE-part reflect diachronous subduction of thinned peri-
of the Schwarzburg unit, Fig. 2) and are typical Gondwanan continental crust down to the stabi-
of the Lower Ordovician strata in central and lity field of metamorphic diamond (c. 120 km;
western Europe. In some localities, these deposits Massonne 1998; Kroner et al. 2007, and references
overstep Lower –Middle Cambrian strata to overlie therein). Top-SE exhumation of high pressure
directly the Neoproterozoic (Cadomian) basement (HP) metamorphic units occurred in a subduction
(Linnemann & Romer 2002). Similar relationships channel (D1 transport) and is preserved in the
are documented in coeval deposits in the Armorican Saxonian Granulitgebirge complex and in parts of
Massif (NW France) and in different parts of the nappe pile of the Erzgebirge complex, and
Iberia (e.g. Graindor 1957; Chauvel & Rabu the nappe complexes of Münchberg, Wildenfels
1988; Chantraine et al. 1994; Fernández-Suárez and Frankenberg (Fig. 4) (Kroner et al. 2007).
et al. 2000; Linnemann et al. 2007a). These de- Finally, regional dextral transpression, rapid
posits, which are accompanied in some localities exhumation and the formation and inversion
by c. 485 Ma rift-related bimodal magmatism (folding) of Variscan flysch basins occurred during
(Sánchez-Garcı́a et al. 2003), are thought to SE–NW oriented D2-transport at c. 340–330 Ma
represent the rift-to-drift transition that heralded (Kroner et al. 2007) (Fig. 4). Related extensive
the development of the Rheic Ocean and the sep- Variscan plutonism occurred in the Bohemian
aration of Avalonia or a related micro-continent or Massif between c. 334– 327 Ma (this paper). The
terrane from the Gondwanan margin (Linnemann Mid-German Crystalline Zone, the suture of the
et al. 2007a). Rheic Ocean, was formed by large-scale oblique
After c. 480 Ma, the Saxo-Thuringian Zone is subduction, collision and exhumation tectonics,
characterized by an extended period (Mid-Upper and strike-slip related slivering (Zeh et al. 2005;
Ordovician–Mid-Devonian) of relative tectonic Kroner et al. 2007).
Fig. 2. Lithological column and magmatic events of the Cadomian Basement and the overlying Palaeozoic strata of the
Thuringian Facies in the Saxo-Thuringian Zone (after Linnemann et al. 2004). 1, immature litharenitic sandstones
(greywackes) and shales; 2, mature sandstones and shales; 3, dark grey and black shales; 4, carbonates, 5, glaciomarine
diamiktite of the Gondwana glaciation in the Late Ordovician (Lederschiefer); 6, level with dated granitoid pebbles; 7,
felsic ash bed; 8, granitoid plutons; 9, rhyolites and porphyroids; 10, mafic volcanic rocks and intrusions. *Traditional
German terms for lithostratigraphic units. Sources of geochronological data: **Pb– Pb TIMS (Linnemann et al. 2000);
***U–Pb SHRIMP (Buschmann et al. 2001).
SW. The geodynamic relationships and the appro-

priate model for the Saxo-Thuringian Zone during
the Variscan Orogeny have been published recently
by Kroner et al. (2007), including also the divi-
sion of Saxo-Thuringian units into autochthonous
and allochthonous domains. The Lausitz Block
belongs to the autochthonous domain of the Saxo-
Thuringian Zone (Fig. 4) and was relatively stable
during Variscan deformation processes (Fig. 4).
The Erzgebirge nappe complex in the SW of the
Elbe Zone is part of the allochthonous domain of
the Saxo-Thuringian Zone and consists of a nappe
pile containing diverse metamorphic rocks rang-
ing from HP/HT to LP/LT units. The Elbe Zone
is a schist belt and represents a part of the wrench-
and-thrust zone and is situated between the
Lausitz Block and the Erzgebirge nappe complex
(Fig. 4) (Kroner et al. 2007).
According to the model of Kroner et al. (2007),
the Variscan Orogeny in the Saxo-Thuringian
Zone is characterized by two regionally-distributed
deformation and tectonic transport events (D1, D2,
Fig. 4). The D1-deformation is restricted to the
allochthonous domain and is related to a top-SW-
directed tectonic transport, which reflects the
exhumation of subducted continental crust along a
subduction channel during the final stage of the
Variscan Orogeny. In the central Bohemian
Massif, rock units that escaped from of the subduc-
tion channel formed a thick pile of metamorphic
rocks. After D1, NW-directed transport of these
over-thickened allochthonous domains occurred
during a regional D2 event. D2 is characterized by
top-NW vergent folding in the southeastern part
of the Saxo-Thuringian Zone and a SE-directed
back-thrusting in its northwestern part. In contrast
to D1, D2 occurs in the allochthonous and the
autochthonous domains as well as in the wrench-
and-thrust zone (Fig. 4). During D2-transport
of the allochthonous nappes of the Erzgebirge
complex towards the NW, dextral strike-slip
between the Erzgebirge nappe complex and the
Fig. 3. Palaeogeography in (a) Lower Silurian
(c. 440 Ma); (b) Lower Devonian (c. 400 Ma); and
autochthonous Lausitz Block occurred, which
(c) Upper Carboniferous (c. 300 Ma) (modified after resulted in the formation of the Elbe Zone. Final
C. R. Scotese: Palaeomap web site: www.scotese.com). exhumation of the deeper parts of the allochthonous
A, Armorica (Brittany, Normandy, Massif Central); B, domain is related to the juxtaposition of the Saxo-
Barrandian; C, Carolina; EA, East Avalonia; EWI, nian Granulitgebirge beneath the wrench-and-thrust
England, Wales, Southern Ireland; F, Florida; I, Iberia; zone (e.g. schist belt of the Elbe Zone) and dextral
IR, Iran; M, Mexican terranes; NF, New Foundland; NS, shear between the Erzgebirge nappe complex
Nova Scotia; PA, Basement of the Alps; RH, Rheno- and the autochthonous domain of the Lausitz
Hercynian; SX, Saxo-Thuringia; TP, ‘Turkish plate’; Block (Linnemann 1994; Kroner et al. 2007).
WA, West Avalonia.
There is broad agreement that the D2-deformation
led to the final crustal architecture of the Saxo-
The Elbe Zone and adjoining areas Thuringian Zone (Fig. 4) and that the schist belt of
the Elbe Zone (Fig. 5) was deformed and over-
The Elbe Zone near Dresden (Figs 4 & 5) is a NW – printed during the dextral strike-slip movements
SE striking schist belt that divides the Lausitz Block (Linnemann 1994; Mattern 1996; Kroner et al.
in the NE from the Erzgebirge nappe complex in the 2007). Estimated distances of strike-slip movement
Fig. 4. Tectonic and geological map of the Saxo-Thuringian Zone in the NE-part of the Bohemian Massif showing units
of Lower Carboniferous and older ages, and the dextral strike-slip movement along the Elbe Zone (modified from
Linnemann & Schauer 1999; Linnemann & Romer 2002; Linnemann et al. 2007a). Subdivision of the Saxo-Thuringian
Zone into autochthonous and allothonous domains and into a wrench-and-thrust Zone is based on Kroner et al. (2007).
Earlier regional Variscan D1 and later regional D2 transport directions are taken from Kroner et al. (2007). Note the
anchorage of the autochthonous domain relative to the allochthonous domain and the wrench-and-thrust zone and the
resulting dextral strike-slip movements along the Elbe Zone during the regional D2 deformation.
Abbreviations and Numbers: Hain, Hainichen; Frkbg, Frankenberg; 1 to 3, Lausitz granitoids complex; 1, Lausitz
anatexite; 2, western Lausitz granitoids; 3, eastern Lausitz granitoids; 4, granitoids of the Leipzig area; 5, granitoids of
the Elbe Zone (Dohna & Laas granodiorites); 6, shear zone-related orthogneisses (Grossenhain gneiss); 7, ‘red’ ortho-
gneisses, anatexites, migmatities (MP– MT unit); 8, gneisses and eclogites with major shear zones (HP– HT unit); 9,
phyllites, garnet phyllites and mica schists (HP– LT unit, MP–LT unit and LP– LT unit); 10, low to high grade ortho-
and para-rocks; 11, Ordovician, Silurian and Devonian volcano-sedimentary rock complexes; 12, Cambro– Ordovician
to Lower Carboniferous volcano-sedimentary rock complexes; 13, Rumburk granite (Lausitz antiform); 14, Granulite
and high grade country rocks of the granulite core; 15, southern phyllite zone (Cambro– Ordovician rock complex); 16,
greywackes, pelites, cherts, volcanic rocks (Altenfeld Fm., Frohnberg Fm., Lausitz Gr., Leipzig Fm.); 17, greywackes,
pelites, cherts, basalts, andesites (Rothstein Fm.); 18, greywackes, pelites, quartzites, basalts (passive margin sequences
and tillites of the Clanzschwitz, Rödern and Weesenstein Gr.); 19, carbonates, sandstones, pelites (Lower to Middle
Cambrian); 20, quartzites and pelites (Skolithos facies; Collmberg Fm., Hainichen–Otterwisch Fm., Dubrau Fm.;
Lower Ordovician, Tremadoc); 21, quartzites, shales, sed. iron ore (Ordovician); 22, Cadomian and Ordovician rocks
affected by the dextral Variscan Blumenau shear zone (Schwarzburg antiform); 23, Lower Graptolite Shale (‘Unterer
Graptolithenschiefer’) and ‘Ockerkalk’ (Silurian); 24, carbonates, sandstones, pelites, diabases (Devonian); 25,
greywackes and pelites (Variscan flysch; Tournai,Visean); 26, Variscan wildflysch with large olistolithes; 27, Variscan
early molasses of Hainichen-Borna (Upper Visean); 28, Variscan early molasses of Doberlug (Upper Visean); 29,
olistolithes of Cambrian to Devonian rock complexes within a wildflysch matrix; 30, acid to basic metamorphic rocks of
the nappe pile remnants of Münchberg and of the Saxon ‘Zwischengebirge’ of Wildenfels and Frankenberg; 31, Permo-
Carboniferous granitoids; 32, Upper Carboniferous rhyolithes; 33, Upper Carboniferous major granitoids dykes.
are in the range of 80–150 km (see Fig. 4, distance time (Fig. 5). Against the Erzgebirge nappe
of D2-transport). The schist belt of the Elbe Zone is complex, the schist belt of the Elbe Zone is bordered
separated in two parts by the overlying Döhlen by the Mid-Saxon Fault (Fig. 5). A major com-
Basin, which is one of the several late-Variscan ponent of the regional dextral shear occurred
molasse basins formed in Permo-Carboniferous along that fault (Pietzsch 1962; Linnemann 1994).
Fig. 5. Geological map of the Elbe Zone based on Pietzsch (1962). Geological cross-section A –B from the NE-part
of the Döhlen Basin and the Meissen Massif modified after Reichel & Schauer (2007). WF, Westlausitz Fault;
LT, Lausitz Thrust; MF, Mid-Saxon Fault; NWSG, Nossen-Wilsdruff Schiefergebirge (Nossen-Wilsdruff schistbelt);
ETSG, Elbtalschiefergebirge (Elbtal schistbelt); TVC, Tharandt Volcanic Complex (Upper Carboniferous). In the map
sample locations are indicated: Elbe 1, sample of an ash bed within the Döhlen Formation in the Döhlen Basin (Elbe
Zone); Klotz 2, sample of a granodiorite from the Lausitz Granitoid Complex (Lausitz Block); MH 5, sample of the
monzonite of the Meissen Massif (Elbe Zone); MH 10, sample of a gneiss from the eastern rim of the Meissen Massif
(Elbe Zone); Mark 1, sample of a granite near the village Markersbach (Elbe Zone).
Towards the Lausitz Block in the NE, the schist belt sigmoidal shape of the Meissen Massif (Fig. 4,
of the Elbe Zone is bordered by two major faults, the inset upper right) and deformation structures along
Westlausitz Fault and the Lausitz Thrust (Fig. 5). the edges of the plutonic complex suggest an intru-
The schist belt in the Elbe Zone contains parts sion into an area of tension in the active dextral
of the Neoproterozoic (Cadomian) basement (Linne- strike-slip regime along the Elbe Zone (Mattern
mann et al. 2007a) and Lower Ordovician to Lower 1996; Wenzel et al. 1997; Kroner et al. 2007).
Carboniferous sedimentary rocks and volcano– The Markersbach Granite intruded into the
sedimentary complexes. The age of the youngest southern part of the schist belt (Fig. 5). This alkali
meta-sediments of the schist belt is poorly con- feldspar granite clearly intruded after dextral shear
strained by badly preserved fossils, but is thought along the Elbe Zone was finished because this
to be Lower Carboniferous (Kurze et al. 1992). pluton cross-cuts all the Variscan structures, includ-
The schist belt of the Elbe Zone was intruded ing faults and cleavage. According to Bonin (2007)
by the granitoids of the Meissen Massif (Fig. 5). this rock with its euhedral dark quartz and pink
This complex consists of plutonic, dioritic to alkali feldspar, mostly mesoperthitic, and late anhe-
mainly monzonitic and granitic rocks, dominated dral mica is a typical A-type granite. Such granites
by hornblende-monzonite. The granitoids show a are commonly emplaced in within-plate, continental
signature typical for I-type granites. The diorite or oceanic settings, or under transtensional regimes
to monzonite intrusions show enriched mantle sig- in postorogenic contexts (Bonin 2007).
natures typical for shoshonitic rocks (Wenzel et al. Adjacent to the southern margin of the Elbe Zone
1997). Nasdala et al. (1999) determined ages of the Tharandt Volcanic Complex (TVC) and the
326 + 6 Ma and 330 + 5 Ma (SHRIMP U –Pb) granite of Niederbobritzsch intruded (Fig. 5). As
for the time of the monzonite intrusion. The there are components of the TVC and the granite
within the earliest c. 330 Ma old Variscan molasse to expose their centre. Prior to the U –Pb analyses,
of the Saxo-Thuringian Zone (Pietzsch 1962; the internal structures of the zircon grains were
Gehmlich 2003), both complexes must have been investigated by cathodoluminescence (CL). Ura-
at the surface before 330 Ma. In addition, neither nium, thorium and lead isotopes from the zircon
the TVC nor the Niederbobritzsch granite show grains were analyzed by laser ablation (LA) using
shear-related structures, what we interpret as both a Thermo-Scientific Element 2 SF-ICP-MS (sector
complexes intruded or extruded, respectively, after field-inductively coupled plasma-mass spectro-
dextral strike-slip movements along the Elbe Zone meter) coupled to a New Wave Research UP-213
finished. It can be inferred that the TVC and the ultraviolet laser system at Goethe University
Niederbobritzsch granite may belong to the same Frankfurt (GFU) (Gerdes & Zeh 2006, 2008). Data
magmatic event as the Markersbach granite, were acquired in time resolved – peak jumping –
marking the age of a final pulse at the very end of pulse counting mode over 810 mass scans, with a
the Variscan orogenic processes. 16 s. background measurement followed by 28 s
The Döhlen Basin post-dates all Variscan sample ablation. Laser spot-sizes varied from 20
structures of the Elbe Zone and adjoining areas. to 30 mm with a typical penetration depth of
It originated as a pull-apart basin, 22 6 km in c. 15 –20 mm. The signal was tuned for maximum
dimension, which opened during post-Variscan sensitivity for Pb and U while keeping oxide pro-
(Late Carboniferous) strike-slip motion along the duction, monitored as 254UO/238U, well below
Elbe Zone (Reichel & Schauer 2007). It is situated 1%. A teardrop-shaped, low volume (,2.5 cm3)
to the SW of Dresden (Fig. 5) and consists of laser cell was used (Frei & Gerdes 2009 and
molasse of the Variscan Orogen with a maximum references therein). This cell enables detection
thickness of c. 700 m (Reichel & Schauer 2007). and sequential sampling of heterogeneous grains
The volcano-sedimentary rocks within the basin (e.g. growth zones) during time resolved data acqui-
overlie unconformably the schist belt of the Elbe sition, due to its response time of ,1 s (time until
Zone, the Meissen Massif and the Erzgebirge maximum signal strength was achieved) and
nappe complex (Figs 4 & 5). Palaeontological data wash-out (,99% of previous signal) time of ,5 s.
suggest a deposition during latest Carboniferous With a depth penetration of c. 0.6 mms21 and a
and Lower Permian times (Roscher & Schneider 0.9 s integration time (¼ 15 mass scans ¼ 1 ratio),
2005). Similar molasse deposits are known from any significant variation of the Pb/Pb and U/Pb in
the Lausitz Block (Weissig Basin). Thus, the base- the mm scale is detectable. Raw data were corrected
ment rocks of the Lausitz Block, the plutonic offline for background signal, common Pb, laser-
rocks of the Meissen Massif, the schist belt of the induced elemental fractionation, instrumental mass
Elbe Zone and the metamorphic rocks of the Erzge- discrimination, and time-dependent elemental frac-
birge nappe complex were completely exhumed in tionation of Pb/U using an in-house MS Excel#
the Upper Carboniferous. spreadsheet program (Gerdes & Zeh 2006). A
common-Pb correction based on the interference-
Samples and methods and background-corrected 204Pb signal and a
model Pb composition (Stacey & Kramers 1975)
All sample locations are situated in the Elbe Zone and was carried out where necessary. The necessity of
adjoining areas (Saxo-Thuringian Zone). Locations the correction was judged on whether the corrected
207
are given in Table 1 and in Figure 5. We collected Pb/206Pb lay outside the internal errors of
one sample of a granodiorite from the Lausitz the measured ratios. The interference of 204Hg
Block (sample Klotz 2) close to the Westlausitz (mean ¼ 110 + 14 cps; counts per second) on
Fault. From the Elbe Zone three samples were mass 204 was estimated using a 204Hg/202Hg ratio
taken. Sample MH 10 is an orthogneiss which of 0.2299 and measured 202Hg. Laser-induced
occurs along the eastern rim of the Meissen Massif. elemental fractionation and instrumental mass
A monzonite (sample MH 05) was taken from the discrimination were corrected by normalization to
Meissen Massif itself, which shape indicates an the reference zircon GJ-1 for each analytical
intrusion into an active dextral shear setting. session. Prior to this normalization, the drift in inter-
Sample Mark 1 is from the Markersbach Granite, elemental fractionation (Pb/U) during 28 s of
which intruded the schist belt of the Elbe Zone and sample ablation was corrected for the individual
crosscuts all tectonic features. Sample Elbe 1 is a analysis. The correction was done by applying a
felsic ash layer from coal seam 5 of the Döhlen linear regression through all measured ratios,
Formation in the Döhlen Basin (Fig. 6). excluding the outliers (+2 standard deviation;
From each sample 2–5 kg of rock material was 2SD), and using the intercept with the y-axis as
collected. Zircons were separated using heavy the initial ratio. The total offset of the measured
liquid, magnetic separation and hand-picking. drift-corrected 206Pb/238U ratio from the ‘true’ ID-
The grains were subsequently mounted and polished TIMS value (0.0983 + 0.0004; ID-TIMS GFU
Table 1. Laser Ablation-SF-ICP-MS U, Pb and Th data of zircon grains from the Cadomian and Variscan basement and from one Upper Carboniferous– Lower Permian
ash bed of the Elbe Zone and Lausitz Block, Saxo-Thuringian Zone, Bohemian Massif (co-ordinates: UTM Zone 33)a
207
Name Pba Ub Pbb Thb/U Isotopic ratiosc Rhod Ages Conc
(spot) (cps) (ppm) (ppm) (%)
206 204 206 238 207 235 207 206 206 238 207 235 207 206
Pb/ Pb Pb/ U 2s Pb/ U 2s Pb/ Pb 2s Pb/ U +2s Pb/ U +2s Pb/ Pb +2s
(%) (%) (%) (Ma) (Ma) (Ma)
TIMING OF DEXTRAL STRIKE-SLIP PROCESSES AND BASEMENT EXHUMATION

Klotz 2 (biotite granodiorite, Lower Cambrian, Lausitz Block, Klotzsche in the north of the city of Dresden, Easting: 54 15448, Northing: 56 66651)
Klotz 2–1 7084 180 16 0.22 1890 0.08797 1.8 0.6841 5.6 0.05640 5.3 0.31 544 10 529 30 468 118 116
Klotz 2–2 9057 213 22 0.51 2715 0.09153 1.5 0.7334 3.4 0.05812 3.1 0.43 565 8 559 19 534 68 106
Klotz 2–3 8293 208 18 0.18 3054 0.08790 1.7 0.7163 2.3 0.05910 1.5 0.73 543 9 548 12 571 33 95
Klotz 2–4 9198 206 21 0.50 1023 0.08770 1.6 0.6799 4.2 0.05623 3.9 0.39 542 9 527 22 461 85 117
Klotz 2–6 6745 171 16 0.25 3608 0.08904 1.5 0.7069 2.4 0.05758 1.9 0.61 550 8 543 13 514 43 107
Klotz 2–7 3891 92 8 0.31 1729 0.08679 2.0 0.6621 10 0.05533 9.8 0.20 537 10 516 51 425 218 126
Klotz 2–8 10124 271 23 0.09 57675 0.08846 1.4 0.7237 2.0 0.05934 1.5 0.68 546 7 553 11 580 32 94
Klotz 2–9 8034 207 17 0.09 21358 0.08855 1.5 0.7167 2.3 0.05870 1.8 0.64 547 8 549 13 556 39 98
Klotz 2–10 1906 46 5 0.46 1315 0.09244 1.8 0.7558 3.2 0.05929 2.6 0.56 570 10 572 18 578 57 99
Klotz 2–11 7868 199 17 0.12 4016 0.08799 1.6 0.7066 2.4 0.05824 1.8 0.65 544 9 543 13 539 40 101
Klotz 2–12 7173 182 15 0.13 2615 0.08650 1.9 0.6816 2.7 0.05715 2.0 0.69 535 10 528 14 497 44 108
Klotz 2–13 3074 77 7 0.28 4466 0.09153 2.1 0.7455 3.3 0.05907 2.6 0.61 565 12 566 19 570 57 99
Klotz 2–14 5322 54 6 6.90 177 0.08872 2.3 0.6918 8.2 0.05655 7.9 0.28 548 13 534 44 474 175 116
Klotz 2–15 2929 73 7 0.40 1871 0.08780 1.8 0.7016 3.5 0.05796 3.0 0.52 543 10 540 19 528 65 103
Klotz 2–16 2612 65 6 0.50 1201 0.09243 1.8 0.7528 3.4 0.05907 2.9 0.52 570 10 570 19 570 62 100
Klotz 2–17 12298 314 27 0.13 6924 0.08793 1.6 0.6993 2.7 0.05769 2.2 0.59 543 9 538 14 518 47 105
Klotz 2–20 7134 153 15 0.61 948 0.08820 1.5 0.6877 4.8 0.05655 4.6 0.30 545 8 531 26 474 102 115
Klotz 2–22 2939 60 6 0.48 860 0.09384 1.9 0.7707 3.5 0.05957 3.0 0.55 578 11 580 20 588 64 98
Klotz 2–23 3617 93 8 0.29 3210 0.08628 2.3 0.6804 6.4 0.05719 6.0 0.36 533 12 527 34 499 132 107
Klotz 2–24 4762 114 10 0.16 3590 0.08818 1.7 0.7056 3.4 0.05804 3.0 0.49 545 9 542 19 531 65 103
Klotz 2–25 3715 91 9 0.50 2142 0.09248 1.5 0.7422 3.2 0.05821 2.8 0.48 570 9 564 18 538 61 106
Klotz 2–26 3337 80 7 0.40 1807 0.08778 1.8 0.7083 3.1 0.05852 2.5 0.58 542 10 544 17 549 55 99
Klotz 2–27 3819 87 8 0.37 1000 0.09271 1.5 0.7560 2.5 0.05914 2.0 0.60 572 9 572 14 572 44 100
Klotz 2–28 10344 153 15 0.83 398 0.08641 2.0 0.7036 3.1 0.05906 2.4 0.63 534 10 541 17 569 53 94
Klotz 2–29 5029 115 10 0.44 1302 0.08776 1.8 0.6628 6.5 0.05477 6.2 0.28 542 10 516 34 403 140 135
Klotz 2–31 7148 178 15 0.21 1252 0.08767 1.7 0.6965 2.9 0.05762 2.3 0.60 542 9 537 15 515 50 105
Klotz 2–32 23586 478 43 0.54 836 0.08699 1.7 0.7086 2.8 0.05908 2.3 0.59 538 9 544 15 570 50 94
Klotz 2–33 4474 116 11 0.43 2569 0.08861 1.4 0.7132 2.0 0.05838 1.3 0.73 547 8 547 11 544 29 101
MH 05 (monzonite, Carboniferous, Visean, Meissen Massif, Elbe Zone, Boxdorf in the north of the city of Dresden, Easting: 54 09338, Northing: 56 65295)
MH 05 – 3 5704 533 33 0.92 932 0.05352 2.2 0.3979 5.2 0.05391 4.7 0.43 336 7 340 15 367 106 91
MH 05 – 7 5097 570 29 0.17 3545 0.05336 1.9 0.3937 2.2 0.05351 1.2 0.85 335 6 337 6 350 26 96
MH 05 – 8 20408 663 55 1.55 301 0.05293 2.0 0.3855 2.8 0.05283 2.0 0.70 332 6 331 8 321 45 103
MH 05 – 11 17585 897 67 1.34 808 0.05262 2.2 0.3958 5.6 0.05456 5.2 0.40 331 7 339 16 394 116 84
MH 05 – 13 5792 634 48 1.67 2140 0.05213 1.9 0.3810 4.4 0.05301 4.0 0.42 328 6 328 13 329 91 100
MH 05 – 18 2842 315 27 2.51 6167 0.05275 4.1 0.3966 5.2 0.05453 3.2 0.79 331 13 339 15 393 72 84
MH 05 – 22 6118 684 39 0.65 2081 0.05464 2.5 0.3955 4.1 0.05250 3.3 0.61 343 8 338 12 307 74 112
MH 05 – 23 6671 726 42 0.68 1630 0.05462 2.2 0.3979 3.5 0.05284 2.8 0.62 343 7 340 10 322 63 107
MH 05 – 24 5611 700 39 1.02 1765 0.05435 2.2 0.3983 4.8 0.05316 4.3 0.45 341 7 340 14 336 98 102
MH 05 – 26 3846 352 20 0.38 870 0.05215 2.1 0.3778 4.6 0.05254 4.1 0.46 328 7 325 13 309 93 106
MH 05 – 29 4330 962 51 0.58 2038 0.05193 4.0 0.3859 5.6 0.05390 3.9 0.72 326 13 331 16 367 88 89
MH 05 – 31 6983 695 43 1.15 796 0.05427 3.7 0.3949 4.9 0.05278 3.2 0.75 341 12 338 14 319 72 107
205
(Continued)
Table 1. Continued
206
207
Name Pba Ub Pbb Thb/U Isotopic ratiosc Rhod Ages Conc
(spot) (cps) (ppm) (ppm) (%)
206 204 206 238 207 235 207 206 206 238 207 235 207 206
Pb/ Pb Pb/ U 2s Pb/ U 2s Pb/ Pb 2s Pb/ U +2s Pb/ U +2s Pb/ Pb +2s
(%) (%) (%) (Ma) (Ma) (Ma)
MH 10 (biotite orthogneiss, Lower Cambrian, Meissen Massif, Elbe Zone, north of the city of Dresden, Easting: 54 14601, Northing: 56 63399)
MH 10 – 3 8026 47 29 1.62 5782 0.3953 2.8 7.452 3.2 0.1367 1.6 0.86 2148 51 2167 29 2186 29 98
MH 10 – 4 5909 306 33 0.49 1304 0.08656 1.7 0.6843 4.3 0.05734 3.9 0.40 535 9 529 18 504 86 106
MH 10 – 5 7213 372 38 0.24 1060 0.08742 1.6 0.6955 3.9 0.05770 3.5 0.40 540 8 536 16 518 78 104
MH 10 – 7 4626 247 27 0.19 3821 0.09868 1.6 0.8205 2.1 0.06031 1.4 0.74 607 9 608 10 615 31 99
MH 10 – 8 15354 529 73 0.82 400 0.1005 1.8 0.8387 3.4 0.06052 2.9 0.53 617 11 618 16 622 62 99
MH 10 – 9 4089 246 26 0.41 2728 0.08874 1.6 0.7198 2.8 0.05883 2.3 0.57 548 8 551 12 561 50 98
MH 10 – 10 4536 201 23 0.71 776 0.08776 2.4 0.6972 7.1 0.05762 6.7 0.33 542 12 537 30 515 148 105
MH 10 – 11 2655 155 17 0.67 4388 0.08807 2.1 0.7184 3.5 0.05916 2.8 0.60 544 11 550 15 573 61 95
MH 10 – 12 52138 200 127 0.94 17630 0.4300 3.9 10.26 4.1 0.1731 1.3 0.95 2306 76 2459 39 2588 22 89
MH 10 – 13 2687 144 19 1.06 4374 0.09726 2.2 0.8000 3.4 0.05966 2.5 0.66 598 13 597 15 591 55 101
MH 10 – 14 11313 564 62 0.48 2972 0.08816 2.6 0.7116 3.4 0.05854 2.2 0.76 545 13 546 14 550 48 99
MH 10 – 16 9437 243 25 0.12 11088 0.08811 4.3 0.6971 6.1 0.05738 4.3 0.71 544 23 537 26 506 95 108
Elbe 1 (felsic ash bed, Upper Carboniferous, Döhlen Basin, Elbe Zone, Elbstolln near Dresden, Easting 54 03464, Northing: 56 55772)
M. HOFMANN ET AL.
Elbe 1–1 2244 122 6 0.45 4287 0.04760 2.1 0.3428 3.9 0.05223 3.2 0.54 300 6 299 10 296 74 101
Elbe 1–4 2005 102 5 1.71 3333 0.04828 3.3 0.3474 6.7 0.05219 5.8 0.50 304 10 303 18 294 133 104
Elbe 1–5 6732 176 10 0.69 4271 0.04830 2.2 0.3384 9.6 0.05081 9.3 0.23 304 7 296 25 232 215 131
Elbe 1–6 4192 205 10 0.44 7953 0.04889 1.5 0.3560 2.6 0.05281 2.1 0.59 308 5 309 7 321 47 96
Elbe 1–9 3553 136 7 0.41 997 0.04841 2.5 0.3560 6.9 0.05334 6.4 0.36 305 7 309 18 343 145 89
Elbe 1–11 2013 96 5 0.59 3788 0.04837 1.8 0.3517 3.5 0.05273 3.0 0.51 305 5 306 9 317 69 96
Mark 1 (granite, Lower Carboniferous, Elbe Zone, near Markersbach, Easting 54 29073, Northing: 56 35602)
Mark 1–10 17185 273 16 0.76 608 0.05068 2.3 0.3671 4.0 0.05253 3.3 0.58 319 7 317 11 309 75 103
Mark 1–13 79884 1190 71 0.23 396 0.05333 2.3 0.3900 2.5 0.05303 1.1 0.90 335 8 334 7 330 25 101
Mark 1–27 609202 1985 182 0.11 75 0.05328 7.2 0.3877 8.8 0.05278 4.9 0.83 335 24 333 25 319 112 105
Mark 1–28 3013 76 4 0.57 5438 0.05264 2.2 0.3868 3.7 0.05330 2.9 0.60 331 7 332 10 341 66 97
Mark 1–32 249445 3010 190 0.08 362 0.05304 4.8 0.3923 5.9 0.05365 3.5 0.81 333 16 336 17 356 78 93
Mark 1–40 114711 2228 231 0.13 4157 0.05274 4.6 0.3809 4.9 0.05238 1.5 0.95 331 15 328 14 302 34 110
Mark 1–41 145046 3151 165 0.10 1085 0.05006 3.6 0.3614 3.8 0.05236 1.4 0.93 315 11 313 10 301 33 105
Mark 1–42 65241 1377 84 0.25 1758 0.05369 2.1 0.3906 2.4 0.05277 1.2 0.87 337 7 335 7 319 27 106
Mark 1–44 333624 2607 194 0.12 96 0.04918 5.7 0.3639 6.8 0.05366 3.7 0.84 309 17 315 19 357 84 87
Mark 1–48 87842 2117 1 0.14 1579 0.05172 2.4 0.3777 2.5 0.05296 0.7 0.95 325 8 325 7 327 17 99
Mark 1–51 82370 904 5 0.28 166 0.05156 2.6 0.3763 3.9 0.05293 2.9 0.67 324 8 324 11 326 66 99
Mark 1–52 102843 1295 5 0.21 298 0.05001 3.5 0.3639 3.8 0.05278 1.4 0.92 315 11 315 10 320 33 98
Mark 1–55 259099 3073 5 0.09 285 0.05318 5.1 0.3841 5.6 0.05238 2.3 0.91 334 17 330 16 302 52 111
a
Within-run background-corrected mean 207Pb signal in counts per second.
b
U and Pb content and Th/U ratio were calculated relative to GJ-1 and are accurate to approximately 10%.
c
Corrected for background, mass bias, laser induced U –Pb fractionation and common Pb (if detectable, see analytical method) using Stacey & Kramers (1975) model Pb composition. 207Pb/235U calculated using
207
Pb/206Pb/(238U/206Pb 1/137.88). 238U/206Pb errors are propagated by quadratic addition of within-run errors (2 standard error) and the reproducibility of GJ-1 (2 standard deviation). 207Pb/206Pb errors were
propagated following Gerdes & Zeh (2008).
d
Rho is the error correlation defined as err206Pb/238U/err207Pb/235U.
See text for more details.
Fig. 6. Lithosection (normal profile) of the Döhlen Basin (Upper Carboniferous– Lower Permian) and location of
the dated ash bed (sample Elbe 1) (modified after Reichel & Schauer 2007). Sample Elbe 1 was taken within the
sequence of coal seam 5. Photograph to the left (by WISMUT GmbH 2008) shows the sample location in the gallery
‘Tiefer Elbstolln’ at 5872 m. BHF, Bannewitz –Hainsberg Formation; NSF, Niederhäslich– Schweinsdorf Formation;
DF, Döhlen Formation; UPF, Unkersdorf–Potschappel Formation; B, Palaeozoic Basement.
Fig. 7. U/Pb concordia plots of zircon grains from samples MH 10, Klotz 2, Mark 1, MH 5 and Elbe 1 (for location of
the samples see Figs 5 & 6 and Table 1). Concordia plots (a) and (b) show the age of the two Cadomian samples (MH 10,
Klotz 2), whereas the plots (c), (d) and (e) show the age of the Variscan samples (Mark 1, MH 05, Elbe 1). All data-point
error ellipses are 2s. MSWDCþE ¼ mean squared weighted deviation of concordance and equivalence. (a) and
(b): The ages of the younger zircon populations are interpreted as the crystallization age of both Cadomian granitoids
(542 + 4 Ma and 543 + 2 Ma, respectively). The older populations are interpreted in terms of inheritance from the
value) of the analyzed GJ-1 grain was typically which was deformed and metamorphosed during
around 3–9%. Reported uncertainties (2s) of the Variscan strike-slip in the Elbe Zone (Figs 4 & 5).
206
Pb/238U ratio were propagated by quadratic From the monzonite of the Meissen Massif
addition of the external reproducibility (2 SD %) (sample MH 05) 12 grains analyzed were concor-
obtained from the standard zircon GJ-1 (n ¼ 12; dant and equivalent with a concordia age of
2 SD c. 1.3%) during the analytical session and 334 + 3 Ma (Figs 7d & 8a–b, Table 1). This
the within-run precision of each analysis (2 SE %; zircon population is the only one that could be
standard error). In the case of the 207Pb/206Pb, found and is interpreted to reflect the crystallization
we used a 207Pb signal-dependent uncertainty pro- age of the monzonite. This age is more precise than,
pagation (Gerdes & Zeh 2008). The accuracy of but overlaps within error with, the U – Pb SHRIMP
the method was verified by analyses of reference age of Nasdala et al. (1999), who published crystal-
zircon 91500 (1064.8 + 4.3 Ma, MSWD of con- lization ages of 326 + 6 Ma and 330 + 5 Ma for
cordance and equivalence ¼ 0.86), Plešovice the monzonite of the Meissen Massif from other
(337.7 + 1.6 Ma, MSWDCþE ¼ 0.84) and Temora localities, and confirms the importance of magma-
(416.6 + 2.5 Ma, MSWDCþE ¼ 0.9). tism of this age in the region.
In the Markersbach Granite (sample Mark 1) 13
Results grains were analyzed (Figs 7c & 8c–d, Table 1) and
gave equivalent and concordant results. These
Results of the U –Pb geochronology are given in grains represent the only population that could be
Table 1 and in Figures 7, 8 and 9. For construction of found in this sample. The calculated concordia age
concordia plots and the calculation of concordia ages is 327 + 4 Ma and interpreted to be the crystalliza-
Isoplot 2.49 (Ludwig 2001) was used. The grano- tion age of the granite.
diorite from the Lausitz Block (sample Klotz 2) Six zircons from the felsic ash bed of the coal
(Fig. 7b, Table 1) exhibits two distinct zircon popu- seam 5 from the Döhlen Basin were analyzed and
lations. The older population yielded a concordia yielded equivalent and concordant results corres-
age of 570 + 4 Ma (7 grains). The younger popu- ponding to the concordia age of 305 + 3 Ma
lation gave a concordia age of 543 + 2 Ma (Figs 7e & 8e –f, Table 1). This age is interpreted to
(21 grains), which is interpreted as the crystalliza- date the volcanic eruption that had formed the ash
tion age of this sample. This lowermost Cambrian bed. Further, it provides an absolute age for the
age of intrusion indicates that the granodiorite is deposition of coal seam 5 in the Döhlen Formation.
part of the Cadomian basement from the southern
part of the Lausitz Block. The older c. 570 Ma popu- Discussion and conclusion
lation is interpreted to reflect inherited zircon grains
consistent with the known age of magmatic activity Our study identifies the granitoids from the Lausitz
in the Cadomian arc (e.g. Linnemann et al. 2007a). Block (sample Klotz 2) and from the Elbe Zone
The orthogneiss of the Elbe Zone (sample MH (sample MH 10) as intruded at c. 542–543 Ma.
10) (Fig. 7a, Table 1) shows similarities to sample They, therefore, belong to the Cadomian orogenic
Klotz 2 concerning the concordia age of its youngest cycle. According to Linnemann et al. (2007a) grani-
zircon population, which yielded an age of toids of this age were generated during an interval of
542 + 4 Ma (7 grains). This age is likewise inter- high heat flow as a result of slap break-off during the
preted to represent the protolith age of the ortho- arc –continent-collision of the Cadomian Orogeny.
gneiss. Three older grains with a concordia age of The Cadomian granodiorite from the Lausitz Block
608 + 6 Ma are interpreted as inherited zircons (sample Klotz 2) is not deformed by the strike-slip
from an older Cadomian magmatic event. Other movements because it lays outside the shear zones
analyzed zircons from the orthogneiss include a along the Elbe Zone. In contrast, the Cadomian
Palaeoproterozoic zircon with a 207Pb/206Pb age orthogneiss within the Elbe Zone (sample MH 10)
of 2186 + 29 Ma and one Neoarchaean zircon obtained its gneissic foliation from the Variscan
with an age of 2588 + 22 Ma (207Pb/206Pb) strike-slip processes along the Elbe Zone.
(Table 1). The crystallization age and the inherited However, the conditions during this shearing event
zircon populations suggests that the orthogneiss is did not affect the igneous zircon grains in
part of the Gondwanan Cadomian basement, the protoliths.
Fig. 7. (Continued) activity of the Cadomian arc magmatism at 608 + 6 Ma and 570 + 4 Ma, respectively. (c) and (d):
Zircon grains of Variscan plutonic rocks of the Elbe Zone. The monzonite of the Meissen Massif (MH 05) results in a
concordia age of 334 + 3 Ma. The Markersbach Granite (sample Mark 1) is slightly younger (327 + 4 Ma).
(e): Concordia plot showing U –Pb ages of 6 zircons of the ash bed in coal seam 5 of the Döhlen Formation (Elbe 1). The
concordia age of 305 + 3 Ma is interpreted as dating the volcanic eruption that formed the felsic ash bed and thus also
reflects the absolute age of sedimentation for coal seam 5 from the Döhlen Formation (Döhlen Basin).
Fig. 8. Cathodoluminescence–SEM–Images of representative zircon grains of different samples, which have

been dated in this study to constrain the timing of the strike-slip and the final tectonic procedures during the Variscan
Orogeny. (a) and (b): Images of zircon grains of sample MH 05 from the Meissen Massif (Elbe Zone) frequently
reveal a uniform or complex zoned core and an wide outer domain with well-developed oscillatory zoning. The
obtained ages from different domains, however, are not distinguishable, suggesting that this structure reflect some
change of the physico-chemical conditions during magma crystallization. The obtained Variscan concordia age of
334 + 3 Ma for the monzonite is interpreted to date an active strike-slip regime. (c) and (d): Zircon grains of sample
Mark 1 generally show dark luminescent CL images with less well developed or wide oscillatory zoning. This reflects
their high U content of up to 3150 ppm. Although the U –Pb spots have been carefully selected it is still remarkable
that no Pb-loss has been observed despite of the high U content. It clearly indicates that no later event has affected
this granite. The Variscan concordia age of 327 + 4 Ma for the granite of Markersbach (Elbe Zone) marks the end of
the strike-slip movements along the Elbe Zone. (e) and (f): Zircon grains of sample Elbe 1 show well-developed
oscillatory zoning and yielded equivalent and concordant ages. The Variscan concordia age of 305 + 3 Ma for the
ash bed within coal seam 5 of the Döhlen Formation in the Döhlen Basin (Elbe Zone) is reflecting a time where all
basement units where exhumed and sedimentation already started.
Fig. 9. Timing of the final pulse of the Variscan Orogeny and general geotectonic setting during the Carboniferous
in the Elbe Zone and adjoining areas (summary) showing all samples with their zircon U/Pb ages presented in
this study and additional data from the literature. Source of geochronological data: *, U–Pb LA-SF-ICP-MS (this
study); **, Pb–Pb TIMS (Gehmlich 2003); ***, U –Pb TIMS (Romer & Rötzler 2001; Rötzler & Romer 2001);
****, U–Pb SHRIMP (Nasdala et al. 1996).
The final pulse of the Variscan Orogeny is that exhumation forced dextral strike-slip pro-
responsible for the formation of the Elbe Zone. cesses along the Elbe Zone during D2-deformation
Due to the D2-deformation (NW-ward transport of and NW-ward transport of allochthonous units.
the allochthonous units Saxonian Granulitgebirge The Meissen Massif intruded at c. 334 Ma (age of
and Erzgebirge nappe complex) in the northern intrusion, Fig. 7d) into a local area of extension in
Bohemian Massif, dextral strike-slip movements the schist belt of the Elbe Zone during ongoing
along today’s Elbe Zone were produced (Fig. 4) dextral shear (Kroner et al. 2007). The sigmoidal
(Kroner et al. 2007). Our geochronological data shape of the Meissen Massif demonstrates that
allow a precise timing of related plutonic, volcano- strike-slip movements along the Elbe Zone were
sedimentary and tectonic processes (Fig. 9). still active at that time (Fig. 4). Dextral shear had
Fossils from the sedimentary rocks in the Elbe terminated before the Markersbach Granite intruded
Zone demonstrate that marine sedimentation at 327 + 4 Ma, because the granite cross-cuts all
occurred until the Lower Carboniferous (Kurze tectonic structures related to the dextral shear. In
et al. 1992). The only published age from marine contrast to the monzonite of the Meissen Massif,
Lower Carboniferous volcano-sedimentary com- the Markersbach Granite itself is undeformed. At
plexes is a Pb– Pb-age of 357 + 2 Ma from a about the same time, the first terrestrial molasse
syn-sedimentary submarine quarz-keratophyre (Early Variscan Molasse of Hainichen) was de-
extrusion (Gehmlich 2003) from a tectonostrati- posited at 330 + 4 Ma along the southern periphery
graphic unit close to the Elbe Zone (Frankenberg of the Saxonian Granulitgebirge overlying the
Zwischengebirge) (Fig. 9). This unit is situated Frankenberg Zwischengebirge (Gehmlich 2003)
between the Saxonian Granulitgebirge and the (Figs 4 & 9). These events in the range of c. 327–
Erzgebirge nappe complex (Fig. 4). From these 330 Ma are related to the final pulse of the Variscan
data we infer that Lower Carboniferous sedimen- Orogeny in the northern Bohemian Massif (Fig. 9).
tation occurs until at least the Tournaisian (Fig. 9). For the time span between c. 327 Ma and the
Peak metamorphism for the Saxonian Granulit- opening of the Döhlen Basin in Upper Carbon-
gebirge is defined at c. 342 Ma (Fig. 9). After that, iferous times, no significant magmatism is recorded.
exhumation of the Saxonian Granulitgebirge into But during this episode the general uplift and base-
upper crustal level occurred until c. 323 Ma ment exhumation to the palaeosurface occurred.
(Romer & Rötzler 2001; Rötzler & Romer 2001). The Döhlen Basin overlies the metamorphic rocks
According to Kroner et al (2007), the beginning of of the Erzgebirge nappe complex, the schist belt of
the Elbe Zone and the Meissen Massif (Figs 4 & 5). The final exhumation and equilibrium of the
In addition, equivalents of the Döhlen Basin overlie Saxo-Thuringian crust in the area of the Elbe Zone
the Cadomian basement of the Lausitz Block took place in the time span from c. 327 Ma until
(Weissig Basin). the onset of Late Carboniferous molasse deposition
This study provides the first precise estimate of at c. 305 Ma in the Döhlen Basin (this study). From
the commencement of sedimentation in the Döhlen the beginning of subduction of continental crust at
Basin. Available palaeontological data could not c. 370 Ma until the opening of the Late
resolve the age of the onset of sedimentation. For Carboniferous molasse basins at c. 305 Ma, it took
example, the palaeoflora only allows to constrain c. 65 Ma. When we interpolate that timing to the
the age of deposition to the transition from Upper formation of the whole of Pangaea, we assume
Carboniferous–Lower Permian. The dated ash bed that for the formation of the supercontinent,
(sample Elbe 1 305 + 3 Ma) from coal seam 5 of c. 65 Ma were needed from the beginning of con-
the Döhlen Formation shows clearly that sedimen- tinental subduction to the start of dispersal rep-
tation in the Döhlen Basin commenced by the Upper resented by general stretching of the lithosphere
Carboniferous. However, the start of sedimentation and opening of molasse basins, as the Döhlen Basin.
in the basin must be earlier, because the underlying
Unkerdorf–Potschappel–Formation unconformably This is a contribution to IGCP 453 ‘Modern and Ancient
overlies the Cadomian basement (Figs 5 & 6). Orogens’ and IGCP 497 (‘The Rheic Ocean: Its Origin,
The time of general uplift, exhumation and denu- Evolution and Correlatives’). The authors benefited from
funding by the IGCP 497, by the Ludwig–Reichenbach–
dation of the Variscan basement lasted a maximum
Gesellschaft e.V., and the Museum of Mineralogy
of c. 22 Ma between the intrusion of the Markers- and Geology (Senckenberg Naturhistorische Sammlungen
bach Granite at c. 327 Ma and the onset of sedimen- Dresden). R. Strachan (Portsmouth, UK) and M. F. Pereira
tation in the Döhlen Basin at c. 305 Ma (Fig. 9). (Evora, Portugal) are thanked very much for critical and
This study defines the very end of Variscan constructive reviews. Special thanks to B. Murphy for a
orogenic processes in central Europe. These pro- lot of useful discussions and numerous corrections, not
cesses where strongly linked to the evolution of only of the language of the manuscript.
the Rheic Ocean, especially to its closure. As the
closure of the Rheic started in the west and pro-
gressed towards the east (Fig. 3), collision of the
References
dispersed Cadomian blocks with the Old Red Con- B ONIN , B. 2007. A-type granites and related rocks:
tinent (Laurussia) and related subduction proces- evolution of a concept, problems and prospects.
ses show chronological differentiations and vary Lithos, 97, 1 –29.
between Early Devonian until Early Carboniferous B USCHMANN , B., N ASDALA , L., J ONAS , P.,
(Kroner et al. 2007). Due to the northward-drift L INNEMANN , U. & G EHMLICH , M. 2001. SHRIMP
of Gondwana (and Cadomia, see Fig. 3) closure of U–Pb dating of tuff-derived and detrital zircons
from Cadomian marginal basin fragments (Neopro-
the Rheic started at c. 400 Ma and triggered sub- terozoic) in the northeastern Saxothuringian Zone
duction of thinned Gondwanan continental litho- (Germany). Neues Jahrbuch Geologie Paläontologie,
sphere in central Europe (i.e. Bohemian Massif) at Monatshefte, 321–342.
c. 370 Ma (Linnemann et al. 2007a; Kroner et al. C HANTRAINE , J., A UVRAY , B., B RUN , J. P., C HAUVEL ,
2007). In the Bohemian Massif (BM, see Fig. 1), J. J. & R ABU , D. 1994. Introduction. In: K EPPIE ,
the allochthonous domains (e.g. Saxonian Granulit- J. D. (ed.) Pre-Mesozoic Geology in France and
gebirge, Erzgebirge nappe complex, see Fig. 4) Related Areas. Springer, Berlin, 75–80.
mark such parts of subducted continental crust. C HAUVEL , J. J. & R ABU , D. 1988. Brioverian in
At least for the northern BM in central Europe, Central Brittany. In: Z OUBEK , V., C ONGE , D.,
K OZOUKHAROV , H. & K RAUTNER , G. (eds) Pre-
subduction-related D1-deformation lasted until cambrian in Younger Fold Belts. J. Wiley & Sons,
c. 340 Ma (Kroner et al. 2007). D2-deformation Chichester, 462– 470.
and related dextral strike-slip movements were E LICKI , O. 1997. Biostratigraphic data of the German
still active at c. 334 Ma represented by the intrusion Cambrian – present state of knowledge. Freiberger
of the Meissen Massif into the Elbe Zone (this Forschungshefte, C 466, 155– 165.
study, Fig. 7d). This means that, for central F ERNÁNDEZ -S UÁREZ , J., G UTIÉRRES -A LONSO , G.,
Europe, it took c. 36 Ma from the beginning of J ENNER , G. A. & T UBRETT , M. N. 2000. New ideas
first subduction of continental lithosphere (c. on the Proterzoic-Early Palaeozoic evolution of NW
370 Ma) up to the point of completed subduction Iberia: insights from U– Pb detrital zircon ages.
Precambrian Research, 102, 185– 206.
of oceanic and ongoing subduction of thinned con- F RANKE , W. 2000. The mid-European segment of
tinental crust and final orogen-related tectonic pro- the Variscides: tectonometamorphic units, terrane
cesses (c. 334 Ma). This D2-deformation was boundaries and plate tectonic evolution. In: F RANKE ,
finished at c. 327 Ma marked by the intrusion of W., H AAK , V., O NCKEN , O. & T ANNER , D. (eds)
the Markersbach Granite (this study, Fig. 7c). Orogenic processes – Quantification and Modelling
in the Variscan Belt of Central Europe. Geological Terranes – Konsequenzen aus einer abgedeckten
Society, London, Special Publications, 179, 35–61. geologischen Karte. Zeitschrift für Geologische
F RANKE , W. & Ż ELAZNIEWICS , A. 2002. Structure and Wissenschaften, 27, 529– 561.
Evolution of the Bohemian Arc. Geological Society, L INNEMANN , U., G EHMLICH , M. ET AL . 2000. From
London, Special Publications, 201, 279–293. Cadomian subduction to Early Palaeozoic rifting: the
F REI , D. & G ERDES , A. 2009. Precise and accurate in situ evolution of Saxo-Thuringia at the margin of Gond-
U–Pb dating of zircon with high sample throughput by wana in the light of single zircon geochronology and
automated LA-SF-ICP-MS. Chemical Geology, 261, basin development (Central European Variscides,
261–270, doi: 10.1016/j.chemgeo.2008.08.025. Germany). London. Geological Society, London,
G EHMLICH , M. 2003. Die Cadomiden und Varisziden des Special Publications, 179, 131–153.
Saxothuringischen Terranes – Geochronologie mag- L INNEMANN , U., M C N AUGHTON , N. J., R OMER , R. L.,
matischer Ereignisse. Freiberger Forschungshefte, G EHMLICH , M., D ROST , K. & T ONK , C. 2004. West
500, 1 –129. African provenance for Saxo-Thuringia (Bohemian
G ERDES , A. & Z EH , A. 2006. Combined U–Pb and Massif): did Armorica ever leave pre-Pangean Gond-
Hf isotope LA-(MC)-ICP-MS analyses of detrital wana? U–Pb-SHRIMP zircon evidence and the
zircons: comparison with SHRIMP and new con- Nd-isotopic record. International Journal of Earth
straints for the provenance and age of an Armorican Sciences, 93, 683–705.
metasediment in Central Germany. Earth and Plane- L INNEMANN , U., G ERDES , A., D ROST , K. & B USCH-
tary Sciences Letters, 249, 47– 61. MANN , B. 2007a. The continuum between Cadomian
G ERDES , A. & Z EH , A. 2008. Zircon formation versus Orogenesis and opening of the Rheic Ocean:
zircon alteration – New insights from combined Constraints from LA-ICP-MS U– Pb zircon dating
U–Pb and Lu–Hf in-situ La-ICP-MS analyses of and analysis of plate-tectonic setting (Saxo-Thuringian
Archean zircons from the Limpopo Belt. Chemical zone, NE Bohemian massif, Germany). In:
Geology, doi: 10.1016/j.chemgeo.2008.07.025. L INNEMANN , U., N ANCE , R. D., K RAFT , P. &
G RAINDOR , M.-J. 1957. Le Briovérien dans le Nord-East Z ULAUF , G. (eds) The Evolution of the Rheic Ocean:
du massif Armoricain. Memoires pour servir a l’expli- From Avalonian-Cadomian Active Margin to
cation de la Carte Géologique detaillée de la France. Alleghenian-Variscan Collision. Boulder, Colorado,
Paris, 1–211. Geological Society of America Special Paper, 423,
K OSSMAT , F. 1927. Gliederung des varistischen Gebirgs- 61–96.
baues. Abhandlungen des Sächsischen Geologischen L INNEMANN , U., N ANCE , R. D., Z ULAUF , G. & K RAFT ,
Landesamtes, 1, 1 –39. P. 2007b. Avalonian-Cadomian Belt, Adjoining
K RONER , U., H AHN , T., R OMER , R. L. & L INNEMANN , Cratons and the Rheic Ocean. In: L INNEMANN , U.,
U. 2007. The Variscan orogeny in the Saxo-Thuringian N ANCE , R. D., K RAFT , P. & Z ULAUF , G. (eds) The
zone-heterogenous overprint of Cadomian/Palaeozoic Evolution of the Rheic Ocean: From Avalonian–
peri-Gondwana crust. In: L INNEMANN , U., N ANCE , Cadomian Active Margin to Alleghenian– Variscan
R. D., K RAFT , P. & Z ULAUF , G. (eds) The Evolution Collision. Boulder, Colorado, Geological Society of
of the Rheic Ocean: From Avalonian-Cadomian America Special Paper, 423, 7– 8.
Active Margin to Alleghenian-Variscan collision. L INNEMANN , U., P EREIRA , F., J EFFRIES , T. E., D ROST ,
Boulder, Colorado, Geological Society of America K. & G ERDES , A. 2008a. The Cadomian Orogeny
Special Paper, 423, 153 –172. and the opening of the Rheic Ocean: the diacrony of
K URZE , M., L INNEMANN , U. & T RÖGER , K.-A. geotectonic processes constrained by LA-ICP-MS
1992. Weesensteiner Gruppe und Altpaläozoikum U– Pb zircon dating (Ossa-Morena and Saxo-
in der Elbtalzone (Sachsen). Geotektonische For- Thuringian Zones, Iberian and Bohemian Massifs).
schungen, 77, 101 –167. Tectonophysics, 461, 21– 43.
L INNEMANN , U.-G. 1994. Geologischer Bau und Struk- L INNEMANN , U., R OMER , R. L. ET AL . 2008b. Introduction
turentwicklung der südlichen Elbezone. Abhandlungen (Chapter 2: ‘The Precambrian’). In: M C C ANN , T. (ed.)
des Staatlichen Museums für Mineralogie und Geolo- The Geology of Central Europe. Geological Society,
gie zu Dresden, 40, 7 –36. London, Special Publications, 21– 102.
L INNEMANN , U. 2003a. Die Struktureinheiten des Sax- L INNEMANN , U., D’L EMOS , R. ET AL . 2008c. Introduc-
othuringikums. In: L INNEMANN , U. (ed.) Das tion (Chapter 3: ‘The Cadomian Orogeny’). In:
Saxothuringikum. Geologica Saxonica, 48/49, 19– 28. M C C ANN , T. (ed.) The Geology of Central Europe.
L INNEMANN , U. 2003b. Sedimentation und geotekto- Geological Society, London, Special Publications,
nischer Rahmen der Beckenentwicklung im Saxo- 103– 154.
thuringikum (Neoproterozoikum – Unterkarbon). In: L UDWIG , K. R. 2001. Users Manual for Isoplot/Ex rev.
L INNEMANN , U. (ed.) Das Saxothuringikum. Geolo- 2.49. Berkeley Geochronology Center Special Publi-
gica Saxonica, 48/49, 71–110. cation, 1, 1–56.
L INNEMANN , U. & R OMER , R. L. 2002. The Cadomian M ARTÍNEZ , S. S., A RENAS , R., G ARCÍA , F. D.,
Orogeny in Saxo-Thuringia, Germany: geochemical M ARTÍNEZ C ATALAN , J. R., G ÓMEZ -B ARREIRO , J.
and Nd– Sr– Pb isotopic characterisation of marginal & P EARCE , J. A. 2007. Careón ophiolithe, NW
basins with constraints to geotectonic setting and pro- Spain: Suprasubduction zone setting for the youngest
venance. Tectonophysics, 352, 33– 64. Rheic Ocean floor. Geology, 35, 53–56.
L INNEMANN , U. & S CHAUER , M. 1999. Die Entstehung M ASSONNE , H.-J. 1998. A new occurrence of microdia-
der Elbezone vor dem Hintergrund der cadomischen monds in quartzofeldspathic rocks of the Saxonian Erz-
und variszischen Geschichte des Saxothuringischen gebirge, Germany, and their metamorphic evolution.
7th International Kimberlite Conference, Abstracts, (ed.) Das Saxothuringikum. Geologica Saxonica,
Cape Town, 552–554. 48/49, 19–28.
M ATTERN , F. 1996. The Elbe Zone at Dresden – a R OSCHER , M. & S CHNEIDER , J. W. 2005. An annotated
Late Palaeozoic pull-apart intruded shear zone. correlation chart for continental Late Pennsylvanian
Zeitschrift der Deutschen Geologischen Geselschaft, and Permian basins and the marine scale. In: L UCAS ,
147, 57–80. S. G. & Z EIGLER , K. E. (eds) The nonmarine
N ANCE , R. D., M URPHY , J. B. & K EPPIE , J. D. 2002. Permian. New Mexico Museum of Natural History
A Cordilleran model for the evolution of Avalonia. and Science Bulletin, 30, 282–291.
Tectonophysics, 352, 11– 31. R ÖTZLER , J. & R OMER , R. L. 2001. P –T –t Evolution
N ANCE , R. D., B RENDAN , M. ET AL . 2008. of Ultrahigh-Temperature Granulites from the Saxon
Neoproterozoic-early Palaeozoic tectonostratigraphy Granulite Massif, Germany. Part I: petrology.
and palaeogeography of the peri-Gondwanan terranes: Journal of Petrology, 42/11, 1995– 2013.
Amazonian v. West African connections. In: E NNIH , S ÁNCHEZ -G ARCÍA , T., B ELLIDO , F. & Q UESADA , C.
H. & L IÉGEOIS , L.-P. (eds) The Boundaries of the 2003. Geodynamic setting and geochemical signatures
West African Craton. Geological Society, London, of Cambrian-Ordovician rift-related igneous rocks
Special Publications, 297, 345–383. (Ossa-Morena Zone, SW Iberia). Tectonophysics,
N ASDALA , L., G RUNER , T., N EMCHIN , A. A., P IDGEON , 365, 233 –255.
R. T. & T ICHOMIROWA , M. 1996. New SCHRIMP ion S ÁNCHEZ M ARTÍNEZ , S., A RENAS , R., D ÍAZ G ARCÍA ,
microprobe measurements on zircons from Saxonian F., M ARTÍNEZ C ATAĹN , J. R., G ÓMEZ -B ARREIRO ,
magmatic and metamorphic rocks. Proceedings of J. & P EARCE , J. A. 2007. Careón ophiolite, NW
the Freiberg Isotope Colloquium, TU Bergakadamie Spain: suprasubduction zone setting for the youngest
Freiberg, 205– 214. Rheic Ocean floor. Geology, 35(1), 53– 56.
N ASDALA , L., W ENZEL , T., P IDGEON , R. T. & K RONZ , A. S COTESE , C. R. & B ARRET , S. F. 1990. Gondwana’s
1999. Internal structures and dating of complex zircons movement over the South Pole during the Palaeozoic:
from Meissen Massif monzonites, Saxony. Chemical evidence from lithological indicators of climate. In:
Geology, 159, 331– 341. M C K ERROW , W. S. & S COTESE , C. R. (eds) Palaeo-
P IETZSCH , K. 1962. Geologie von Sachsen. VEB zoic Palaeogeography and Biogeography. Geological
Deutscher Verlag der Wissenschaften, Berlin. Society, London, Memoir, 12, 75–85.
R EICHEL , W. & S CHAUER , M. 2007. Das Döhlener S TACEY , J. S. & K RAMERS , J. D. 1975. Approximation
Becken bei Dresden – Geologie und Bergbau. of terrestrial lead isotope evolution by a two-stage
Bergbaumonographie. Landesamt für Umwelt und model. Earth and Planetary Science Letters, 26,
Geologie Sachsen. 207–221.
R OBARDET , M. 2002. Alternative approach to the Variscan T ORSVIK , T. H. & R EHNSTRÖM , E. F. 2002. The Torn-
Belt in southwestern Europe: preorogenic paleobio- quist Sea and Baltica-Avalonia docking. Tectonophy-
geographicalo constraints. In: M ARTINEZ C ATALAN , sics, 362, 67–82.
J. R., H ATCHER , R. D. J., A RENAS , R. & D IAZ W ENZEL , T., M ERTZ , D. F., O BERHÄNSLI , R. &
G ARCIA , F. (eds) Variscan-Appalachian Dynamics: B ECKER , T. 1997. Age, geodynamic setting, and
The Building of the Late Palaeozoic Basement. mantle enrichment processes of a K-rich intrusion
Geological Society of America Special Paper, 364, from the Meissen massif (northern Bohemian massif)
1 –15. and implications for related occurrences from the mid-
R OMER , R. L. & R ÖTZLER , J. 2001. P– T– t Evolution of European Hercynian. Geologische Rundschau, 86,
Ultrahigh-Temperature Granulites from the Saxon 556–570.
Granulite Massif, Germany. Part II: geochronology. Z EH , A., G ERDES , A., W ILL , T. M. & M ILLAR , I. L.
Journal of Petrology, 42(11), 2015–2032. 2005. Provenance and Magmatic-Metamorphic Evol-
R OMER , R. L., L INNEMANN , U. & G EHMLICH , M. 2003. ution of a Variscan Island-arc Complex: constraints
Geochronologische und isotopengeochemische Rand- from U– Pb dating, Petrology, and Geospeedometry
bedingungen für die cadomische und variszische of the Kyffhäuser Crystalline Complex, Central
Orogenese im Saxothuringikum. In: L INNEMANN , U. Germany. Journal of Petrology, 46, 1393– 1420.
Variscan intra-orogenic extensional tectonics in the Ossa –Morena
Zone (Évora– Aracena – Lora del Rı́o metamorphic belt, SW Iberian
Massif ): SHRIMP zircon U – Th– Pb geochronology
M. FRANCISCO PEREIRA1*, MARTIM CHICHORRO2, IAN S. WILLIAMS3,
JOSÉ B. SILVA4, CARLOS FERNÁNDEZ5, MANUEL DÍAZ-AZPÍROZ6,
ARTURO APRAIZ7 & ANTONIO CASTRO8
1
Departamento de Geociências, Centro de Geofı́sica de Évora, Universidade
de Évora, Apt. 94, 7001-554 Évora, Portugal
2
Centro de Investigação em Ciência e Engenharia Geológica,
Universidade Nova de Lisboa, Portugal
3
Research School of Earth Sciences, The Australian National University,
Canberra, ACT, 0200, Australia
4
Departamento de Geologia, Faculdade de Ciências, Universidade de Lisboa,
Edifı́cio C3, Campo Grande, Lisboa, Portugal
5
Departamento de Geodinámica y Paleontologı́a, Facultad Ciencias Experimentales,
Universidad de Huelva, Campus Carmen, 21071 Huelva, Spain
6
Departamento Ciencias Ambientales, Universidad Pablo Olavide, 41013 Sevilla, Spain
7
Geodinamika Saila, Zientzia eta Teknologia Fak., Euskal Herriko
Unibertsitatea, Apt. 644, 48080 Bilbo, Spain
8
Departamento de Geologı́a, Universidad de Huelva, E21819 La Rábida, Huelva, Spain
*Corresponding author (e-mail: mpereira@uevora.pt)
Abstract: Following a Middle–Late Devonian (c. 390– 360 Ma) phase of crustal shortening and
mountain building, continental extension and onset of high-medium-grade metamorphic terrains
occurred in the SW Iberian Massif during the Visean (c. 345– 326 Ma). The Évora–Aracena–
Lora del Rı́o metamorphic belt extends along the Ossa– Morena Zone southern margin from
south Portugal through the south of Spain, a distance of 250 km. This major structural domain is
characterized by local development of high-temperature–low-pressure metamorphism (c. 345 –
335 Ma) that reached high amphibolite to granulite facies. These high-medium-grade metamorphic
terrains consist of strongly sheared Ediacaran and Cambrian –early Ordovician (c. 600–480 Ma)
protoliths. The dominant structure is a widespread steeply-dipping foliation with a gently-plunging
stretching lineation generally oriented parallel to the fold axes. Despite of the wrench nature of this
collisional orogen, kinematic indicators of left-lateral shearing are locally compatible with an
oblique component of extension. These extensional transcurrent movements associated with
pervasive mylonitic foliation (c. 345–335 Ma) explain the exhumation of scarce occurrences of
eclogites (c. 370 Ma). Mafic-intermediate plutonic and hypabyssal rocks (c. 355–320 Ma),
mainly I-type high-K calc-alkaline diorites, tonalites, granodiorites, gabbros and peraluminous
biotite granites, are associated with these metamorphic terrains. Volcanic rocks of the same chemi-
cal composition and age are preserved in Tournaisian–Visean (c. 350–335 Ma) marine basins
dominated by detrital sequences with local development of syn-sedimentary gravitational collapse
structures. This study, supported by new U– Pb zircon dating, demonstrates the importance of intra-
orogenic transtension in the Gondwana margin during the Early Carboniferous when the Rheic
ocean between Laurussia and Gondwana closed, forming the Appalachian and Variscan mountains.
Introduction of intra-continental extensional tectonics during

the terminal stages of, and immediately after, conti-
Recent studies of ancient and modern examples of nental accretion (late- to post-orogenic processes).
orogenic belts have demonstrated the importance The main effects of such crustal thinning processes
216 M. F. PEREIRA ET AL.
are distributed over vast areas affected by large- results have been interpreted through comparisons
scale orogen-parallel strike-slip tectonics. High with NW Iberian Massif correlatives, which help
heat flow due to asthenospheric upwelling and us to better understand the crustal growth processes
extensive emplacement of igneous rocks are charac- that led to the amalgamation of Pangaea in the
teristic of such thermally weakened lithosphere. Lower –Middle Mississipian. We discuss the tec-
Ductile deformation on major detachments favours tonic model of intra-continental extension and
the ascent and emplacement of metamorphic onset of high-grade metamorphism that occurred
complexes and the rapid exhumation of medium- at this time, contemporaneously with large-scale
high-pressure metamorphic rocks (Sandiford & orogen-parallel tectonic (transcurrent) movements
Powell 1986). In the upper crust, strike-slip and in the northern Gondwana margin.
normal faulting controls the generation of basins
with the development of gravitational collapse
structures (Rey et al. 2001). The Pangaea amalga- Geological background
mation with the sequential closure of the Rheic
ocean and assembly of continental blocks in a multi- The geological processes related to the formation of
phase Middle and Late Palaeozoic orogenesis Pangaea during the Tournaisian– Visean are recog-
(Variscan Orogeny; Matte 2001) have been high- nizable within the SW Iberian Massif regions
lighted as a major feature of the pre-Permian geody- (Fig. 1), in the western part of the European Varis-
namic history of the Iberian Massif (e.g. Ribeiro can chain. The Ossa–Morena Zone experienced
et al. 1990). During the last two decades several extension involving crustal deformation and meta-
contrasting plate tectonic models have been pro- morphism of deep crust, generation of metamorphic
posed to explain the remnants of the Avalonia– complexes, the emplacement of voluminous
Gondwana collisional margin exposed in the magmas and the development of marine basins.
western Iberian Massif. Exhaustive searching for The Ossa– Morena Zone is composed of conti-
the geodynamic significance of tectonothermal and nental crustal rocks with a complex but continuous
sedimentary events related to the evolution of stratigraphy. The geological record started in the
this Variscan oblique convergence margin (Ossa – Ediacaran with sedimentation and magmatism
Morena Zone/South– Portuguese Zone suture related to a magmatic arc (c. 560 Ma; Quesada
zone) has lead to several conflicting ideas and 1990; Schäfer et al. 1993; Eguı́luz et al. 2000;
sparked ongoing controversy (Silva et al. 1990; Pereira et al. 2006b, 2008), then evolved to a
Crespo-Blanc & Orozco 1991; Fonseca & Ribeiro rifting process, active from the Cambrian to the
1993; Giese et al. 1994; Castro et al. 1996a; Matte Early Ordovician (c. 540–480 Ma; Liñán &
2001; Simancas et al. 2005; Onèzime et al. 2003; Quesada 1990; Sanchez-Garcia et al. 2003;
Silva & Pereira 2004; Dı́az Azpı́roz et al. 2005, Robardet & Gutierrez-Marco 2004; Chichorro
2006; Pereira et al. 2007; Azor et al. 2008; Pin et al. in press). The record of tectonothermal
et al. 2008; Rosas et al. 2008). The same happens processes for the period of approximately 25 Ma in
with the regional geodynamic models that try to the Middle–Late Devonian (c. 395–360 Ma) that
link this Variscan suture with the rootless suture of predates the Tournaisian–Visean evolution is poorly
the NW Iberian Massif (Matte 2001; Robardet known. On contrary, Devonian sedimentation is
2002; Simancas et al. 2005; Ribeiro et al. 2007; well-known and includes flysch (Terena syncline;
Martinez Catalan et al. 2007). Pereira & Oliveira 1996a) and platformal (Valle
The aim of the present study was to contribute to and Cerrón del Hornillo synclines; Robardet &
a better understanding of the geodynamic evolution Gutiérrez Marco 2004) sequences.
of this segment of the European Variscan chain The Évora–Aracena –Lora del Rı́o metamorphic
during the Visean, a period of poorly-known belt discontinuously extends along the southern
intra-orogenic extension. Based on structural, geo- margin of the Ossa–Morena Zone from south
chemical and geochronological data now available Portugal through the south of Spain, a distance of
from studies of high-grade metamorphic terrains 250 km. The belt includes three major exposures of
exposed in Portugal (Évora-Chichorro 2006; metamorphic complexes (high-medium-grade meta-
Pereira et al. 2007) and Spain (Aracena-Dı́az morphic terrains surrounded by lower-grade meta-
Azpı́roz 2006; Dı́az Azpı́roz et al. 2006; and Lora morphic rocks). These metamorphic complexes
del Rı́o-Apraiz 1998; Apraiz & Eguiluz 2002). are characterized by the development of high-
The results obtained are the key to constraining temperature –low-pressure metamorphic conditions
the onset of peak metamorphism and the duration that reached upper amphibolite to granulite facies
of ductile deformation that accompanied the exhu- during the Visean (c. 345–335 Ma), affecting
mation of medium to high-pressure rocks, and also strongly sheared Ediacaran and Cambrian –early
the local gravitational collapse recognizable within Ordovician protoliths (c. 540– 480 Ma). The main
the Tournaisian –Visean sedimentary basins from structure consists of a dominant mylonitic foliation
the Ossa –Morena Zone (Pereira et al. 2006a). Our with gently plunging stretching lineations generally
VARISCAN INTRA-OROGENIC EXTENSIONAL TECTONICS IN THE OSSA-MORENA ZONE 217
Fig. 1. Location of the Évora– Aracena– Lora del Rı́o metamorphic belt in the SW Iberian Massif.
oriented parallel to the fold axes. Kinematic indi- metagreywackes, micaschists, paragneisses and
cators of extension with a coeval left-lateral com- interbedded black metacherts, amphibolites and
ponent (transtension) seem to be locally linked to felsic gneisses (Série Negra; Carvalhosa 1965),
the exhumation of eclogites. an Early– Middle Cambrian igneous (felsic-
The recognition of 20-km-wide MORB-derived dominated) complex with marbles, interbedded
amphibolites and phyllites of unknown age in felsic and mafic metavolcanics, felsic gneisses and
between the Ossa –Morena Zone (Gondwana) and micaschists, and a Late Cambrian –Early Ordovi-
South Portuguese Zone (Laurussia) (Bard & cian igneous (mafic-dominated) complex mainly
Moine 1979; Munhá 1986; Fonseca & Ribeiro composed of amphibolites with micaschists, quart-
1993; Quesada et al. 1994; Castro et al. 1999; zites, metatuffs and calc-silicate rocks (Carvalhosa
Dı́az Azpı́roz et al. 2005, 2006) has been used to 1999; Chichorro 2006; Pereira et al. 2007, 2008;
define a Variscan suture zone (Pulo do Lobo Zone; Chichorro et al. 2008).
Quesada 1991; Quesada et al. 1994). Evidence for Three main units overprint this lithostratigraphy
subduction of oceanic crust (Rheic Ocean?) (Pereira et al. 2003, 2007): a southern low- to
beneath the Ossa –Morena Zone is accepted by the medium-grade unit (greenschist-upper amphibolite
majority of authors but fundamental questions facies; Montemor-o-Novo shear zone) and a north-
regarding the interactions of deformation, meta- ern medium-grade unit (amphibolite facies; Évora
morphism and magmatism related to this particular medium-grade metamorphic terrains) represent the
stage of the SW Iberian Massif geodynamic hanging-wall of a 25 –45 km-wide and 75 km-long
evolution still remain unsolved: Why are there no central high-grade unit (Évora high-grade metamor-
significant testimonies of early Variscan foliation, phic terrains). Orogen-parallel movements were
regional metamorphism, arc magmatism and responsible for the partial exhumation of a central
fore-arc sedimentation related with the supposed high-grade unit with anatectic granitoids, migmati-
Middle–Upper Devonian subduction/collision tic gneisses and diatexites (high-amphibolite facies
(c. 390 –370 Ma)? and transitional between the amphibolite and granu-
lite facies). The huge volume of melt produced
Évora massif during uplift and decompression of this high-grade
unit is indicated by numerous granitoids emplaced
The Évora massif (Fig. 2), located in the south- along the boundary with the hanging-wall low- to
westernmost part of the Ossa –Morena Zone, medium-grade units (Pereira et al. 2007).
extends from Montemor-o-Novo to Évora (Portu- The structure is characterized by a moderately
gal; Carvalhosa 1983; Pereira et al. 2003, 2007). to steeply dipping mylonitic foliation and a
There the lithostratigraphy includes Ediacaran weakly to moderately dipping stretching lineation.
Fig. 2. Schematic representation of the structure and metamorphism distribution in the Èvora Massif (adapted from
Pereira et al. 2007). Location of samples analysed for U– Pb SHRIMP zircon geochronology.
This mylonitic fabric is locally folded with fold axis 1969; Dupuy et al. 1979; Quesada et al. 1994;
parallel to the stretching lineation. Different mineral Castro et al. 1996b) dated at c. 340–332 Ma (Azor
assemblages are found associated with the same et al. 2008); and a northern domain separated
fabric, revealing significant variations in the meta- into a northernmost low- to medium-grade unit
morphic conditions without any noticeable change (greenschist-amphibolite facies) and a southernmost
in the regional strain regime. Application of geo- high-grade unit (upper amphibolite–granulite
thermobarometers to garnet-rich amphibolites from facies). Both units consist of aluminous and calc-
the southern low- to medium-grade unit revealed magnesium series derived from Ediacaran (Série
an increase in temperature and pressure from Negra) and Cambrian protoliths, respectively. The
c. 375 –425 8C at 4.5–5.25 kbar to c. 475 –525 8C high-grade unit is composed of pelitic and calc-
at 6.6–7.5 kbar (Pereira et al. 2007). A regional silicate gneisses, marbles, amphibolites, migmatites
episode of albitization has been interpreted as a con- and granulites (Ediacaran and Early Cambrian pro-
sequence of pressure collapse followed by dynamic toliths) that were metamorphosed at high tempera-
plagioclase recrystallization associated with trans- ture and low pressure. The complex structure of
current shearing (Chichorro 2006). the high-grade unit of the Aracena metamorphic
There has been a strong telescoping of the meta- belt is the result of up to four phases of deformation
morphic gradient, as is typical of extensional (Dı́az Azpı́roz 2006). The pervasive phase was
regimes. The metamorphic isograds along the responsible for the generation of a crenulation and
major detachment that separates the southern unit mylonitic foliation, which is the main planar
from the high-grade unit are telescoped, passing fabric observed in the field. Structural analysis has
from the biotite to the sillimanite zone over a revealed that this main foliation was originally sub-
distance of few hundred metres. Paragneisses, horizontal, although it was later folded during the
which record the first stages of local melting, differ- last deformation phases (Dı́az Azpı́roz 2006).
entiation and the highest strain gradient, are charac- Quartz c-axis fabrics associated with the main foli-
terized by well-developed mylonitic structures and ation show characteristic patterns indicative of axial
textures (c. 650– 7508C, 2–4 kbar; Chichorro flattening with a vertical maximum shortening axis
et al. 2004). (Dı́az Azpı́roz 2006; Dı́az Azpı́roz et al. 2006).
The contact of the northern medium-grade unit Chocolate tablet boudinage of calc-silicate rock
with the core high-grade unit is marked by an layers interspersed within marbles supports this con-
increase of temperature (with local partial melting) clusion. Locally, shear zones have top-to-the-north
and shearing (local mylonitization). Low pressure and top-to-the-NE displacements. Therefore, the
and medium- to high-temperature metamorphism is evolution of the high-grade unit of the Aracena
characterized by the presence of sillimanite-garnet- metamorphic belt during the second deformation
cordierite-rich paragneisses (sillimanite zone) and phase is best described as due to the extensional
andalusite-rich mica schists (andalusite zone) collapse of a previously thickened crust.
(Carvalhosa 1999). The highest temperatures were Metamorphic isograds are telescoped at the
reached within the high-grade unit with diatexites northern boundary of the high-grade unit, sug-
showing mineral paragenesis consisting of plagio- gesting that this could be an original extensional
clase, K– feldspar, sillimanite and cordierite, and contact (Florindo & Quesada 1984). Extension pre-
associated with progressive partial melting with dated and partly coincided with the temperature
dehydration reactions of muscovite and biotite. A peak recorded in the metamorphic assemblages
regional well-developed late post-mylonitic replace- and with migmatization and intrusion of abundant
ment of biotite-andalusite-silimanite-cordierite by noritic and granitic plutons. Garnet-cordierite-
muscovite suggests slow cooling. The subsequent sillimanite-rich migmatites experienced maximum
decrease of temperature produced low-grade assem- metamorphic conditions estimated at 9008C and
blages with local replacement of biotite by chlorite 4 –6 kbar (El-Biad 2000; Dı́az Azpı́roz 2006).
(Pereira et al. 2007). Peak temperatures recorded in granulites with cor-
dierite, garnet, spinel, sillimanite and K –feldspar
Aracena massif reached 950–9758C at a pressure of 5 kbar (Patiño
Douce et al. 1997; El-Biad 2000). Migmatites
The Aracena massif (Bard 1969; Florido & Quesada have yielded a Visean age of c. 330–320 Ma
1984; Crespo-Blanc & Orozco 1991) is 2–10 km (Rb –Sr biotite, feldspar; Castro et al. 1999),
wide and extends for 120 km from Aroche to which is the supposed to be related to the exten-
Almadén de la Plata (Spain). This massif (Fig. 3) sional deformation. 40Ar/39Ar ages of the Ace-
has been divided into two domains (Castro et al. buches amphibolites range from c. 340–330 Ma
1996a, b, 1999; Dı́az Azpı́roz 2006; Dı́az Azpı́roz (Fonseca et al. 1993; Castro et al. 1999). The
et al. 2006): a southern domain that includes the eastern extension of the Aracena metamorphic
Acebuches MORB-derived amphibolites (Bard complex, the Almadén de la Plata core (Abalos
Fig. 3. Schematic representation of the structure and metamorphism distribution in the Aracena Massif (adapted from
Dı́az Azpı́roz et al. 2006).
et al. 1991), also shows a prograde metamorphic approximate age of extension, as given by meta-
evolution reaching in the highest-grade areas low morphic overgrowths on zircons in high-grade
pressure–high temperature mineral parageneses rocks, is Visean (c. 340 Ma, SHRIMP U –Pb;
with garnet, cordierite and sillimanite (peak of Ordóñez-Casado 1998). Three tectonic units that
metamorphism at 600 –700 8C and 3 –4 kbar). represents three structural levels were here defined
(Fig. 4) and attest to a prograde metamorphism
Lora del Rı́o massif under low-pressure conditions (Apraiz & Eguiluz
2002): the upper low-medium grade unit (Los
The Lora del Rı́o massif, which is approximately Miradores unit) in the north is represented by
12 km wide and 20 km long, crops out at the south- felsic tuffs, rhyolites, limestones, gabbros, slates
eastern end of the Ossa –Morena Zone, where it is and arkoses, micaschists and pelitic gneisses
partially covered by Cenozoic sediments of the Gua- (greenschist-upper amphibolite facies); the inter-
dalquivir basin (north of Seville, Spain; Fabriès mediate low-medium grade unit (Hueznar unit)
1963; Apraiz 1998; Apraiz & Eguiluz 2002). The with slates, greywackes, metarhyolites and
Fig. 4. Schematic representation of the structure and metamorphism distribution in the Lora del Rı́o Massif (adapted
from Apraiz & Eguiluz 2002).
micaschists (greenschist-upper amphibolite facies); high-grade unit suggests a prograde metamorphic

and the lower high-grade unit (Lora del Rı́o meta- evolution through the kyanite stability field with
morphic core) made of metatexites, amphibolites, the presence of garnet and kyanite (650–750 8C
amphibolitic gneisses, graphite-rich micaschists, and 10–12 kbar, Apraiz 1998); these relics of
diatexites, leucogranulites and anatectic granites higher pressure conditions suggest previous crustal
(upper amphibolite to granulite facies). Both upper thickening. Stretching lineations strike predomi-
unit and intermediate unit rocks are derived from nantly N– S, and kinematic criteria indicate a
Early–Middle Cambrian rocks, and the lower unit top-to-the-north normal shear sense for the northern
may also include Ediacaran protoliths. The meta- boundary of the high-grade unit. Isobaric or decom-
morphism of the upper structural unit increases pression heating is interpreted to be coeval with
downwards to high-grade amphibolite facies con- normal slip on the main detachments that bound
ditions close to the detachment that marks its limit the lower high-grade unit (c. 750– 8008C and 4–
with the intermediate unit. At the detachment, 6 kbar, with growth of cordierite and sillimanite;
partial melting is associated with the growth of silli- Apraiz 1998).
manite and K –feldspar, indicating that these rocks
reached temperatures of c. 650 –7508C and press- Tournaisian-Visean magmatism and
ures of c. 3– 5 kbar (Apraiz 1998). The intermediate marine basins
unit records maximum metamorphic conditions
estimated as c. 650 –700 8C and 3–5 kbar (Apraiz Diorites, tonalites, granodiorites and gabbros
1998). The occurrence of leucogranulites with (c. 350–320 Ma) represent I –type, high–K calc-
garnet, kyanite, K –feldspar and rutile in the lower alkaline magmatism (Santos et al. 1987, 1990;
Moita et al. 2005a; Tornos et al. 2005) associated reflect the influence of gravitational processes
with the Évora –Aracena –Lora del Rio meta- during sedimentation.
morphic belt. Less representative norites of bonini-
tic affinity (Castro et al. 1996a; El Hmidi 2000; Dı́az
Azpı́roz et al. 2006) and peraluminous biotite SHRIMP U– Th – Pb geochronology
granite are also present (Chichorro 2006; Pereira Sample preparation and analytical methods
et al. 2008).
The Beja mafic-ultramafic rocks (c. 350 Ma; Pin Five samples were collected from the Évora Massif
et al. 1999, 2008) are banded gabbros with tholeiitic for U– Pb dating of zircon: two paragneisses from
to calc-alkaline affinity, cumulate textures and mag- the Ediacaran Série Negra, BSC-1 (Universal Trans-
matic foliation (Santos et al. 1987). This intrusive verse Mercator co-ordinates– 29SNC707609) and
episode was followed by another magmatic event SEC-1 (29SNC742677); two of the orthogneisses
at c. 330 Ma (Jesus et al. 2006) that produced the from the Cambrian igneous-sedimentary com-
calc-alkaline Cuba gabbros and diorites. plexes, ALC-10 (29SNC734568) and VLV-3
Further to the north (present co-ordinates), pera- (29SNC855655) and one of an anatectic biotite
luminous I-type, calc-alkaline granites (Arraiolos granite ARL-6 (29SNC762664), probably formed
biotite granite; Chichorro 2006) and tonalites and by partial melting of the Série Negra sediments
gabbros (Hospitais tonalites and Alto de São and Cambrian igneous-sedimentary complexes.
Bento leucogranites, with a crystallization/cooling The zircon grains show complex structures, most
age of c. 320 Ma; Moita 2007) show intrusive consisting of a core surrounded by a much
relations with the Évora high-grade and medium- younger overgrowth. The ages of the cores and
grade metamorphic terrains. The mafic-intermediate their tectonic implications have been reported by
intrusions are thought to be derived from the mantle Pereira et al. (2008). Here we report the ages of
based on their isotopic signature (Santos et al. 1987; the zircon metamorphic/hydrothermal overgrowths
Pin et al. 1999; Moita et al. 2005a; Moita 2007), and their significance.
while the granites contain a mixture of mantle and Zircon grains were extracted, then mounted in
crustal components. epoxy resin with zircon standards SL13
The same chemical compositions and ages are (U ¼ 238 ppm) and TEMORA (206Pb*/238U ¼
preserved in Tournaisian– Visean (c. 350 – 0.06683), using procedures described by Pereira
335 Ma) volcanic rocks associated with marine et al. (2008). The polished mount was photographed
sedimentation in a relatively shallow-water plat- and imaged by SEM cathodoluminescence (CL) to
form environment (Santos et al. 1987; Chichorro document the internal growth zoning of the grains,
2006; Pereira et al. 2007). These marine basins, prior to SHRIMP analysis at the ANU using a pro-
dominated by detrital sequences with syn- cedure similar to that described by Williams &
sedimentary gravitational collapse structures, are Claesson (1987). Monazite recovered from VLV-3
preserved in the southwestern Ossa –Morena Zone was mounted separately with monazite standard
in Portugal (Pereira et al. 2006a) and record an WB.T.329 from Thompson Mine, Manitoba
early period of mixed siliciclastic-carbonate plat- (206Pb*/238U ¼ 0.3152, U 2100 ppm), and was
form sedimentation, which was transitional to pre- analysed using a procedure similar to that described
dominantly siliciclastic deposits. The Toca de by Williams et al. (1996). The Th-related isobar
Moura volcano-sedimentary complex exposed to which potentially interferes with 204Pb was
the north of the Beja gabbros consists of basalts reduced to insignificance by using moderate
(pillow-lavas and hyaloclastites), microdiorites, energy filtering.
andesites, rhyolites, felsic-intermediate tuffs and The plotted and tabulated analytical uncertain-
shales (Gonçalves 1985; Santos et al. 1987; ties are 1s precision estimates. Uncertainties in
Pereira et al. 2006a). To the north, the Cabrela the calculated mean ages are 95% confidence
volcano-sedimentary complex is exposed. A limits (ts, where t is the Student’s t multiplier)
siliciclastic-carbonate sequence (calcic-turbidites) and, for the mean 206Pb/238U ages, include the
near the base passes upwards into a predominantly uncertainty in the Pb/U calibration (c. 0.3–0.5%).
siliciclastic sequence with shales, greywackes Ages were calculated using the constants rec-
and interbedded conglomerates, with associated ommended by the IUGS Subcommission on Geo-
felsic tuffs, volcanic breccias, rhyodacites, rhyo- chronology (Steiger & Jäger 1977). Common Pb
lites and subordinate andesites (Ribeiro 1983; corrections assumed a model common Pb compo-
Chichorro 2006). sition appropriate to the age of each spot
Shales from both basins contain miospores of (Cumming & Richards 1975). Best estimates of
late Tournaisian to late Visean age (Pereira et al. the individual ages (Inferred age, Tables 1 & 2)
2006a). Slumps, growth-faults and olistoliths of were calculated from the radiogenic 206Pb/238U
limestone (of Fammenian and Fransnian ages) (common Pb correction based on 207Pb). Wherever
VARISCAN INTRA-OROGENIC EXTENSIONAL TECTONICS IN THE OSSA-MORENA ZONE
Table 1. Zircon SHRIMP data from the paragneisses SEC-1 and BSC-1
Apparent age (Ma)

204 204 208 208 206 207
Analysis Type Pb* U Th Th/ Pb/ Pb/ + f Pb*/ + Pb*/ + Pb*/ + Pb*/ + 208/ + 206/ + 207/ + Inf/ +
206 206 232 238 206
(ppm) (ppm) (ppm) U ppb Pb 206
Pb Pb Th U Pb 232 238 206 Age§
Paragneiss SEC-1
1.1 UZO 32 448 108 0.24 4 1.5E-04 6.5E-05 0.00266 0.0856 0.0033 0.02612 0.00113 0.07330 0.00108 0.05692 0.00146 521.1 22.3 456.0 6.5 488.6 57.6 455.5 6.5
2.1 UZO 25 523 2 0.00 5 5.3E-04 1.7E-04 0.00373 – – – – 0.05219 0.00065 0.05491 0.00154 – – 327.9 4.0 408.4 63.8 327.1 4.0
4.1 UZO 42 812 34 0.04 2 5.2E-05 2.9E-05 0.00095 0.0187 0.0023 0.02501 0.00309 0.05564 0.00050 0.05283 0.00176 499.4 61.1 349.0 3.1 321.5 77.3 349.3 3.1
5.1 UZO 15 277 28 0.10 6 4.3E-04 2.2E-04 0.00751 – – – – 0.05843 0.00186 0.05144 0.00279 – – 366.1 11.3 260.7 129.4 367.1 11.0
6.1 UZO 27 390 40 0.10 32 1.0E-03 1.9E-04 0.02304 – – – – 0.07356 0.00095 0.05573 0.00338 – – 457.6 5.7 441.7 141.0 457.8 5.9
8.1 UZO 21 349 36 0.10 16 8.9E-04 2.0E-04 0.01621 0.1327 0.0108 0.07448 0.00626 0.05847 0.00094 0.06022 0.00360 1451.9 118.1 366.3 5.8 611.6 134.9 363.5 5.6
9.1 UZO 30 546 8 0.02 25 1.6E-03 3.2E-04 0.01605 – – – – 0.05997 0.00055 0.05999 0.00204 – – 375.5 3.3 603.2 75.5 372.8 3.4
6.2 UZO 38 668 13 0.02 34 2.2E-03 3.8E-04 0.01715 – – – – 0.06228 0.00165 0.05607 0.00209 – – 389.5 10.0 455.3 85.1 388.7 10.0
2.2 UZO 47 958 9 0.01 11 1.4E-04 9.0E-05 0.00436 – – – – 0.05446 0.00034 0.05295 0.00127 – – 341.9 2.1 326.8 55.5 342.0 2.1
2.3 UZO 53 951 5 0.00 27 9.2E-04 2.5E-04 0.00973 – – – – 0.06160 0.00077 0.05427 0.00113 – – 385.3 4.7 382.2 47.5 385.4 4.7
4.2 UZO 13 297 7 0.02 3 2.1E-04 1.3E-04 0.00473 – – – – 0.04959 0.00086 0.05570 0.00236 – – 312.0 5.3 440.4 97.2 310.8 5.3
11.1 UZO 16 337 2 0.01 2 8.6E-05 1.5E-04 0.00224 – – – – 0.05147 0.00097 0.05236 0.00158 – – 323.5 5.9 301.0 70.1 323.8 6.0
Paragneiss BSC-1
1.1 UZO 11 224 4 0.02 4 9.1E-04 6.3E-04 0.00646 – – – – 0.05601 0.00060 0.05538 0.00200 – – 351.3 3.6 427.5 82.5 350.5 3.7
3.1 LUZO 7 153 5 0.03 13 3.1E-03 5.9E-04 0.03270 – – – – 0.05363 0.00164 0.05312 0.00382 – – 336.8 10.0 333.7 171.9 336.8 10.0
*, Radiogenic.
f, Correction for common Pb –Fraction of total 206Pb (206PbT) that is common 206Pb (206Pbc).
§
, Best estimated of the age of the analysed zircon. (See text for explanation.) Uncertainties one standard error. UZO, Unzoned overgrowth; LUZO, Light unzoned overgrowth.
223
224
Table 2. Zircon SHRIMP data from the orthogneisses ALC-10 and VLV-3
Apparent age (Ma)

204 204 208 208 206 207
Analysis Type Pb* U Th Th/ Pb/ Pb/ + f Pb*/ + Pb*/ + Pb*/ + Pb*/ + 208/ + 206/ + 207/ + Inf +
206 206 232 238 206
Orthogneiss ALC-10
1.1 UZ(r) 19 376 15 0.04 4 3.4E-05 3.6E-05 0.00391 – – – – 0.05425 0.00078 0.05179 0.00160 – – 340.6 4.7 276.2 72.4 341.2 4.8
2.1 UZ(r) 14 272 10 0.04 2 2.6E-04 1.5E-04 0.00339 – – – – 0.05457 0.00111 0.05524 0.00194 – – 342.5 6.8 421.8 80.5 341.7 6.8
3.1 UZ(r) 27 535 10 0.02 4 8.1E-05 7.7E-05 0.00258 – – – – 0.05455 0.00049 0.05206 0.00141 – – 342.4 3.0 288.2 63.2 342.9 3.0
4.1 UZO 33 665 28 0.04 3 1.5E-04 1.4E-04 0.00170 – – – – 0.05400 0.00052 0.05385 0.00129 – – 339.1 3.2 364.6 54.8 338.8 3.2
5.1 LUZ(r) 11 230 8 0.03 4 4.5E-04 2.5E-04 0.00658 – – – – 0.05244 0.00084 0.05174 0.00263 – – 329.5 5.2 274.0 120.9 330.0 5.2
M. F. PEREIRA ET AL.
Orthogneiss VLV
1.1 UZO 87 1374 334 0.24 337 4.1E-03 2.4E-04 0.07419 0.0860 0.0099 0.02287 0.00266 0.06471 0.00066 0.05744 0.00431 457.1 52.6 404.2 4.0 508.3 174.0 403.0 3.7
2.1 UZO 53 722 207 0.29 46 7.1E-04 1.4E-04 0.01760 – – – – 0.07462 0.00072 0.05437 0.00208 – – 464.0 4.3 386.6 88.3 465.0 4.4
3.1 UZ 97 1278 351 0.27 6 7.1E-05 4.2E-05 0.00127 0.0841 0.0023 0.02368 0.00068 0.07737 0.00064 0.05733 0.00108 473.1 13.4 480.4 3.8 504.3 42.0 480.0 3.8
4.1 UZ 54 666 219 0.33 1 2.1E-05 1.8E-05 0.00038 0.1066 0.0019 0.02637 0.00068 0.08134 0.00042 0.05837 0.00084 526.0 13.5 504.1 2.5 543.7 31.7 503.4 2.5
6.1 UZO 193 3343 470 0.14 1907 9.7E-03 2.2E-04 0.16493 – – – – 0.06121 0.00164 0.05508 0.00295 – – 383.0 10.0 415.6 124.5 382.7 9.9
7.1 UZO 66 1317 89 0.07 19 3.6E-04 8.1E-05 0.00548 – – – – 0.05460 0.00045 0.05277 0.00127 – – 342.7 2.7 318.7 55.8 342.9 2.8
8.1 UZO 70 1293 100 0.08 14 2.6E-04 7.2E-05 0.00395 – – – – 0.05873 0.00051 0.05502 0.00151 – – 367.9 3.1 413.2 62.4 367.4 3.2
9.1 UZO 48 984 51 0.05 3 4.0E-05 3.6E-05 0.00105 – – – – 0.05340 0.00035 0.05405 0.00110 – – 335.4 2.2 373.3 46.3 335.0 2.2
10.1 UZO 70 984 196 0.20 16 1.3E-04 1.0E-04 0.00477 – – – – 0.07355 0.00068 0.05511 0.00148 – – 457.5 4.1 416.7 61.3 458.1 4.1
11.1 UZO 56 1059 105 0.10 60 9.9E-04 1.9E-04 0.02080 – – – – 0.05690 0.00040 0.05321 0.00137 – – 356.8 2.5 337.9 59.5 357.0 2.5
12.1 CZ 64 850 244 0.29 92 1.4E-03 9.9E-05 0.02900 – – – – 0.07634 0.00047 0.05706 0.00144 – – 474.2 2.8 493.9 56.6 473.9 2.8
13.1 UZC 90 950 926 0.97 14 2.2E-04 1.1E-04 0.00391 0.3011 0.0048 0.02509 0.00044 0.08125 0.00047 0.05624 0.00184 501.0 8.6 503.6 2.8 461.8 74.1 504.2 2.7
13.2 LCZC 9 102 60 0.59 43 5.0E-03 6.1E-04 0.09968 – – – – 0.08089 0.00131 0.05728 0.00824 – – 501.5 7.8 502.2 352.4 501.4 7.4
14.1 UZC 85 918 819 0.89 11 1.7E-04 3.2E-05 0.00305 0.2647 0.0023 0.02408 0.00034 0.08112 0.00029 0.05715 0.00072 481.0 6.7 502.8 1.7 497.4 28.1 502.9 1.7
14.2 CZ 61 639 698 1.09 11 2.5E-04 7.4E-05 0.00455 0.3287 0.0040 0.02411 0.00036 0.08017 0.00054 0.05627 0.00133 481.5 7.0 497.2 3.2 462.9 53.3 497.7 3.2
15.1 UZr 37 460 105 0.23 2 5.7E-05 3.7E-05 0.00102 0.0710 0.0021 0.02591 0.00080 0.08315 0.00056 0.05741 0.00171 517.1 15.8 514.9 3.4 507.4 66.8 515.0 3.5
16.1 UZC 86 908 797 0.88 10 1.5E-04 3.1E-05 0.00276 0.2756 0.0052 0.02598 0.00053 0.08278 0.00058 0.05761 0.00075 518.3 10.4 512.7 3.5 515.1 29.0 512.7 3.5
*, Radiogenic.
§
, Best estimated of the age of the analysed zircon. (See text for explanation.) Uncertainties one standard error. UZO, Unzoned overgrowth; UZ, Unzoned; CZ, Concentric zoned; UZC, Unzoned core; LCZC, Light
concentric zoned core; LUZ, Light unzoned; (r), Recrystallization front.
possible, the analyses tabulated and plotted were that might help explain the scatter as due to either
corrected for common Pb using 208Pb, meaning Pb loss or inheritance of radiogenic Pb. Interpret-
that radiogenic 208Pb/206Pb and 208Pb/232Th were ation of the analyses must remain inconclusive. If
not independently determined. Some analyses Pb loss was the dominant effect, the primary age
(those with U contents .2500 ppm) required a of the overgrowths is c. 380 Ma; if inheritance of
small matrix correction to the Pb/U ratios, which radiogenic Pb predominated, the overgrowths
was applied using the relationship determined by might be as young as c. 320 Ma. Given that unsup-
Williams & Hergt (2001). ported radiogenic Pb is relatively rare in zircon
(Williams et al. 1984), an age of c. 380 Ma is con-
SHRIMP U – Th– Pb results sidered to be the more likely.
BSC-1 was a medium-grained paragneiss
The results of the zircon U –Th –Pb analyses are (Ediacaran sedimentary protolith) consisting of
listed in Tables 1–3 (monazite in Table 4) and CL quartz, biotite, muscovite, plagioclase and feldspar
images of representative zircons in Figures 6 and 7. aligned parallel to the regional foliation. Albite por-
The data are plotted on a concordia diagrams in phyroblasts included aligned and folded graphitic
Figure 8, and a compilation of all inferred ages is inclusions, and had tails of new-grown oligoclase.
illustrated as a relative probability density distri- Blades of muscovite and biotite were oriented paral-
bution (Dodson et al. 1988) in Figure 9. The lel to dynamically recrystallized quartz ribbons that
zircon grains are complexly structured, most con- defined the foliation. Only 11 zircon grains were
sisting of a core surrounded by a much younger recovered from the sample. Most of those consisted
overgrowth. The ages of the cores and their tectonic of a detrital zircon core partly overgrown by a
implications have been reported by Pereira et al. thin (,10 mm), discontinuous, nodular layer of
(2008). Here we report the ages of the zircon moderately luminescent, irregularly zoned zircon
metamorphic/hydrothermal overgrowths and their (Fig. 5). Three overgrowths were thick enough to
significance. analyse with a 10 mm probe. Two of those had
very similar compositions, moderate U contents
Série Negra paragneisses (224 and 153 ppm) and very low Th/U (0.02 and
0.03), consistent with growth during metamorph-
SEC-1 (Fig. 6) was a foliated paragneiss (Ediacaran ism. The 206Pb/238U ages of these overgrowths
sedimentary protolith) with a mosaic texture con- were also very similar, c. 340 Ma. Their weighted
sisting of alternating mm-wide bands of dynami- mean age, 348.9 + 3.5 (s) Ma, is dominated by
cally recrystallized quartz and feldspar, ribbons of the more precise analysis. With only two analyses,
polygonal quartz and bands with biotite, quartz, the 95% confidence limits on this result are very
feldspar, sillimanite, zircon, apatite, rutile and large, c. 45 Ma. In contrast, the third overgrowth
opaque minerals (mainly graphite). Most of the det- was much thicker (c. 20 mm), strongly luminescent
rital zircon grains recovered were partially over- and faceted. It had much lower U (9 ppm) and
grown by a very thin (,20 mm) nodular layer of higher Th/U. It also had a much higher age,
weakly luminescent zircon commonly rich in 748 + 48 (s) Ma. This is interpreted as an over-
inclusions that have yet to be identified (Fig. 5). growth that was present on a detrital grain, and
Only the thickest of the overgrowths were accessi- indeed close inspection showed that it was over-
ble to analysis using the 10 mm diameter SHRIMP grown in turn by tiny (c. 2 mm) nodules of zircon
primary probe and accurate targeting to avoid resembling the young overgrowths. The best esti-
overlap with the host zircons was very difficult. mate of the age of metamorphism of paragneiss
The twelve areas analysed, with a single exception BSC-1 is 350 + 45 (95% c.l.) Ma.
(analysis 1.1) had moderate U contents (280 –
960 ppm) and low to very low Th/U (0.1–0.004), Cambrian orthogneisses
consistent with growth during metamorphism
(Williams & Claesson 1987; Heaman et al. 1990). Orthogneiss ALC-10 (Fig. 7) was an igneous rock
With two exceptions (8.1, 9.1) the analyses were deformed under upper amphibolite facies meta-
concordant within analytical uncertainty, but there morphic conditions. It had a foliation defined by
was a very large range in their radiogenic aligned individual biotite and amphibole crystals
206
Pb/238U apparent ages. Two analyses (1.1, 6.1), and an alternation of continuous layers of biotite
both of which probably overlapped the old detrital aggregates with bands of heterogeneously dynami-
host grain, gave Ordovician (c. 455 Ma) apparent cally recrystallized quartz and feldspar. Rare fibro-
ages that are considered to have no geological sig- lite sillimanite was present. CL imaging revealed
nificance. The remaining 206Pb/238U ages ranged that the great majority of the zircon grains had
from c. 390 –310 Ma, and there was no clustering weak concentric growth zoning consistent with pre-
towards either the upper or lower end of the range cipitation from magma of felsic to intermediate
Table 3. Zircon SHRIMP data from the biotite granite ARL-6
226
Apparent age (Ma)
204 204 208 208 206 207
Analysis Type Pb* U Th Th/ Pb/ Pb/ + f Pb*/ + Pb*/ + Pb*/ + Pb*/ + 208/ + 206/ + 207/ + Inf +
206 206 232 238 206
Biotite granite ARL-6

1.1 CZIO 37 682 212 0.31 1 4.4E-05 1.9E-05 0.00080 0.0946 0.0022 0.01659 0.00041 0.05453 0.00037 0.05305 0.00074 332.6 8.1 342.3 2.3 330.8 32.0 342.2 2.3
2.1 CZ 14 257 123 0.48 2 1.5E-04 6.1E-05 0.00277 0.1448 0.0049 0.01608 0.00056 0.05316 0.00041 0.04837 0.00154 322.4 11.2 333.9 2.5 117.2 76.7 334.8 2.5
3.1 CZIO 17 301 166 0.55 1 9.5E-05 8.0E-05 0.00173 0.1705 0.0042 0.01678 0.00044 0.05427 0.00048 0.05258 0.00155 336.4 8.8 340.7 2.9 310.5 68.5 341.0 2.9
4.1 DBZ 108 1684 1671 0.99 1 1.7E-05 1.3E-05 0.00032 0.3072 0.0022 0.01698 0.00016 0.05483 0.00026 0.05351 0.00046 340.3 3.1 344.1 1.6 350.4 19.7 344.0 1.6
5.1 DBZ 171 2777 2295 0.83 4 2.8E-05 2.2E-05 0.00051 0.2487 0.0019 0.01650 0.00015 0.05481 0.00025 0.05360 0.00049 330.7 3.1 344.0 1.5 354.2 21.0 343.9 1.5
7.1 DBZ 165 2741 2684 0.98 102 8.1E-04 4.5E-05 0.01482 0.2750 0.0031 0.01481 0.00019 0.05273 0.00026 0.05365 0.00082 297.2 3.8 331.2 1.6 356.2 34.9 331.0 1.6
8.2 CZIO 26 522 40 0.08 0 7.4E-06 7.9E-06 0.00014 0.0248 0.0012 0.01742 0.00084 0.05319 0.00035 0.05388 0.00083 349.1 16.7 334.1 2.2 365.9 35.0 333.8 2.2
13.1 UZO 217 4281 335 0.08 34 1.6E-04 1.9E-05 0.00299 0.0043 0.0008 0.00307 0.00060 0.05573 0.00017 0.05332 0.00046 62.0 12.1 349.6 1.0 342.6 19.8 349.8 1.1
22.1 CZO 17 293 163 0.56 2 1.5E-04 1.0E-04 0.00269 0.1711 0.0058 0.01648 0.00058 0.05367 0.00044 0.05347 0.00185 330.3 11.5 337.0 2.7 348.7 80.3 336.9 2.6
23.1 CZC 9 162 87 0.54 2 2.5E-04 1.1E-04 0.00463 0.1636 0.0059 0.01633 0.00071 0.05371 0.00106 0.05031 0.00229 327.4 14.2 337.3 6.5 209.1 108.9 338.5 6.5
25.1 CZC 33 633 165 0.26 0 3.4E-06 5.8E-06 0.00006 0.0799 0.0025 0.01623 0.00054 0.05294 0.00052 0.05363 0.00096 325.4 10.8 332.6 3.2 355.7 40.8 332.3 3.2
M. F. PEREIRA ET AL.
*, Radiogenic.
f, Correction for common Pb–Fraction of total 206Pb (206PbT) that is common 206Pb (206Pbc).
§
, Best estimated of the age of the analysed zircon. (See text for explanation.) Uncertainties one standard error. CZ, Concentric zoned; CZO, Concentric zoned overgrowth; CZIO, Concentric zoned intermediary
overgrowth; UZO, Unzoned overgrowth; CZC, Concentric zoned core; DBZ, Dark banded zoned.
Table 4. Monazite SHRIMP data from the orthogneiss VLV-3
Apparent age (Ma)

204 208 208 206 207
Analysis Pb* U Th Th/U Pb/ + f Pb*/ + Pb*/ + Pb*/ + Pb*/ + 208/ + 206/ + 207/ + Inf +
206 206 232 238 206
(ppm) (ppm) (ppm) Pb 206
1.1 703 2020 44061 21.8 6.79E-05 9.44E-05 0.0011 6.696 0.093 0.01591 0.00035 0.05183 0.00082 0.0533 0.0019 319.1 7.1 325.8 5.0 339.5 80.9 325.6 5.0
2.1 525 1864 31780 17.0 4.46E-05 3.64E-05 0.0007 5.273 0.057 0.0159 0.00037 0.05142 0.00093 0.0513 0.0012 318.9 7.4 323.2 5.7 255.3 56.3 323.8 5.7
3.1 437 4281 17145 4.0 2.69E-05 2.31E-05 0.0004 1.256 0.013 0.01606 0.00036 0.05122 0.00086 0.0519 0.0013 322.0 7.1 322.0 5.3 281.3 57.7 322.3 5.3
4.1 316 1057 19404 18.4 7.01E-05 7.02E-05 0.0011 5.730 0.102 0.01587 0.00047 0.05082 0.00109 0.0522 0.0018 318.2 9.4 319.6 6.7 294.4 80.8 319.8 6.7
5.1 411 1263 25473 20.2 2.00E-05 2.00E-05 0.0003 6.275 0.079 0.01594 0.00037 0.05122 0.00092 0.0521 0.0013 319.6 7.5 322.0 5.6 287.9 59.9 322.3 5.7
6.1 339 1043 21430 20.6 1.33E-04 7.55E-05 0.0021 6.291 0.084 0.01564 0.00037 0.05109 0.00088 0.0513 0.0020 313.7 7.5 321.2 5.4 252.1 94.5 321.7 5.4
*, Radiogenic.
§
, Best estimated of the age of the analysed zircon. (See text for explanation.) Uncertainties one standard error.
Fig. 5. I–II, III–IV and V– VI are cross sections through western (Èvora Massif), middle (Aracena Massif) and eastern
(Lora del Rio Massif) sectors of the investigated metamorphic belt of the SW Iberian Massif.
composition. Dating of such grains by Chichorro and very low Th/U (0.02–0.04). Such textureless
et al. (2008), showed that the igneous protolith areas are a common feature of zircon from high-
was of Cambrian age. Many of the zircon grains grade metaigneous rocks, and have been interpreted
had patches, embayments and overgrowths of as the product of thermally-activated solid-state
more luminescent zircon with no visible internal recrystallization (Hoskin & Black 2000). In the
texture. In contrast to the zoned portions of the present case, the distinctively low Th/U in the
grains, which had moderate to high Th contents recrystallized areas indicates that one of the trace
(150 –600 ppm) and moderate to high Th/U (0.2– elements preferentially expelled during the recrys-
1.2), five luminescent areas analysed for the tallization was Th. The five U –Th –Pb isotopic ana-
present study had low to very low Th (28–8 ppm) lyses all yielded concordant Mississippian apparent
Fig. 6. CL images of representative zircons with analytical sites and their resulting ages indicated, from two
samples of Ediacaran Serie Negra metasediments from the Evora Massif. Analysis spots and ages are listed in Table 1,
first and last columns, respectively.
ages, indicating that another trace element expelled as medium- to coarse-grained (100–200 mm diam-
during the recrystallization was radiogenic Pb. In eter) euhedral to subhedral prisms with simple oscil-
fact, all five 206Pb/238U ages were the same within latory growth zoning, consistent with precipitation
analytical uncertainty, with a well-defined weighted from the melt phase of a magma. Many of the
mean age of 339.7 + 5.5 (95% c.l.) Ma. This is the grains, however, had complex textures, including
best estimate of the age of the recrystallization and convolute zoning and textureless patches, in this
the metamorphism that caused it. case more weakly luminescent than the remainder
VLV-3 was a fine- to medium-grained orthog- of the grain. These features are indicative of perva-
neiss consisting of aligned individual biotite and sive solid-state recrystallization. Some grains also
amphibole crystals, and an alternation of continuous had very distinct, nodular, weakly luminescent,
layers of biotite aggregates with bands of heteroge- weakly zoned or unzoned overgrowths.
neously dynamically recrystallized quartz and feld- U –Th –Pb isotopic analyses of 10 igneous grains
spar. Apatite, zircon and monazite were accessory showed a range of moderate to high U contents
phases. As in ALC-10, most of the zircon occurred (100 –1280 ppm) and moderate to high Th/U
Fig. 7. CL images of representative zircons with analytical sites and their resulting ages indicated, from two Cambrian
metaigneous rocks and Mississipian biotite granite from the Évora Massif. Analysis spots and ages are listed in
Tables 2 and 3, first and last columns, respectively.
(0.22–1.1). The U –Pb isotopic compositions were single felsic magma. These grains also all have the
all concordant within analytical uncertainty, but same 206Pb/238U age within analytical uncertainty
there was a small range in radiogenic 206Pb/238U (MSWD ¼ 2.1), 337.1 + 4.3 (95% c.l.) Ma. This
ages (c. 515– 475 Ma). Most of the scatter is due is the best estimate of the crystallization age of
to two low values (analyses 3.1 and 12.1). To elim- the granite.
inate the scatter entirely requires the omission of
two more analyses, one high and one low; Discussion
however this is hard to justify. A more objective
result is to pool all analyses in the upper group, The significance of different zircon growths
giving a weighted mean 206Pb/238U age of and distribution of ages
504.9 + 5.4 (95% c.l.) Ma, the uncertainty taking
the minor excess scatter (MSWD ¼ 3.2) into The term ‘metamorphic zircon’ is applied to zircon
account. The igneous protolith of the orthogneiss formed under metamorphic conditions by a range of
was of Cambrian age (Chichorro et al. 2008). different postulated processes such as: (1) sub-
Isotopic analyses of 8 overgrowths and recrystal- solidus nucleation and crystallization by diffusion
lized areas showed a narrower range in U content of Zr and Si released by metamorphic breakdown
(720–3340 ppm) and Th/U (0.05– 0.28), but a reactions of major silicates or other Zr-bearing min-
much larger range in 206Pb/238U age (c. 465 – erals (Fraser et al. 1997); (2) precipitation from
335 Ma). The four ‘oldest’ areas had relatively Zr-saturated silicate anatectic melts (Vavra et al.
high Th/U (0.14–0.28) and the four ‘youngest’ 1996; Williams et al. 1996; Watson 1996), direct
had low Th/U (0.05 –0.10), consistent with the crystallization from a zircon-saturated aqueous
observation in sample ALC-10 that recrystallization fluid (hydrothermal fluids) which can be aqueous
was accompanied by the preferential expulsion of metamorphic fluids (Williams et al. 1996), fluids
Th and Pb. The difference in the case of the zircon evolved from a magma during the final stages of
from VLV-3 is that the expulsion of radiogenic Pb crystallization (Hoskin 2005) or low temperature
was incomplete and a partial memory of the fluids, which precipitate authigenic zircon (Arm-
primary age was retained. The best estimate of the strong et al. 1995); and (3) solid state annealing
metamorphic age is provided by the isotopic ana- (Vavra et al. 1996) or recrystallization of protolith
lyses of six monazite grains, all of which gave the zircon (Black et al. 1986; Friend & Kinny 1995;
same concordant age within analytical uncertainty, Pan 1997; Bowring & Williams 1999).
322.8 + 6.8 (95% c.l.) Ma. All the external overgrowths and recrystallized
areas that were dated for the present study,
Mississipian biotite granite whether on detrital zircon grains from Ediacaran
paragneisses, on igneous zircon from Cambrian
ARL-6 was a medium-grained, weakly foliated ana- orthogneisses or on inherited zircon from an anatec-
tectic granite composed of plagioclase, K – tic granite, have very low Th/U (,0.1). This feature
feldspar, quartz and biotite, with zircon as the main is typical of the zircon precipitated during the meta-
accessory phase. The zircon population was mor- morphism or partial melting of a peraluminous rock
phologically very heterogeneous. In addition to (Williams & Claesson 1987; Heaman et al. 1990;
euhedral grains with simple oscillatory growth Williams 2001). There are significant differences,
zoning, there were large, CL-dark, prismatic grains however, between the zircon precipitated in these
with indistinct banded zoning and a large number different environments. Seven analyses (4 from
of grains composed of a core surrounded by a orthogneiss VLV-3, 1 from orthogneiss ALC-10
weakly zoned or unzoned overgrowth. SHRIMP and 2 from biotite granite ARL-6) have relatively
U– Pb analyses of the cores (Pereira et al. 2008) high U contents (522–4281 ppm) consistent with
showed that they were a Neoproterozoic– Cambrian zircon precipitated from trace element rich partial
inherited component. In contrast, the grains without melts. These overgrowths are very well defined,
cores have very similar apparent ages (c. 350 – generally unzoned and pyramidal and tend to have
330 Ma). Nevertheless, there is a significant range developed preferentially on rapidly growing pyra-
which correlates with grain morphology and com- midal crystal faces.
position (Fig. 8). Three dark banded grains and a The two analysed overgrowths on old inherited
dark unzoned overgrowth had high U contents zircon cores from the biotite granite probably rep-
(1685–4280 ppm) and a wide range of Th/U resent the same high temperature processes associ-
(0.07–0.99). These contribute most of the scatter. ated with mylonitization recorded in the Serie
In contrast, the simply zoned euhedral grains have Negra paragneisses. These thick new additions
lower and more uniform U (160 –680 ppm) and, resemble the overgrowths developed on detrital
with one exception, more uniform Th/U (0.26– zircon during the high temperature/low pressure
0.55), consistent with having crystallized from a metamorphism of the Cooma Complex in eastern
Australia that culminated in the production of an (5 analyses) and 337.1 + 4.3 Ma for the emplacement
anatectic granite (Williams 2001). of biotite granite ARL-6 (7 analyses). These
The thin, discontinuous, nodular, inclusion rich results are slightly but significantly older than the
(spongy) overgrowths developed on the detrital crystallization age of the monazite from orthogneiss
zircon grains from the Serie Negra paragneisses VLV-3, 322.8 + 6.8 Ma. This Middle– Upper
are quite different, and do not appear to be associ- Mississipian age overlaps the ages obtained for the
ated with high-grade partial melting. The morpho- Hospitais tonalites (323.5 + 5.2 Ma; Ar– Ar amphi-
logy of the inclusion-rich overgrowths suggests bole) and the Alto de São Bento leucogranites
precipitation from intergranular Zr saturated fluid (325 + 9 Ma; Rb–Sr whole rock – biotite) inter-
films. Such fluid films, with low Th contents, might preted to be related with crystallization/cooling
be produced during the early stages of partial (Moita 2007).
melting related to prograde high temperature/low The results of this study are consistent with a dia-
pressure metamorphism, or might reflect the chronous and variable spatial distribution of mul-
inhibition of partial melting, and consequently the tiple zircon growth and recrystallization related to
inhibition of new zircon growth, by the relatively an extended thermal history (discrete thermal
low Al2O3 contents (16% , Al2O3 , 18%) of pulses) concentrated in the Early Carboniferous,
these immature Ediacaran metasediments. This within the interval c. 350–335 Ma (contains 56%
hypothesis suggests a relationship between high- of the analyses and is statistically significant).
grade solid –solid metamorphic reactions and con- These thermal events are associated with the devel-
sequent melting conditions. An alternative is that opment of ductile deformation (widespread myloni-
the new zircon growth was linked to hydrothermal tic fabrics) related to transcurrent movements.
fluid flow driven by temperature gradients around Finally, the obtained ages of 388– 367 Ma from
thermal anomalies (Buick & Cartwright 1994). sample SEC-1 seem to be related with metamorphic
Fluids would have been generated in the Serie overgrowths. Apparent ages of 368 + 21 Ma (2s,
Negra paragneisses by the metamorphic biotite MSWD ¼ 1.6) were also obtained for the Alcáçovas
dehydration and the melting reaction that produced felsic gneisses and interpreted to represent an early
peritetic cordierite. Variscan metamorphic overprinting (Cordani et al.
Local retrogression of the metamorphic assem- 2006). Similar Middle–Upper Devonian ages were
blages (also associated with mylonitization) is found in eclogites and blueschist from Safira
indicated by intense alteration of cordierite and and Alvito (371 + 17 Ma; Sm –Nd, whole-rock –
the replacement of biotite by chlorite (Ribeiro garnet and Ar –Ar, glaucophane; Moita et al.
1994; Chichorro 2006). The presented evidence 2005b).
suggests that hydrothermal fluids may have played
a major role in the precipitation of the new zircon.
As hydrothermal zircon can be used to date the The significance of Variscan
infilling fluid events and the consequent metaso- intra-continental extensional tectonics
matic interactions between water and rock (Hoskin in the SW Iberian Massif
2005), the results obtained here suggest that in the
Évora Massif, the transport and deposition of Several tectonic models have been proposed to
metals in this gold province was probably concomi- explain the geological evolution of the Variscan
tant with high-grade metamorphic activity, with a orogeny in the SW Iberian Massif. Bard et al.
maximum remobilization around the age of (1973) proposed a geological scheme with a north-
340 Ma (Visean). directed (present-day co-ordinates) subduction
Four analyses from the orthogneisses were made zone, marked by the exposure of the Ossa–Morena
on cross-cutting mantles with morphologies typical and South Portuguese suture zone. The hypothesis
of thermally activated solid-state recrystallization of a SW Iberian Massif suture zone was reinforced
(Black et al. 1986; Varva et al. 1996; Hoskin & by the identification of MORB-derived amphibolites
Black 2000; Bowring & Williams 1999). The along this boundary (unknown age Beja–Acebuches
observed difference in Th/U from the igneous amphibolites; Bard & Moine 1979; Dupuy et al.
grains indicates that the high-grade conditions per- 1979; Quesada 1991; Crespo-Blanc & Orozco
sisted for an extended period of time. 1991; Fonseca & Ribeiro 1993; Quesada et al.
The distribution of 206Pb/238U ages is illustrated on 1994; Castro et al. 1996b).
a relative probability diagram (Dodson et al. 1988) in The widespread Tournaisian (c. 350 Ma)
Figure 9. The main peak at 342 Ma contains nearly calc-alkaline igneous rocks from the southwestern
32% of the analyses and is statistically significant. Ossa –Morena, which were also interpreted to
This is the principal metamorphic age obtained be arc-related (Beja gabbros and Toca da Moura–
from several of the samples; 339.7 + 5.5 Ma for the Odivelas volcanics; Santos et al. 1987, 1990),
metamorphic imprint on orthogneiss ALC-10 favoured this plate convergence tectonic setting.
A more complex structural scenario was proposed

by Araújo et al. (2005) who postulated the existence
of obducted oceanic crust within the Ossa –Morena
zone. These ophiolitic nappes, composed of MORB-
derived amphibolites and serpentinized peridotites,
occur to the north of the supposed suture zone and
magmatic arc. According to these authors, such dis-
membered oceanic rocks were transported on the
top of, and/or imbricated in, a tectonic mélange
(Moura phyllonitic accretionary complex). This
contractional deformation was, by their interpret-
ation, responsible for the exhumation of high- to
medium-pressure rocks (Safira –Viana do Alentejo
eclogites and Moura blueschists; Araújo et al.
2005; Ribeiro et al. 2007) during the Middle–Late
Devonian (c. 390 –370 Ma). Meanwhile, Rosas et
al. (2008) admit extension to explain the exhuma-
tion of high-pressure rocks after the earlier episode
of contraction deformation.
A major feature of the SW Ossa –Morena Zone is
the extensive record of a high-grade thermal event
with a prograde history that reached temperatures
of crustal anatexis under low-pressure conditions
(Abalos et al. 1991; Eguiluz et al. 2000; Dı́az
Azpı́roz 2006; Dı́az Azpı́roz et al. 2006). This high-
temperature/low-pressure metamorphism locally
overprinted high- to medium-pressure mineral
assemblages (Apraiz 1998; Apraiz & Eguiluz
2002; Chichorro 2006; Pereira et al. 2007) observed
in small outcrops (metre-scale boudins) with as yet
undeciphered relationships to their deformed
host rocks.
The Évora –Aracena –Lora del Rio metamorphic
belt is a discontinuous linear metamorphic zone
characterized by strong ductile shearing associa-
ted with high-medium-temperature/low-pressure
mineral assemblages with characteristic isobarically
cooled P –T paths (Fig. 9). The high-grade meta-
morphism appears to be related to thermal relax-
ation, extensional unroofing and rapid exhumation
due to the activity of major transcurrent faults
(c. 350 –335 Ma). Structural elements formed Fig. 8. Modified Tera –Wasserburg concordia diagrams
during ductile conditions show the progressive char- for zircon overgrowths from: Série Negra paragneiss,
Cambrian orthogneisses and Mississipian biotite granite
acter of shear deformation (strong mylonitization) at
of the Évora Massif.
higher and lower temperatures, with orogen-parallel
tectonic transport.
Our data emphasise the existence of an import- pressure rocks surrounded by a strong mylonitic
ant Tournaisian –Visean intra-orogenic extension fabric were probably rapidly exhumed along these
(transtension) of continental lithosphere in the SW major structures that gave place to uplift of meta-
Iberian Massif, with strong across-strike thermal morphic complexes and melt generation and intru-
gradients. sion. The dynamics associated with mantle
A strong heat source probably induced partial thermal perturbations and the consequent gravity
melting in the upper mantle and hot material collapse of the crust still need to be identified.
moved upward into shallow levels where melting The thermally anomaly responsible for the
was promoted in the continental crust. The strong Mississippian extensional collapse has been attribu-
ductile deformation is attributed to the influence of ted to astenospheric up-welling apparently in con-
transcurrent shear zones, which in places penetrated nection with a lithosphere slab-window (Castro
down to deep crustal levels. The high-medium et al. 1996a, b; Dı́az Azpı́roz et al. 2006). Another
Fig. 9. (a) Schematic reconstruction of the geodynamic scenario of the SW Iberian Massif in the Visean (c. 342 Ma)
before the shortening caused by the dominant Variscan wrench regime (the final result was the steepening of
earlier structures as shown in cross-sections of Figs 2, 3 & 4); (b) Summary of the P –T – t path for the Évora– Aracena–
Lora del Rı́o metamorphic belt (data compiled from Apraiz & Eguiluz 2002; Moita et al. 2005b; Dı́az Azpı́roz et al.
2006; Pereira et al. 2007); (c) Relative probability density distribution with compilation of all inferred ages obtained
from zircon overgrowths (time scale of Gradstein et al. 2004).
interpretation proposes that the heat source to the Gondwana continental margin involved in
trigger lithosphere extension was a kilometre-scale the Variscan collision (Martinez Catalán et al.
layered mid-crustal mafic intrusion probably associ- 1996). They consist of unknown age metasediments
ated with a sub-crustal thermal perturbation akin to a (Cambrian –Ordovician) intruded by early Ordovi-
mantle plume (Simancas et al. 2003). cian felsic and mafic igneous rocks that underwent
Finally, mantle convection slowed down, ther- subduction and high-pressure metamorphism in
mal gradients attenuated, and extension probably the course of early Variscan events (Martinez
stopped at c. 320 Ma. Catalán et al. 1996, 2007).
Comparison of our dating results with those The Malpica –Tui Complex lithostratigraphy
obtained for the NW Iberian Massif indicate com- and structure looks like that of the Évora Massif.
mon geological characteristics. An upper Igneous (mafic dominated)– Sedimentary
The NW Iberian Massif (Galicia-Trás-os- Complex composed of MORB-derived amphibo-
Montes Zone; Martinez Catalan et al. 2003) lites and phyllites with interbedded carbonates and
preserves relicts of the Variscan suture related cherts overlies a probably older Igneous (felsic
with the convergence between Gondwana and dominated) –Sedimentary Complex with calc-
Laurussia (Martinez Catalán et al. 2007). Here alkaline felsic orthogneisses, schists and para-
the allochthonous complexes of Cabo Ortegal, gneisses. Contrary to the Évora Massif here,
Órdones, Malpica-Tui, Bragança and Morais evidence for early Variscan subduction seems to
constitute a nappe pile of different tectonic units. be well preserved by variably retrograde high-
The basal units are considered to derive from pressure metamorphic fabrics in blueschists and
eclogites (Rodriguez 2005). The age of c. 370 Ma isotopic constraints on provenance and post-
obtained for the Évora Massif eclogites and blues- depositional alteration of the Witwatersrand Super-
chists (Sm–Nd, garnet and Ar– Ar, amphibole, group, Extended Abstracts, Centennial Geocongress
Moita et al. 2005b) coincide with the eclogite-facies (1995). Geological Society of South Africa, 1086.
A ZOR , A., R UBATTO , D., S IMANCAS , J. F., G ONZÁLEZ
metamorphic event recorded in the Malpica –Tui L ODEIRO , F., M ARTÍNEZ P OYATOS , D., M ARTÍN
Complex at 365 + 1 Ma (Ar– Ar, phengite; P ARRA , L. M. & M ATAS , J. 2008. Rheic Ocean ophio-
Rodriguez et al. 2003), interpreted to be related to litic remnants in southern Iberia questioned by
exhumation. The tangential deformation respon- SHRIMP U –Pb zircon ages on the Beja– Acebuches
sible for the emplacement of nappes occurred by a amphibolites. Tectonics, 25, TC5006.
combination of thrusts and normal faults developed B ARD , J. P. 1969. Le métamorphisme régional progressif
under amphibole-facies at 350 –340 Ma. The perva- des Sierra de Aracena en Andalousie occidentale
sive greenschist-facies mylonitic fabric parallel (Espagne). Unpublished PhD dissertation, Université
with the older ones indicate orogen-parallel tectonic de Montpellier, France.
B ARD , J. P. & M OINE , B. 1979. Acebuches amphibolites
transport until the Visean (330 + 1 Ma; Ar –Ar, in the Aracena hercynian metamorphic belt (southwest
phengite; Rodriguez et al. 2003). Spain): geochemical variations and basaltic affinities.
Variscan intra-continental extensional tectonics Lithos, 12, 271 –282.
related with orogen-parallel transport seems to be B ARD , J. P., C APDEVILA , R., M ATTE , PH . & R IBEIRO , A.
more important than previously recognized in 1973. Geotectonic model for the Iberian Variscan
the Iberian Massif. This tectonic context was pro- Orogen. Nature, 241, 50–52.
bably caused by complex interactions between B LACK , L. P., W ILLIAMS , I. S. & C OMPSTON , W. 1986.
lithospheric plates in convergent settings, when Four zircon ages from one rock: the history of a
any process affecting subduction/collision rates 3930 Ma-old granulite from Mount Sones, Enderby
Land, Antarctica. Contributions to Mineralogy and
and/or tectonic consequences of triple junction Petrology, 94, 427–437.
migration causes collapse of thickened continental B OWRING , S. A. & W ILLIAMS , I. S. 1999. Priscoan
crust. Additional influence from sub-lithosphere (4.00–4.03 Ga) orthogneisses from northwestern
mantle dynamics cannot be disregarded. Canada. Contributions to Mineral Petrology, 134,
3– 16.
This work is a contribution to the projects: CGL 2004- B UICK , I. S. & C ARTWRIGHT , I. 1994. Fluid-rock inter-
06808-C04-02-BTE, ‘Estudio geoquı́mico, tectónico y action across a green-schist-to granulite facies tran-
experimental de los procesos de reciclaje cortical y interac- sition, Reynolds range, central Australia: implications
ción manto-corteza’ and IGCP 497: ‘The Rheic Ocean: for regional-scale fluid flow in LP/HT orogenic belts.
Its origin, evolution and correlatives’. The authors thank Mineralogical Magazine, 58, 130– 131.
C. Quesada and J. Allen for their constructive revisions, C ARVALHOSA , C. 1965. Contribuição para o conheci-
and P. Moita for providing sample VLV. mento geológico da região entre Portel e Ficalho
(Alentejo). Memória dos Serviços Geológicos de
Portugal, 11.
References C ARVALHOSA , A. 1983. Esquema geológico do Maciço
de Évora. Comunicações dos Serviços Geológicos de
A BALOS , B. V., G IL I BARGUCHI , I. & E GUILUZ , L. 1991. Portugal, 69, 201–208.
Structural and metamorphic evolution of the Almadén C ARVALHOSA , A. 1999. Carta Geológica de Portugal,
de la Plata Core (Seville, Spain) in relation to syn- Noticia Explicativa da Folha 36-C (Arraiolos). Insti-
metamorphic shear between the Ossa– Morena and tuto Geológico e Mineiro, Lisboa, Portugal, scale
South Portuguese zones of the Iberian Variscan fold 1:50 000.
belt. Tectonophysics, 191, 365–387. C ASTRO , A., F ERNÁNDEZ , C. ET AL . 1996a. Triple-
A PRAIZ , A. 1998. Geologı́a de los macizos de Lora del junction migration during Paleozoic plate conver-
Rı́o y Valuengo (Zona de Ossa Morena). Evolución gence: the Aracena metamorphic belt, Hercynian
tectonometamórfica y significado geodinámico. massif, Spain. Geologische Rundschau, 85, 180–185.
Unpublished PhD dissertation, Universidad del Paı́s C ASTRO , A., F ERNÁNDEZ , C., DE LA R OSA , J. D.,
Vasco. M ORENO -V ENTAS , I. & R OGERS , G. 1996b. Signifi-
A PRAIZ , A. & E GUILUZ , L. 2002. Hercynian tectono- cance of MORB-derived Amphiboles from the
thermal evolution associated with crustal extension Aracena Metamorphic Belt, Southwest Spain.
and exhumation of the Lora del Rı́o metamorphic Journal of Petrology, 37, 235–260.
core complex (Ossa–Morena Zone, Iberian Massif, C ASTRO , A., F ERNÁNDEZ , C., E L -H MIDI , H., E L -B IAD ,
SW Spain). Geologische Rundschau, 91, 76–92. M., D ÍAZ , M., DE LA R OSA , J. & S TUART , F. 1999.
A RAÚJO , A., F ONSECA , P., M UNHÁ , J., M OITA , P., Age constraints to the relationships between magma-
P EDRO , J. & R IBEIRO , A. 2005. The Moura phyllonitic tism, metamorphism and tectonism in the Aracena
complex: an accretionary complex related with obduc- metamorphic belt, southern Spain. International
tion in the southern Iberia Variscan suture. Geodina- Journal of Earth Sciences, 88, 26– 37.
mica Acta, 18, 375– 388. C HICHORRO , M. 2006. Estrutura do Sudoeste da Zona de
A RMSTRONG , R. A., F ANNING , C. M., E LDRIDGE , C. S. Ossa-Morena: Área de Santiago de Escoural –
& F RIMMEL , H. E. 1995. Geochronological and Cabrela (Zona de Cisalhamento de Montemor-o-Novo,
Maciço de Évora). Unpublished PhD dissertation, (Espagne). Éssai sur le métamorphisme des roches
Universidade de Évora, Portugal. éruptives basiques. Unpublished PhD dissertation,
C HICHORRO , M., P EREIRA , M. F., A PRAIZ , A. & S ILVA , Université de Nancy, France.
J. B. 2004. High temperature– Low pressure tectonites F LORIDO , P. & Q UESADA , C. 1984. Estado actual de con-
from the Boa Fé Fault Zone (Évora Massif, Ossa ocimientos sobre el Macizo de Aracena. Cuadernos
Morena Zone, Portugal): evidences for transtensional Laboratorio Geológico de Laxe, 8, 257 –277.
tectonics. Geogaceta, 34, 43–46. F ONSECA , P. & R IBEIRO , A. 1993. Tectonics of the
C HICHORRO , M., P EREIRA , M. F., D ÍAZ -A ZPIROZ , M., Beja-Acebuches ophiolite: a major suture in the
W ILLIAMS , I. S., F ÉRNANDEZ , C., P IN , C. & S ILVA , Iberian Variscan Foldbelt. Geologische Rundschau,
J. B. 2008. Cambrian ensialic rift-related magmatism 82, 440 –447.
in the Ossa–Morena Zone (Évora–Aracena meta- F RASER , G., E LLIS , D. & E GGINS , S. M. 1997. Zirconium
morphic belt, SW Iberian Massif): Sm– Nd isotopes abundance in granulite-facies minerals, with impli-
and SHRIMP zircon U– Th–Pb geochronology. Tecto- cations for zircon geochronology in high-grade rocks.
nophysics, 461, 91–113. Geology, 25, 607–610.
C ORDANI , U. G., N UTMAN , A. P., A NDRADE , A. S., F RIEND , C. R. L. & K INNY , P. D. 1995. New evidence for
S ANTOS , J. F., A ZEVEDO , M. R., M ENDES , M. H. & protolith ages of Lewisian granulites, northwest Scot-
P INTO , M. S. 2006. New U–Pb SHRIMP zircon ages land. Geology, 23, 1027– 1030.
for pre-variscan orthogneisses from Portugal and G IESE , U., H OEGEN VON , R., H OYMANN , K.-H. &
their bearing on the evolution of the Ossa–Morena W ALTER , R. 1994. The Palaeozoic evolution of the
Tectonic Zone. Anais da Academia Brasileira de Ossa Morena Zone and its boundary to the South Por-
Ciencias, 78(1), 133 –149. tuguese Zone in SW Spain: geological constraints and
C RESPO -B LANC , A. & O ROZCO , M. 1991. The boundary geodynamic interpretation of a suture in the Iberian
between the Ossa–Morena and South Portuguese Variscan orogen. Neues Jahrbuch fur Geologie und
zones (southern Iberian Massif): a major suture in the Palaontologie, Abhandlungen, 192, 383– 412.
European Hercynian chain. Geologische Rundschau, G ONÇALVES , F. 1985. Contribuição para o conhecimento
80, 691–702. geológico do complexo volcano-sedimentar da Toca
C UMMING , G. L. & R ICHARDS , J. R. 1975. Ore lead da Moura (Alcácer do Sal). Memoria da Academia de
isotope ratios in a continuously changing earth. Earth Ciências de Lisboa, 26, 263–267.
and Planetary Science Letters, 28, 155– 171. H EAMAN , L. M., B OWINS , R. & C ROCKET , J. 1990. The
D ÍAZ A ZPIÍROZ , M. 2006. Evolución tectono- chemical composition of igneous zircon studies: impli-
metamórfica del domı́nio de alto grado da la banda cations for geochemical tracer studies. Geochimica et
metamórfica de Aracena, Laboratorio Xeolóxico de Cosmochimica Acta, 54, 1597–1607.
Laxe, Coruña. Nova Terra, 30, 858. H OSKIN , P. W. O. 2005. Trace-element composition of
D ÍAZ A ZPÍROZ , M. & F ERNÁNDEZ , C. 2005. Kinematic hydrothermal zircon and the alteration of Hadean
analysis of the southern Iberian shear zone and tectonic zircon from the Jack Hills, Australia. Geochimica et
evolution of the Acebuches metabasites (SW Variscan Cosmochimica Acta, 69(3), 637–648.
Iberian Massif). Tectonics, 24, TC3010, doi: 10.1029/ H OSKIN , P. W. O. & B LACK , L. P. 2000. Metamorphic
2004TC001682. zircon formation by solid-state recrystallization of pro-
D ÍAZ A ZPÍROZ , M., F ERNÁNDEZ , C., C ASTRO , A. & tolith igneous zircon. Journal of Metamorphic
E L -B IAD , M. 2006. Tectonometamorphic evolution Geology, 18, 423–439.
of the Aracena metamorphic belt (SW Spain) resulting J ESUS , A., M UNHÁ , J., M ATEUS , A., T ASSINARI , C. &
from ridge-trench interaction during Variscan plate N UTMAN , A. P. 2007. The Beja Layered Gabbroic
convergence. Tectonics, 25(1), TC1001, 10.1029/ Sequence (Ossa–Morena Zone, Southern Portugal):
2004TC001742. geochronology and geodynamic implications. Geo-
D ODSON , M. H., C OMPSTON , W., W ILLIAMS , I. S. & dinamica Acta, 20(3), 139–157.
W ILSON , J. F. 1988. A search for ancient detrital L IÑÁN , E. & Q UESADA , C. 1990. Part V Ossa-Morena
zircons in Zimbabwean sediments. Journal of the Geo- Zone: 2. Stratigraphy. In: D ALLMEYER , R. D. &
logical Society, London, 145, 977–983. M ARTÍNEZ -G ARCÍA , E. (eds) Pre-Mesozoic Geology
D UPUY , C., D OSTAL , J. & B ARD , J. P. 1979. Trace of Iberia. Springer, Berlin, 229– 266.
element geochemistry of Paleozoic amphibolites M ARTINEZ -C ATALAN , J. R., A RENAS , R. & D ÍEZ
from SW Spain. Tschermaks Minerologische und B ALDA , M. A. 2003. Large extensional structures
Petrographische Mitteilungen, 26, 87– 93. developed during emplacement of a crystalline thrust
E GUILUZ , L., G IL I BARGUCHI , J. I., A BALOS , B. & sheet: the Mondoñedo nappe (NW Spain). Journal of
A PRAIZ , A. 2000. Superposed Hercynian and Cado- Structural Geology, 25, 1815– 1839.
mian orogenic cycles in the Ossa-Morena Zone and M ARTINEZ -C ATALAN , J. R., A RENAS , R. ET AL . 2007.
related areas of the Iberian Massif. Geological Space and time in the tectonic evolution of the north-
Society America Bulletin, 112, 1398– 1413. western Iberian Massif. Implications for the compre-
E L -B IAD , M. 2000. Generacı́on de granitoides en hension of the Variscan belt. In: H ATCHER , R. D.,
ambientes geológicamente contrastados del Macizo C ARLSON , M. P., M C B RIDE , J. H. & M ARTINEZ -
Ibérico: Limitaciones experimentales entre 2 y 15 C ATALAN , J. R. (eds) 4-D Framework of Continental
kbar. Unpublished PhD dissertation, Universidad de Crust. Geological Society of America Memoir, 200.
Huelva, Spain. M ATTE , P. 2001. The Variscan collage and orogeny (480–
F ABRIÈS , J. 1963. Les formations cristallines et métamor- 290 Ma) and the tectonic definition of the Armorica
phiques du Nord-Est de la Province de Seville microplate: a review. Terra Nova, 13, 122–128.
M OITA , P. 2007. Granitóides no SW da zona de Ossa– link between the Ossa– Morena Ediacaran basins and
Morena (Montemor-o-Novo-Évora). Petrogénese e the West African craton. Geological Society, London,
processos geodinâmicos. Unpublished PhD thesis, Special Publications, 297, 385– 408.
Universidade de Évora. P IN , C., P AQUETTE , J.-L. & F ONSECA , P. 1999. 350 Ma
M OITA , P., S ANTOS , J. F. & P EREIRA , M. F. 2005a. Tona- (U–Pb zircon) igneous emplacement age and Sm-Nd
lites from the Hospitais Massif (Ossa-Morena Zone, isotopic study of the Beja gabbroic complex (S Portu-
SW Iberian Massif, Portugal) II: geochemistry and gal), XV Reunión Geologı́a del Oeste Peninsular,
petrogenesis. Geogaceta, 37, 55–58. Badajoz, Spain. Journal of Conference Abstracts,
M OITA , P., F ONSECA , P., T ASSINARI , C., A RAUJO , A. & 4(3), 1019.
P ALACIOS , T. 2005b. Phase equilibria and geochronol- P IN , C., F ONSECA , P. E., P AQUETTE , J.-L., C ASTRO , P. &
ogy of Ossa–Morena eclogites. Actas da XIV Semana M ATTE , P. 2008. The ca. 350 Ma Beja Igneous
de Geoquı́mica/VII Congresso de geoquı́mica dos Complex: A record of transcurrent slab break-off in
Paı́ses de Lı́ngua Portuguesa, 2, 463– 466. the Southern Iberia Variscan Belt? Tectonophysics,
M UNHÁ , J., O LIVEIRA , J. T., R IBEIRO , A., O LIVEIRA , V., 461, 356 –377.
Q UESADA , C. & K ERRICH , R. 1986. Beja –Acebuches Q UESADA , C. 1990. Precambrian successions in SW
Ophiolite: characterization and geodynamic signifi- Iberia: their relationship to ‘Cadomian’ orogenic
cance. Maleo, 2, 1 –3. events. In: D’L EMOS , R. S., S TRACHAN , R. A. &
O NÈZIME , J., C HARVET , J., F AURE , M., B OURDIER , J. & T OPLEY , G. G. (eds) The Cadomian Orogeny. Geo-
C HAUVET , A. 2003. A new geodynamic interpretation logical Society, London, Special Publications, 551,
for the South Portuguese Zone (SW Iberia) and the 353–362.
Iberian Pyrite Belt genesis. Tectonics, 22, 10–27. Q UESADA , C. 1991. Geological constraints on the Paleo-
O RDOÑEZ -C ASADO , B. 1998. Geochronological studies zoic tectonic evolution of tectonostratigraphic terranes
of the Pre-Mesozoic basement of the Iberian Massif: in the Iberian Massif. Tectonophysics, 185, 225 –245.
the Ossa Morena Zone and Allochthonous Complexes Q UESADA , C., F ONSECA , P. E., M UNHÁ , J. M.,
within the Central-Iberian Zone. PhD thesis, ETH O LIVEIRA , J. T. & R IBEIRO , A. 1994. The Beja-
(12.940), Zurich. Acebuches Ophiolite (southern Iberia Variscan fold
P AN , Y. 1997. Zircon- and monazite-forming meta- belt): geological characterization and geodynamic
morphic reactions at Manitouwadge, Ontario. Cana- significance. Boletı́n Geológico y Minero, 105, 3 –49.
dian Mineralogist, 35, 105–118. R EY , P., V ANDERHAEGHE , O. & T EYSSIER , C. 2001.
P ATIÑO D OUCE , A. E. A., C ASTRO , A. & E L -B IAD , M. Gravitational collapse of the continental crust: defi-
1997. Thermal evolution and tectonic implications of nition, regimes and modes. Tectonophysics, 342,
spinel-cordierite granulites from the Aracena meta- 435–449.
morphic belt, Southwest Spain. GAC/MAC Annual R IBEIRO , A. 1983. Relações entre formações do Devónico
meeting, Ottawa, Canada. superior e o Maciço de Évora na região de Cabrela
P EREIRA , M. F., S ILVA , J. B. & C HICHORRO , M. 2003. (Vendas Novas). In: R IBEIRO , A., G ONÇALVES , F.,
Internal Structure of the Évora High-grade Terrains S OARES DE A NDRADE , A. & O LIVEIRA , V. (eds)
and the Montemor-o-Novo Shear Zone (Ossa– Guia das excursões do bordo sudoeste da Zona de
Morena Zone, Portugal). Geogaceta, 33, 79–82. Ossa-Morena. Comunicações dos Serviços Geológicos
P EREIRA , Z., O LIVEIRA , V. & O LIVEIRA , J. T. 2006a. de Portugal, 69, 267– 269.
Palynostratigraphy of the Toca da Moura and Cabrela R IBEIRO , A., Q UESADA , C. & D ALLMEYER , R. D. 1990.
Complexes, Ossa– Morena Zone, Portugal. Geody- Geodynamic evolution of the Iberian Massif. In:
namic implications. Review of Palaeobotany & Paly- D ALLMEYER , R. D. & M ARTÍNEZ G ARCÍA , E. (eds)
nology, 139, 227–240. Pre-Mesozoic Geology of Iberia. Berlin-Heidelberg,
P EREIRA , M. F., C HICHORRO , M., L INNEMANN , U., Springer-Verlag, 399– 409.
E GUILUZ , L. & S ILVA , J. B. 2006b. Inherited arc sig- R IBEIRO , A., M UNHÁ , J. ET AL . 2007. Geodynamic evol-
nature in Ediacaran and Early Cambrian basins of the ution of the SW Europe Variscides. Tectonics, 26,
Ossa–Morena Zone (Iberian Massif, Portugal): paleo- TC6009.
geographic link with European and North African R OBARDET , M. 2002. Alternative approach to the
Cadomian correlatives. Precambrian Research, 144, Variscan Belt in southwestern Europe: Preorogenic
297– 315. palaeobiogeographical constraints. In: M ARTINEZ -
P EREIRA , M. F., S ILVA , J. B., C HICHORRO , M., M OITA , C ATALAN , J. R., H ATCHER , R. D., J R ., A RENAS , R.
P., S ANTOS , J. F., A PRAIZ , A. & R IBEIRO , C. 2007. & D IAZ G ARCIA , F. (eds) Variscan– Appalachian
Crustal growth and deformational processes in the Dynamics: The Building of the Late Palaeozoic Base-
Northern Gondwana margin: Constraints from the ment. Geological Society of America, Special Papers,
Évora Massif (Ossa-Morena Zone, SW Iberia, Portu- 364, 1 –15.
gal). In: L INNEMANN , U., N ANCE , R. D., K RAFT , P. R OBARDET , M. & G UTÍERREZ -M ARCO , J. C. 2004. The
& Z ULAUF , G. (eds) The Evolution of the Rheic Ordovician, Silurian and Devonian sedimentary rocks
Ocean: From Avalonian–Cadomian Active Margin of the Ossa–Morena Zone (SW Iberian Peninsula,
to Alleghenian– Variscan Collision. Special Paper Spain). Journal of Iberian Geology, 30, 73–92.
Geological Society of America, 423, 333– 358. R ODRIGUEZ , J. 2005. Recristaizacion y deformacion de
P EREIRA , M. F., C HICHORRO , M., W ILLIAMS , I. S. & litologias supracorticales sometidas a metamorfismo
S ILVA , J. B. 2008. Zircon U –Pb geochronology of de alta presion (Complejo Malpica-Tuy, NO del
paragneisses and biotite granites from the SW Iberian Macizo Ibérico). Laboratório Xeológico de Laxe,
Massif (Portugal): Evidence for a paleogeographic Serie Nova Terra, 29.
R ODRIGUEZ , J., C OSCA , M. A., G IL I BARGUCHI , J. I. & H ADI , H. 2005. The tectonic frame of the Variscan-
D ALLMEYER , R. D. 2003. Strain partitioning and pres- Alleghanian orogen in Southern Europe and Northern
ervation of 40Ar/39Ar ages during Variscan exhuma- Africa. Tectonophysics, 398, 181 –198.
tion of a subducted crust (Malpica-Tuy Complex, S TEIGER , R. H. & J ÄGER , E. 1977. Subcommision on geo-
SW Spain). Lithos, 70, 111– 139. chronology: convention on the use of decay constants
R OSAS , F. M., M ARQUES , F. O., B ALLÈVRE , M. & T AS- in geo- and cosmochronology. Earth and Planetary
SINARI , C. 2008. Geodynamic evolution of the SW Science Letters, 36, 359– 362.
Variscides: Orogenic collapse shown by new tectono- T ORNOS , F., C ASQUET , C. & R ELVAS , M. R. S. J. 2005.
metamorphic and isotopic data from western Ossa– 4: Transpressional tectonics, lower crust decoupling
Morena Zone, SW Iberia. Tectonics, 27, TC6008. and intrusion of deep mafic sills: a model for the
S ANCHEZ -G ARCIA , T., B ELLINDO , F. & Q UESADA , C. unsual metallogenesis of SW Iberia. Ore Geology
2003. Geodynamic setting and geochemical signatures Reviews, 27, 133– 163.
of Cambrian– Ordovician rift-related igneous rocks V AVRA , G., G EBAUER , D., S CHMID , R. & C OMPSTON ,
(Ossa– Morena Zone, SW Iberia). Tectonophysics, W. 1996. Multiple zircon growth and recrystallisation
365, 233– 255. during polyphase Late Carboniferous to Triassic
S ANDIFORD , M. & P OWELL , R. 1986. Deep crustal meta- metamorphism in granulites of the Ivrea Zone
morphism during continental extension: modern and (Southern Alps): an ion microprobe (SHRIMP) study.
ancient examples. Earth and Planetary Science Contributions to Mineralogy and Petrology, 122,
Letters, 79, 151–158. 337– 358.
S ANTOS , J. F., M ATA , J., G ONÇALVES , F. & M UNHÁ , J. M. W ATSON , E. B. 1996. Dissolution, growth and survival of
1987. Contribuição para o conhecimento geológico- zircons during crustal fusion: kinetic principles, geo-
petrológico da região de Santa Suzana: o Complexo logic models and implications for isotopic inheritance.
Vulcano-sedimentar de Toca da Moura. Comuni- Geological Society of America Special Paper, 315,
cações Serviços Geológicos Portugal, 73, 29– 40. 43–56.
S ANTOS , J. F., S OARES DE A NDRADE , A. & M UNHÁ , W ILLIAMS , I. S. 2001. Response of detrital zircon and
J. M. 1990. Magmatismo orogénico Varisco no limite monazite, and their U-Pb isotopic systems, to regional
meridional da Zona de Ossa-Morena. Comunicações metamorphism and host-rock partial melting, Cooma
dos Serviços Geológicos de Portugal, 76, 91– 124. Complex, southeastern Australia. Australian Journal
S CHÄFER , H. J., G EBAUER , D., N ÄGLER , T. F. & of Earth Sciences, 48, 557– 580.
E GUILUZ , L. 1993. Conventional and ion-microprobe W ILLIAMS , I. S. & C LAESSON , S. 1987. Isotopic evidence
U-Pb dating of detrital zircons of the Tentudı́a Group for the Precambrian provenance and Caledonian meta-
(Serie Negra, SW Spain): implications for zircon sys- morphism of high grade paragneisses from the Seve
tematics, stratigraphy, tectonics and the Precam- Nappes, Scandinavian Caledonides, II Ion microprobe
brian/Cambrian boundary. Contributions to Mineral zircon U-Th-Pb. Contributions to Mineralogy and
Petrology, 113, 289–299. Petrology, 97, 205–217.
S ILVA , J. B. & P EREIRA , M. F. 2004. Transcurrent conti- W ILLIAMS , I. S. & H ERGT , J. M. 2001. U–Pb dating of
nental tectonics model for the Ossa-Morena Zone Neo- Tasmanian dolerites: a cautionary tale of SHRIMP
proterozoic-Paleozoic evolution, SW Iberian Massif, analysis of high-U zircon. In: W OODHEAD , J. D.,
Portugal. Geologische Rundschau, 93, 886–896. H ERGT , J. M. & N OBLE , W. P. (eds) Beyond 2000:
S ILVA , J. B., O LIVEIRA , J. T. & R IBEIRO , A. 1990. South New frontiers in Isotope Geoscience. University of
Portuguese Zone, structural outline. In: D ALLMEYER , Melbourne, Melbourne, 185– 188.
R. D. & M ARTÍNEZ G ARCÍA , E. (eds) Pre-Mesozoic W ILLIAMS , I. S., B UICK , I. S. & C ARTWRIGHT , I.
Geology of Iberia. Berlin-Heidelberg, Springer- 1996. An extended episode of early Mesopro-
Verlag, 348– 362. terozoic metamorphic fluid flow in the Reynolds
S IMANCAS , F., C ARBONELL , R. ET AL . 2003. Crustal Range, central Australia. Journal of Metamorphic
structure of the transpressional Variscan orogen of Geology, 14, 29–47.
SW Iberia: SW Iberia deep seismic reflection profile W ILLIAMS , O. S., C OMPSTON , W., B LACK , L. P.,
(IBERSEIS). Tectonics, 22(6), 1062, doi: 10.1029/ I RELAND , T. R. & F OSTER , J. J. 1984. Unsupported
2002TC001479. radiogenic Pb in zircon: a cause of anomalously high
S IMANCAS , J. F., T AHIRI , A., A ZOR , A., G ONZÁLEZ Pb– Pb, U– Pb and Th– Pb ages. Contributions to
L ODEIRO , F., M ARTÍNEZ P OYATOS , D. J. & E L Mineralogy and Petrology, 88, 322–327.
Palaeozoic palaeogeography of Mexico: constraints from
detrital zircon age data
R. DAMIAN NANCE1*, J. DUNCAN KEPPIE2, BRENT V. MILLER3,
J. BRENDAN MURPHY4 & JAROSLAV DOSTAL5
1
Department of Geological Sciences, 316 Clippinger Laboratories, Ohio University,
Athens, Ohio 45701 USA
2
Departamento de Geologı́a Regional, Instituto de Geologı́a, Universidad Nacional
Autonoma de México, 04510, México D.F., México
3
Department of Geology and Geophysics, Texas A&M University,
College Station, Texas 77843, USA
4
Department of Earth Sciences, St. Francis Xavier University, Antigonish, Nova Scotia,
Canada, B2G 2W5
5
Department of Geology, St. Mary’s University, Halifax, Nova Scotia, Canada, B3H 3C3
*Corresponding author (e-mail: nance@ohio.edu)
Abstract: Detrital zircon age populations from Palaeozoic sedimentary and metasedimentary
rocks in Mexico support palinspastic linkages to the northwestern margin of Gondwana (Amazo-
nia) during the late Proterozoic –Palaeozoic. Age data from: (1) the latest Cambrian-Pennsylvanian
cover of the c. 1 Ga Oaxacan Complex of southern Mexico; (2) the ?Cambro-Ordovician to Triassic
Acatlán Complex of southern Mexico’s Mixteca terrane; and (3) the ?Silurian Granjeno Schist of
northeastern Mexico’s Sierra Madre terrane, collectively suggest Precambrian provenances in: (1)
the c. 500– 650 Ma Brasiliano orogens and c. 600–950 Ma Goias magmatic arc of South America,
the Pan-African Maya terrane of the Yucatan Peninsula, and/or the c. 550 –600 Ma basement that
potentially underlies parts of the Acatlán Complex; (2) the Oaxaquia terrane or other c. 1 Ga base-
ment complexes of the northern Andes; and (3) c. 1.4–3.0 Ga cratonic provinces that most closely
match those of Amazonia. Exhumation within the Acatlán Complex of c. 440 –480 Ma granitoids
prior to the Late Devonian–early Mississippian, and c. 290 Ma granitoids in the early Permian,
likely provided additional sources in the Palaeozoic. The detrital age data support the broad corre-
lation of Palaeozoic strata in the Mixteca and Sierra Madre terranes, and suggest that, rather than
representing vestiges of Iapetus or earlier oceanic tracts as has previously been proposed, both were
deposited along the southern, Gondwanan (Oaxaquia) margin of the Rheic Ocean and were
accreted to Laurentia during the assembly of Pangaea in the late Palaeozoic.
Introduction dominated by Palaeozoic siliciclastic and oceanic

rocks tectonically juxtaposed against the Oaxaquia
The geology of Mexico is dominated by a collage of terrane along major, north– south dextral faults
suspect terranes, the majority of which constitute of late Palaeozoic age (e.g. Elı́as-Herrera & Ortega-
part of the North American Cordillera having been Gutiérrez 2002; Dowe et al. 2005).
accreted to the southwestern margin of Laurentia The metasedimentary rocks of the Mixteca and
during the Mesozoic –Cenozoic (e.g. Keppie Sierra Madre terranes have long been recognized to
2004). However, several Mexican terranes were preserve an important record of Palaeozoic ocean
accreted to Laurentia during the late Palaeozoic opening and closure with links to the adjacent
amalgamation of Pangaea and record Palaeozoic Oaxaquia terrane (e.g. Yañez et al. 1991; Ramı́rez-
histories that constrain continental reconstructions Ramı́rez 1992). But whether this ocean was Iape-
for Pangaea assembly. These terranes (Fig. 1) tus, the Rheic Ocean, or some other oceanic tract,
include Oaxaquia, a c. 1 Ga crustal block that remains controversial. Yañez et al. (1991) proposed
underlies much of central Mexico and is overlain that the Acatlán Complex documented a Devonian
by a thin veneer of unmetamorphosed Palaeozoic episode of Laurentia–Gondwana collision, an event
strata (Ortega-Gutiérrez et al. 1995), and the they linked to the Acadian belt of the Appalachian
Mixteca and Sierra Madre terranes, which are orogen. On the basis of revised geochronology,
240 R. D. NANCE ET AL.
Fig. 1. Simplified tectonic map of Middle America showing location of Mexico’s Oaxacan Complex and Novillo
Gneiss (Oaxaquia terrane), Acatlán Complex (Mixteca terrane) and Granjeno Schist (Sierra Madre terrrane).
Figure modified after Keppie (2004).
however, Ortega-Gutiérrez et al. (1999) later reas- Palaeozoic detrital zircon data
signed this event to the Late Ordovician –early Silur-
ian and proposed that the collision recorded closure Palaeozoic U –Pb detrital zircon age data are pre-
of the Iapetus Ocean, a view subsequently developed sently available in Mexico from three main sources
for the Acatlán Complex by Talavera-Mendoza et al. (Fig. 1): (1) the sedimentary cover of the c. 1 Ga
(2005) and Vega-Granillo et al. (2007). Ramı́rez- Oaxacan Complex in southern Mexico, the largest
Ramı́rez (1992) similarly advocated a marginal exposed portion of the Oaxaquia terrane (Gillis
basin of the Iapetus Ocean as the depositional et al. 2005); (2) metasedimentary rocks of the
setting for the metasedimentary rocks of the Sierra largely Palaeozoic Acatlán Complex, which forms
Madre terrane. A linkage to the Iapetus Ocean, the basement of southern Mexico’s Mixteca terrane
however, is not supported by new age data that con- (Sánchez-Zavala et al. 2004; Talavera-Mendoza
strain the high-grade collisional metamorphism of et al. 2005; Keppie et al. 2006a; Murphy et al.
the Acatlán Complex to the latest Devonian– early 2006; Vega-Granillo et al. 2007; Grodzicki et al.
Mississippian (Middleton et al. 2007). Instead, 2008; Hinojosa-Prieto et al. 2008; Morales-Gámez
these data have been used to link the Acatlán et al. 2008); and (3) the Palaeozoic Granjeno Schist
Complex to the southern (Gondwanan) margin of of northeastern Mexico’s Sierra Madre terrane,
the Rheic Ocean (e.g. Nance et al. 2006, 2007a; which is tectonically juxtaposed against the second
Keppie et al. 2008a), a view consistent with the con- largest exposure of the Oaxaquia terrane, the
tinental reconstructions of Keppie & Ramos (1999). c. 1 Ga Novillo Gneiss (Nance et al. 2007b).
Discriminating between these mutually exclu- These ages have been determined for samples of
sive models is essential if the Palaeozoic strati- variable size using a variety of protocols and
graphic and tectonothermal history of Mexico is to analytical techniques, including conventional,
be placed within a broader plate tectonic frame- isotope dilution and thermo-ionization mass spec-
work, and the country’s important role in continen- trometry (ID-TIMS), laser ablation-inductively
tal reconstructions for the assembly of Pangaea is to coupled plasma mass spectrometry (LA-ICPMS),
be resolved. In this article, we examine the problem and sensitive high-resolution ion microprobe
from the perspective of sedimentary provenance by (SHRIMP). For comparative purposes, therefore,
using available detrital zircon data from Palaeozoic the data are best depicted using relative age
strata to identify potential cratonic source areas and, probability plots (Ludwig 2003) normalized for
thereby, constrain the palinspastic restoration of the sample size. These plots are constructed by:
those terranes that record Palaeozoic sedimentary (1) calculating a normal distribution for each analy-
histories. sis based on the reported age and uncertainty; (2)
PALAEOZOIC PALAEOGEOGRAPHY OF MEXICO 241
summing the probability distributions of all accepta- include three formations that are separated by
ble analyses into a single curve; and (3) dividing the shear zones (Centeno-Garcia & Keppie 1999): (1)
area under the curve by the number of analyses the latest Cambrian –Ordovician Tiñu Formation,
(Gehrels et al. 2006). The plots were constructed which contains shallow-marine fauna of Gondwa-
at the same scale from original data using an nan affinity (Robison & Pantoja-Alor 1968;
Excel macro available on the website of the Laser- Landing et al. 2007); (2) the Mississippian Santiago
Chron Center at the University of Arizona (www. Formation with midcontinent (US) brachiopod
geo.arizona.edu/alc). For consistency, the normal- fauna (e.g. Navarro-Santillan et al. 2002); and (3)
ized relative age probability plots were constructed the Pennsylvanian Ixtaltepec Formation. Gillis
using the 206Pb/238U age for young (,1.0 Ga) et al. (2005) report ID-TIMS and LA-ICPMS
zircons and the 206Pb/207Pb age for older U –Pb detrital zircon ages from each of these three
(.1.0 Ga) grains, and analyses with an error formations (Fig. 2).
.10%, and older ages with .20% discordance or
.5% reverse discordance, were not used. Tiñu Formation. The latest Cambrian –earliest
Ordovician Tiñu Formation comprises up to 200 m
of limestone, shale and siltstone interpreted to rep-
Oaxacan Complex resent a transgressive sequence of wave-dominated
The Oaxacan Complex of the Oaxaquia terrane shelf sediments (Sánchez-Zavala et al. 1999;
(Fig. 1) comprises granulite facies paragneisses Landing et al. 2007). The basal unit is a coarse
and within-plate, rift-related AMCG (anorthosite- arkosic sandstone, detrital zircons from which
mangerite-charnockite-granite) suites that have yield a normalized relative age population (Fig. 2a)
yielded U –Pb emplacement ages of c. 1000–1020 with a series of peaks between c. 995 Ma and
and c. 1100–1160 Ma (Keppie et al. 2001, 2003; 1225 Ma, and a single older zircon at c. 1450 Ma
Solari et al. 2003), and orthogneisses with a proto- (Gillis et al. 2005). The youngest detrital zircon
lith age of 1350 Ma that were migmatized at gives an LA-ICPMS age of c. 960 Ma.
c. 1106 Ma. These rocks underwent polyphase Santiago Formation. The early Mississippian
deformation and granulite facies metamorphism Santiago Formation comprises some 200 m of
at c. 980 –1005 Ma, cooling through 500 8C brachiopod-rich limestone, shale and minor sand-
(40Ar/39Ar on hornblende) at c. 975 Ma (Solari stone (Pantoja-Alor & Robison 1967; Navarro-
et al. 2003). A shallow-level, calc-alkaline granitoid Santillan et al. 2002) interpreted to have been
pluton intruded the northern Oaxacan Complex at deposited in either a shallow marine (Sánchez-
c. 917 Ma (Ortega-Obregon et al. 2003). Zavala et al. 1999) or deeper shelf –slope (Flores
Further north, inliers of the Oaxaquia terrane de Dios Gonzalez et al. 1998) environment.
include the Huiznopala and Novillo gneisses Detrital zircons from a calcareous sandstone at the
(Fig. 1), the latter juxtaposed against the Palaeozoic base of the formation yield a normalized relative
Granjeno Schist of the Sierra Madre terrane along a age population (Fig. 2b) with peaks at c. 470 and
north–south dextral shear zone of late Palaeozoic 990 Ma (Gillis et al. 2005). The youngest detrital
age (Dowe et al. 2005). The Huiznopala Gneiss zircon gives an ID-TIMS age of c. 460 Ma
comprises layered paragneisses and arc-related (mid-Ordovician; Gradstein et al. 2004).
orthogneisses that have yielded ages of c. 1150 –
1200 Ma (Lawlor et al. 1999), and an anorthosite- Ixtaltepec Formation. The fossiliferous early –
gabbro complex emplaced during granulite facies middle Pennsylvanian Ixtaltepec Formation com-
metamorphism at c. 1000 Ma. Pegmatites that give prises c. 425 m of shale with interbedded limestone
an age of 988 + 3 Ma post-date ductile deformation. and sandstone (Pantoja-Alor & Robison 1967;
The Novillo Gneiss likewise comprises metasedi- Flores de Dios Gonzalez et al. 1998) that are inter-
mentary rocks intruded by gabbro-anorthosite, preted to be shallow marine (Sánchez-Zavala et al.
granite and amphibolite that have yielded U –Pb 1999). Detrital zircons from a sandstone near the
emplacement ages of c. 1010– 1035 and c. 1115 – base of the formation yield a normalized relative
1235 Ma (Cameron et al. 2004). These rocks under- age population (Fig. 2c) with a peak at c. 360 Ma
went polyphase deformation and granulite facies and a series of peaks in the interval c. 990–
metamorphism at 990 + 5 Ma, followed by post- 1185 Ma (Gillis et al. 2005). The youngest detrital
tectonic anorthositic pegmatite emplacement at zircons give ID-TIMS and LA-ICPMS ages of
978 + 13 Ma. At c. 546 Ma, the Novillo Gneiss c. 340 Ma (Mississippian; Gradstein et al. 2004).
was intruded by a mafic dyke swarm with
plume-related geochemical affinities (Keppie et al. Acatlán Complex
2006b).
Unmetamorphosed Palaeozoic strata that non- To the west, the Oaxacan Complex is faulted against
conformably overlie the Oaxacan Complex the mainly Palaeozoic rocks of the Acatlán Complex
Fig. 2. Normalized relative age distribution plots of U–Pb detrital zircon analyses from the sedimentary cover of
the Oaxacan Complex: (a) latest Cambrian-Ordovician Tiñu Formation; (b) early Mississippian Santiago Formation;
and (c) early–middle Pennsylvanian Ixtaltepec Formation (data from Gillis et al. 2005). Shaded bars outline the age
ranges characteristic of the Ordovician megacrystic granitoids of the Acatlán Complex (c. 440–480 Ma: Sánchez-
Zavala et al. 2004; Talavera-Mendoza et al. 2005; Miller et al. 2007; Vega-Granillo et al. 2007) and the
Mesoproterozoic basement of the Oaxaquia terrane (c. 920 –1350 Ma: Keppie et al. 2001, 2003; Solari et al. 2003;
Cameron et al. 2004). n ¼ sample size.
of the Mixteca terrane (Fig. 1). The north –south (Petlalcingo Suite) that are thought to represent an
Caltepec fault zone that defines the terrane boundary active margin clastic wedge (Keppie et al. 2006a)
between these two complexes (Fig. 3) is a dextral and comprise pelites, psammites and minor mafic
transpressive structure that was active at c. 276 Ma units (Magdalena protolith and Chazumba For-
and is overstepped by early Permian (Leonardian) mation) that were metamorphosed in the amphibo-
red beds of the Matzitzi Formation (Elı́as-Herrera lite facies and locally pervasively migmatized in
& Ortega-Gutiérrez 2002). the mid-Jurassic.
The Acatlán Complex is made up of several
tectonic assemblages that range in age from Pre-Silurian high-grade metasedimentary rocks
?Cambro–Ordovician to Triassic and record major (Piaxtla Suite). LA-ICPMS U –Pb detrital zircon
tectonothermal events of Devono–Carboniferous, age data have been obtained from several high-
Permo– Triasssic and Jurassic age (Fig. 4). For grade metasedimentary units of the Acatlán
recent reviews, see Nance et al. (2006, 2007a) and Complex sampled within the Piaxtla Suite (Fig. 5).
Keppie et al. (2008a). For the purpose of this High-grade metasedimentary rocks reportedly
paper, however, the principal units of the complex intruded by the Esperanza Granitoids have been
can be grouped into five major packages (Keppie sampled for detrital zircons at Mimilulco (Xayaca-
et al. 2008a): (1) Pre-Silurian high-grade metasedi- tlán Formation; Talavera-Mendoz et al. 2005),
mentary rocks (Piaxtla Suite) that are thought to San Francisco de Asis (Asis Lithodeme; Murphy
represent an allochthonous, ?Cambro–Ordovician et al. 2006) and Santa Cruz Organal (quartzite;
rift-drift assemblage and include locally eclogitic Vega-Granillo et al. 2007) near the northern
mafic-ultramafic rocks, serpentinites, blueschists, margin of the Acatlán Complex (Fig. 3). They
migmatites and amphibolite facies metasedimentary have also been sampled at Ixcamilpa (Ixcamilpa
units (Xayacatlán Formation, Asis Unit and Ixca- blueschist; Talavera-Mendoza et al. 2005) near the
milpa blueschists) that are intruded by a suite of western margin of the complex.
highly deformed K –feldspar/quartz megacrystic At Mimilulco, detrital zircons from a chloritoid–
granitoid bodies (Esperanza Grantoids) of Ordovi- phengite –garnet psammitic schist of the Xayacatlán
cian (c. 480 –440 Ma) age (e.g. Talavera-Mendoza Formation interbedded with retrogressed eclog-
et al. 2005; Murphy et al. 2006; Miller et al. 2007; ites yield a normalized relative age population
Vega-Granillo et al. 2007) and were exhumed in (Fig. 5a) with peaks at c. 870, 990, 1120 and
the Late Devonian –early Mississippian (Middleton 1390 Ma (Talavera-Mendoza et al. 2005). At
et al. 2007); (2) pre-Silurian low-grade metasedi- San Francisco de Asis, detrital zircons from a
mentary rocks that are thought to represent a quartz-rich metapsammite of the Asis Lithodeme
?Cambro–Ordovician rift-passive margin clastic interlayered with garnet-bearing metapelites and
sequence (El Rodeo Formation and Amate, amphibolites yield a normalized population
Canoas, Huerta, El Epazote and Las Calaveras (Fig. 5b) with major peaks at c. 920, 1160 and
units) and are associated with bimodal, rift-related 1200 Ma (Murphy et al. 2006). Older zircons give
igneous rocks and intruded by Ordovician minor peaks at c. 1490 and 1555 Ma. At Santa
megacrystic granitoids (e.g. Keppie et al. 2007, Cruz Organal, detrital zircons from a garnet-
2008b; Miller et al. 2007; Hinojosa-Prieto et al. phengite quartzite yield age probabilities (Fig. 5c)
2008; Grodzicki et al. 2008; Morales-Gámez et al. with a peak at c. 720 Ma and a series of peaks
2008; Ramos-Arias et al. 2008); (3) Devonian– between c. 945 and 1125 Ma (Vega-Granillo et al.
Carboniferous low-grade metasedimentary rocks 2007). Older zircon peaks occur at c. 1450, 1510
that include phyllite, quartzite and minor mafic and 1675 Ma. At Ixcamilpa, zircons from a chlorite-
volcanic rocks (Cosoltepec Formation, Salida and phengite schist interbedded with blueschists (Ixca-
Coatlaco units), the deposition of which was milpa blueschist) yield a very different normalized
coeval with exhumation of the high pressure rocks relative age population (Fig. 5d) with major peaks
(Grodzicki et al. 2008; Morales-Gámez et al. at c. 475, 540 and 600 Ma, a broad spectrum of
2008); (4) Carboniferous –Permian low-grade meta- smaller peaks between c. 730 and c. 1230 Ma, and
sedimentary rocks that form a continental-shallow older peaks at c. 1820 and 1950 Ma (Talavera-
marine succession dominated by slate, sandstone, Mendoza et al. 2005).
conglomerate and limestone (Tecomate, Olinalá The youngest detrital zircons in the four units
and Patlanoaya formations), the deposition of give ages of c. 800, 705, 670 and 450 Ma (Late
which was coeval with Permo –Triassic arc magma- Ordovician; Gradstein et al. 2004), respectively,
tism (Torres et al. 1999; Malone et al. 2002) and and place maximum age constraints on the timing
which are locally overthrust by the broadly coeval of their deposition. The youngest zircon cluster in
(c. 288 Ma), arc-related Totoltepec pluton (Yañez the Ixcamilpa sample, which Talavera-Mendoza
et al. 1991; Keppie et al. 2004a); and (5) Permian – et al. (2005) consider to provide a more reliable
Triassic high-grade metasedimentary rocks maximum age constraint, peaks at c. 477 Ma
244
R. D. NANCE ET AL.
Fig. 3. Simplified geological map of the Acatlán Complex, southern Mexico (modified from Ortega-Gutiérrez et al. 1999 & Keppie et al. 2008a). See Figure 1 for location.
Fig. 4. Simplified tectonostratigraphic column for the Acatlán Complex based on the geochronological data of
Keppie et al. (2004a, b, 2006a), Sánchez-Zavala et al. (2004), Talavera-Mendoza et al. (2005), Middleton et al. (2007),
Miller et al. (2007), Vega-Granillo et al. (2007) and Morales-Gámez et al. (2008). Arrows indicate present structural
positions of the units but not the nature of their emplacement. Figure modified from Nance et al. (2006). C, Cambrian;
O, Ordovician; S, Silurian; D, Devonian; C, Carboniferous; P, Permian; Tr, Triassic; J, Jurassic; A, away; T, towards.
(Early Ordovician; Gradstein et al. 2004). The sig- Ordovician megacrystic granitoids. ID-TIMS,
nificant difference in youngest zircon age between SHRIMP and LA-ICPMS crystallization and inheri-
the northern and western units suggests that the tance ages have been determined for 15 megacrystic
former provide constraints on the deposition of the granitoid bodies. The granitoids have yielded a
host rock into which the megacrystic granitoids lower intercept age of 440 + 14 Ma (Ortega-
were emplaced, whereas the latter dates a volcani- Gutiérrez et al. 1999), LA-ICPMS ages ranging
clastic protolith whose deposition was broadly from 440 + 15– 478 + 5 Ma (Sánchez-Zavala
coeval with granitoid intrusion. et al. 2004; Talavera-Mendoza et al. 2005;
Fig. 5. Normalized relative age probability plots of U–Pb detrital zircon analyses from pre-Silurian high-grade
metasedimentary units of the Acatlán Complex (Piaxtla Suite): (a) Xayacatlán Formation at Mimilulco
(Talavera-Mendoza et al. 2005); (b) Asis Lithodeme at San Francisco de Asis (Murphy et al. 2006); (c) quartzite from
Santa Cruz Organal (Vega-Mendoza et al. 2007); and (d) Ixcamilpa blueschist at Ixcamilpa (Talavera-Mendoza et al.
2005). Shaded bars outline the age ranges characteristic of the Ordovician megacrystic granitoids of the Acatlán
Complex (c. 440– 480 Ma: Sánchez-Zavala et al. 2004; Talavera-Mendoza et al. 2005; Miller et al. 2007;
Vega-Granillo et al. 2007) and the Mesoproterozoic basement of the Oaxaquia terrane (c. 920–1350 Ma: Keppie et al.
2001, 2003; Solari et al. 2003; Cameron et al. 2004). n ¼ sample size.
Vega-Granillo et al. 2007), a SHRIMP age of The Canoas Unit is dominated by low-grade
467 + 16 Ma, and a single precise ID-TIMS age psammites and pelites that record three phases of
of 461 + 2 Ma (Miller et al. 2007), all of which penetrative deformation. Detrital zircons from a
are interpreted to date their crystallization. Two sample of metapsammite show a normalized rela-
additional granitoids were interpreted by Talavera- tive age population (Fig. 7c) with peaks at c. 960,
Mendoza et al. (2005) to have Mesoproterozoic 1160 and 1510 Ma (Grodzicki et al. 2008). The
(1043 + 50 and 1149 + 6 Ma) crystallization youngest detrital zircon, which records an age of
ages. However, this interpretation is contentious c. 453 Ma (Late Ordovician; Gradstein et al.
since the granitoids contain significant Mesoproter- 2004), provides a maximum age constraint for the
ozoic inheritance in precisely this age range. timing of deposition.
LA-ICPMS ages of 464 + 4 Ma for a granite peg- The Huerta Unit comprises greenschist facies
matite and 452 + 6 and 447 + 3 Ma for foliated interbedded psammites and pelites that closely
granite dykes have been reported by Morales- resemble those of the Canoas Unit and likewise
Gámez et al. (2008), and a younger LA-ICPMS record three phases of penetrative deformation
age of 372 + 8 Ma has been reported for a leucogra- (Malone et al. 2002). A semipelite from the Huerta
nite (Vega-Granillo et al. 2007). Collectively, the Unit yielded detrital zircons that show a normalized
intrusive age of the granitoids centres at c. 460 Ma relative age population (Fig. 7d) with major peaks at
(Fig. 6a), whereas the normalized relative age popu- c. 510, 570 and 630 Ma, a series of smaller peaks
lation for xenocrystic zircons peaks at c. 565, 1015 between c. 950 and 1220 Ma, and minor peaks at
and 1185 Ma (Fig. 6b). c. 1680 and 1995 Ma (Keppie et al. 2006a). The
youngest detrital zircon records an age of c. 455 Ma.
Pre-Silurian low-grade metasedimentary The El Epazote and Las Calaveras units are
rocks. LA-ICPMS U –Pb detrital zircon age data greenschist facies meta-volcanosedimentary assem-
have been obtained from low-grade, ?Cambro– blages in the northern Acatlán Complex (Hinojosa-
Ordovician siliciclastic rocks of the Acatlán Prieto et al. 2008). The El Epazote Unit comprises
Complex near the villages of Olinalá (El Rodeo metapelites, metavolcaniclastics and fine-grained
Formation and Canoas Unit) in the southwestern metapsammites that are in tectonic contact with
part of the complex, Xayacatlán (Amate Unit) megacrystic granitoids, one of which (the La Noria
and La Huerta (Huerta Unit) in the east-central granite) has yielded a weighted mean 206Pb/238U
part, and near La Noria (Las Calaveras and El SHRIMP age of 467 + 16 Ma (Miller et al. 2007).
Epazote units) in the northern Acatlán Complex The Las Calaveras Unit comprises quartzites and
(Fig. 3). metagreywackes that are intruded by granite. Both
The El Rodeo Formation consists of greens- units record correlative deformational histories
chist facies mafic, conglomeratic and volcaniclastic involving three phases of deformation that are also
rocks that are intruded by an Early Ordovician recorded in the Ordovician La Noria granite
(476 + 8 Ma) granitoid and unconformably over- (Hinojosa-Prieto et al. 2008). Detrital zircons from
lain by the middle Permian Olinalá Formation. a volcaniclastic epidote-chlorite schist of the El
Detrital zircons from a metavolcaniclastic sand- Epazote Unit yield a normalized relative age
stone of the El Rodeo Formation yield a normalized population with a series of peaks in the intervals
relative age population (Fig. 7a) that defines a broad c. 440–650 and c. 975–1280 Ma, and a minor
peak between c. 990 and c. 1500 Ma centred at peak at c. 1800 Ma (Fig. 7f). Zircons from a meta-
c. 1170 Ma (Talavera-Mendoza et al. 2005). The greywacke of the Las Calaveras Unit yield a normal-
youngest zircon at c. 950 Ma or the youngest ized relative age probability with major peaks at
zircon cluster at c. 988 Ma, and the c. 476 Ma age c. 470, 1040 and 1225 Ma, and a minor peak at
of the intrusive granite, provide maximum and c. 1760 Ma (Fig. 7e). The youngest concordant
minimum constraints, respectively, on the for- zircons at c. 486 Ma (Early Ordovician; Gradstein
mation’s depositional age. et al. 2004) and c. 452 Ma (Late Ordovician; Grad-
The Amate Unit comprises low-grade arkoses, stein et al. 2004), respectively, and youngest zircon
psammites and pelites that are intruded by mafic clusters at c. 506 Ma (late Cambrian; Gradstein
dykes dated at 442 + 1 Ma (Keppie et al. 2008b) et al. 2004) and 466 Ma (Middle Ordovician; Grad-
and granitoid dykes in which the youngest concor- stein et al. 2004), provide maximum limits on the
dant 206Pb/238U zircon ages are 452 + 6 and depositional ages of the two units.
447 + 3 Ma (Morales-Gámez et al. 2008). A felds-
pathic metapsammite from this unit yielded detrital Devonian –Carboniferous low-grade metasedimen-
zircons with a normalized relative age population tary rocks. LA-ICPMS U–Pb detrital zircon age
that shows a series of peaks between c. 950 and data have been obtained from low-grade, Devonian–
c. 1300 Ma (Fig. 7b). The youngest concordant Carboniferous siliciclastic rocks near the village of
zircon gives an age of c. 900 Ma. Mimilulco (Cosoltepec Formation) in the northern
Fig. 6. Normalized relative age probability plot of U–Pb zircon analyses from Ordovician granitoids of the Acatlán
Complex: (a) compilation of all zircon data and (b) compilation of xenocrystic zircon analyses (all those with ages
550 Ma). Data from Sánchez-Zavala et al. (2004), Talavera-Mendoza et al. (2005), Miller et al. (2007), Vega-Granillo
et al. (2007) and Morales-Gámez et al. (2008). Shaded bar outlines the age range characteristic of the Mesoproterozoic
basement of the Oaxaquia terrane (c. 920–1350 Ma: Keppie et al. 2001, 2003; Solari et al. 2003; Cameron et al. 2004).
n ¼ sample size.
Fig. 7. Normalized relative age probability plots of U– Pb detrital zircon analyses from pre-Silurian low-grade
metasedimentary units of the Acatlán Complex: (a) El Rodeo Formation at Olinalá (Talavera-Mendoza et al. 2005);
(b) Amate Unit at Xayacatlán (Morales-Gámez et al. 2008); (c) Canoas Unit near Olinalá (Grodzicki et al. 2008);
(d) Huerta Unit at La Huerta (Keppie et al. 2006a); (e) El Epazote Unit near La Noria (Hinojosa-Prieto et al. 2008); and
(f) Las Calaveras Unit near La Noria (Hinojosa-Prieto et al. 2008). Shaded bars outline the age ranges characteristic of
the Ordovician megacrystic granitoids of the Acatlán Complex (c. 440–480 Ma: Sánchez-Zavala et al. 2004;
Talavera-Mendoza et al. 2005; Miller et al. 2007; Vega-Granillo et al. 2007) and the Mesoproterozoic basement of the
Oaxaquia terrane (c. 920– 1350 Ma: Keppie et al. 2001, 2003; Solari et al. 2003; Cameron et al. 2004). n ¼ sample size.
Acatlán Complex, at Cosoltepec (Cosoltepec For- a within-plate geochemical signature (Grodzicki

mation) and Xayacatlán (Salada Unit) in the east- et al. 2008). The unit is exposed near Olinalá
central part, and near Olinalá (Coatlaco Unit) in where it is in tectonic contact with the pre-Silurian
the southwestern part of the complex (Fig. 3). low-grade Canoas Unit, but records only two
The Cosoltepec Formation comprises pelites, phases of penetrative deformation compared to the
quartzites and minor mafic volcanic rocks that three or more phases present in the Canoas Unit,
record three phases of penetrative deformation suggesting an originally unconformable relation-
(e.g. Malone et al. 2002). As originally mapped ship. Detrital zircons from a quartzite of the
(Ortega-Gutiérrez 1975), the formation included Coatlaco Unit (Grodzicki et al. 2008) yield a nor-
all low-grade pelite-psammite assemblages with malized relative age distribution with peaks in the
this record of deformation and formed the most interval c. 310– 370 Ma, major peaks at c. 535 and
widely distributed unit within the complex (Fig. 3). 575 Ma, a series of peaks between c. 780 and
However, detrital zircon data (Talavera-Mendoza c. 1200 Ma, and numerous minor peaks in the
et al. 2005; Keppie et al. 2006b; Morales-Gámez broad interval c. 1760–2670 Ma (Fig. 8d). The
et al. 2008) have since revealed the existence of youngest zircon at c. 308 Ma (Pennsylvanian;
polydeformed pelite-psammite units of at least two Gradstein et al. 2004) and the youngest zircon
separate ages. The term Cosoltepec Formation is cluster at c. 357 Ma (late Devonian; Gradstein
used here only for pelite-psammite units of et al. 2004) provide maximum constraints on the
Devonian– Carboniferous age, like those of the timing of the unit’s deposition.
type-area at Cosoltepec.
Detrital zircons from a quartzite of the Cosole- Carboniferous–Permian low-grade metasedimen-
pec Formation sampled near Mimilulco (Talavera- tary rocks. LA-ICPMS U –Pb detrital zircon age
Mendoza et al. 2005) yield a normalized relative data for Carboniferous–Permian units of the
age population with peaks at c. 340 and 410 Ma, a Acatlán Complex have been obtained from low-
series of major peaks in the interval c. 500 – grade siliciclastic rocks near the village of Acatlán
700 Ma, and minor peaks at c. 985 and 2190 Ma (Tecomate Formation) in the centre of the complex
(Fig. 8a). Zircons from a quartzite of the Cosoltepec and at Olinalá (Olinalá Formation) near its western
Formation sampled in the type-area (Talavera- margin (Fig. 3). The Patlanoaya Formation, which
Mendoza et al. 2005) yield a similar population is exposed near the northern margin of the
with a peak at c. 395 Ma, a series of major peaks complex, has not been sampled for detrital zircons
in the interval c. 540 –620 Ma, and a series of but comprises fossiliferous shales, sandstones,
minor peaks in the intervals c. 1000–1360 and conglomerates and limestones that range in age
c. 1700–1960 Ma (Fig. 8b). Maximum age con- from latest Devonian (Strunian) to early Permian
straints on the timing of deposition of the formation (Leonardian) (Vachard et al. 2000; Vachard &
are provided by the youngest detrital zircons in each Flores de Dios 2002).
sample, which record ages of c. 340 Ma (Mississip- The Tecomate Formation comprises low-grade,
pian; Gradstein et al. 2004) and c. 375 Ma (late pelitic and volcanosedimentary psammitic rocks,
Devonian; Gradstein et al. 2004), respectively. conglomerates and marbles that, in the type
The youngest zircon cluster in each sample record section south of the town of Acatlán (Fig. 3), non-
an age of c. 410 Ma (early Devonian; Gradstein conformably overlie a deformed megacrystic grani-
et al. 2004). toid of presumed Ordovician age (Sánchez-Zavala
The Salida Unit consists of greenschist facies et al. 2004). The formation is mildly to strongly
psammites, pelites and thin tholeitiic mafic slices deformed, recording up to two phases of penetrative
tectonically juxtaposed against the pre-Silurian low- deformation (e.g. Malone et al. 2002), and is
grade Amate Unit in the vicinity of Xayacatlán stratigraphically similar to the essentially unde-
(Fig. 3). Detrital zircons from a metapsammite in formed Patlanoaya Formation. Two marble hor-
this unit (Morales-Gámez et al. 2008) yield a nor- izons towards the top of the type section have
malized relative age distribution with a series of yielded conodonts with age ranges close to the Car-
peaks in the interval c. 355 –625 Ma, a peak at boniferous–Permian and early– middle Permian
895 Ma, and a series of minor peaks in the intervals boundaries, respectively (Keppie et al. 2004a).
c. 1070–1300 and c. 1765–1985 Ma (Fig. 8c). An East of the town of Acatlán, the formation is over-
older limit on the time of deposition of the Salada thrust by the Totoltepec pluton, the early Permian
Unit is provided by the c. 350 Ma age (early Missis- ID-TIMS U –Pb age of which (287 + 2 Ma,
sippian; Gradstein et al. 2004) of the youngest con- Yañez et al. 1991; 289 + 1 Ma, Keppie et al.
cordant detrital zircon. 2004b) matches the SHRIMP U – Pb ages of
The Coatlaco Unit consists of interbedded quart- granitoid pebbles sampled from a conglomerate
zite and locally pillowed, greenschist facies meta- of the Tecomate Formation west of Cosoltepec
basalts that show sub-alkaline tholeiitic affinity and (c. 280–310 Ma; Keppie et al. 2004a).
Fig. 8. Normalized relative age probability plots of U– Pb detrital zircon analyses from Devonian –Carboniferous
low-grade metasedimentary units of the Acatlán Complex: (a) Cosoltepec Formation at Mimiluco (Talavera-Mendoza
et al. 2005); (b) Cosoltepec Formation at Cosoltepec (Talavera-Mendoza et al. 2005); (c) Salida Unit at Xayactalán
(Morales-Gámez et al. 2008); and (d) Coatlaco Unit near Olinalá (Grodzicki et al. 2008). Shaded bars outline the
age ranges characteristic of the Ordovician megacrystic granitoids of the Acatlán Complex (c. 440–480 Ma:
Sánchez-Zavala et al. 2004; Talavera-Mendoza et al. 2005; Miller et al. 2007; Vega-Granillo et al. 2007) and the
Mesoproterozoic basement of the Oaxaquia terrane (c. 920–1350 Ma: Keppie et al. 2001, 2003; Solari et al. 2003;
Detrital zircon age data are available for seven (174 + 1 Ma) age (Keppie et al. 2004b). The grada-
low-grade recrystallized sandstones of the Teco- tionally underlying Magdalena Migmatite is a
mate Formation sampled across the type section similar lithological unit with additional calcsilicate
south of Acatlán (Sánchez-Zavala et al. 2004). Col- and marble lenses, and was pervasively migmatized
lectively, these yield a normalized relative age dis- during the same Jurassic (c. 175–170 Ma; Keppie
tribution with a prominent peak at c. 465 Ma, a et al. 2004b) tectonothermal event. Both units are
broad Mesoproterozoic population in the interval interpreted to represent portions of a clastic wedge
c. 1000–1300 Ma, and a minor peak at c. 1520 Ma deposited ahead of thrust faults developed along
(Fig. 9a). No systematic variation in detrital age the active margin of the palaeo-Pacific during the
populations exists with stratigraphic level, but indi- Permo –Traissic (Keppie et al. 2006a).
vidual sandstones are dominated by either Ordovi- Data are available for two samples of the Magda-
cian (c. 460– 480 Ma) or Mesoproterozoic– early lena Migmatite palaeosome, both collected south of
Neoproterozoic (c. 900 –1300 Ma) zircons, presum- Magdalena, and three samples of the Chazumba For-
ably in response to varying drainage patterns. The mation, one collected from the structural base of the
dominance of Carboniferous –Permian granitoid succession, near Magdalena, and the others from the
cobbles in the conglomerate west of Cosoltepec structural top to the north (Talavera-Mendoza et al.
further documents the influence of local drainage 2005; Keppie et al. 2006a). The two Magdalena
on the age populations of the formation’s detrital Migmatite samples – a biotite–hornblende schist
zircons. Older age limits for the deposition of the and a biotite–garnet– amphibole schist – yield nor-
Tecomate Formation are provided by the youngest malized relative age populations with peaks (1) in
concordant zircon at c. 450 Ma and the youngest the intervals c. 315–370, 500– 650, 800–1200 Ma
zircon cluster at c. 460 Ma (Late Ordovician; and a c. 1865 Ma (Fig. 10a; Talavera-Mendoza
Gradstein et al. 2004). et al. 2005); and (2) at c. 310, 890 and 1000 Ma, in
The Olinalá Formation is a marine succession in the interval c. 1100–1200 Ma, and at c. 1590 Ma
the southwestern Acatlán Complex that, like the (Fig. 10b; Keppie et al. 2006a), respectively. The
Tecomate Formation, unconformably overlies older youngest detrital zircons at c. 303 Ma (late Penn-
metasedimentary units. The formation is floored by sylvanian; Gradstein et al. 2004) and c. 244 Ma
a massive conglomerate containing quartzite and (earliest Middle Triassic; Gradstein et al. 2004),
schist detritus, and includes fossiliferous black respectively, and youngest zircon cluster at
shales, estuarine sandstones and locally stromatolitic c. 317 Ma (earliest Pennsylvanian; Gradstein et al.
limestones that have been dated as middle Permian 2004), provide maximum limits on the depositional
on the basis of ammonoids, brachiopods and age of the protolith.
fusilinids (González-Arreola et al. 1994; Vachard The three Chazumba Formation samples – a
et al. 2004). biotite–sillimanite schist from the base of the suc-
Detrital zircons from a quartz-rich calcareous cession and two biotite–muscovite –garnet schists
sandstone collected from the type section of the from near the top – yield normalized relative age
Olinalá Formation, east of the town of Olinalá populations with peaks (1) at c. 275, 300 and
(Talavera-Mendoza et al. 2005), yield a normalized 740 Ma, in the intervals c. 850–1000 and 1095–
relative age distribution with a major peak at 1230 Ma, and at c. 1465 Ma (Fig. 10c; Talavera-
c. 300 Ma, a minor peak at c. 825 Ma, and a broad Mendoza et al. 2005); (2) at c. 300 Ma, 590 Ma and
Mesoproterozoic population in the interval in the intervals c. 900– 1000 and 1120– 1240 Ma
c. 1100–1350 Ma (Fig. 9b). The youngest concor- (Fig. 10d; Talavera-Mendoza et al. 2005); and (3)
dant zircon at c. 290 Ma and the youngest zircon at c. 240 Ma and in the interval c. 925–1130 Ma
cluster at c. 297 Ma (early Permian; Gradstein (Fig. 10e; Keppie et al. 2006a), respectively. The
et al. 2004) provide maximum constraints on the youngest detrital zircons at c. 274 Ma (mid-Permian;
formation’s depositional age. Gradstein et al. 2004), c. 249 Ma (Early Triassic;
Gradstein et al. 2004) and c. 239 Ma (Middle
Permian-Triassic high-grade metasedimentary Triassic; Gradstein et al. 2004), respectively, and
rocks. LA-ICPMS U –Pb detrital zircon age data youngest zircon clusters at c. 275 Ma (mid-Permian;
have been obtained from both of the high-grade Gradstein et al. 2004) and 301 Ma (late Pennsylva-
metasedimentary units of the Permian–Triassic nian; Gradstein et al. 2004), provide maximum age
Petlalcingo Suite (Magdalena Migmatite and Cha- constraints on the formation’s deposition.
zumba Formation) near the village of Magdalena
in the eastern Acatlán Complex (Fig. 3). The Granjeno Schist
Chazumba Formation consists of repeatedly defor-
med, amphibolite facies psammites and pelites that The Granjeno Schist of northeastern Mexico’s
were metamorphosed during the Jurassic and Sierra Madre terrane (Fig. 1) comprises repeat-
contain several mafic-ultramafic lenses of Jurassic edly deformed, pelitic metasedimentary and
Fig. 9. Normalized relative age probability plots of U– Pb detrital zircon analyses from Carboniferous–Permian
low-grade metasedimentary rocks of the Acatlán Complex: (a) Tecomate Formation south of Acatlán (Sánchez-Zavala
et al. 2004) and (b) Olinalá Formation at Olinalá (Talavera-Mendoza et al. 2005). Shaded bars outline the age
ranges characteristic of the Totoltepec pluton (280– 310 Ma: Yañez et al. 1991; Keppie et al. 2004a, b) and the
Ordovician megacrystic granitoids (c. 440– 480 Ma: Sánchez-Zavala et al. 2004; Talavera-Mendoza et al. 2005; Miller
et al. 2007; Vega-Granillo et al. 2007) of the Acatlán Complex, and the Mesoproterozoic basement of the Oaxaquia
terrane (c. 920–1350 Ma: Keppie et al. 2001, 2003; Solari et al. 2003; Cameron et al. 2004). n ¼ sample size.
Fig. 10. Normalized relative age probability plots of U– Pb detrital zircon analyses from Permian– Triassic high-grade
metasedimentary rocks of the Acatlán Complex (Petlalcingo Suite) near Magdalena: (a) Magdalena protolith
(Talavera-Mendiza et al. 2005); (b) Magdalena protolith (Keppie et al. 2006a); (c) base of Chazumba Formation
(Talavera-Mendiza et al. 2005); (d) top of Chazumba Formation (Talavera-Mendiza et al. 2005); and (e) Chazumba
Formation (Keppie et al. 2006a). Shaded bars outline the age ranges characteristic of the Totoltepec pluton
(280–310 Ma: Yañez et al. 1991; Keppie et al. 2004a, b) and the Ordovician megacrystic granitoids (c. 440– 480 Ma:
Sánchez-Zavala et al. 2004; Talavera-Mendoza et al. 2005; Miller et al. 2007; Vega-Granillo et al. 2007) of the Acatlán
Complex, and the Mesoproterozoic basement of the Oaxaquia terrane (c. 920– 1350 Ma: Keppie et al. 2001, 2003;
Solari et al. 2003; Cameron et al. 2004). n ¼ sample size.
Fig. 11. Normalized relative age probability plots of U –Pb detrital zircon analyses (a) from the Granjeno Schist of the
Sierra Madre terrane (data from Nance et al. 2007); (b) from all samples of the pre-Silurian low-grade units of the
Acatlán Complex (Talavera-Mendoza et al. 2005; Keppie et al. 2006a; Grodzicki et al. 2008; Hinojosa-Prieto et al.
2008; Morales-Gámez et al. 2008); and (c) from all samples of the Devonian-Carboniferous low-grade units of the
Acatlán Complex (Talavera-Mendoza et al. 2005; Grodzicki et al. 2008; Morales-Gámez et al. 2008). Shaded bars
outline the age ranges characteristic of the Ordovician megacrystic granitoids of the Acatlán Complex (c. 440 –480 Ma:
Sánchez-Zavala et al. 2004; Talavera-Mendoza et al. 2005; Miller et al. 2007; Vega-Granillo et al. 2007) and the
Mesoproterozoic basement of the Oaxaquia terrane (c. 920–1350 Ma: Keppie et al. 2001, 2003; Solari et al. 2003;
256
R. D. NANCE ET AL.
Fig. 12. U –Pb detrital zircon ages from (meta)sedimentary rocks of the Oaxaquia, Mixteca (Acatlán Complex) and Sierra Madre (Granjeno Schist) terranes, the Oaxaquia
basement (Oaxacan Complex and Novillo Gniess), and the megacrystic granitoids of the Acatlán Complex (see text for sources) compared to the cratonic age provinces of eastern
Laurentia (Cawood et al. 2001, 2007), Baltica (Gower et al. 1990; Roberts 2003), the Amazon craton (Sadowski & Bettencourt 1996; Tassinari & Macambira 1999; Santos et al. 2000;
Santos 2003) and the West Africa craton (Rocci et al. 1991; Boher et al. 1992; Potrel et al. 1998; Hirdes & Davis 2002). Duplicate samples (Cosoltepec, Magdalena and Chazumba) are
listed in same order as presented in text. Among the detrital zircons, red bars identify zircon age ranges, whereas red dots show ages of individual zircons or zircon pairs.
metavolcaniclastic rocks that enclose lenses of derived (Gillis et al. 2005). Similarly, the youngest
serpentinite-metagabbro. This low-grade assem- (c. 340–360 Ma) zircons in the Ixtaltepec For-
blage has been correlated with the pre-Silurian mation record ages that coincide with the high-
units (formerly the Cosoltepec Formation) of the pressure metamorphism and exhumation of the
Acatlán Complex on the basis of lithologic similarity pre-Silurian high-grade units (Piaxtla Suite) of the
(Ortega-Gutiérrez 1978; Ramı́rez-Ramı́rez 1992), Acatlán Complex (c. 346 Ma; Middleton et al.
deformational history (Dowe et al. 2005) and det- 2007) and so may also have been derived locally.
rital zircon ages (Nance et al. 2007b) and, like However, a more distal source in the magmatic arc
the Acatlán Complex, is tectonically juxtaposed above the subduction zone from which the Piaxtla
against c. 1 Ga granulite facies gneisses (Novillo Suite was exhumed cannot be excluded.
Gneiss), the unmetamorphosed middle Silurian sedi-
mentary cover of which contains fauna of Gondwa- Acatlán Complex: Pre-Silurian high-grade
nan affinity (Boucot et al. 1997; Stewart et al. 1999). metasedimentary rocks
LA-ICPMS U –Pb detrital zircon age data
obtained from a single sample of phyllite from the Based on their detrital zircon age populations
Granjeno Schist yielded a normalized relative (Fig. 5), two separate sets of pre-Silurian high-
age population (Fig. 11a) with major peaks in the pressure rocks (Piaxtla Suite) have been sampled
intervals c. 450 –650 and 880 –1330 Ma (Nance from the Acatlán Complex – those that are intruded
et al. 2007b). The youngest detrital zircon at by, and so host, the Ordovician megacrystic grani-
433 + 14 Ma (early Silurian; Gradstein et al. toids (Xayacatlán Formation and Asis Lithodeme)
2004) and youngest zircon cluster at c. 453 Ma and those that broadly coincide with granitoid
(Late Ordovician; Gradstein et al. 2004) indicate a emplacement (Ixcamilpa blueschist). The first set
maximum depositional age for the Granjeno Schist. shows normalized relative age probabilities
(Fig. 5a –c) that closely match that of the latest
Cambrian –Ordovician Tiñu Formation overlying
the Oaxacan Complex (Fig. 2a). None of these
Zircon Provenance units contains Palaeozoic zircons and all are domi-
Oaxacan Complex nated by detrital zircons of Mesoproterozoic–
early Neoproterozoic age, the likely provenance
Deposition of the Tiñu, Santiago and Ixtaltepec for- for which is the c. 920–1350 Ma basement of the
mations, which form the sedimentary cover to the Oaxacan Complex. In addition, with the exception
Oaxacan Complex, spans an interval of almost of the late Neoproterozoic peak, the relative age fre-
200 Myr (latest Cambrian –middle Pennsylvanian). quencies of these units closely match the xenocrys-
Yet they are characterized by similar detrital zircon tic age populations of the megacrystic granitoids
age spectra dominated by zircons of Mesoprotero- (Fig. 6b). These correspondences in zircon ages
zoic –early Neoproterozoic (c. 980 –1080 Ma and/ suggest that the Oaxacan Complex provided an
or 1100–1240 Ma) age (Fig. 2). These data important source for both the sedimentary protoliths
suggest that the cover sequence records very little of the Piaxtla Suite and the Ordovician granitoid
change in provenance for much of the Palaeozoic. magmas. Potential cratonic sources for those
The most obvious source for these zircons is the zircons lying outside the age range represented in
underlying basement rocks of the Oaxacan the Oaxacan Complex are present in eastern Lauren-
Complex (Gillis et al. 2005), the granulite facies tia, Baltica and the Amazon craton but, contrary to
gneisses and AMCG suites of which have yielded the eastern Laurentian provenance proposed by
ages of c. 980– 1020 and 1100– 1265 Ma (Keppie Talavera –Mendoza et al. (2005), the presence of
et al. 2001, 2003; Solari et al. 2003). The Palaeozoic both xenocrystic and detrital zircons with ages in
detrital zircon clusters that peak at c. 470 Ma in the the ranges c. 700–900 Ma (which is not preserved in
Santiago and Ixtaltepec formations match the age of eastern Laurentia and Baltica but present in South
the megacrystic granitoids in the Acatlán Complex America) and c. 1500–1600 Ma (which is not pre-
(c. 440–480 Ma) from which they were likely served in eastern Laurentia but present in the
Fig. 12. (Continued) Purple bars and dots identify xenocrystic zircons in the megacrystic granitoids of the Acatlán
Complex. Horizontal light blue bars outline the age ranges characteristic of the Totoltepec pluton (280–310 Ma: Yañez
et al. 1991; Keppie et al. 2004a, b) and the Ordovician megacrystic granitoids (c. 440– 480 Ma: Sánchez-Zavala et al.
2004; Talavera-Mendoza et al. 2005; Miller et al. 2007; Vega-Granillo et al. 2007) of the Acatlán Complex, and the
Mesoproterozoic basement of the Oaxaquia terrane (c. 920– 1350 Ma: Keppie et al. 2001, 2003; Solari et al. 2003;
Cameron et al. 2004). Horizontal yellow bars highlight c. 700–850, 1500– 1600 and 2250–2400 Ma age ranges most
pertinent to the identification of cratonic source areas. DEV–CARB, Devonian– Carboniferous low-grade units;
CARB– P, Carboniferous –Permian low-grade units; ?SIL, possibly Silurian; S–M, Sierra Madre terrane.
Amazon craton) favours an Amazonian source with their emplacement (Canoas, Huerta, El
(Fig. 12). Epazote and Las Calaveras units).
In contrast, the Ixcamilpa blueschist, while con- The El Rodeo Formation and Amate Unit show
taining Mesoproterozoic detrital zircons that match normalized relative age probabilities (Fig. 7a– b)
the age range of the Oaxacan Complex (Fig. 5d), is that closely match those of the older set of high-
dominated by Palaeozoic and Neoproterozoic grade metasedimentary rocks (Xayacatlán For-
zircons with ages that peak at c. 475, 540, 600 and mation, Asis Lithodeme and Quartzite; Fig. 5a–c),
730 Ma (Talavera –Mendoza et al. 2005). The as well as those of the latest Cambrian– Ordovician
youngest of these peaks overlaps the emplacement Tiñu Formation of the Oaxacan Complex (Fig. 2a)
age range of the Ordovician megacrystic granitoids, and the xenocrystic population within the megacrys-
the magmatic episode of which the sample’s proto- tic granitoids (Fig. 6b). In each case, the age spectra
lith may have been part. However, the remaining lack Palaeozoic detrital zircons and are dominated
clusters cannot presently be sourced in either the by those of Mesoproterozoic–early Neoproterozoic
Acatlán or Oaxacan complexes unless the presence age like that of the Oaxacan Complex, which likely
of Neoproterozoic xenocrystic zircons in the Ordo- provided the source. This striking similarity in the
vocian megacrystic granitoids (Fig. 6b) indicates detrital zircon signatures suggests that the pre-
the existence of Neoproterozoic basement beneath Silurian siliciclastic units of the Acatlán and
the Acatlán Complex that may once have been Oaxacan complexes are broadly correlative and rep-
exposed. The likelihood, however, is that these resent portions of a shelf succession deposited on
ages indicate a significant change in provenance. the margin of Oaxaquia (Xayacatlán Formation,
Rocks of this age are present in Baltica only in Asis Lithodeme and Quartzite) that were later sub-
the Timanides on its northern margin (e.g. Roberts jected to varying grades of metamorphism.
& Siedlecka 2002), and were largely missing in With the exception of a cluster of Ordovician
eastern Laurentia prior to the Silurian arrival of (c. 450–475 Ma) detrital zircons likely derived
the peri-Gondwanan terranes (e.g. van Staal et al. from the megacrystic granitoids, the age spectrum
1998), as indicated by the almost complete of the Canoas Unit (Fig. 7c) strongly resembles
absence of detrital zircons of this age in those of the El Rodeo Formation and Amate Unit,
Neoproterozoic–Ordovician strata on the eastern suggesting a similar, largely local provenance.
Laurentian margin (Cawood & Nemchin 2001; Older zircons in the Canoas Unit, like those of the
Cawood et al. 2007). If deposition of the Ixcamilpa El Rodeo Formation, include ages in the range
Unit is as old as c. 477 Ma, a local source for the c. 1500–1600 Ma, which favours an Amazonian
Neoproterozoic–Cambrian detrital zircons may source (Fig. 12).
have existed in peri-Gondwanan Avalonia and Car- The detrital zircon signatures of the Huerte, El
olinia, since these terranes are inferred to have lain Epazote and Las Calaveras units (Fig. 7d–f), on
adjacent to Oaxaquia until the opening of the the other hand, are closely similar to that of the Ixca-
Rheic Ocean in the Early Ordovician (Keppie milpa blueschist (Fig. 5d), showing a broad age span
2004). However, the zircons are considered more dominated by Ordovician (c. 450–490 Ma) and/or
likely to have come from either the Pan-African late Neoproterozoic –Cambrian (c. 510–630 Ma)
basement of the Maya terrane (Fig. 1) beneath the zircons with less conspicuous Mesoproterozoic–
Yucatan Peninsula (Krogh et al. 1993a), or the Bra- early Neoproterozoic populations. Sources for the
siliano orogens and Goias magmatic arc of South Ordovician and Mesoproterozoic zircons exist
America (e.g. Brito Neves et al. 1999; Pimental both locally and in eastern Laurentia. However, a
et al. 2000). Potential sources for the older (Palaeo- late Neoproterozoic –Cambrian source was essen-
proterozoic and Archaean) zircons can likewise be tially absent in eastern Laurentia prior to the accre-
found in the age provinces of Amazonia (Fig. 12). tion of the peri-Gondwanan terranes in the Silurian.
Hence, likely sources for the three age populations
include the Ordovician megacrystic granitoids,
Acatlán Complex: Pre-Silurian low-grade the Pan-African Maya terrane and/or the Brasiliano
metasedimentary rocks belts of South America, and the Oaxacan Complex,
although the possibility of nearby peri-Gondwanan
The detrital zircon age populations of the pre- terranes cannot be excluded. In addition, all three
Silurian low-grade units of the Acatlán Complex units contain a representation of ages in the range
show striking similarities to those of the high-grade c. 750–950 Ma, the only viable source for which
metasedimentary rocks, the data again suggesting is the Goias magmatic arc (Fig. 12), which flanks
the presence of two groups of assemblages – those the Amazon craton in the Tocantins Province of
that pre-date the Ordovician megacrystic granitoids Brazil (e.g. Pimental et al. 2000). Possible sources
(El Rodeo Formation and Amate Unit) and those for the Palaeoproterozoic and Archaean zircons in
whose deposition was broadly contemporaneous these units also exist within the Amazon craton.
Acatlán Complex: Devonian – Carboniferous additional source of zircons of this age. No such
low-grade metasedimentary rocks arc exists within the Acatlán Complex, but evidence
of an arc of this age is preserved in the Ouachita
All of the low-grade Devono–Carboniferous rocks orogenic belt on the southern margin of Laurentia
sampled within the Acatlán Complex – the Cosolte- (e.g. Viele & Thomas 1989).
pec Formation (Fig. 8a, b), the Salida Unit (Fig. 8c)
and the Coatlaco Unit (Fig. 8d) – show similar det- Acatlán Complex: Carboniferous– Permian
rital zircon age spectra (Talavera-Mendoza et al. low-grade metasedimentary rocks
2005; Morales-Gámez et al. 2008), suggesting a
similar provenance for each of these assemblages. Detrital zircons in the two sampled Carboniferous–
Like those of the pre-Silurian Ixcamilpa blueschist Permian low-grade units of the Acatlán Complex –
(Fig. 5d), Huerta Unit (Fig. 7d) and El Epasote the Tecomate and Olinalá formations (Fig. 9) – share
Unit (Fig. 7e), each of the Devonian –Carboniferous a broad Mesoproterozoic (c. 950–1300 Ma) age
units shows a spectrum with a broad span of ages population most likely derived from the Oaxacan
that is dominated by late Neoproterozoic (c. 540– Complex, but differ in their Palaeozoic age popu-
640 Ma) zircons with only subordinate populations lations. The Tecomate Formation is dominated
of Mesoproterozoic–early Neoproterozoic age like by an Ordovician (c. 450– 480 Ma) population of
that of the Oaxacan Complex. zircons that matches the age of the megacrystic gran-
The importance of late Neoproterozoic zircon itoids, whereas the Olinalá Formation lacks this age
ages again points to sources in the Maya terrane group and contains, instead, a population of Permo-
and/or the Brasiliano belts (Keppie et al. 2006a), Carboniferous (c. 280–310 Ma) zircons that
although a local or locally recycled source cannot matches the TIMS U –Pb age of the Totoltepec
be excluded given the matching xenocrystic popu- pluton (c. 290 Ma; Yañez et al. 1991; Keppie et al.
lations in the Ordovician megacrystic granitoids 2004a) and the SHRIMP U –Pb ages of granitoid
(Fig. 6b). Given the maximum depositional age of pebbles sampled from a conglomerate of the
these units, however, the abundance of late Neo- Tecomate Formation west of Cosoltepec (c. 280–
proterozoic detrital zircons could also reflect 310 Ma; Keppie et al. 2004a). In addition, detrital
increasing proximity to the previously accreted peri- zircon age spectra from individual horizons of
Gondwanan terranes of the eastern Laurentian the Tecomate Formation differ significantly, some
margin (Grodzicki et al. 2008), but the units also being dominated by Mesoproterozoic zircons,
contain a significant population of older Neoproter- whereas others are dominated by zircons of Ordovi-
ozoic (c. 750 –900 Ma) zircons, a source for which cian age (Sánchez-Zavala et al. 2004).
is absent in Laurentia and present only in the A provenance for the Mesoproterozoic zircons in
Goias magmatic arc of Amazonia (Fig. 12). the granulite facies Oaxacan Complex rather than the
The Coatlaco Unit, which records a structural megacrystic granitoids, which record inheritance of
history like that of the low-grade Permo- similar age, is supported by the presence of pebbles
Carboniferous units, nevertheless contains detrital of blue-quartz-bearing gneiss in the Tecomate For-
zircons whose age spectrum is closely similar to mation, and the existence in the age-equivalent
those of the structurally more complex Cosoltepec Patlanoaya Formation of abundant granulite facies
Formation and Salida Unit, particularly with grains such as mesoperthite and rutile-bearing
respect to the broad age-span of the spectrum quartz (Sánchez-Zavala et al. 2004). As these
and the importance of detrital zircons of late authors suggest, the data imply that the formation
Neoproterozoic– Cambrian (c. 530 –630 Ma) and was sourced predominantly locally and that vari-
Neoproterozoic (c. 770 –940 Ma) age (Fig. 8d), ations in drainage pattern played an important role
both of which suggest sources in Amazonia (Fig. 12). in controlling the local provenance. Given that its
Relatively few of the Palaeozoic detrital zircons deposition post-dates the assembly of Pangaea,
in the Devonian –Carboniferous units match the age older detrital zircons in the Tecomate Formation
of the megacrystic granitoids, the bulk showing could have been sourced in either Amazonia, Lauren-
younger (c. 340 –410 Ma) ages, like those of the tia or West Africa. However, most show populations
Pennsylvanian Ixtaltepec Formation overlying the with ages in the range c. 1500–1600 Ma, which
Oaxacan Complex. These ages broadly coincide favours an Amazonian source (Fig. 12). The presence
with the exhumation of the pre-Silurian high-grade of a c. 825 Ma population in the Olinalá Formation
units (Piaxtla Suite) within the Acatlán Complex likewise favours a South American provenance.
(c. 346 Ma; Middleton et al. 2007) and so could Hence, the data suggest that the two formations
have been derived locally. However, subduction of are not only of similar age, but also shared a
the high-grade units would have been associated similar, largely local provenance sourced in the
with the development of a Devono –Carboniferous Oaxacan and Acatlán complexes, with possible
magmatic arc, which could have provided an minor contributions of South American origin.
Acatlán Complex: Carboniferous –Permian Neoproterozoic (at c. 590, 740 and 850– 1000 Ma)
high-grade metasedimentary rocks and Late Palaeozoic (at c. 275 and 300 Ma).
However, almost all the zircons in the sample
The relative age probabilities of detrital zircons collected by Keppie et al. (2006a) fall in the range
in the Carboniferous –Permian high-grade units of c. 925–1130 Ma, with a single Triassic grain at
the Acatlán Complex – the Magdalena Migmatite c. 240 Ma. Once again, a likely provenance for the
and the Chazumba Formation – show significant youngest zircon grains lies in the Permo –Triassic
variation, even within the same unit. This variation magmatic arc (Torres et al. 1999) represented by
may reflect the creation of new, lithologically the Totoltepec pluton, whereas those of Mesoproter-
complex source areas as a result of the onset of con- ozoic age were probably sourced from either the
vergent tectonics along the palaeo-Pacific margin. Oaxacan or Acatlán complexes. Older detrital
Detrital zircons from the palaeosome of the zircons are few, but again match age provinces in
Magdalena Migmatite show a broadly similar span both Amazonia and eastern Laurentia. However,
of age populations but quite variable normalized the Neoproterzoic zircon ages best match sources
relative age frequencies (Fig. 10a, b). Both in the Brasiliano belts and Goiás magmatic arc of
samples of the protolith contain significant popu- South America (Fig. 12).
lations of Mesoproterozoic, early Neoproterozoic
(c. 850 –920 Ma), Cambrian (c. 520 –530 Ma) and Granjeno Schist
late Palaeozoic zircons (c. 310 –370 Ma), but
whereas that sampled by Keppie et al. (2006a) is Detrital zircons from the Granjeno Schist (Nance
dominated by c. 1100 –1200 Ma ages like those of et al. 2007b) show a relative age distribution
the older components of the Oaxacan Complex, (Fig. 11a) that resembles those of various units
relatively few zircons of this age occur in the within the Acatlán Complex, with two major popu-
sample collected by Talvera-Mendoza et al. lations in the early Neoproterozoic– Mesoprotero-
(2005). In both cases, the zircons were likely zoic (c. 880–1330 Ma) and late Neoproterozoic –
sourced from the Oaxacan Complex or recycled early Palaeozoic (c. 450–650 Ma). However, with
from older units of the Acatlán Complex, although respect to the pre-Permian low-grade units
a South American provenance is suggested by the (Fig. 11b, c), the most striking match is with those
early Neoproterozoic zircons, the only obvious of pre-Silurian age (formerly the Cosoltopec For-
source for which lies in the Goiás magmatic arc of mation) with which the schist has been traditionally
central Brazil (e.g. Pimental et al. 2000). The correlated (Ortega-Gutiérrez 1978). Despite the
latter sample also contains an important late Neo- slightly younger maximum depositional age of the
proterozoic (c. 650 Ma) population and shows a schist (c. 435 Ma), this correlation is further sup-
broader range of ages that extend into the Palaeopro- ported by its lithologic association (Ramı́rez-
terozoic. Talavera-Mendoza et al. (2005) link the Ramı́rez 1992) and deformational history (Dowe
late Palaeozoic, late Neoproterozoic– Cambrian et al. 2005).
and Palaeoproterozoic zircons to the Alleghanian Potential provenances for the Mesoproterozoic
orogen, the rifting of the Iapetus Ocean and the detrital zircons exist within the Grenville Belt of
Trans-Hudson orogen, respectively, and so argue southern Laurentia and in belts of similar age in Ama-
for a Laurentian provenance for these detrital popu- zonia. However, the most likely provenance is the
lations. This is a plausible source, given the deposi- adjacent Novillo Gneiss, which has yielded ages of
tional age of the Magdalena protolith, but better c. 990–980, c. 1035–1010 and c. 1235–1115 Ma
matches for the Cambrian– Neoproterozoic zircons (Cameron et al. 2004) that are almost identical to
exist within the Ribeira Belt and Borborema Pro- those of the Oaxacan Complex (Fig. 12). Sources
vince of South America (e.g. Machado et al. 1996; for the Neoproterzoic detrital zircons can be found
Brito Neves et al. 2003) and the late Palaeozoic in the Maya terrane and in the Brasiliano belts of
zircons match the age of locally derived boulders Amazonia, but also occur in the rifted margin of
of the Totoltetec granite in the Tecomate Formation Laurentia and in the Appalachian peri-Gondwanan
(Keppie et al. 2004a). Sources for the older zircons terranes, some of which may have been accreted to
in both samples similarly match age provinces in Laurentia prior to the schist’s deposition (e.g. van
both the Amazon craton and Laurentia (Fig. 12). Staal et al. 1998). Sources for the early Palaeozoic
The detrital zircon populations in the Chazumba zircons exist within the Taconian belt of the Appala-
Formation resemble those of the Magdalena proto- chians and the Arequipa–Antofalla terrane on the
lith, and again show similar age ranges but differing southwestern margin of Amazonia, but the most
normalized relative age frequencies (Fig. 10c). The likely source is the c. 480–440 Ma megacrystic gran-
two samples collected by Talavera-Mendoza et al. itoids of the Acatlán Complex (Nance et al. 2007b).
(2005) show similar spectra dominated by peaks Sources for the older zircons exist in both Amazonia
in the Mesoproterozoic (at c. 1100–1250 Ma), and eastern Laurentia (Fig. 12).
Palaeogeographic significance El Rodeo, Amate and Canoas units) and in the

Permo-Carboniferous Tecomate and Olinalá for-
The detrital zircon signatures of Palaeozoic sedimen- mations. However, they are present in all other
tary and metasedimentary rocks in the Oaxaquia, units and constitute an important population in the
Mixteca and Sierra Madre terranes show striking pre-Silurian Ixcamilpa blueschists and Huerta, El
similarities despite the wide range in their likely Epazote and Las Calaveras units, the Devonian –
depositional ages (Fig. 12). The most important Carboniferous low-grade units, and the protolith of
zircon population is of Mesoproterozoic (c. 950– the Magdalena Migmatite. Zircons of this age also
1300 Ma) age. A provenance of this age provided occur in the Granjeno Schist and as inheritance in
virtually all of the detrital zircons in the sedimentary the megacrystic granitoids. Potential proximal
cover of the Oaxacan Complex from the latest sources of zircons of this age may exist in the base-
Cambrian to the Pennsylvanian. Such a provenance ment of the Acatlán Complex and are available in
was also a major source in almost all units of the the Maya terrane of the Yucatan Peninsula
Acatlán Complex into the Triassic, and dominates (Fig. 1). However, prior to the Jurassic opening of
the age distribution in the Granjeno Schist. Potential the Gulf of Mexico, this basement was contiguous
sources for detrital zircons of this age exist within with that of the Suwannee terrane in Florida (e.g.
the Grenville Belt of eastern Laurentia (e.g. Tollo Pindell et al. 2000; Dickinson & Lawson 2001),
et al. 2004; Cawood et al. 2007), in the ‘Grenville’ the accretion of which did not occur until the Alle-
massifs of the northern Andes (e.g. Cediel et al. ghanian at c. 300 Ma (e.g. Heatherington et al.
2003), and in the Sunsas Belt of the Amazon craton 1996). Hence, this source is unlikely to have been
(e.g. Boger et al. 2005). However, the most likely available prior to the late Palaeozoic assembly of
source is exposed portions of the basement of the Pangaea. More distal sources of such zircons exist
Oaxaquia terrane, represented today by the in the Brasiliano belts of South America (e.g.
Oaxacan Complex and Novillo Gneiss, available Brito Neves et al. 2003) and, following the Silurian
U –Pb ages for which span the interval c. 920– closure of the Iapetus Ocean (e.g. van Staal et al.
1350 Ma (Keppie et al. 2003; Cameron et al. 1998), in the accreted peri-Gondwanan terranes of
2004). That this basement source, which also eastern Laurentia. In those (meta)sedimentary
includes other c. 1 Ga inliers in Mexico (Huiznopala units whose deposition pre-dates the Silurian, how-
Gneiss and Guichicovi Gneiss) with similar age ever, a Gondwanan (peri-Amazonian) source would
ranges (Lawlor et al. 1999; Weber & Köehler appear to be required.
1999), is of Gondwanan rather than Laurentian A peri-Amazonian source is also suggested by
affinity as indicated by the Gondwanan fauna of detrital zircons of c. 750–950 Ma (early–middle
the early Palaeozoic cover sequences of both the Neoproterozoic) age, which are present in virtually
Oaxacan Complex and the Novillo Gneiss (e.g. all Palaeozoic units of the Mixteca and Sierra
Stewart et al. 1999; Landing et al. 2007). The signifi- Madre terranes (Fig. 12). Provenances of this age
cant number of studied (meta)sedimentary units that range, which post-dates the Grenvillian orogeny
contain Ordovician detrital zircons, which match the (e.g. Rivers 1997), are absent in eastern Laurentia,
c. 440– 480 Ma age of the megacrystic granitoids of Baltica and West Africa. However, ages in this
the Acatlán Complex, and Permo–Carboniferous det- range are characteristic of the c. 600–930 Ma
rital zircons, which match the age of the c. 290 Ma Goias magmatic arc of central Brazil (Pimental
Totoltepec pluton, further attests to the importance et al. 2000).
of local sources throughout the Palaeozoic. Likewise, the populations of mid-Proterozoic
However, the detrital zircon signatures of the and older (,1300 Ma) detrital zircons in both the
(meta)sedimentary rocks also contain important Acatlán Complex and the Granjeno Schist most
age populations that are not represented by local sou- closely match the age provinces of the Amazon
rces and so bear upon the palaeogeography of the craton (Fig. 12), suggesting an Amazonian prove-
Oaxaquia, Mixteca and Sierra Madre terranes at the nance. This is indicated by the presence of detrital
time of their deposition. Chief among these are popu- zircons with ages in the range c. 1300 –1900 Ma
lations of c. 500–650 Ma (late Neoproterozoic– (absent in West Africa), c. 1900–2250 Ma ages
Cambrian), c. 750– 950 Ma (early Neoproterozoic) (absent in Baltica) and c. 1500–1600 and 2000–
and .1300 Ma (Mesoproterozoic and older) age. 2100 Ma ages (absent in Laurentia), and by the
Detrital zircons of late Neoproterozoic– paucity of detrital zircons with ages in the range
Cambrian age are absent in the latest Cambrian– c. 2250– 2550 Ma (absent in Amazonia). In
Ordovician, Mississippian and Pennsylvanian support of this, Bream et al. (2004) have shown in
sedimentary cover of the Oaxacan Complex, in the southern Appalachians that basements of Ama-
metasedimentary rocks of the Acatlán Complex zonian and Laurentian affinity can be distinguished
that pre-date emplacement of the megacrystic gran- on the basis of zircon inheritance. Whereas zircon
itoids (Xayacatlán Formation and Asis, Quartzite, inheritance from a Laurentian source is typically
no older than c. 1300–1350 Ma, inheritance from an granitoids) are attributed to east-directed subduction
Amazonian source is distinguished by the frequency beneath a Laurentia-derived arc terrane floored by
of 1600–2100 Ma and, to a lesser degree, 2700– Grenville basement that is accreted to Laurentia in
2900 Ma age ranges, both of which are well rep- the early Ordovician; (2) blueschist metamorphism
resented in the detrital zircon spectra from the (Ixcamilpa blueschists) and renewed arc magma-
Mixteca and Sierra Madre terranes. Cawood et al. tism (c. 440–460 Ma megacrystic granitoids) are
(2007) have likewise shown that sedimentary attributed to east-directed subduction beneath an
basins of Laurentian provenance in the southern ensimatic arc that is accreted to Laurentia in the
Appalachians are essentially devoid of detrital late Ordovician–early Silurian; and (3) thrusting
zircons much older than those of Grenville age. of these previously accreted elements over the
Cosoltopec Formation (following its deposition at
Discussion the leading edge of South America) is attributed to
west-directed subduction beneath Laurentia and
Available detrital zircon data for the Cambrian the tectonic transfer of these elements from the
through Permo–Triassic (meta)sedimentary rocks upper plate to the lower plate during Laurentia-
of the Oaxaquia, Mixteca and Sierra Madre terranes South America collision in the late Pennsylvanian.
suggests a preponderance of local and/or Amazo- The oceanic domains whose closure is recorded by
nian sources throughout the Palaeozoic. This con- each of these collisional events are not specified,
clusion concurs with that of Gillis et al. (2005) for but from the perspective of the eastern Laurentian
the sedimentary cover of the Oaxacan Complex, margin, they would correspond to the Baie Verte
with that of Talavera-Mendoza et al. (2005) and seaway, and the Iapetus and Rheic oceans, respect-
Keppie et al. (2006a) for the low-grade Devonian – ively (van Staal et al. 1998; Martı́nez Catalán
Carboniferous and high-grade Permian–Triassic et al. 2002). According to their model, therefore,
units of the Acatlán Complex, and with that of the Acatlán Complex records the closure of three
Nance et al. (2007b) for the Granjeno Schist. separate oceanic tracts.
However, it is contrary to the view put forward by However, we would argue that the detrital zircon
Talavera-Mendoza et al. (2005), and adopted by signatures of pre-Silurian units in the Acatlán
Vega-Granillo et al. (2007), with regard to the prove- Complex, and the Ixcamilpa blueschist in particular,
nance of the pre-Silurian high-grade units of the are more consistent with local and Amazonian
Acatlán Complex. These authors interpret the detrital (rather than Laurentian) sources given that: (1)
zircon signatures of the Xayacatlán and Ixcamilpa c. 800–950 Ma ages are rare in Laurentia but charac-
units to indicate a Laurentian provenance for this teristic of the Goias magmatic arc of central Brazil
component of the Acatlán Complex, and likewise (Pimental et al. 2000); (2) c. 1150–1300 Ma ages
proposed a Laurentian source for the El Rodeo For- can be found in the Oaxacan Complex (Solari et al.
mation, the depositional age constraints for which 2003); (3) c. 480 Ma ages occur in the megacrystic
are the same as those for the Xayacatlán Formation. granitoids (Sánchez-Zavala et al. 2004; Talavera-
Their interpretation is based on the existence in the Mendoza et al. 2005); and (4) c. 600 and 700 Ma
Xayacatlán and El Rodeo formations of c. 800 – ages are typical of the peri-Gondwana terranes and
950 Ma and .1260 Ma age populations, which are rare in Laurentia until the accretion of these terranes
rare or absent in the neighbouring Oaxacan in the Silurian (van Staal et al. 1998). In addition,
Complex, and c. 1150–1300 Ma ages, which are several of the pre-Silurian units, including the El
rare in Amazonia but common in the southwestern Rodeo Formation, contain detrital zircons with
Grenville Province of North America. They also ages in the range c. 1500 –1600 Ma, which can be
link the Ordovician (c. 480 Ma), Neoproterozoic linked to Amazonia but not to Laurentia (Fig. 12).
(c. 600 and 710 Ma) and Palaeoproterozoic If the pre-Silurian high-grade units (Piaxtla
(c. 1800 Ma) zircon populations in the Ixcamilpa Suite) of the Acatlán Complex are of Amazonian
blueschist to age provinces in eastern Laurentia. rather than Laurentian provenance and their eclogite
Accordingly, for the Acatlán Complex, these facies metamorphism, which has been dated as
authors propose contrasting (Laurentian and Amazo- late Devonian –Mississippian (Middleton et al.
nian) provenances for the pre-Silurian high-grade 2007, Elı́as-Herrera et al. 2007), is divorced from
units (Piaxtla Suite) and Devonian– Carboniferous the Ordovician emplacement of the megacrystic
low-grade units (Cosoltepec Formation). granitoids, a much simpler model for the Palaeozoic
Based on this interpretation, Talavera-Mendoza evolution of the Acatlán Complex than that pro-
et al. (2005) proposed a complex model for the tec- posed by Talavera-Mendoza et al. (2005) can be
tonothermal evolution of the Acatlán Complex in advanced (Fig. 13). This model (e.g. Nance et al.
which: (1) eclogite facies metamorphism (Xayaca- 2006, 2007a), which can be used to link the Oaxa-
tlán Formation) and rifted arc magmatism (El quia, Mixteca and Sierra Madre terranes, is analo-
Rodeo Formation and c. 470 –480 Ma megacrystic gous to that originally proposed for the Acatlán
Fig. 13. Plate tectonic interpretation of the Acatlán Complex showing: (a) late Devonian–Mississippian subduction
and exhumation of the pre-Silurian high-grade rocks of the Acatlán Complex (Piaxtla Suite), an oceanic and rift-related
continental margin assemblage (Tehuitzingo ophiolite from Proenza et al. 2004), onto the Gondwanan continental
rise deposits of the low-grade pre-Silurian and Devono-Carboniferous pelite-psammite assemblages during closure
of the Rheic Ocean between Oaxaquia and Laurentia; (b) early Permian development of the Totoltepec arc as a result of
subduction of the palaeo-Pacific following the amalgamation of Pangaea, and deposition and deformation of the
Carboniferous-Permian low-grade units of the Acatlán Complex coincident with dextral strike-slip faulting that
juxtaposed the Acatlán and Oaxacan complexes; and (c) Middle Triassic development of the clastic wedge represented
by the Permian-Triassic high-grade units of the Acatlán Complex (Petlalcingo Suite) in response to continued
subduction of the palaeo-Pacific. S, south; A, away; T, toward (modified from Nance et al. 2007b).
Complex by Ortega-Gutierrez et al. (1999), but tholeiitic (Keppie et al. 2008b) affinity, are inter-
views this complex (and the Granjeno Schist), not preted to record the development of a rift/passive
as a remnant of the Iapetus Ocean, but as a vestige margin on the southern (Oaxaquia) flank of the Rheic
of its immediate successor, the Rheic Ocean, that Ocean in the ?Cambro– Ordovician. The broadly
was subsequently overprinted by convergent tec- coeval megacrystic granitoids geochemically strad-
tonics at the margin of the palaeo-Pacific. dle the volcanic arc/within plate fields (Reyes-Salas
Accordingly, the pre-Silurian units, which 2003; Miller et al. 2007; Hinojosa-Prieto et al.
include quartzites, pelites and mafic rocks of both 2008); however, they are synchronous with the
MORB (Meza-Figueroa et al. 2003) and continental intrusion of mafic rocks and an Ordovician,
tholeiitic rift-related, dyke swarm (Morales-Gámez

et al. 2008). Thus the magmatism is bimodal and
likely formed a rift-related suite that developed
along the northern margin of Gondwana during the
opening of the Rheic Ocean. Analogous megacrys-
tic granitoid magmatism is associated with the
opening of the Rheic Ocean in the Central Iberian
Zone of Spain (e.g. Valverde-Vaquero et al. 2000).
The low-grade pelite-psammite units of the
Acatlán Complex (Huerta Unit, Cosoltepec For-
mation and Salida Unit) and the Granjeno Schist,
on the other hand, are interpreted to be parts of an
Ordovician-Mississippian continental rise prism
deposited on oceanic lithosphere along the northern
margin of Gondwana, represented by the Oaxaquia
terrane, following ocean opening. This interpret-
ation is based on the distal nature of the continental-
derived detritus in these units (Dowe et al. 2005;
Keppie et al. 2006a) and the presence of MORB
and OIB basalts within the low-grade pelite-
psammite units of the Acatlán Complex (Keppie
et al. 2007) and tectonic lenses of serpentinite-
metagabbro in the Granjeno Schist (Ramı́rez-
Ramı́rez 1992), both of which are likely to be of
oceanic origin.
The subsequent high-grade metamorphism of
pre-Silurian units in the Acatlán Complex and
their emplacement onto low-grade pre-Silurian
and Devono– Carboniferous rocks is taken to docu-
ment subduction and exhumation of the leading
edge of Gondwana with the closure of the Rheic
Ocean in the Late Devonian –Mississippian. Term-
inal collision at this time is consistent with the
appearance of midcontinent (US) brachiopod
fauna in the Mississippian rocks of the Oaxaquia
Fig. 14. (Continued) Oaxaquia, the Acatlán Complex

(Acatlán continental rise and Piaxtla Suite) and Sierra
Madre terrane (Granjeno continental rise) relative to
Gondwana (South America– Africa), North America
(Laurentia) and Baltica. (a) Silurian (c. 420 Ma)
palaeogeographic map showing location of pre-Silurian
low-grade units of the Acatlán Complex (Mixteca
terrane) and the Granjeno Schist of the Sierra Madre
terrane as parts of a continental rise flanking Oaxaquia
and South America; (b) Devonian –Carboniferous
palaeogeographic map showing subduction and
exhumation following the amalgamation of Pangaea of
the leading edge of Oaxaquia/Amazonia beneath the
southern margin of Lauentia. Loss of Devono–
Carboniferous magmatic arc as a result of subduction
erosion (Keppie et al. 2008a) would reverse polarity of
subduction; (c) Permian–Triassic palaeogeographic map
showing formation of a Permo– Triassic magmatic arc
and dextral strike-slip fault system, and subsequent
Fig. 14. Continental reconstructions of the margins of development of a clastic wedge, in response to
the Rheic Ocean (modified from Keppie 2004 and subduction of the palaeo-Pacific following the
Keppie et al. 2007, 2008a) showing locations of amalgamation of Pangaea. þ, South Pole.
terrane (Stewart et al. 1999; Navarro-Santillan et al. by Mesoproterozoic (c. 950–1300 Ma), late
2002). Neoproterozoic –Cambrian (c. 500– 700 Ma), Ordo-
Following the amalgamation of Pangaea and the vician (c. 440–480 Ma) and Permo– Carboniferous
tectonic juxtapositioning of the Acatlán Complex (c. 290 Ma) ages, with additional early –middle
and Granjeno Schist with the Oaxaquia terrane by Neoproterozoic (c. 750– 950 Ma) and mid-
dextral strike slip, the development of a Permo– Proterozoic and older (c. 1300–2200 Ma) signa-
Triassic magmatic arc is recorded in the Acatlán tures. These ages are interpreted to indicate
Complex by the emplacement of the Totoltepec sources dominated by local (Oaxaquia terrane and
pluton and the deposition of the arc-related Teco- Acatlán Complex) and Gondwanan (Amazonia,
mate, Olinalá and Patlanoaya formations. This arc Brasiliano and Tocantins) provenances, and support
is considered to have developed in response to sub- the continental reconstruction of Keppie & Ramos
duction of the palaeo-Pacific on the southwestern (1999), which places the Oaxaquia terrane adjacent
(present co-ordinates) flank of Pangaea following to northwestern South America throughout the
Pangaea assembly. Continued convergent tectonics Palaeozoic. We therefore consider the Mixteca and
along this margin during the Permo –Triassic led Sierra Madre terranes to preserve vestiges, not of
to the development of the Chazumba–Magdalena the Iapetus Ocean as traditionally advocated, but
clastic wedge, and was followed in the Jurassic of the southern margin of the Rheic Ocean. The pre-
by hotspot activity and the development of the Mag- Silurian sedimentary and magmatic history of the
dalena Migmatite coeval with the opening of the Mixteca terrane is interpreted to record the initial
Gulf of Mexico. rifting of this ocean, whereas deposition of its
In this much simpler scenario, the Acatlán pre-Silurian and Devono–Carboniferous pelite-
Complex and Granjeno Schist were deposited psammite units, and the Granjeno Schist of the
along the southern Rheic margin of the Oaxaquia Sierra Madre terrane, chronicle the development
terrane (Fig. 14a), which lay adjacent to northwes- of a continental rise prism on its Gondwanan (Oax-
tern South America throughout the Palaeozoic as aquia) margin following ocean opening. Closure of
proposed in the continental reconstructions of the Rheic Ocean during the Late Devonian –
Keppie & Ramos (1999). Hence, rather than record- Mississippian is thought to be documented in the
ing the closure of a succession of oceanic tracts as Mixteca terrane by the subduction-related eclogite
early as the Ordovician, the Mixteca and Sierra facies metamorphism and exhumation of the pre-
Madre terranes are considered to have been accreted Silurian high-grade units of the Acatlán Complex
to Laurentia only with the late Palaeozoic closure (Piaxtla Suite), while the deposition and subse-
of the Rheic Ocean and the assembly of Pangaea quent deformation of its Permo–Carboniferous
(Fig. 14b). The subsequent development of a and Permo– Triassic units are interpreted to reflect
dextral strike-slip regime and Permian magmatic convergent tectonics on the palaeo-Pacific margin
arc is considered to record subduction of the palaeo- following the amalgamation of Pangaea.
Pacific, which led to thrusting and the deposition of
a Permo –Triassic clastic wedge (Fig. 14c). The authors are indebted to P. Cawood and F. McDowell
This model places the Oaxaquia, Mixteca and for their constructive reviews, and to G. Gehrels and the
Sierra Madre terranes on the lower plate during University of Arizona LaserChron Center for making
available on-line the analysis tool for constructing normal-
closure of the Rheic Ocean (Fig. 13a), consistent
ized relative age distribution plots. Funding for this project
with the absence of a Devono– Carboniferous arc, was provided by NSF grant (EAR-0308105) and an Ohio
and on the upper plate with respect to subduction University Baker Award to RDN, Conacyt and PAPIIT
of the palaeo-Pacific (Fig. 13b, c), consistent with (IN103003) grants to JDK, NSF grant (EAR-0308437) to
the existence of Permian arc. However, Keppie BVM, and NSERC Discovery and Research Capacity
et al. (2008a) advocate an alternative explanation Development grants to JBM and JD. For introducing us
for the missing Devono–Carboniferous arc, in to the Acatlán Complex and for continued discussions
which its absence is attributed to subduction and support, we are especially indebted to F. Ortega-
erosion beneath the Gondwanan margin. If so, the Gutiérrez. This paper is a contribution to IGCP Projects
453 and 497.
polarity of subduction would be reversed, placing
southern Laurentia on the lower plate consistent
with models for the tectonic evolution of the Oua- References
chita orogen (Viele & Thomas 1989).
B OGER , S. D., R AETZ , A., G ILES , D., E TCHART , E. &
F ANNING , C. M. 2005. U– Pb age data from the
Conclusions Sunsas region of Eastern Bolivia, evidence for the
allochthonous origin of the Paragua block. Precam-
Detrital zircon age spectra in the Palaeozoic (meta)- brian Research, 139, 121– 146.
sedimentary rocks of the Oaxaquia, Mixteca and B OHER , M., A BOUCHAMI , W., M ICHARD , A. N.,
Sierra Madre terranes of Mexico are dominated A LBAREDE , F. & A RNDT , N. 1992. Crustal growth
in West Africa at 2.1 Ga. Journal of Geophysical Oaxaca Terrane of southern Mexico: deformation
Research, 97, 345–369. near a triple junction. Tectonophysics, 301, 231 –242.
B OUCOT , A. J., B LODGETT , R. B. & S TEWART , J. H. D ICKINSON , W. R. & L AWTON , T. F. 2001. Carboni-
1997. European Province Late Silurian brachiopods ferous to Cretaceous assembly and fragmentation of
from the Ciudad Victoria area, Tamaulipas, northeast- Mexico. Geological Society of America Bulletin, 113,
ern Mexico. In: K LAPPER , G., M URPHY , M. A. & 1142– 1160.
T ALENT , J. A. (eds) Paleozoic Sequence Stratigraphy, D OWE , D. S., N ANCE , R. D., K EPPIE , J. D., C AMERON ,
Biostratigraphy, and Biogeography: Studies in Honor K. L., O RTEGA -R IVERA , A., O RTEGA -G UTIÉRREZ ,
of J. Granville (‘Jess’) Johnson. Geological Society F. & L EE , J. W. K. 2005. Deformational history
of America Special Paper, 321, 273–293. of the Granjeno Schist, Ciudad Victoria, Mexico:
B REAM , B. R., H ATCHER , R. D., M ILLER , C. F. & constraints on the closure of the Rheic Ocean?
F ULLAGAR , P. D. 2004. Detrital zircon ages and Nd International Geology Review, 47, 920–937.
isotopic data from the southern Appalachian crystal- E LÍAS -H ERRERA , M. & O RTEGA -G UTIÉRREZ , F.
line core, Georgia, South Carolina, North Carolina, 2002. Caltepec fault zone: an Early Permian dextral
and Tennessee: new provenance constraints for transpressional boundary between the Proterozoic
part of the Laurentian Margin. In: T OLLO , R. P., C OR- Oaxacan and Paleozoic Acatlán complexes, southern
RIVEAU , L., M C L ELLAND , J. & B ARTHOLOMEW , M. Mexico, and regional implications. Tectonics, 21(3),
J. (eds) Proterozoic Evolution of the Grenville Orogen 10.1029/200TC001278.
in North America. Geological Society of America E LÍAS -H ERRERA , M., M ACÍAS -R OMO , C., O RTEGA -
Memoir, 197, 459–476. G UTIÉRREZ , F., S ÁNCHEZ -Z AVALA , J. L., I RIONDO ,
B RITO N EVES , B. B., C AMPOS N ETO , M. C. & F UCK , R. A. & O RTEGA -R IVERA , A. 2007. Conflicting
1999. From Rodinia to Western Gondwana: an stratigraphic and geochronologic data from the
approach to the Brasiliano – Pan African cycle and Acatlán Complex: ‘Ordovician’ granites intrude meta-
orogenic collage. Episodes, 22, 155–199. morphic and sedimentary rocks of Devonian–Permian
B RITO N EVES , B. B., P ASSARELLI , C. R., B ASEI , age. Eos Transactions of the American Geophysical
M. A. S. & S ANTOS , E. J. 2003. U– Pb zircon ages Union, 88(23), Joint Assembly Supplement, Abstract
of some classic granites of the Borborema Province. T41A-12.
In: Short Papers – IV South American Symposium F LORES DE D IOS G ONZALÉZ , L. A., V ACHARD , D. &
on Isotope Geology, Salvador, Brazil, August 24–27, B UITRON -S ÁNCHEZ , B. E. 1998. The Tiñu, Santiago
2003 [CD-ROM], 158–159. Ixtaltepec and Yododene Formations, Oaxaca State:
C AMERON , K. L., L OPEZ , R., O RTEGA -G UTIÉRREZ , F., sedimentological, stratigraphic and paleogeographic
S OLARI , L. A., K EPPIE , J. D. & S CHULZE , C. reinterpretations. In: P ROGRAM , I. G. C. (ed.)
2004. U –Pb constraints and Pb isotopic compositions Laurentia– Gondwanan Connections Before Pangea.
of leached feldspars: constraints on the origin and Program and Abstracts. Instituto de Geologı́a,
evolution of Grenvillian rocks from eastern and Universidad Nacional Autonoma de México, Oaxaca
southern Mexico. In: T OLLO , R. P., C ORRIVEAU , L., City, Oaxaca Mexico, 16.
M C L ELLAND , J. & B ARTHOLOMEW , M. J. (eds) G EHRELS , G., V ALENCIA , V. & P ULLEN , A. 2006.
Proterozoic Tectonic Evolution of the Grenville Detrital zircon geochronology by laser ablation
Orogen in North America. Geological Society of multicollector ICPMS at the Arizona LaserChron
America Memoir, 197, 755–770. Center. In: O LSZEWSKI , T. (ed.) Geochronology:
C AWOOD , P. A. & N EMCHIN , A. A. 2001. Paleogeo- Emerging Opportunities. Paleontology Society
graphic development of the East Laurentian margin; Papers, 12, 67–76.
constraints from U– Pb dating of detrital zircons in G ILLIS , R. J., G EHRELS , G. E., R UIZ , J. & F LORES DE
the Newfoundland Appalachians. Geological Society D IOS G ONZALÉZ , L. A. 2005. Detrital zircon prove-
of America Bulletin, 113, 1234–1246. nance of Cambrian-Ordovician and Carboniferous
C AWOOD , P. A., M C C AUSLAND , P. J. A. & D UNNING , strata of the Oaxaca terrane, southern Mexico. Sedi-
G. R. 2001. Opening Iapetus: constraints from the mentary Geology, 182, 87–100.
Laurentian margin of Newfoundland. Geological G ONZÁLEZ -A RREOLA , C., V ILLASEÑOR -M ARTINEZ ,
Society of America Bulletin, 113, 443–453. A. B. & C ORONA -E SQUIVEL , R. 1994. Permian
C AWOOD , P. A., N EMCHIN , A. A., S TRACHAN , R., fauna of the Los Arcos Formation, Municipality of
P RAVE , T. & K RABBENDAM , M. 2007. Sedimentary Olinalá, State of Guerrero, Mexico. Revista Mexicana
basin and detrital zircon record along East Laurentia de Ciencias Geológicas, 2(2), 214–221.
and Baltica during assembly and breakup of Rodinia. G OWER , C. F., R YAN , A. B. & R IVERS , T. 1990.
Journal of the Geological Society, London, 164, Mid-Proterozoic Laurentia–Baltica: an overview of
257– 275. its geological evolution and summary of the contri-
C EDIEL , F., S HAW , R. P. & C ÁCERES , C. 2003. Tectonic butions by this volume. In: G OWER , C. F., R IVERS ,
assembly of the Northern Andean Block. In: T. & R YAN , A. B. (eds) Mid-Proterozoic Laurentia–
B ARTOLINI , C., B UFFLER , R. T. & B LICKWEDE , J. Baltica. Geological Association of Canada Special
(eds) The Circum-Gulf of Mexico and the Caribbean: Paper, 38, 1– 20.
Hydrocarbon Habitats, Basin Formation, and Plate G RADSTEIN , F. M., O GG , J. G. & S MITH , A. G. 2004. A
Tectonics. American Association of Petroleum Geol- Geologic Time Scale. Cambridge University Press,
ogists Memoir, 79, 815– 848. Cambridge, 1–585.
C ENTENO -G ARCIA , E. & K EPPIE , J. D. 1999. Latest G RODZICKI , K. R., N ANCE , R. D., K EPPIE , J. D., D OSTAL ,
Paleozoic–early Mesozoic structures in the central J. & M URPHY , J. B. 2008. Structural, geochemical and
geochronological analysis of metasedimentary and K EPPIE , J. D., D OSTAL , J., N ANCE , R. D., M ILLER , B. V.,
metavolcanic rocks of the Coatlaco area, Acatlán O RTEGA -R IVERA , A. & L EE , J. K. W. 2006b.
Complex, southern Mexico. Tectonophysics, 461, Circa 546 Ma plume-related dykes in the 1 Ga
311–323. Novillo Gneiss (east-central Mexico): evidence of
H EATHERINGTON , A. L., M UELLER , P. A. & N UTMAN , the initial separation of Avalonia. Precambrian
A. P. 1996. Neoproterozoic magmatism in the Suwan- Research, 147, 342– 353.
nee terrane: implications for terrane correlation. In: K EPPIE , J. D., D OSTAL , J. & E LIÁS -H ERRERA , M. 2007.
N ANCE , R. D. & T HOMPSON , M. D. (eds) Avalonian Ordovician-Devonian oceanic basalts in the Cosolte-
and Related Peri-Gondwanan Terranes of the Circum- pec Formation, Acatlán Complex, southern Mexico:
North Atlantic. Geological Society of America Special vestiges of the Rheic Ocean? In: L INNEMANN , U.,
Paper, 304, 257–268. N ANCE , R. D., K RAFT , P. & Z ULAUF , G. (eds) The
H INOJOSA -P RIETO , H. R., N ANCE , R. D., K EPPIE , J. D., Evolution of the Rheic Ocean: From Avalonian-
D OSTAL , J., O RTEGA -R IVERA , A. & L EE , J. W. K. Cadomian Active Margin to Alleghenian-Variscan
2008. Ordovician and Late Paleozoic-Early Mesozoic Collision. Geological Society of America Special
tectonothermal history of the La Noria area, northern Paper, 423, 477– 487.
Acatlán Complex, southern Mexico: record of con- K EPPIE , J. D., D OSTAL , J., M URPHY , J. B. & N ANCE ,
vergence in the Rheic and paleo-Pacific Oceans. Tecto- R. D. 2008a. Synthesis and tectonic interpretation of
nophysics, 461, 324– 342. the westernmost Paleozoic Variscan orogen in
H IRDES , W. & D AVIS , D. W. 2002. U –Pb geochronology southern Mexico: from rifted Rheic margin to active
of Paleoproterozoic rocks in the southern part of Pacific margin. Tectonophysics, 461, 277 –290.
the Kedougou-Kéniéba inlier, Senegal, West Africa: K EPPIE , J. D., D OSTAL , J. ET AL . 2008b. Ordovician rift
evidence for diachronous accretionary development tholeiites in the Acatlán Complex, southern Mexico:
of the Eburnian province. Precambrian Research, evidence of rifting on the southern margin of the
118, 83–99. Rheic Ocean. Tectonophysics, 461, 130–156.
K EPPIE , J. D. 2004. Terranes of Mexico revisited: a 1.3 K ROGH , T. E., K AMO , S. L. & B OHOR , B. F. 1993a. Fin-
billion year odyssey. International Geology Review, gerprinting the K/T impact site and determining the
46, 765–794. time of impact by U/Pb dating of single shocked
K EPPIE , J. D. & R AMOS , V. S. 1999. Odyssey of ter- zircons from distal ejecta. Earth and Planetary
ranes in the Iapetus and Rheic Oceans during the Science Letters, 119, 425– 429.
Paleozoic. In: K EPPIE , J. D. & R AMOS , V. A. (eds) L ANDING , E., W ESTROP , S. R. & K EPPIE , J. D. 2007.
Laurentia-Gondwana Connections Before Pangea. Terminal Cambrian and lowest Ordovician succession
Geological Society of America Special Paper, 336, of Mexican West Gondwana: biotas and sequence stra-
267–276. tigraphy of the Tiñu Formation. Geological Magazine,
K EPPIE , J. D., D OSTAL , J., O RTEGA -G UTIÉRREZ , F. & 144, 909–9372.
L OPEZ , R. 2001. A Grenvillian arc on the margin of L AWLOR , P. J., O RTEGA -G UTIÉRREZ , F., C AMERON ,
Amazonia: evidence from the southern Oaxacan K. L., O CHOA -C AMARILLO , H., L OPEZ , R. &
Complex, southern Mexico. Precambrian Research, S AMPSON , D. E. 1999. U– Pb geochronology, geo-
112, 165– 181. chemistry, and provenance of the Grenvillian Huizno-
K EPPIE , J. D., D OSTAL , J., C AMERON , K. L., S OLARI , pala Gneiss of eastern Mexico. Precambrian Research,
L. A., O RTEGA -G UTIÉRREZ , F. & L OPEZ , R. 2003. 94, 73– 99.
Geochronology and geochemistry of Grenvillian L UDWIG , K. J. 2003. Isoplot 3.00. Berkeley Geochronol-
igneous suites in the northern Oaxacan Complex, ogy Center Special Publication, 4, 70.
southern Mexico: tectonic implications. Precambrian M ACHADO , N., V ALLADARES , C. S., H EILBRON , M. &
Research, 120, 365 –389. V ALERIANO , C. M. 1996. U/Pb geochronology of
K EPPIE , J. D., S ANDBERG , C. A., M ILLER , B. V., the Central Ribeira Belt. Precambrian Research, 79,
S ÁNCHEZ -Z AVALA , J. L., N ANCE , R. D. & P OOLE , 347– 361.
F. G. 2004a. Implications of latest Pennsylvanian to M ALONE , J. W., N ANCE , R. D., K EPPIE , J. D. &
Middle Permian paleontological and U– Pb SHRIMP D OSTAL , J. 2002. Deformational history of part of
data from the Tecomate Formation to re-dating the Acatlán Complex: Late Ordovician– Early Silur-
tectonothermal events in the Acatlán Complex, ian and Early Permian orogenesis in southern
southern Mexico. International Geology Review, 46, Mexico. Journal of South American Earth Sciences,
745–753. 15, 511– 524.
K EPPIE , J. D., N ANCE , R. D. ET AL . 2004b. Mid-Jurassic M ARTÍNEZ C ATALÁN , J. R., H ATCHER , R. D., J R .,
tectonothermal event superposed on a Paleozoic geo- A RENAS , R. & D ÍAZ G ARCÍA , F. (eds) 2002.
logical record in the Acatlán Complex of southern Variscan-Appalachian dynamics: The building of the
Mexico: hotspot activity during the breakup of late Paleozoic basement. Geological Society of
Pangea. Gondwana Research, 7, 239– 260. Amererica Special Paper, 364, 312.
K EPPIE , J. D., N ANCE , R. D., F ERNÁNDEZ -S UÁREZ , J., M EZA -F IGUEROA , D., R UIZ , J., T ALAVERA -M ENDOZA ,
S TOREY , C. D., J EFFRIES , T. E. & M URPHY , J. B. O. & O RTEGA -G UTIERREZ , F. 2003. Tectonometa-
2006a. Detrital zircon data from the eastern Mixteca morphic evolution of the Acatlán Complex eclogites
terrane, southern Mexico: evidence for an Ordovician- (southern Mexico). Canadian Journal of Earth
Mississippian continental rise and a Permo-Triassic Sciences, 40, 27– 44.
clastic wedge adjacent to Oaxaquia. International M IDDLETON , M., K EPPIE , J. D., M URPHY , J. B., M ILLER ,
Geology Review, 48, 97– 111. B. V., N ANCE , R. D., O RETEGA -R IVERA , A. & L EE ,
J. K. W. 2007. P– T–t constraints on exumation fol- O RTEGA -G UTIÉRREZ , F. 1978. El Gneiss Novillo y rocas
lowing subduction in the Rheic Ocean from eclogitic metamorficas asociadas en los canones del Novillo y la
rocks in the Acatlán Complex of southern Mexico. Peregrina, area Ciudad Victoria, Tamaulipas. Univer-
In: L INNEMANN , U., N ANCE , R. D., K RAFT , P. & sidad Nacional Autónoma de México, Instituto de
Z ULAUF , G. (eds) The Evolution of the Rheic Geologı́a. Revista, 2, 19– 30.
Ocean: From Avalonian–Cadomian active margin to O RTEGA -G UTIÉRREZ , F., R UIZ , J. & C ENTENO -
Alleghenian– Variscan collision. Geological Society G ARCÍA , E. 1995. Oaxaquia, a Proterozoic micro-
of America Special Paper, 423, 489–509. continent accreted to North America during the late
M ILLER , B. V., D OSTAL , J., K EPPIE , J. D., N ANCE , R. D., Paleozoic. Geology, 23, 1127– 1130.
O RTEGA -R IVERA , A. & L EE , J. W. K. 2007. Ordovi- O RTEGA -G UTIÉRREZ , F., E LIAS -H ERRERA , M., R EYES -
cian calc-alkaline granitoids in the Acatlán Complex, S ALAS , M., M ACIAS -R OMO , C. & L ÓPEZ , R. 1999.
southern Mexico: geochemical and geochronological Late Ordovician-Early Silurian continental collision
data and implications for the tectonics of the Gondwa- orogeny in southern Mexico and its bearing on
nan margin of the Rheic Ocean. In: L INNEMANN , U., Gondwana-Laurentia connections. Geology, 27,
N ANCE , R. D., Z ULAF , G. & K RAFT , P. (eds) The 719–722.
Evolution of the Rheic Ocean: From Avalonian– O RTEGA -O BREGÓN , C., K EPPIE , J. D. ET AL . 2003. Geo-
Cadomian active margin to Alleghenian –Variscan chronology and geochemistry of the 917 Ma,
collision. Geological Society of America Special calc-alkaline Etla granitoid pluton (Oaxaca, southern
Paper, 423, 465–475. México): evidence of post-Grenvillian subduction
M ORALES -G ÁMEZ , M., K EPPIE , J. D. & N ORMAN , M. along the northern margin of Amazonia. International
2008. The Ordovician– Silurian rift-passive margin Geology Review, 45, 596–610.
on the Mexican margin of the Rheic Ocean overlain P ANTOJA -A LOR , J. & R OBISON , R. 1967. Paleozoic sedi-
by Carboniferous-Permian periarc rocks: evidence mentary rocks in Oaxaca, Mexico. Science, 157,
from the eastern Acatlán Complex, southern Mexico. 1033– 1035.
Tectonophysics, 461, 291–310. P IMENTAL , M. M., F UCK , R. A. & G IOIA , S. M. C. 2000.
M URPHY , J. B., K EPPIE , J. D., N ANCE , R. D., M ILLER , The Neoproterozoic Goiás magmatic arc, central
B. V., M IDDLETON , M., F ERNÁNDEZ -S UAREZ , J. & Brazil: a review and new Sm–Nd isotopic data.
J EFFRIES , T. E. 2006. Geochemistry and U–Pb proto- Revista Brasileira de Geociências, 30, 35– 39.
lith ages of eclogitic rocks of the Ası́s Lithodeme, P INDELL , J., K ENNAN , L. & B ARRETT , S. 2000. Putting
Piaxtla Suite, Acatlán Complex, southern Mexico: it all together again. AAPG Explorer, 21(10), 34–37.
tectonothermal activity along the southern margin of P OTREL , A., P EUCAT , J. & F ANNING , C. M. 1998.
the Rheic Ocean. Journal of the Geological Society, Archean crustal evolution of the West African
London, 163, 683– 695. Craton: example of the Amsaga Area (Reguibat
N ANCE , R. D., M ILLER , B. V., K EPPIE , J. D., M URPHY , Rise). U– Pb and Sm–Nd evidence for crustal growth
J. B. & D OSTAL , J. 2006. The Acatlán Complex, and recycling. Precambrian Research, 90, 107–117.
southern Mexico: record of Pangea assembly to P ROENZA , J. A., O RTEGA -G UTIÉRREZ , F., C AMPRUBÍ ,
breakup. Geology, 34, 857–860. A., T RITLLA , J., E LÍAS -H ERRERA , M. & R EYES -
N ANCE , R. D., M ILLER , B. V., K EPPIE , J. D., M URPHY , S ALAS , M. 2004. Paleozoic serpentinite-enclosed
J. B. & D OSTAL , J. 2007a. Vestige of the Rheic chromitites from Tehuitzingo (Acatlán Complex,
Ocean in North America: the Acatlán Complex of southern Mexico): a petrological and mineralogical
southern México. In: L INNEMANN , U., N ANCE , study. Journal of South America Earth Sciences, 16,
R. D., Z ULAF , G. & K RAFT , P. (eds) The Evolution of 649–666.
the Rheic Ocean: From Avalonian–Cadomian active R AMÍREZ -R AMÍREZ , C. 1992. Pre-Mesozoic geology of
margin to Alleghenian– Variscan collision. Geological Huizachal-Peregrina Anticlinorium, Ciudad Victoria,
Society of America Special Paper, 423, 437–452. Tamaulipas, and Adjacent Parts of Eastern Mexico.
N ANCE , R. D., F ERNÁNDEZ -S UÁREZ , J., K EPPIE , J. D., PhD thesis, University of Texas, Austin, 1–317.
S TOREY , C. & J EFFRIES , T. E. 2007b. Provenance of R AMOS -A RÍAS , M. A., K EPPIE , J. D., O RTEGA -R IVERA ,
the Granjeno Schist, Ciudad Victoria, México: detrital A. & L EE , J. W. K. 2008. Extensional Late Paleozoic
zircon U –Pb age constraints and implications for the deformation on the western margin of Pangea, Patla-
Paleozoic paleogeography of the Rheic Ocean. In: noaya area, Acatlán Complex, southern Mexico. Tecto-
L INNEMANN , U., N ANCE , R. D., Z ULAF , G. & nophysics, 461, 60– 76.
K RAFT , P. (eds) The Evolution of the Rheic Ocean: R EYES -S ALAS , A. M. 2003. Mineralogia y petrologia de
From Avalonian-Cadomian active margin to los granitoides Esperanza del compljo Acatlán, sur
Alleghenian-Variscan collision. Geological Society de México. PhD thesis, Universidad Nacional Auton-
of America Special Paper, 423, 453–464. oma de México, Mexico City.
N AVARRO -S ANTILLAN , D., S OUR -T OVAR , F. & R IVERS , T. 1997. Lithotectonic elements of the Grenville
C ENTENO -G ARCIA , E. 2002. Lower Mississippian Province: review and tectonic implications. Precam-
(Osagean) brachiopods from the Santiago Formation, brian Research, 86, 117–154.
Oaxaca, Mexico: stratigraphic and tectonic impli- R OBERTS , D. 2003. The Scandinavian Caledonides: event
cations. Journal of South American Earth Sciences, chronology, palaeogeographic settings and likely
15, 327–336. modern analogues. Tectonophysics, 365, 283– 299.
O RTEGA -G UTIÉRREZ , F. 1975. The Pre-Mesozoic R OBERTS , D. & S IEDLECKA , A. 2002. Timanian oro-
geology of the Acatlán area, south Mexico. PhD genic deformation along the northeastern margin of
thesis, Leeds University, UK. Baltica, northwest Russia and northeast Norway, and
Avalonian– Cadomian connections. Tectonophysics, T OLLO , R. P., C ORRIVEAU , L., M C L ELLAND , J. &
352, 169– 184. B ARTHOLOMEW , M. J. (eds) 2004. Proterozoic tec-
R OBISON , R. & P ANTOJA -A LOR , J. 1968. Tremadocian tonic evolution of the Grenville Orogen in North
trilobites from Nochixtlan region, Oaxaca, Mexico. America. Geological Society of America Memoir,
Journal of Paleontology, 42, 767–800. 197, 1– 830.
R OCCI , G., B RONNER , G. & D ESCHAMPS , M. 1991. T ORRES , R., R UÍZ , J., P ATCHETT , P. J. & G RAJALES -
Crystalline basement of the West African craton. N ISHIMURA , J. M. 1999. Permo-Triassic continental
In: D ALLMEYER , R. D. & L ÉCORCHÉ , J. P. (eds) arc in eastern Mexico; tectonic implications for recon-
The West African Orogens and Circum-Atlantic structions of southern North America. In: B ARTOLINI ,
Correlatives. Springer-Verlag, New York, 31– 61. C., W ILSON , J. L. & L AWTON , T. F. (eds) Mesozoic
S ADOWSKI , G. R. & B ETTENCOURT , J. S. 1996. Mesopro- Sedimentary and Tectonic History of North-central
terozoic tectonic correlations between eastern Lauren- Mexico. Geological Society of America Special
tia and the western border of the Amazon craton. Paper, 340, 191– 196.
Precambrian Research, 76, 213–227. V ACHARD , D. & F LORES DE D IOS , A. 2002. Discovery of
S ÁNCHEZ -Z AVALA , J. L., C ENTENO -G ARCIA , E. & latest Devonian/earliest Mississippian microfossils in
O RTEGA -G UTIÉRREZ , F. 1999. Review of Paleozoic San Salvador Patlanoaya (Puebla, Mexico): biogeo-
stratigraphy of Mexico and its role in the graphic and geodynamic consequences. Compte
Gondwana– Laurentia connections. In: R AMOS , Rendu Geoscience, 334, 1095– 1101.
V. A. & K EPPIE , J. D. (eds) Laurentian –Gondwanan V ACHARD , D., F LORES DE D IOS , A., B UITRÓN , B. E. &
Connections Before Pangea. Geological Society of G RAJALES , M. 2000. Biostratigraphie par fusu-
America Special Paper, 336, 211 –226. lines des calcaires Carbonifères et Permiens de San
S ÁNCHEZ -Z AVALA , J. L., O RTEGA -G UTIÉRREZ , F., Salvador Patlanoaya (Puebla, Mexique). Geobios, 33,
K EPPIE , J. D., J ENNER , G. A., B ELOUSOVA , E. & 5– 33.
M ACIÁS -R OMO , C. 2004. Ordovician and Mesoproter- V ACHARD , D., F LORES DE D IOS , A. & B UITRÓN , B.
ozoic zircons from the Tecomate Formation and Esper- 2004. Guadalupian and Lopingian (Middle and Late
anza granitoids, Acatlán Complex, southern Mexico: Permian) deposits from México and Guatemala, a
local provenance in the Acatlán and Oaxacan com- review with new data. Geobios, 37, 99–115.
plexes. International Geology Review, 46, 1005–1021. VAN S TAAL , C. R., D EWEY , J. F., M AC N IOCAILL , C. &
S ANTOS , J. O. S. 2003. Geotectônica dos Escudos das M C K ERROW , W. S. 1998. The Cambrian-Silurian
Guianas e Brasil-Central. In: B IZZI , L. A., S CHOB- tectonic evolution of the Northern Appalachians
BENHAUS , C., V IDOTTI , R. M. & G ONÇALVES , J. H. and British Caledonides; history of a complex, west
(eds) Geologia, Tectônica e Recursos Minerais do and southwest Pacific-type segment of Iapetus. In:
Brasil, Texto, Mapas e SIG. Brasilia, Seriço Geológico B LUNDELL , D. & S COTT , A. C. (eds) Lyell: The past
do Brasil – CPRM, 169–226. is the key to the present. Geological Society, London,
S ANTOS , J. O. S., H ARTMANN , L. A., G AUDETTE , H. E., Special Publications, 143, 199–242.
G ROVES , D. I., M C N AUGHTON , N. J. & F LETCHER , V ALVERDE -V AQUERO , P. & D UNNING , G. R. 2000.
I. R. 2000. A new understanding of the provinces of New U–Pb ages from Early Ordovician magmatism
the Amazon Craton based on integration of field in Central Spain. Journal of the Geological Society,
mapping and U– Pb and Sm–Nd geochronology. London, 157, 15–26.
Gondwana Research, 3, 453– 488. V EGA -G RANILLO , R., T ALAVERA -M ENDOZA , O.,
S OLARI , L. A., K EPPIE , J. D., O RTEGA -G UTIÉRREZ , F., D IANA M EZA -F IGUEROA , D., R UIZ , J., G EHRELS ,
C AMERON , K. L., L OPEZ , R. & H AMES , W. E. 2003. G. E. & L ÓPEZ -M ARTÍNEZ , M. 2007. Pressure-
990 and 1100 Ma Grenvillian tectonothermal events temperature-time evolution of Paleozoic high-pressure
in the northern Oaxacan Complex, southern Mexico: rocks of the Acatlán Complex (southern Mexico):
roots of an orogen. Tectonophysics, 365, 257–282. implications for the evolution of the Iapetus and
S TEWART , J. H., B LODGETT , R. B., B OUCOT , A. J., Rheic Oceans. Geological Society of America Bulletin,
C ARTER , J. L. & L OPEZ , R. 1999. Exotic Paleozoic 119, 1249– 1264.
strata of Gondwanan provenance near Ciudad Victoria, V IELE , G. W. & T HOMAS , W. A. 1989. Tectonic Synthesis
Tamaulipas, Mexico. In: K EPPIE , J. D. & R AMOS , of the Ouachita Orogenic Belt. In: H ATCHER , R. D.,
V. A. (eds) Laurentia–Gondwana Connections T HOMAS , W. A. & V IELE , G. W. (eds) The Geology
Before Pangea. Geological Society of America of North America. The Appalachian –Ouachita
Special Paper, 336, 227 –252. Orogen in the United States, F-2, 695–728.
T ALAVERA -M ENDOZA , O., R UÍZ , J., G EHRELS , G. E., W EBER , B. & K ÖEHLER , H. 1999. Sm–Nd, Rb –Sr and
M EZA -F IGUEROA , D. M., V EGA -G RANILLO , R. & U– Pb geochronology of a Grenville Terrane in
C AMPA -U RANGA , M. F. 2005. U–Pb geochronology Southern Mexico: origin and geologic history of the
of the Acatlán Complex and implications for the Paleo- Guichicovi Complex. Precambrian Research, 96,
zoic paleogeography and tectonic evolution of 245– 262.
southern Mexico. Earth and Planetary Science Y AÑEZ , P., R UIZ , J., P ATCHETT , P. J., O RTEGA -
Letters, 235, 682– 699. G UTIÉRREZ , F. & G EHRELS , G. 1991. Isotopic
T ASSINARI , C. C. G. & M ACAMBIRA , M. J. B. 1999. studies of the Acatlán Complex, southern Mexico:
Geochronological provinces of the Amazon Craton. implications for Paleozoic North American tectonics.
Episodes, 22, 174– 182. Geological Society of America Bulletin, 103, 817–828.
Pre-Carboniferous, episodic accretion-related, orogenesis along the
Laurentian margin of the northern Appalachians
CEES R. VAN STAAL1*, JOSEPH B. WHALEN2, PABLO VALVERDE-VAQUERO3,
ALEXANDRE ZAGOREVSKI2 & NEIL ROGERS2
1
Geological Survey of Canada, 625 Robson Street, Vancouver, British Columbia, Canada
2
Geological Survey of Canada, 601 Booth Street, Ottawa, Ontario, K1A 0E8, Canada
3
Instituto Geologico y Minero de España (IGME), La Calera 1,
Tres Cantos (Madrid), 28760, Spain
*Corresponding author (e-mail: Cees.vanStaal@NRCan-RNCan.gc.ca)
Abstract: During the Early to Middle Palaeozoic, prior to formation of Pangaea, the Canadian and
adjacent New England Appalachians evolved as an accretionary orogen. Episodic orogenesis
mainly resulted from accretion of four microcontinents or crustal ribbons: Dashwoods, Ganderia,
Avalonia and Meguma. Dashwoods is peri-Laurentian, whereas Ganderia, Avalonia and Meguma
have Gondwanan provenance. Accretion led to a progressive eastwards (present co-ordinates)
migration of the onset of collision-related deformation, metamorphism and magmatism. Volumi-
nous, syn-collisional felsic granitoid-dominated pulses are explained as products of slab-breakoff
rather than contemporaneous slab subduction. The four phases of orogenesis associated with accre-
tion of these microcontinents are known as the Taconic, Salinic, Acadian and Neoacadian oroge-
nies, respectively. The Ordovician Taconic orogeny was a composite event comprising three
different phases, due to involvement of three peri-Laurentian oceanic and continental terranes.
The Taconic orogeny was terminated with an arc–arc collision due to the docking of the active
leading edge of Ganderia, the Popelogan –Victoria arc, to an active Laurentian margin (Red
Indian Lake arc) during the Late Ordovician (460–450 Ma).
The Salinic orogeny was due to Late Ordovician– Early Silurian (450–423 Ma) closure of the
Tetagouche– Exploits backarc basin, which separated the active leading edge of Ganderia from
its trailing passive edge, the Gander margin. Salinic closure was initiated following accretion of
the active leading edge of Ganderia to Laurentia and stepping back of the west-directed subduction
zone behind the accreted Popelogan –Victoria arc. The Salinic orogeny was immediately followed
by Late Silurian– Early Devonian accretion of Avalonia (421– 400 Ma) and Middle Devonian–
Early Carboniferous accretion of Meguma (395– 350 Ma), which led to the Acadian and Neoaca-
dian orogenies, respectively. Each accretion took place after stepping-back of the west-dipping
subduction zone behind an earlier accreted crustal ribbon, which led to progressive outboard
growth of Laurentia. The Acadian orogeny was characterized by a flat-slab setting after the
onset of collision, which coincided with rapid southerly palaeolatitudinal motion of Laurentia.
Acadian orogenesis preferentially started in the hot and hence, weak backarc region. Subsequently
it was characterized by a time-transgressive, hinterland migrating fold-and-thrust belt antithetic to
the west-dipping A– subduction zone. The Acadian deformation front appears to have been closely
tracked in space by migration of the Acadian magmatic front. Syn-orogenic, Acadian magmatism
is interpreted to mainly represent partial melting of subducted fore-arc material and pockets of fluid-
fluxed asthenosphere above the flat-slab, in areas where Ganderian’s lithosphere was thinned by
extension during Silurian subduction of the Acadian oceanic slab. Final Acadian magmatism
from 395– c. 375 Ma is tentatively attributed to slab-breakoff.
Neoacadian accretion of Meguma was accommodated by wedging of the leading edge of
Laurentia, which at this time was represented by Avalonia. The Neoacadian was devoid of any
accompanying arc magmatism, probably because it was characterized by a flat-slab setting through-
out its history.
Introduction by using the critical relationships established in the

Canadian, and to a lesser extent, the immediately
This paper focuses on the style and nature of the adjacent parts of the New England segment of the
relationships between deformation, metamorphism northern Appalachians (Fig. 1). The Alleghenian
and magmatism generated during the Taconic, orogenic overprint in the Canadian segment of the
Salinic, Acadian and Neoacadian orogenies, mainly Appalachians is relatively small and we are
272
C. R. VAN STAAL ET AL.
Fig. 1. Geology of the Canadian and adjacent New England Appalachians with the geographical distribution of the major tectonic elements discussed in text. A, Arisaig Group;
AC, Ackley granite; B, Burgeo batholith; BB, Badger basin; BBF, Bamford Brook fault; BBL, Baie Verte Brompton Line; BVOT, Baie Verte oceanic tract; CCF, Cobequid–
Chedabucto fault; CF, Cabot fault; CL, Chain Lakes Massif; CO, Cookson Group; DBL, Dog Bay Line; FO, Fournier Group; F, Fogo Island pluton; GBF, Green Bay fault;
GRUB, Gander River ultrabasic belt; HH, Hodges Hill pluton; LBOT, Lushs Bight oceanic tract; MP, Mount Peyton pluton; RBF, Rocky Brook –Millstream fault system;
RF, Restigouche fault; RIL, Red Indian Line; SGB, St. George batholith; SM, South Mountain batholith; U, Utopia granite; VA, Victoria arc; TP, Tally Pond Group.
PRE-CARBONIFEROUS, EPISODIC ACCRETION-RELATED, OROGENESIS ALONG 273
intimately familiar with its geology, some of us within arc or back-arc settings prior to accretion
having worked here for .25 years. In addition, (e.g. Swinden et al. 1997; van Staal et al. 1998;
this segment of the northern Appalachians has rela- Barr et al. 2002). Details concerning spatial-temporal
tively large and regionally extensive databases of groupings, interpretation methods/approach and
radiometric and fossil ages as well as high-quality data sources are given in Appendix 1. The reader is
magmatic rock geochemical data, while the grade referred to these data sources for much more com-
of regional metamorphism is generally low. Also prehensive interpretations of specific portions of
available is a wide set of seismic reflection and refrac- our dataset.
tion data acquired during Lithoprobe, which imaged A problem encountered in any orogenic research
the structural architecture at depth and allows its is the well-established classification of orogenies on
linking with crustal structures mapped at surface the basis of time. We know from modern plate tec-
(e.g. van der Velden et al. 2004, and references tonics that tectonic events are commonly markedly
therein). Hence, the vital linkages between defor- diachronous and additionally, spatially and kinema-
mation, metamorphism, magmatism, exhumation tically unrelated orogenic events, which overlap in
and syn-orogenic sedimentation, essential to under- time but not in space, may be juxtaposed fortui-
stand the dynamic processes responsible for orogen- tously (e.g. van Staal et al. 1998). Hence, as much
esis, are relatively well preserved, yet the constraints as possible, we grouped and separated structures,
provided by these processes rarely have been inte- metamorphism and associated magmatic rocks on
grated together. The latter particularly applies to the basis of their inferred relationship to the specific
the relationships between magmatism and tectonic tectonic events that have been recognised.
processes responsible for the Appalachian orogenic The data discussed in this paper indicate that the
events (see below). The details of the tectonic Appalachians and closely-related British Caledo-
elements, plate tectonic setting and plate interactions nides (Fig. 1) were an accretionary orogen prior to
involved have been discussed in several recent papers Alleghenian– Variscan orogenesis. During the early
(e.g. van Staal et al. 1998, 2007; van Staal 2005, to middle Palaeozoic, the roughly east –west oriented
2007; Lissenberg et al. 2005b; Zagorevski et al. Appalachian Laurentian margin, situated close to the
2006, 2007a, c; Valverde-Vaquero et al. 2006; Lin equator, progressively expanded southwards (van
et al. 2007; Hibbard et al. 2007) and the reader is Staal et al. 1998) due to episodic accretions of
referred to these for more details. ribbons of suprasubduction zone oceanic and conti-
Although plutonic rocks make up one quarter of nental lithosphere that were present in the Iapetus
the exposure in the Canadian and adjacent New and to a less extent the Rheic oceans (van Staal
England Appalachians and occur in all the tectonos- et al. 1998; Hibbard 2000; van Staal 2005, 2007;
tratigraphic terranes discussed herein, understanding van Staal & Hatcher 2010). Hence, large parts of
of their petrogenesis generally has been either poor or this orogen are polyorogenic (Fig. 2). For the sake
equivocal (cf. Currie 1995). The subject of their tec- of simplicity we will discuss the orientation of large
tonic implications commonly tended to raise more structural features, such as Laurentia’s Appalachian
questions than provide answers. In this study, we margin, in present co-ordinates rather than use Laur-
have attempted to interpret the observed spatial- entia’s true orientation during the Palaeozoic as
temporal changes and/or repetitions in plutonic rock determined by palaeomagnetism. Accretion gener-
compositions within the tectonic context, to a large ally led to aerially restricted, episodic orogenesis
extent deduced from other less-equivocal lines of with a progressive eastwards migration of the locus
geological evidence. This approach not only pro- of collision-related deformation, metamorphism
vides new insights into understanding granitoid and magmatism (Fig. 2). The eastward migration
petrogenesis in the Appalachians, but it may also of the locus of orogenesis was followed by uplift
represent a ‘granitoid geochemical template’ that and cratonization of the accreted terranes in the
could aid in understanding the tectonic evolution orogen’s hinterland. This process is particularly
of orogens elsewhere that are also characterized evident in Newfoundland by deposition of an uncon-
by abundant plutonism but with less well-defined formable cover of terrestrial, Laurentia-derived
tectonic constraints. The vast majority of the clastic rocks (equivalent to the Old Red Sandstone
over 5400 analyses employed herein are from pluto- in the British Isles). Deposition of the terrestrial
nic rocks, because our evidence indicates that rocks and cratonization of the accreted terranes
the most voluminous plutonic episodes occurred generally becomes younger eastwards (Fig. 3) and
in syn- to post-collisional settings, and as such mimics the eastward time-transgressive migration
mainly reflect the accretion-related processes and of the site of accretion. This eastward shift is due to
provide information concerning the various crustal- progressive stepping back of the subduction zone
and mantle materials juxtaposed and/or involved behind the accreted terranes (see below) and explains
during these events. In contrast, much of the the progressive, time-transgressive uplift of the
volcanic rock record is thought to have formed latter. Some of the accreted ribbons were composite
274
C. R. VAN STAAL ET AL.
Fig. 2. Map of the Canadian and adjacent New England Appalachians showing the geographical distribution of the various orogenic belts and associated Silurian–Early
Carboniferous plutonism. Note that the Acadian starts in the part of Ganderia where Salinic orogenesis was absent. Over time it progressively overprints first Salinic and then
the Taconic orogenic belts.
PRE-CARBONIFEROUS, EPISODIC ACCRETION-RELATED, OROGENESIS ALONG
Fig. 3. Summary of the stratigrapic and tectonic relationships of the various tectonic elements present in the Canadian and adjacent New England Appalachians. All elements west
of the Gander margin are mainly based on relationships established in Newfoundland. Gander margin, Avalonia and Meguma are principally based on relationships in maritime
Canada and adjacent Maine, where these elements and their mutual relationships are best preserved. Note that rifting and separation from Gondwana becomes progressively younger
from Ganderia to Meguma. Modified from van Staal et al. (2007) and van Staal (2007). AG, Arisaig Group; AOB, Annieopsquotch ophiolite belt; BBF, Bamford Brook fault;
BVOT, Baie Verte oceanic tract; C, Coaker porphyry; CSP mélange, Cold Spring Pond mélange; DP, Dunn Point volcanics; HH, Hodges Hill pluton; LB, Loon Bay pluton;
LBOT, Lushs Bight oceanic tract; LL, Long Lake volcanics; LOT, Liberty–Orrington thrust (Maine); Met. Sole, metamorphic sole; MG, magmatic gap; MP, Mount Peyton pluton;
275
RA, Robert Arm Group; RLA, Red Indian Lake arc; SH/BP, Sops Head/Boones Point complex (mélange); SM, South Mountain batholith; TP, Tally Pond volcanics;
WR, White Rock Formation.
276 C. R. VAN STAAL ET AL.
due to collision with suprasubduction zone oceanic or microcontinent at c. 505 Ma, while its leading edge
continental lithospheric terranes present within was still an active margin (511 –505 Ma Tally Pond-
Iapetus, prior to their docking with Laurentia. For Long Lake phase of Penobscot arc; Rogers et al.
example, the closure of the Penobscot back-arc 2006) (Figs 1 & 3). The latter suggest that slab roll-
basin led to Tremadoc (486– 479 Ma) obduction of back forces may have been involved in Ganderia’s
back-arc ophiolites onto the Gander margin (Fig. 3) separation from Gondwana. Slab rollback-induced
and hence to orogenesis, while the latter was still extension continued and culminated in rifting of
proximal to Gondwana in high southerly latitudes the Penobscot arc and opening of a backarc basin
(van Staal 1994; Liss et al. 1994; Zagorevski et al. at 500– 495 Ma, which divided Ganderia into two
2007a). This phase of orogenesis, which took place blocks: an active leading edge (Penobscot arc) and
near the Gondwanan margin far from Laurentia, is a passive trailing edge. The latter is known as the
not dealt with here. Gander margin (van Staal 1994). The Penobscot
The accretionary stage of the Appalachian backarc basin temporarily closed between 485 and
orogen was terminated by the arrival of the bulk 479 Ma, which produced the short-lived Penobscot
of Gondwana during the Carboniferous, which led orogeny (Zagorevski et al. 2007a). Penobscot oro-
to formation of the supercontinent Pangaea. The genesis mainly led to obduction of back-arc ophio-
part of the northern Appalachians exposed on land lites onto the Gander margin (Colman-Sadd et al.
in Canada and northern New England, in contrast 1992) and temporary arc shut-off (Fig. 3). After
to the Appalachians further to the south, largely 478 Ma a new arc was resurrected on Ganderia’s
escaped the penetrative effects of terminal leading edge, which is known as the Popelogan –
continent–continent collision-related deformation, Victoria arc (van Staal et al. 1998). This arc rifted
metamorphism and magmatism. The remnants of too, which led to the opening of the wide
the Alleghenian collision zone are mainly situated Tetagouche– Exploits back-arc basin. Ganderia’s
offshore below the modern Atlantic margin, where active leading edge was accreted to Laurentia at
they were strongly overprinted by Mesozoic exten- 455–450 Ma (van Staal 1994; Zagorevski et al.
sion accompanying opening of the Atlantic Ocean. 2007c), while the passive trailing edge was accreted
at 430–423 Ma, producing the Salinic orogeny
(see below).
Northern Appalachian microcontinents Avalonia (Fig. 1) is a collage of several fault-
bounded Neoproterozoic, largely juvenile
The northern Appalchians comprise the remnants arc-related volcanic-sedimentary belts of low meta-
of at least four accreted ribbon-shaped crustal morphic grade and associated plutonic rocks that
blocks or microcontinents: Dashwoods, Ganderia, experienced a complicated and long-lived Neopro-
Avalonia and Meguma (van Staal 2007). The accre- terozoic tectonic history before deposition of an
tion of these crustal ribbons to Laurentia is the main overstepping Cambrian– Early Ordovician shale-
focus of this paper and hence, below we present a rich platformal sedimentary succession (Fig. 3)
brief overview of the tectonic setting and prove- (O’Brien et al. 1996; Landing 1996; Kerr et al.
nance of these four crustal blocks. 1995). Palaeomagnetic data indicate that Avalonia
Dashwoods, which is the most inboard crustal resided at a high southerly latitude near Gondwana
block, has a Laurentian provenance. Isotopic data from the Middle Cambrian to the end of the Early
and strong zircon inheritance in Ordovician arc plu- Ordovician (Johnson & van der Voo 1986; van der
tons and volcanic rocks suggest it is probably under- Voo & Johnson 1985; MacNiocaill 2000; Hamilton
lain by North American Grenvillian basement (van & Murphy 2004) following a more intermediate lati-
Staal et al. 2007). Fossils and palaeomagnetic data tude position during the Late Neoproteozoic
suggest it was never separated by a wide ocean (c. 580 Ma; McNamara et al. 2001). Fossils also
basin from the Humber passive margin of Laurentia show strong links to Gondwana (e.g. Fortey &
(¼ Humber Zone of Williams 1979) (van Staal et al. Cocks 2003), but previously proposed connections
1998; Waldron & van Staal 2001, and references to NW Africa are inconsistent with a wide range
therein). of geological arguments (e.g. Landing 1996). This
Ganderia, which underlies most of the central led Murphy et al. (2002) to propose an alternative
core of the northern Appalachians (Fig. 1), was a position opposite the Neoproterozoic northern
Late Neoproterozoic –Early Cambrian arc terrane margin of Amazonia, in proximity to but not con-
that was probably built upon the Amazonian nected yet with Ganderia (van Staal et al. 1996;
margin of Gondwana (van Staal et al. 1996). The Rogers et al. 2006). Avalonia has a very different
composition and nature of Middle Cambrian volca- geological evolution during the Palaeozoic than
nic rocks extruded at its trailing edge (White et al. Ganderia (Fig. 3), indicating they were two differ-
1994; Schultz et al. 2008), suggest that Ganderia ent, unrelated terranes, at least after the onset of
rifted off Amazonia and became an isolated the Palaeozoic.
Bimodal volcanic rocks were extruded sporadi- tract (LBOT) and correlatives, the Tremadoc
cally during the Middle Cambrian in Avalonia (490 –483 Ma) Baie Verte oceanic tract (BVOT),
and attributed to transtension (Murphy et al. the Arenig (480 –464 Ma) Annieopsquotch accre-
1996). They are probably related to the onset of tionary tract (AAT) and remnants of dismembered
extension that culminated in Avalonia’s departure Middle Ordovician ophiolitic rocks formed in the
from Gondwana during the Early Ordovician at Tetagouche –Exploits back-arc basin (van Staal
c. 475 Ma (Prigmore et al. 1997; van Staal et al. et al. 2003; Valverde-Vaquero et al. 2006). All
1998). This event is characterized by deposition of accreted oceanic terranes have suprasubduction
a transgressive arenitic cover sequence both in zone geochemical signatures and either formed
Gondwana (Armorican quartzite) and Avalonia during upper plate extension and spreading related
(e.g. Stiperstones quartzite in England). Avalonia to subduction initiation and creation of infant arcs
thus rifted-off Gondwana c. 30 Ma after Ganderia, (van Staal et al. 1998, 2007; Lissenberg et al.
which is an additional argument that Avalonia and 2005b) or back-arc basin opening (Rogers &
Ganderia were two separate microcontinents van Staal 2003; Zagorevski et al. 2006; Valverde-
during the Early Palaeozoic. Vaquero et al. 2006). The LBOT, BVOT and
Meguma represents the most outboard terrane AAT formed close to Laurentia and are related to
preserved in the Canadian Appalachians and is Taconic convergence (Zagorevski et al. 2006,
exposed on land only in southern Nova Scotia 2007c; van Staal et al. 2007), whereas the
(Fig. 1). However, its regional extent is much Tetagouche –Exploits ophiolitic rocks formed
larger and its rocks have been traced offshore by while Ganderia was still at high southerly latitudes
an impressive set of geophysical and well data near Gondwana (Liss et al. 1994).
from the southernmost part of the Grand Banks SE
of Newfoundland across the Scotian shelf, and the
Gulf of Maine to southernmost Cape Cod (see Taconic Orogeny
inset Fig. 1) (Hutchinson et al. 1988; Keen et al.
1991; Pe-Piper & Jansa 1999). Meguma overall The Taconic orogeny recently has been redefined by
had a different Palaeozoic geological evolution than van Staal et al. (2007) to encompass all the orogenic
Avalonia (Fig. 3). The oldest exposed rocks in events that took place in the peri-Laurentian realm
Meguma comprises a thick (,10 km) Cambrian – between the Late Cambrian and Late Ordovician
Early Ordovician turbiditic sandstone–shale sequ- (495 –450 Ma) and the reader is referred to this
ence, which was deposited on the continental rise and related papers (Lissenberg et al. 2005a, b;
and/or slope to outer shelf of a Gondwanan passive Lissenberg & van Staal 2006; Zagorevski et al.
margin. These rocks are disconformably overlain by 2006, 2007c) for more detailed information on
Silurian rift-related bimodal volcanic rocks (442 – the tectonic processes responsible for orogene-
438 Ma) of the White Rock Formation (Schenk sis. It encompasses three dynamically distinct
1997; Keppie & Krogh 2000; MacDonald et al. orogenic events, referred to as Taconic 1, 2 and 3
2002), which may mark the onset of final rifting (Figs 3, 4 & 5).
and departure of Meguma from Gondwana. A com-
bination of detrital zircon, sedimentological and Taconic 1
sparse fossil data support an original provenance
along the West African continental margin (e.g. Taconic 1 represents short-lived, Late Cambrian
Krogh & Keppie 1990; Schenk 1997), possibly in (c. 495 Ma) west-directed obduction of an oceanic
a position separating Avalonia from West Africa infant arc terrane, the Lushs Bight oceanic tract
(Waldron et al. 2009). Fossil evidence suggests (LBOT) onto the peri-Laurentian Dashwoods micro-
that during the Late Silurian, Meguma was close continent in Newfoundland (Fig. 4) (Waldron & van
to Avalonia and/or Baltica and probably had Staal 2001; van Staal et al. 2007). Equivalents of
moved to Laurentia (Bouyx et al. 1997) indepen- Dashwoods and the LBOT have also been reco-
dently from Gondwana. gnized or postulated to exist in the subsurface
Gondwana is generally inferred to have accreted of southern New England, Maine and Quebec
to Laurentia not earlier than the end of the Early (Laird et al. 1993; Karabinos et al. 1998; Huot
Carboniferous. et al. 2002; Gerbi et al. 2006a). Evidence for
LBOT’s Late Cambrian obduction in Newfound-
Accreted oceanic terranes land is principally indicated by: (1) the c. 495 Ma
age of the metamorphic aureole beneath the
In addition to the accreted crustal ribbons, the Appa- ophiolitic St. Anthony complex (Jamieson 1988;
lachian margin of Laurentia grew by accretion G. Dunning, pers. comm. 2004); (2) the c. 493 Ma
of several oceanic terranes (Fig. 3). They are the markedly crustal-contaminated mafic dikes that
Cambrian (510 –500 Ma) Lushs Bight oceanic cut and intruded into metamorphic tectonites of
Fig. 4. Cambrian– Middle Ordovician tectonic evolution of the Humber margin and outboard peri-Laurentian terranes.
Modified from van Staal et al. (2007). (a) Subduction initiation (using an abandoned spreading ridge, see van Staal et al.
2007) and rapid hinge retreat of the east-dipping (present co-ordinates) Dashwoods plate is responsible for formation of
the Lushs Bight oceanic tract infant arc terrane. (b) Stepping-back of the subduction zone in the Taconic (Humber)
Seaway produces the Baie Verte oceanic tract and Notre Dame arc and led to the Taconic 2 arc-continent collision. The
onset of collision slowed down convergence in the Taconic seaway, which is thought to have been responsible for
initiation of west-directed subduction outboard of Dashwoods. The latter led to formation of the Annieopsquotch
ophiolite belt (Lissenberg et al. 2005b).
the LBOT late synkinematically (Szybinski 1995; Notre Dame arc (see below) after its assembly
Swinden et al. 1997); and (3) the c. 488 Ma age of with the overlying ophiolitic rocks during Taconic
the ensialic Cape Ray granodiorite (Dube et al. 2 (Fig. 4).
1996; Whalen et al. 1997b), which cuts the ophioli- Evidence for ductile deformation and meta-
tic Long Range ultramafic –mafic complex, a corre- morphism related to Taconic 1 is relatively cryptic.
lative of the LBOT in southern Newfoundland, and Szybinski (1995) documented c. 493 Ma dykes that
underlying Mischief mélange (Hall & van Staal had intruded late syn-kinematically into chlorite-
1999). The mélange (Fig. 3) contains large knockers rich shear zones. In addition, foliated and flattened
of gabbroic and serpentinized ultramafic rocks, Cambrian basalts of the Sleepy Cove Group of
including dunite and harzburgite, in a predomi- the LBOT in Notre Dame Bay were cut by the
nantly pelitic to semi-pelitic matrix that generally c. 507 Ma Twillingate trondhjemite (Elliot et al.
lacks any coherent stratification (Fig. 6a). The 1991; Dewey 2002), which itself was also foliated
zone of mélange can be traced at least for c. before the emplacement of the c. 473 Ma crosscut-
100 km along strike (Fox & van Berkel 1988; ting Moreton’s Harbour dikes (Williams et al.
Brem et al. 2007). The mélange was strongly 1976), indicating that the LBOT locally has pre-
deformed, metamorphosed and injected by Ordovi- served evidence for a complex structural history
cian plutonic rocks of the 1st and 2nd stages of the prior to and during its accretion to Dashwoods.
Fig. 5. (a) Full-scale Middle Ordovician collision of Dashwoods with the Humber margin (Taconic 2). This phase of the
collision led to significant thickening of Dashwoods and east-directed thrusting at its boundary with the Annieopsquotch
ophiolite belt (AOB). The latter led to closure of the AOB, accretion of parts of the Red Indian Lake arc and hence
formation of the AAT. (b) Late Ordovocian Taconic 3 arc–arc collision (Zagorevski et al. 2007c). Accretion of the
Popelogan –Victoria arc (leading edge of Ganderia) closed the main Iapetan oceanic tract.
In northern Vermont, barroisite in garnet (present co-ordinates) peri-Laurentian arc (Notre

amphibolite that forms part of an ophiolitic sliver Dame arc in Newfoundland; Ascot arc in Quebec),
at Belvidere Mountain yielded 39Ar/40Ar-ages of containing both ensimatic and ensialic segments,
505 –490 Ma (Laird et al. 1993), suggesting that with the Humber margin and obduction of suprasub-
this phase of relatively high-pressure metamorph- duction zone oceanic lithosphere of the intervening
ism and ophiolite emplacement may also be Taconic (Humber in Newfoundland) seaway
related to Taconic 1. The much localized nature of (Fig. 4) (Waldron & van Staal 2001; van Staal
ductile deformation and metamorphism associated et al. 2007; Hibbard et al. 2007; Brem et al.
with this oceanic infant arc –microcontinent col- 2007). The continental part of the arc was built
lision, suggests that Taconic 1 was a relatively soft upon Dashwoods in Newfoundland and upon equiv-
event, mainly characterized by mélange develop- alent continental ribbons elsewhere (e.g. Chain
ment beneath the overriding infant arc and develop- Lakes massif in southern Quebec and adjacent
ment of low-temperature shear zones. Ductile Maine) (Karabinos et al. 1998; Gerbi et al.
deformation and metamorphism were probably 2006a). Lower Ordovician arc plutons, formed
mainly restricted to a relatively narrow subduction during east-directed subduction of the Taconic
channel at depth, remnants of which appears to seaway, intrude into continental basement and its
have been rarely exhumed and/or preserved. LBOT suprastructure in Newfoundland (Hall &
van Staal 1999; van Staal et al. 2007), probably
Taconic 2 severely masking structures formed during
Taconic 1. Closure of the Taconic seaway was
Taconic 2 is the main Ordovician orogenic phase in initiated following choking of the Taconic 1 subduc-
the Appalachians and was due to dextral oblique tion zone after entrance of Dashwoods into the
collision of an Early Ordovician west-facing trench (Fig. 4), which forced subduction to step
Fig. 6. (a) Slab of Taconic 1 mélange (c. 1 m in length), that occurs structurally below the Cambrian ophiolitic
Long Range ultramafic –mafic complex in southwestern Newfoundland. Slab contains a clast of sandstone (left) and
gabbro (right). Pelitic to semi-pelitic matrix was deformed and metamorphosed to sillimanite grade, and injected by
c. 488 Ma and younger granitoid veins. (b) Amphibolite facies ultramylonite of granodiorite and tonalite along the
moderately to shallowly NW-dipping Hungry Mountain thrust. Hand lens for scale. Thrust was active during Middle
Ordovician magmatism (van Staal et al. 2007). Inset shows protolith of mylonite. (c) Amphibolite tectonite with pods
(red) of relict and partly retrogressed eclogite in thrust panel at the base of the Hungry Mountain thrust zone. Scale at
right of photograph. (d) Panel of tilted and weakly foliated Ordovician pillow basalt in the AAT, structurally below the
Hungry Mountain–Lobster Cove thrust system. Rocks locally contain pumpellyite.
back in the remaining oceanic seaway that separated sedimentary rocks to high pressures and tempera-
Dashwoods from Laurentia (van Staal et al. 2007). tures, for example, garnet–clinopyroxene granulite
Remnants of the Taconic 2-related arc can be traced facies conditions at 467–462 Ma (Pehrsson et al.
along the length of the orogen from the southern 2003; van Staal et al. 2007). Such burial of upper
Appalachians into the British Isles, where its accre- plate rocks is explained by widening of the sub-
tion is known as the Grampian orogeny (Dewey & duction channel towards the arc hinterland, such
Shackleton 1984; van Staal et al. 1998). that progressively more of the channel’s hanging
Taconic 2-related arc magmatism is most exten- wall became dragged down and buried (Fig. 8).
sively exposed in Newfoundland, where it is known Deformation of the upper plate also led to
as the 1st stage of the Notre Dame arc (van Staal SE-directed thrusting at the boundary (Fig. 6b) of
et al. 2007), which lasted from 489 –475 Ma. Arc Dashwoods with the Annieopsquotch accretionary
magmatism in Dashwoods has a continental signa- tract (Figs 3 & 4). Tectonic transport in Dashwoods
ture, although the degree of contamination by conti- thus was accommodated by bi-vergent thrusting
nental crust varies along its length, and comprises (Thurlow et al. 1992; Lissenberg et al. 2005a;
both volcanic and plutonic rocks. Geochemical Lissenberg & van Staal 2006). Deformation led to
plots (Fig. 7a, b), together with extended element development of a strong metamorphic foliation
plots, showing well-developed negative Nb ano- (S1) in the buried arc volcanic and sedimentary
malies and relatively flat REE patterns (van Staal rocks, marked by alignment of minerals such as sil-
et al. 2007, their fig. 13) indicate that the mafic limanite, biotite, hornblende and gedrite. S1 was
volcanic and plutonic rocks exhibit island arc refolded into steep orientations by upright or
tholeiite characteristics. steeply inclined, tight to isoclinal F2 folds forming
Stratigraphic and other age constraints suggest a composite S1/S2 transposition foliation during
that arc –continent collision had started, with little terminal collision. The interlimb angle of F2 folds
diachroneity between 480 and 470 Ma along the appears to become tighter and the S2 transposition
full length of the mountain chain, with most foliation better developed on approaching the
collision-related ductile deformation finished by now subvertical Baie Verte–Brompton Line
c. 455 Ma (Friedrich et al. 1999; Chew et al. 2003; (¼ Taconic suture, Williams & St. Julien 1982) in
Castonguay et al. 2001, 2007; van Staal et al. central Newfoundland (van der Velden et al.
2007). Collision-related deformation, mainly 2004). Dashwoods rocks in the high strain zone,
accommodated by west-directed thrusting and immediately east of the suture, locally show
folding (e.g. Tremblay & Pinet 1994; Waldron marked retrogression to greenschist facies phyllo-
et al. 2003), locally seems to have continued until nites and were intruded by syn-tectonic 459–
450 –445 Ma in parts of New England (e.g. Ratcliffe 455 Ma S-type muscovite granite and pegmatite
et al. 1998, 1999) and Quebec (Sacks et al. 2004), (Brem et al. 2007). Such granitic rocks are very
possibly due to convergence driven by subduction rare or absent elsewhere in Dashwoods and its
rollback of old oceanic lithosphere trapped in Notre Dame arc and are obviously spatially and tem-
re-entrants in the Laurentian margin. Blueschists porally associated with the final increments of
and eclogites in the Piedmont of the southern Appa- Taconic deformation localized in the sheared
lachians, Vermont, southern Quebec and the west of rocks present immediately east of the boundary
Ireland, which mainly involved rocks of the conti- with the adjacent Humber margin in central New-
nental margin, yielded ages between 470 and foundland. The S-type melts probably formed by
459 Ma (Laird et al. 1993; Chew et al. 2003; melting of underthrusted Laurentian sediments,
Miller et al. 2006). These high-pressure meta- consistent with the presence of xenocrystic Gren-
morphic rocks thus formed during A –subduction ville and older zircons typical of Laurentia, and
of the leading edge of the Laurentian margin to were channelled into a progressively narrowing
blueschist and locally eclogite facies depths after and steepening shear zone that must have accom-
the start of arc –continent collision. Combined, the modated significant uplift of the arc after peak
available age data indicate little diachroneity in metamorphism. In many places the fault zone also
the time of collision along the full length of the accommodated significant dextral transcurrent
Laurentia’s Appalachian margin. translation (Brem 2007; Brem et al. 2007). Investi-
Taconic 2 led to intense ductile deformation and gations in southern Quebec (e.g. Castonguay et al.
metamorphism, locally reaching granulite facies, in 2001, 2007) and Gaspé (Sacks et al. 2004; Pinchivy
rocks involved in arc thickening and tectonic under- et al. 2003) revealed a comparable Taconic move-
plating in Newfoundland (van Staal et al. 2007) and ment picture involving both orthogonal west-
New England (Karabinos et al. 1998; Gerbi et al. directed thrusting and oblique –dextral tectonic
2006a). Thickening of the Notre Dame arc is mani- transport near the Baie Verte– Brompton line.
fested by the rapid burial of Lower Ordovician Although evidence for ophiolite obduction,
(489 –475 Ma) arc volcanic and associated Taconic thrusting and mélange formation is
Fig. 7. 1st (489– 475 Ma) and 2nd stage (470–455 Ma) Notre Dame Arc plutonism in western Newfoundland.
Mafic samples (,60 wt% SiO2) plotted on: (a) La –Y –Nb diagram (Cabanis & Lecolle 1989); and (b) Th– Zr– Nb
diagram (Wood 1980). Felsic samples (.60 wt% SiO2) plotted on: (c) Rb–Y þ Nb diagram (Pearce 1996) for
2nd stage granitoids (1st stage granitoids were omitted because most exhibited evidence of Rb-mobility, see van
Staal et al. 2007). Beside the diagrams, information on number (N) of samples plotted, sample total (mafic þ felsic)
and % samples plotted in diagram represent of group total is given. Abbreviations: CAB, calc-alkaline basalt;
VAT, volcanic arc tholeiite; N- and E-MORB, normal and enriched mid-ocean-ridge basalt; OIB, ocean island
basalt; BAB, backarc basin basalt; CON, continental tholeiite; VAG, volcanic-arc granite; syn-, post-COLG,
syn-/post-collisional granite; WPG, within-plate granite; and ORG, ocean-ridge granite. See Appendix 1 for
data sources.
Fig. 8. Proposed tectonic evolution of the A– subduction channel beneath Dashwoods. Progressive shallowing of the
Humber margin slab caused progressive widening of the subduction channel, which led to underthrusting of upper plate,
Notre Dame arc (phase 1) rocks to great depths. The widening of subduction channel may also have involved rocks of
the down-going plate, producing a complex kinematic system of sliding lower and upper plate rocks.
widespread in the rocks of the Laurentian margin suture zone of the Baie Verte area of northern
(e.g. Tremblay & Pinet 1994; Waldron et al. 1998, Newfoundland and structural windows east of the
2003), dating using both the 39Ar/40Ar and U– Pb suture in southern Quebec that van Staal et al.
methods (Cawood et al. 1994; Castonguay et al. (2009) and Castonguay et al. (2001, 2007) respect-
2001, Brem 2007) revealed surprisingly little evi- ively measured Taconic metamorphic ages. Alterna-
dence for Taconic metamorphism. Most meta- tively, large strike-slip movements could have
morphism is either Salinic or Acadian. Castonguay excised segments with high-grade Taconic meta-
et al. (2007) suggested that the scarcity of Taconic morphism and juxtaposed parts of the Laurentian
ages west of the Baie Verte –Brompton line in margin that structurally were never deeply buried
southern Quebec was mainly due to nearly complete with high-grade upper plate rocks. For example,
recrystallization and resetting of Taconic micas rocks deposited in first- or second-order re-entrants
during superimposed Silurian (Salinic) and Devo- in the Laurentian margin may have escaped deep
nian (Acadian) orogenesis in Quebec. On the other Taconic-related structural burial, because A –
hand, Brem (2007) replicated the results of subduction ceased or slowed down significantly
Cawood et al. (1994) in the metamorphic rocks of shortly after arrival of promontories at the trench.
the Humber margin of central Newfoundland Such early choking of the A– subduction channel
(internal Humber Zone of Williams 1979) and was inferred to be the main cause for initiation of
failed to detect any evidence for Taconic meta- west-directed subduction at c. 480 Ma, immediately
morphism, despite careful in-situ dating of tiny outboard of the Notre Dame arc in the Iapetus Ocean
zircon and monazite inclusions in garnets grown (van Staal et al. 2007). This led to formation of the
in Salinic amphibolite-facies pelitic rocks immedi- 480–473 Ma suprasubduction zone oceanic litho-
ately west of the Baie Verte–Brompton line. sphere (e.g. Annieopsquotch and Lloyds River
Taconic collision-related metamorphism thus ophiolites, Lissenberg et al. 2005b; Zagorevski
seems to have been restricted mainly to rocks of et al. 2006) (Fig. 4b) and associated arc sequences,
the overriding plate and underplated Laurentian such as the Roberts Arm Group in Newfoundland
sediments, but appears generally weak or absent in and Boil Mountain complex in Maine (Gerbi et al.
metamorphic tectonites developed in rocks of the 2006b). The Annieopsquotch and Lloyds River
immediately adjacent Humber margin (van Staal ophiolites were partially subducted westwards
et al. 2007). The lack of evidence for this event in beneath Dashwoods shortly after their formation at
the adjoining metamorphic tectonites of the c. 468 Ma (Lissenberg et al. 2005b). This event
Humber margin is puzzling, particularly consider- established the east-facing Annieopsquotch accre-
ing the relatively large database of radiometric tionary tract (Fig. 5a) (AAT, van Staal et al. 1998;
ages. Hence, it is unlikely to be due to inadequate van der Velden et al. 2004). The AAT progressively
dating. One possible solution is that Taconically expanded eastwards over time due to accretion
thickened and deeply buried parts of the Humber and underplating of oceanic and arc lithosphere
margin were later (post-peak Taconic metamorph- that was originally situated slightly outboard of
ism) overthrust and hidden beneath rocks of the the Dashwoods microcontinent (Zagorevski et al.
Notre Dame arc and/or adjacent fore-arc ophiolites 2006, 2007b, c). Amphibolite and eclogite assem-
of the Baie Verte oceanic tract (BVOT; van Staal blages in the most inboard panels of the AAT
et al. 2007). It is in such rocks, exposed in the (Fig. 6c) indicate that these rocks were subducted
to relatively deep depths (Lissenberg & van Staal arc), previously accreted to Dashwoods during
2006; Zagorevski et al. 2007d). Taconic 2 during closure of the Lloyds River
Middle Ordovician Taconic 2 deformation and backarc basin (Lissenberg et al. 2005b; Zagorevski
metamorphism were accompanied by voluminous et al. 2006) with a west-facing peri-Gondwanan arc
granitoid magmatism (van Staal et al. 2007). This (Popelogan –Victoria arc), which was built on the
is the 2nd stage of Notre Dame arc magmatism leading edge of Ganderia (van Staal et al. 1998;
(Figs 3 & 7b). Its mafic end-members exhibit Zagorevski et al. 2007a, c) (Fig. 5b). Taconic 3
mainly calc-alkaline to arc tholeiite characteristics led to assembly of arc volcanic and associated sedi-
but also include non-arc-like N- and E-MORB com- mentary rocks into a series of west-dipping struc-
positions (Fig. 7a, b). Almost equal proportions of tural panels at the leading edge of the AAT. The
mafic and felsic compositions attest to the major structures and other evidence supporting these two
role played by mantle-derived magmas in this processes are described in more detail by Zagor-
2nd stage of Notre Dame arc magmatism. Felsic evski et al. (2007b, c). The presence of Laurentian-
compositions exhibit volcanic-arc granite (VAG) derived detrital zircon populations and Notre Dame
signatures, plotting mainly to the left or below the arc tonalite clasts in the late Ordovician (450 –
area of overlap between VAG and post-collisional 440 Ma) sandstones (Fig. 9a) overlying the Victoria
granites (post-COLG) (Fig. 7c). Such low-Rb and arc in Newfoundland (McNicoll et al. 2001;
Y þ Nb contents indicate derivation from mainly O’Brien 2003) confirmed that these two opposing
primitive depleted sources. However, the range of arcs were assembled at this time and hence the
1Nd(T) values (214 to þ1) and infracrustal d18O main tract of the Iapetus Ocean had closed by
(VSMOW) values (þ5 to þ10‰) exhibited by this stage.
this magmatism (Whalen et al. 1997a) indicate it Mélanges (Figs 3 & 9b; Dunnage and Sops
formed within Precambrian Laurentian crust. Head-Boones Point mélanges), west-dipping,
Characteristics of 2nd stage Notre Dame arc brittle-ductile, sinistral oblique reverse faults, tight
magmatism, such as the presence of non-arc-like overturned folds and localized low-grade metamor-
athenospheric mantle-derived magmas, support a phism with stilpnomelane- and/or phengite-bearing
model in which progressive removal of the bulk of greenschist, pumpellyite–actinolite and prehnite-
the arc’s mantle lithosphere during the final stages pumpellyite facies assemblages (Fig. 6d) character-
of collision and breakoff of the oceanic part of the ize the style of tectonometamorphism of the rocks
downgoing Taconic seaway slab induced rapid accreted as a result of Taconic 3 in the AAT (e.g.
upwelling of the asthenosphere, which caused large- Bostock 1988; Zagorevski et al. 2006, 2007b, c).
scale melting of the arc’s infrastructure (Fig. 5a). Such metamorphic assemblages, particularly the
Melting led to a relatively short-lived (464– presence of pumpellyite and phengite, suggest rela-
461 Ma), but a very voluminous tonalite bloom, tively low geothermal gradients typical of accretion-
which overlapped with D2 (van Staal et al. 2007) ary complexes formed above subduction zones.
and was rapidly followed by uplift and exhumation
of the arc’s infrastructure. The latter was already
exhumed and exposed at the surface by the end of
the Caradoc (c. 453 Ma). Collision-related Taconic Salinic orogeny
magmatism thus was mainly confined to the site of
the original arc. The Salinic orogeny has only recently been recog-
nized as a significant tectonic event, although its
Taconic 3 imprint on the geology of the Ganderian core of
the northern Appalachians had been recognized as
Rocks involved in Taconic 3 mainly occur in a early as the 1960s (e.g. Boucot 1962). Structural
narrow belt situated along the Red Indian Line investigations in conjunction with extensive radio-
(RIL, Williams et al. 1988), which is the principal metric age dating of critical rock units highlighted
Iapetan suture in the northern Appalachians the regional importance of this orogenic event,
(Figs 1, 3 & 5), along which rocks with Laurentian which has been documented in Newfoundland
and Gondwanan provenances have been juxtaposed. (Dunning et al. 1990), New Brunswick (van Staal
The Red Indian Line (RIL) is largely covered in & de Roo 1995) and Maine (West et al. 1992;
New Brunswick and New England by younger Hibbard 1994). Salinic orogenesis is kinematically
rocks; hence the nature of Taconic 3 is largely and dynamically distinct from the predominantly
based on our research in Newfoundland. Rocks Early–Middle Devonian Acadian orogeny (van
affected by Taconic 3 were deformed and metamor- Staal 2007). Structural evidence suggests that
phosed to low grade conditions as a result of a Salinic orogenesis was due to sinistral oblique
Caradoc (460 –450 Ma) collision between an east- convergence, whereas the obliquity of the Acadian
facing Middle Ordovician peri-Laurentian (RIL appears predominantly dextral (Holdsworth 1994;
Fig. 9. (a) Poorly sorted matrix-supported conglomerate (olistostrome) in the Upper Ordovician Point Leamington
Formation, Badger Group, a few hundred metres south of the Red Indian Line, SE of Cottrel’s Cove, Notre Dame Bay,
Newfoundland. Inset: detail of a tonalite cobble clast typical of the Notre Dame arc, indicating that the latter was uplifted
due to Taconic 3 arc–arc collision and supplied detritus to the Badger basin. The latter in part is a new, post-Taconic
basin associated with the downgoing Salinic slab (oceanic lithosphere of the Tetagouche–Exploits basin). (a) Zagorevski
for scale. (b) Dunnage mélange with clasts of basalt and sandstone in a scaly cleaved shaly matrix. Hammer for scale.
Hibbard 1994; van Staal 1994; van Staal & de Caledonides; hence, the coeval Scandian and Salinic
Roo 1995). orgenies are both dynamically related to collision
Salinic orogenesis was mainly due to a mid- between Laurentia and an enlarged and composite
Silurian (430–422 Ma) collision between the Baltica (Valverde-Vaquero et al. 2006; van Staal
Gander margin and composite Laurentia following & Hatcher 2009). The Tetagouche –Exploits
terminal closure of the intervening, wide Teta- back –arc basin had started to open during the
gouche–Exploits backarc basin (Fig. 10a). This Early Ordovician (c. 475 Ma) as a result of rifting
basin separated its passive Gander margin (Fig. 3) of the Popelogan–Victoria arc (van Staal et al.
from Ganderia’s leading edge, the Popelogan– 1998, 2003; Valverde-Vaquero et al. 2006). Back-
Victoria arc (van Staal 1994, 2007). Ganderia had arc spreading responsible for the Tetagouche –
already accreted to Baltica during the Ashgill Shel- Exploits back-arc basin lasted 15– 20 Ma and led
vian orogeny at its eastern extremity in the British to a significant oceanic basin, estimated on
Fig. 10. (a) Late Ordovician–Silurian closure of the Tetagouche– Exploits back-arc basin, which is the principal cause
of the Salinic orogeny; (b) Silurian closure of the Acadian seaway that separated Ganderia and Avalonia, which led to
the Acadian orogeny. The early stages of the Acadian (421–417 Ma) were localized preferentially in the back-arc and/
or intra-arc basins. Note that most of the fore-arc is assumed to have been subducted to a position beneath the coastal arc.
Late stages of the Acadian mainly involve antithetic retro-arc thrusting towards the orogen’s hinterland and dextral
strike-slip on steep, orogen-parallel faults. (c) Accretion of Meguma, which is interpreted to have been accompanied by
wedging and breakoff of the downgoing Rheic slab. A new west-dipping subduction zone was probably established
outboard of Meguma, necessary to accommodate convergence and Alleghenian collision with Gondwana. Wedging and
lower crustal architecture in (b) and (c) is largely based on seismic interpretations (Keen et al. 1991; van der Velden
et al. 2004).
palaeontological and palaeomagnetic evidence to 2007) that spatially overlaps the Taconic-related
have been in the order of 1000–1500 km arc and collision-related magmatism (Figs 1 &
wide. The Japan Sea is a close modern analogue 10a; van Staal et al. 2007). Mafic and felsic plutonic
(van Staal et al. 1991, 2003). Closure of the rocks are present in equal proportions and all
back-arc basin started at c. 450 Ma, immediately samples exhibit well-developed negative Nb
after the Popelogan– Victoria arc had accreted to anomalies on extended element plots (see fig. 6b
Laurentia at c. 455 Ma (van Staal 1994; van Staal of Whalen et al. 2006). Geochemical plots for
et al. 2008). Closure was initiated following these magmatic rocks (Fig. 13) indicate that mafic
stepping-back of the west-dipping subduction zone components exhibit calc-alkaline arc signatures
that had rimmed Laurentia during Taconic 3 and felsic components exhibit VAG signatures,
(Fig. 10a; van Staal et al. 1998). Subduction of the which are on average more Rb and Y þ Nb
backarc’s oceanic lithosphere led to sporadic rich than 2nd stage Notre Dame arc tonalitic
Late Ordovician –Early Silurian (445 –435 Ma) arc rocks (Fig. 7c). In their 1Nd(T) values these
magmatism in Newfoundland (Fig. 3; 3rd phase of plutons exhibit both contaminated and juvenile sig-
Notre Dame arc, Whalen et al. 2006; van Staal natures (Whalen et al. 2006). Combined, these
characteristics suggest this magmatism is Deformation in the Brunswick subduction complex

subduction-related. was closely monitored by tectonism in the adjacent
Equivalents of this magmatism (mainly volcanic Matapedia fore-arc basin, which exhibits several
rocks) occur also in northern New Brunswick stages of localized Llandoverian deformation
(c. 430 Ma volcanic rocks of the Weir Formation, (Fig. 11d, e) and sedimentary onlap onto exhumed
Wilson et al. 2008) and adjacent Quebec parts of the accretionary wedge (van Staal & de
(c. 433 Ma volcanic rocks of the Pointe aux Trem- Roo 1995; Dimitrov et al. 2004; van Staal et al.
bles and Lac Raymond formations, David & 2008). As in New Brunswick, ample evidence is
Gariepy 1990). In northern Maine, this phase of preserved for Llandoverian deformation with
arc magmatism is represented by the c. 443 Ma SE-directed tectonic transport in the same tectonic
Attean pluton in the Chain Lakes massif (Gerbi setting in Newfoundland. Fossil-bearing olistos-
et al. 2006b) and the coeval Quimby volcanic tromes (e.g. Joey’s Cove Mélange), which had
rocks (Moench & Aleinikoff 2003). Correlatives developed in front of seafloor-breaching thrusts,
also occur further south in the Bronson Hill arc. were partly overridden and deformed into mélange
by the same faults (van der Pluijm 1986; P. F.
Salinic accretionary phase (445 – 430 Ma) Williams et al. 1988). Elsewhere, Llandovery
thrust-related deformation is tightly constrained by
Closure of the Tetagouche –Exploits backarc basin syn-tectonic intrusions and/or stratigraphy. The
led to punctuated latest Ordovician–Early Silurian intrusives cut the enclosing fault-related rocks
accretion of seamounts and isolated back-arc such as mélange, but were also deformed by sub-
crustal ribbons (e.g. van Staal et al. 2003, 2008) sequent increments of the responsible progressive
into a progressively expanding subduction complex deformation (Zagorevski et al. 2007b). The regional
(Fig. 10a), which is exceptionally well-preserved in extent of Llandoverian Salinic deformation is not
northern New Brunswick (van Staal et al. 2001, well known, as it was gradually followed by more
2003, 2008), but poorly-preserved and/or largely widespread deformation during the Wenlock
covered by younger Siluro–Devonian sedimentary (Figs 11c & 12a), which produced, at least in part,
rocks in Newfoundland and Maine. Growth of the similar structures (O’Brien 2003).
subduction complex and progressive stepping back
of the subduction zone was concomitant with an
eastward migration of relatively deep marine Salinic collision-related orogenesis
basin(s) situated in the arc-trench gap (Matapedia- (430– 422 Ma)
Badger basin) and its adjacent hinterland during the
Late Ordovician –Early Silurian (Elliot et al. 1989; Salinic deformation became regional in extent near
Williams et al. 1995; van Staal & de Roo 1995; the Llandovery–Wenlock boundary (c. 429 Ma),
van Staal 2007; Pollock et al. 2007). Zeolite, now also affecting rocks of the Gander margin fore-
prehnite–pumpellyite, pumpellyite–actinolite, greenland, mainly underlain by Cambrian –Ordovician
schist and blueschist facies metamorphism (Richter rocks of the Gander Group and correlatives else-
& Roy 1974; van der Pluijm 1986; van Staal et al. where, and rocks in the Laurentian hinterland. Pro-
1990, 2008; O’Brien 2003; Wilson 2003) accom- gressive widening of the Salinic deformation zone
panied deformation of the rocks in the subduction probably reflects the onset of full-scale collision
complex and adjacent fore-arc basin. between the Gander margin and Laurentia, which
Deformation during this pre-collisional stage of sutured along the Dog Bay Line (DBL) in New-
the Salinic was mainly restricted to the arc-trench foundland (Williams et al. 1993, Valverde-Vaquero
gap region and led to predominantly east- or south- et al. 2006) and the Bamford Brook-Liberty-
directed ductile-brittle thrusting, localized folding Orrington fault system (Figs 1 & 9a) in New Bruns-
and mélange formation (Fig. 11a, b). Tectonic wick and adjacent New England (Ludman et al.
transport direction in Newfoundland locally varied 1993; van Staal 2005, 2007; Hibbard et al. 2006).
from SE to south or SW (e.g. O’Brien et al. 1997; The suture separates two distinct Middle Ordovi-
O’Brien 2003), which possibly reflects the buttres- cian –Silurian sequences (Fig. 3), which are particu-
sing effects of relatively rigid salients in the modi- larly obvious in Newfoundland where Wenlock
fied Laurentian margin, such as the Notre Dame terrestrial rocks (Botwood Group) are separated
Bay flexure. Deformation in the Brunswick sub- from coeval marine sediments (Indian Island
duction complex in New Brunswick was mainly Group), each with a distinctly different provenance
due to underthrusting and underplating, which led (Pollock et al. 2007). The Lower Silurian prove-
to a general, non-coaxial deformation that was nance of the marine foreland sedimentary rocks
accompanied by some of the best developed high- east of the suture is Ganderian, whereas rocks
pressure low-temperature metamorphic tectonites deposited in the arc-trench gap west of the suture
in the Appalachians (van Staal et al. 1990, 2008). have a predominantly Laurentian provenance and
Fig. 11. (a) Uppermost Ordovician– Lower Silurian mélange situated structurally beneath the ophiolitic Fournier
block, Brunswick subduction complex, northern New Brunswick. Pen for scale. (b) Close-up of mélange matrix. (c)
Nearly isoclinal pre-Devonian, Salinic folds in Early Llandovery calcareous mudstones of the Matapedia Group (part of
Matapedia fore-arc basin associated with the Brunswick subduction complex) near New Brunswick–Maine border.
Bottom of photo is c. 1.5m wide. Photo courtesy of R. Wilson of the New Brunswick Geological Survey. (d) Angular
unconformity between folded (Salinic) Lower Silurian (Llandovery–Wenlock boundary) limestones of the La Veille
Formation and upper Silurian calcareous conglomerate of the Simpsons Field Formation (Dimitrov et al. 2004),
Limestone Point, northern New Brunswick. S. McCutcheon of the New Brunswick Geological Survey outlines bedding
(S) and cleavage (Sc) in the La Veille Formation. Hammer parallel to trace of bedding in calcareous conglomerate of
Simpsons Field Formation. Cleavage overprints unconformity and is Acadian. (e) Same outcrop and rocks (broken
also include volcanic rocks. The lithological differ- dextral transpressive deformation (Williams et al.
ences between the Silurian rocks east and west of the 1993). Several lines of evidence (e.g. metamorph-
suture are more subtle in Maine and parts of New ism) suggest that the earliest motion on the DBL
Brunswick, because the fore-arc basin remained and mélange formation involved Silurian reverse
largely marine and was probably simple and movement, which was probably synthetic with the
sloped (Dickinson & Seely 1979) such that its sedi- SE-directed Silurian thrust faults that deform the
ments could spill over from the fore-arc basin into adjacent Badger Group and associated olistostromes
the adjacent trench (Fredericton trough) immedi- (Williams et al. 1993; Valverde-Vaquero et al.
ately east of the suture. A distinctive feature of the 2006; Pollock et al. 2007). The DBL and associated
Fredericton trough is that Silurian volcanic rocks transpressive structures were cut by the late
are absent, while they are locally present in the Silurian –Early Devonian (424 –411 Ma) phases of
fore-arc basin. the Mount Peyton intrusive suite (Dickson et al.
The Bamford Brook-Liberty-Orrington fault 2007). In general, the rocks on both sides of the
system that marks the suture in northern New DBL underwent a significant phase of Silurian
England was originally a major SE-directed thrust deformation and metamorphism prior to the Devo-
that had emplaced Cambrian –Ordovician meta- nian (e.g. Dunning et al. 1990). Gander Group
morphic rocks of the Miramichi and Liberty Orring- rocks east of the DBL in eastern Newfoundland
ton belts above Lower Silurian rocks of the were strongly folded, foliated and metamorphosed
Fredericton Trough (e.g. Ludman et al. 1993; up to migmatite grade at c. 425 Ma (d’Lemos
Tucker et al. 2001; van Staal et al. 2003). Thrust et al. 1997; Scofield & d’Lemos 2000) during
movements and penetrative ductile deformation in a regional sinistral transpressive deformation
the Fredericton trough took place before intrusion (Holdsworth 1994). Silurian deformation and meta-
of the Upper Silurian, 422 + 3 Ma Pocomoonshine morphism, likewise tightly constrained by U –Pb
gabbro (West et al. 1992) and 418 + 1 Ma Lincoln ages of synkinematic or cross-cutting plutons, mig-
syenite (West et al. 2003) in Maine. In New Bruns- matites and high-grade metamorphic minerals, has
wick, the Bamford Brook fault and strongly folded also been documented in Cambrian –Ordovician
turbidites of the Fredericton trough were cut by rocks in southern and southwestern Newfoundland
the 414– 411 Ma Pokiok Batholith (McLeod et al. (Dunning et al. 1990; van Staal et al. 1994).
2003). Furthermore, the 420 + 5 Ma Mohannes Salinic orogenesis led to significant uplift
granite (Fyffe & Bevier 1992), which intruded the (Figs 11d & 12b) and locally rapid exhumation of
Ordovician rocks of the Cookson Group (St. Croix rocks that were incorporated in the hinterland of
belt) that underlies the Fredericton trough, was the orogen, both in the northern New England
either synchronous or postdated the earliest defor- (Maine, New Brunswick and Quebec) and New-
mation in the poly-deformed Cookson Group. The foundland Appalachians (van Staal & de Roo
latter is a correlative of the Gander Group in 1995; van Staal et al. 2007, 2008). Unconformable
eastern Newfoundland (van Staal & Fyffe 1995) deposition of upper Llandovery– Wenlock terrestrial
and hence forms part of the Gander margin clastic rocks above Orovician tectonites, originally
(Fig. 3, van Staal 1994). Unpublished U –Pb ages buried and metamorphosed during Ordovician
(V. McNicoll, pers. comm.) revealed that Ordovi- Taconic shortening of the Notre Dame arc in New-
cian rocks of the Miramichi highlands in the foundland, took place during and immediately
hanging wall of the Bamford Brook fault were after probably faulting-associated late Llandovery
metamorphosed locally up to migmatite grade, (c. 430 Ma) uplift and unroofing (van Staal et al.
during the Silurian. In addition, a detailed struct- 2007). An excellent example is represented by
ural analysis by Park & Whitehead (2003) docu- the regionally extensive Lloyd’s River-Hungry
mented early east-vergent Silurian thrusting and Mountain-Lobster Cove fault system (Figs 1 & 3),
localized recumbent folding in the Fredericton which originally was a major SE-directed Ordovi-
trough, followed by steepening of the structures cian thrust system (Fig. 6b) that accommodated
prior to- or coeval with Early Devonian dextral underthrusting and accretion of ophiolites of the
transpressive deformation. AAT (Lissenberg et al. 2005a; Lissenberg & van
The early movement history accommodated by Staal 2006; Zagorevski et al. 2006) to Laurentia.
the DBL in Newfoundland is poorly known, This fault system was reactivated during the Silurian
although it involved mélange formation prior to as a SE-directed reverse fault (Dean & Strong 1978),
Fig. 11. (Continued) formation) as (f) unconformably overlain with erosional contact by red conglomerate of the Upper
Llandovery (C5– C6) Weir Formation. Red beds were folded and cleaved together with the underlying calcareous
sandstones during both the Salinic and Acadian orogenies. (f) Folded lower Llandovery calcareous sandstones and
shales, which were earlier deformed (early Salinic) into broken formation. Bottom of photo is c. 2m wide.
Fig. 12. (a) Narrow, Salinic SE-directed shear zone (thrust) near Baie Verte, northern Newfoundland. SE is to the
right. Hammer for scale. (b) Greenschist facies (cold) mylonite with boudinaged quartz veins marking the Lobster Cove
fault near Lobster Harbour, Sunday Cove Island, Newfoundland. Blue ballpoint for scale. Shearbands (poorly visible in
photograph) and other shear sense indicators indicate that thrust movement was to the SE over the underlying Silurian
(Wenlock) red beds. (c) Folded Silurian red beds in the immediate footwall of the Lobster Cove fault. A. Zagorevski for
scale. (d) Pristine, cross-bedded red beds in footwall of Lobster Cove fault. (e) Large southerly-overturned F2 recumbent
fold in an apophyse of the 418 + 2 Ma Rose Blanche granite. The granite veins probably intruded during F2 (van Staal
et al. 1992; Valverde-Vaquero et al. 2000) because they cut co-planar south-verging tighter F2 isoclines (out of view),
S1 (layer-parallel) and the S2 axial plane foliations in the sillimanite schists and gneisses, east of Port aux Basques,
such that unmetamorphosed Wenlock red beds of felsic compositions to that west of the RIL but its
(Fig. 12c, d) and volcanic rocks now locally occur mafic rocks are limited to transitional calc-alkaline
in the footwall of the fault, whereas uplifted and to tholeiitic basalt compositions (Fig. 14). This mag-
unroofed Cambrian and Ordovician metamorphic matism, like that west of the RIL, is attributed to
rocks occur in the hanging wall (Fig. 12b). Detritus breakoff of the Salinic slab (Fig. 10a). However,
of the hanging wall occurs in the footwall’s Silurian the magmatism occurring east of the RIL and west
clastic rocks. Hence, the red beds date the defor- of the DBL received only input from depleted
mation associated with Silurian reverse faulting mantle sources, which contrasts with coeval
(Zagorevski et al. 2007b). magmatism occurring west of the RIL. The latter
Widening of the deformation zone into the probably was situated immediately above the
retro-arc hinterland of the orogen also led to signifi- upwelling asthenosphere that replaced the sinking
cant reactivation of Taconic structures in the Salinic slab, whereas the former was situated
Humber margin of Laurentia and renewed imbrica- near the eastern boundary of the area affected by
tion and metamorphism in a relatively narrow belt slab-breakoff. In NW New Brunswick and Gaspé,
near the Baie Vverte Brompton Line (Taconic Quebec, most equivalent Silurian magmatism is
suture), both in Newfoundland (Cawood et al. obscured by younger cover sequences. Our dataset
1994; Brem et al. 2007) and Quebec (Castonguay only includes a small number of mafic rock
et al. 2007). Salinic recrystallization and meta- analyses, which exhibit transitional calc-alkaline
morphism were apparently so penetrative in this to continental tholeiitic basalt compositions
belt that it may have destroyed nearly all evidence (Fig. 15a, b).
for earlier Taconic metamorphism. Strain localiz-
ation in this belt may be due to a combination of
fabric and thermal softening; the latter possibly Acadian orogeny
related to heat generated during slab-breakoff
and/or lithospheric thinning prior to this retro-arc The Acadian orogeny is generally attributed to
deformation. accretion of Avalonia to Laurentia (e.g. Bird &
Temporal-compositional changes in voluminous Dewey 1970; Bradley 1983), a causative mechan-
433 –425 Ma magmatism west of the RIL have been ism retained herein. Points of contention remain
documented by Whalen et al. (2006) and interpreted the start and duration of Acadian orogenesis, the
in terms of a slab-breakoff model of the oceanic location of the suture and the polarity of subduction
segment attached to the downgoing Tetagouche- (Bradley 1983; Robinson et al. 1998; Tucker et al.
Exploits back-arc lithosphere (Salinic slab). The 2001; van Staal 2007). The discussions particularly
compositional spectrum includes calc-alkaline to hinge on the position of the suture between Avalonia
tholeiitic basalts and also non-arc like mafic compo- and Laurentia, the leading edge of which at this
sitions and felsic magmatism includes both VAG stage was represented by Salinic-accreted Ganderia.
and WPG type suites (Fig. 13). In their 1Nd(T) Several lines of evidence suggest that the suture lies
values most of these plutons exhibit juvenile (þ1 along the Hermitage Bay–Dover fault in New-
to þ6) signatures, though there are some slightly foundland and Caledonia fault in New Brunswick
negative values. In general, these geochemical fea- (Fig. 1) (Samson et al. 2000; Barr et al. 1998,
tures substantiate that breakoff of the Salinic slab 2003; van Staal et al. 2004; van Staal 2005; Lin
led to sequential tapping of a combination of asthe- et al. 2007). The latter fault is sinistrally offset by
nospheric, lithospheric and crustal sources both the Oak Bay fault (Fig. 1) to continue south of
proximal to the RIL and in the hinterland of the Grand Manan Island beneath the Gulf of Maine
orogen. Silurian magmatism overlapped with defor- and to reappear on land as the Bloody Bluff fault
mation, is locally associated with high-grade meta- in Massachusetts (Hibbard et al. 2006). Part of the
morphism and migmatites and appears to become problem previously identifying with this suture
younger towards the foreland. Silurian magmatism was that Ganderia and Avalonia both have a Neo-
east of the RIL and west of the DBL in west-central proterozoic arc-like basement, which could be sep-
Newfoundland includes parts of the voluminous arated mainly on the basis of detailed isotopic
Mt. Peyton, Fogo (Currie 2003) and Hodges Hill (Kerr et al. 1995; Whalen et al. 1996a; Samson
suites (O’Brien 2003). It exhibits a similar spectrum et al. 2000; Potter et al. 2008) and tectonothermal
Fig. 12. (Continued) south coast of Newfoundland, which form part of the Gander margin. These Acadian folds occur
all along Newfoundland’s south coast and are interpreted to have formed during the early stages of collision between
Avalonia (lower plate) and Laurentia’s leading edge (represented by Ganderia at this stage). S. Colman-Sadd of the
Newfoundland Geological Survey for scale. (f) Silurian mafic dyke complex (sheeted?) in the Mascarene basin,
southern New Brunswick. P. Valverde-Vaquero for scale.
Fig. 13. 3rd Stage (445– 435 Ma) Notre Dame Arc plutonism and Salinic slab break-off related magmatism (433–
423 Ma) in western Newfoundland. Diagrams (a, b, c) and abbreviations as in Figure 7. See Appendix 1 for data sources.
studies (Barr & White 1996). More significantly, the situated west of the arc (van Staal 2007); (2) tectonic
Early Palaeozoic tectonic histories of these two ter- slivers, characterized by 420–416 Ma subduction-
ranes are markedly different (Fig. 3) (van Staal et al. related fore-arc high pressure –low temperature
2004; van Staal 2007; Lin et al. 2007). metamorphism, occur immediately to the east of
The polarity and timing of the Acadian orogeny the coastal arc in southern New Brunswick, very
(Fig. 10b) is best constrained by the following three close to the Ganderia–Avalonia suture (White
observations: (1) geometry of Silurian arc and et al. 2006); and (3) Silurian shallow water shelf
backarc magmatism (442 –423 Ma) on the trailing sediments (Arisaig Group) deposited on Avalonia
edge of Ganderia facing Avalonia. Backarc is were overlain by a thick sequence of foreland
Fig. 14. Silurian plutons emplaced between the Red Indian Line (RIL) and the Dog Bay Line (DBL) in west-central
Newfoundland. Diagrams (a, b, c) and abbreviations as in Figure 7. See Appendix 1 for data sources.
basin deposits during the latest Silurian –Early Barr et al. 2002; Miller & Fyffe 2002; Lin et al.
Devonian, due to tectonic loading of Avalonia 2007) and beyond into Massachusetts (Hepburn
(Waldron et al. 1996a). Avalonia, thus, was the 2008). They have been recently named the Kingston
lower plate and subduction was to the west or NW arc terrane (Barr et al. 2002; White et al. 2006; van
beneath Laurentia (Fig. 10b). The Silurian Staal 2007), but were previously referred to as the
Acadian arc- and back-arc volcanic rocks coastal (volcanic) arc (Bradley 1983), a name we
(Fig. 12f) can be traced from southern Newfound- prefer because the arc- and back-arc volcanic
land (e.g. La Poile Group and Burgeo granite) rocks define a belt that closely follows the Atlantic
through Cape Breton Island (e.g. Money Point coast of eastern North America (Fig. 2).
Group –Sarach Brook suite and correlatives) to The timing of tectonic loading of Avalonia’s
southern New Brunswick (Kingston and Mascarene leading edge suggests that the Avalonia–Laurentia
terranes) (Barr & Jamieson 1991; Price et al. 1999; collision had started during the Late Silurian at
Fig. 15. Silurian magmatism in Gaspé (Quebec) and New Brunswick. Diagrams (a, b, c, d) and abbreviations as in
Figure 7. See Appendix 1 for data sources.
c. 421 Ma in maritime Canada, which is virtually Bocabec plutons, Fyffe et al. 1999). The latter, and
coeval with inversion of back-arc and/or intra-arc consanguineous phases nearby, yielded ages of
basins, such as the Mascarene (Fig. 12f) and La 423–421 Ma (McLaughlin et al. 2003), suggesting
Poile basins in New Brunswick and Newfoundland, a very rapid transition from back-arc and/or
respectively (O’Brien et al. 1991; Fyffe et al. 1999) intra-arc extension to compression following the
following shut-off of coastal arc volcanism. onset of collision.
Deformed and foliated Mascarene back-arc basin The large collection of new radiometric ages
rocks in southern New Brunswick, which include indicates that Salinic- and Acadian-related arc mag-
volcanic rocks as young as 423 + 1 Ma (van matism partly overlapped in time during the Early
Wagoner et al. 2002), were cut by Upper Silurian Silurian (Fig. 10a, b). Hence, while the coastal arc
phases of the St. George batholith (e.g. Utopia and was in the early stages of being constructed upon
Ganderia’s trailing edge, the Gander margin’s and associated thrust or reverse faults have also
leading edge was in the final stages of being sub- been mapped along Newfoundland’s south coast in
ducted beneath Laurentia (van Staal 2007). It is Lower Palaeozoic rocks of the Gander margin
thus feasible that pre-Acadian transtension-related (Colman-Sadd 1980; Blackwood 1985; van Staal
deformation in the coastal arc (e.g. Doig et al. et al. 1994). Acadian folds deform an earlier layer-
1990; McLeod et al. 2001) induced by oblique subparallel metamorphic foliation, which could be an
duction of the narrow oceanic seaway that separated early Acadian or Salinic structure (van Staal et al.
Ganderia and Avalonia (Valverde-Vaquero et al. 1992). Development of the recumbent or overturned
2006) overlaps in time, although not in space folds and associated shear zones partly overlapped
(Fig. 2) with the waning stages of Salinic orogenesis with high-grade metamorphism, dated at c. 418–
in the other parts of Ganderia (e.g. Hibbard 1994). 414 Ma (Burgess et al. 1995; van Staal et al.
More problematic is defining the time span of the 1994), and were intruded late syn-kinematically by
Acadian orogeny, because the retro-arc deformation Upper Silurian (421 –418 Ma) biotite + muscovite
front (see below) and associated foreland basin pro- granites such as the Gaultois (c. 421 Ma, Piasecki
gressively migrated into the Laurentian hinterland 1988; Dunning et al. 1990) and Rose Blanche
during the Devonian (Bradley et al. 2000). It plutons (c. 418 Ma, van Staal et al. 1992; Valverde-
reaches its final limit at the Appalachian structural Vaquero et al. 2000). Correlative Gander margin
front in the early Middle Devonian at a time when rocks in southern New Brunswick (St. Croix belt)
Meguma had already started to collide with Lauren- also may contain early southerly overturned struc-
tia (Neoacadian, see below). tures (Fyffe et al. 1991), although the dominant
Acadian structural transport direction here and in
adjacent Maine appears to be towards the NW
Late Silurian (421 – 416 Ma), early (Tucker et al. 2001; Castonguay et al. 2003). Such
Acadian orogenesis a movement picture is consistent with the asymme-
try of structures developed during the latest Silurian
Structures formed during the early stages of the inversion of the La Poile and Mascarene intra-arc
Acadian occur close to the suture in the region and/or back-arc basins in Newfoundland and New
mainly underlain by rocks of the coastal arc- Brunswick/Maine, respectively (O’Brien et al.
backarc system (Figs 1 & 2), that is, near the Atlan- 1991; Fyffe et al. 1999; Tucker et al. 2001). Early
tic coast. Metamorphic tectonites related to Acadian Acadian shortening thus appears to have produced
subduction have been preserved in a narrow belt bi-vergent structures, recording both SE and
underlain by the Pocologan metamorphic suite in NW-directed structural transport and was preferen-
southern New Brunswick (White et al. 2006). Corre- tially localized in the back-arc region.
latives may be hidden beneath the seafloor of the Early Acadian metamorphism varied regionally,
Gulf of St. Lawrence, offshore Newfoundland’s but Gander margin rocks in southern Newfoundland
south coast, but the overall scarcity of identifiable and New Brunswick (St. Croix belt) generally
Palaeozoic arc-trench gap rocks between the display Barrovian to Buchan style metamorphism
coastal arc and Avalonia either suggests major sub- (Colman-Sadd 1980; Fyffe et al. 1991), although
duction erosion of the fore-arc block following locally more deeply buried rocks have been
establishment of a flat-slab during the Early Devo- exhumed (e.g. Burgess et al. 1995). Metamorphic
nian (Murphy et al. 1999, see below) shortly after isograds generally seem to dip shallowly to the
the onset of collision (Fig. 10b), or to structural exci- NW or north and high-grade rocks have been juxta-
sion related to orogen-parallel transcurrent faulting. posed locally with low-grade rocks along major
Subduction erosion or widening of the subduction faults, suggesting significant and complex post-
channel may have abraded and dragged down metamorphic movements along some of the
most of the fore-arc’s crust in the hanging wall of belt-bounding faults.
the subduction zone into the mantle to a position
below the arc. Such underthrusting of crustal rocks
may be important during subsequent Acadian mag- Early Devonian (416– 400 Ma) late
matism (e.g. Kay et al. 2005, see below). Consider- Acadian orogenesis
ing the lack of preservation of arc-trench gap rocks
in general, it is not surprising that structures syn- Acadian orogenesis continued uninterrupted into
thetic with west-directed Acadian subduction are the Early Devonian. Structural analyses suggest it
relatively rare, although they have been imaged at involved a combination of orogen-normal shorten-
depth on the seismic profile crossing the Acadian ing and dextral movements along orogen-parallel
suture in southeastern Newfoundland (van der structures (e.g. Hibbard 1994; van Staal & de Roo
Velden et al. 2004). South- or SE-vergent early 1995; Malo & Kirkwood 1995). The structures
Acadian recumbent or overturned folds (Fig. 12e) formed during crustal shortening include upright
or steeply inclined folds and high-angle reverse (c. 425 Ma) from correlative Gander margin rocks
faults. The latter are generally west-vergent, along strike (d’Lemos et al. 1997; Scofield &
hinterland-directed structures that progressively d’Lemos 2000). The western tectonic boundary of
become younger towards the NW in northern New the Meelpaeg allochthon, the Cape Ray Fault-
England and adjacent New Brunswick (Tucker Victoria Lake shear zone, is generally a major
et al. 2001; Wilson et al. 2004), which is well con- ductile, moderately to shallowly east-dipping,
strained by migration of a retro-arc foredeep over thrust-sense shear zone. Locally, where the fault
time (Bradley et al. 2000; Bradley & Tucker zone had steepened significantly, it also has accom-
2002). Early Devonian reverse faults that overprint modated dextral oblique motion (Dubé et al. 1996;
Salinic structures have also been observed in Valverde-Vaquero & van Staal 2001; van der
central Newfoundland (Rogers et al. 2005; Velden et al. 2004). Mylonites yielded biotite and
Valverde-Vaquero & van Staal 2001; Zagorevski muscovite cooling ages between 400 and 390 Ma
et al. 2007a, b), but the link between migration of (Burgess et al. 1995; Dubé et al. 1996), indicating
the Acadian deformation front and an associated that the high-grade rocks were mostly exhumed by
retro-arc foredeep cannot be established or tested the end of the Early Devonian. These thrust-related
in Newfoundland. Devonian sedimentary rocks are structures were cut and stitched by Middle Devo-
rarely preserved and are exposed only in a few iso- nian (c. 386 Ma) granite (Dubé et al. 1996), support-
lated areas. Movements on the final Acadian defor- ing other evidence that Acadian orogenesis was
mation front in westernmost Newfoundland are mainly confined to the Early Devonian.
tightly constrained by deformed foreland basin Acadian deformation reoriented and modified
rocks of the Lochkovian (416 –411 Ma) Clam the earlier established Taconic and Salinic structural
Bank Formation and undeformed Emsian (407 – architecture significantly. Shallow east-dipping
397 Ma) conglomerate of the nearby Red Island thrusts cut and offset major west-dipping structures
Road Formation, which was deposited immediately at depth, such that it produced a wedge geometry,
in front of the Round Head thrust (Stockmal & which was imaged on seismic profiles (van der
Waldron 1993; Burden et al. 2002; Quinn et al. Velden et al. 2004). On the surface, Taconic and/
2004). Acadian deformation lasted longer in Quebec or Salinic folds were refolded (Fig. 11c), whereas
and continued into the Middle/Upper Devonian the predominantly shallowly west-dipping fault
(380–370 Ma) (Bradley et al. 2000). This late stage panels were folded into steep or westerly overturned
of Acadian deformation overlapped in age with structures and/or cut by new east-dipping low-angle
Neoacadian orogenesis in Meguma (see below). reverse faults (Zagorevski et al. 2007b). Regional
A large upper amphibolite-facies metamorphic shear associated with west-directed transport
nappe (Meelpaeg allochthon) is exposed in southern locally rotated and overturned older Salinic-
and south-central Newfoundland. Its extent and generated, west-dipping fault panels into markedly
geometry at depth was largely constructed on the east-dipping overturned panels. Overturning also
basis of seismic profiles (van der Velden et al. rotated the regional cleavage into a new orientation
2004). Rocks of the allochthon reached peak meta- such that it now consistently has a shallower dip
morphism at 418 –414 Ma (van Staal et al. 1994; than the fault panels (Dean & Strong 1978; Lafrance
Burgess et al. 1995; Valverde-Vaquero et al. & Williams 1992).
2000) and must have travelled a fair distance and The progressive hinterland migration of the
for a considerable time interval, because they are Acadian deformation front was interpreted by
generally emplaced above low-grade rocks and Murphy et al. (1999) as being due to establishment
locally were even transported above unmetamor- of a west-dipping ‘flat-slab’, analogous to the
phosed Lower Devonian sediments (van Staal Laramide in the western USA and the present-day
et al. 2005). Attainment of Acadian high-grade Andes in central Chile and Argentina (Kay &
metamorphism in these Gander margin rocks so Abruzzi 1996) (Fig. 10b). The cause of the inferred
shortly after the onset of the Acadian orogeny Acadian flat-slab setting is likely related in some
suggests that these rocks were already hot, either fashion to prolonged subduction of Avalonia
as a result of being situated in the Acadian back- beneath Laurentia, the reason of which is conten-
arc (see below) prior to collision and establishment tious. However, Avalonia’s A – subduction and
of the flat-slab or were already metamorphosed to resultant widening of the Acadian deformation zone
high-grade conditions during the immediately towards the hinterland during the Early Devonian
preceding Salinic orogeny and remained buried coincides with rapid southerly palaeolatitudinal
and relatively hot until their involvement in the motion of Laurentia (van Staal et al. 1998, their
Acadian. The latter is consistent with the presence fig. 8). Hence, construction of the wide Acadian
of a well-developed pre-418 Ma S1 metamorphic mountain belt may be to some extent analogous
layering or schistosity (van Staal et al. 1992, to formation of the Andes and North American
1994) and age constraints on Salinic metamorphism Cordillera as a result of the Mesozoic and younger
westward (trenchward) motion of the America’s tholeiitic and non-arc features (Fig. 15a, b). Felsic
following the diachronous opening of Atlantic samples (Fig. 15c) form two sample clusters, one
Ocean. within the VAG field and the other within the
Acadian convergence probably was strongly par- WPG. Most of the latter Y þ Nb-enriched samples
titioned, because the reverse hinterland-directed are from the c. 418 Ma Mt. Elizabeth complex,
movements were coeval with dextral transcurrent which Whalen (1990) documented as being com-
motion on steeply dipping faults (King & Barr prised of contemporaneous I-, A- and mafic-type
2004) near the suture with Avalonia (Holdsworth suites. The spectrum and proportions of mafic and
1994; Doig et al. 1990; Nance & Dallmeyer 1993; felsic compositions is remarkably similar to both
Park et al. 1994). Progressive steepening of the western and west-central Newfoundland Silurian
Acadian reverse faults over time probably facili- magmatism (Figs 13 & 14) and, for this reason,
tated transcurrent motion in the hinterland as well we interpret these 423–416 Ma magmatic rocks as
(e.g. van Staal & de Roo 1995). a final, post-kinematic pulse of magmatism related
to breakoff of the Salinic slab (cf. Whalen et al.
2006) Hence, if correct, they are not related to the
Late Silurian-Middle Devonian Acadian orogeny.
(Acadian) magmatism The Upper Silurian (423 –416 Ma) magmatic
rocks that occur east of the Fredericton trough in
Late Silurian (423 –416 Ma) magmatism is volumi- southern New Brunswick, adjacent Maine and east
nous both in Newfoundland, New Brunswick, Gaspé or south of the DBL in Newfoundland are spatially
and Maine (Figs 1 & 2) and represents a high-flux associated with the Acadian suture and are related to
event, although we will argue below that not the early stages of the Acadian collision between
all rocks of this age are related to the Acadian Avalonia and the leading edge of Laurentia (rep-
orogeny. Upper Silurian magmatic rocks occur in resented at this stage by the accreted Ganderia),
two geographically separated belts, which are both since they generally formed syn- to post-kinemati-
represented by volcanic and plutonic rocks. The cally with respect to the associated early Acadian
most westerly belt is spatially associated with the structures and metamorphism in this part of the
part of Ganderia where Salinic structures and meta- Canadian Appalachians (see above). The relative
morphism are well-developed, whereas the easterly abundance of isotopic and geochemical data in
belt is generally not. In Newfoundland, the westerly Newfoundland (see below) revealed that this
belt is represented by such igneous rocks as the magmatic suite can be subdivided into distal and
Late Silurian phases of the 424 + 2 Ma Mt. Peyton proximal subgroups based on their nearness to the
intrusive suite (Dickson et al. 2007), 422 + 2 Ma DBL. This subdivision probably correlates with
Fogo Island pluton or batholith (Aydin 1995), the the different crustal evolutions experienced by the
422 + 2 Ma Port Albert bimodal dikes (Elliot inboard (distal group) and outboard (proximal) seg-
et al. 1991), 423þ3/22 Ma Stony Lake volcanic ments of the Gander margin during the Ordovician
rocks (Dunning et al. 1990) and other related volca- and Silurian prior to Acadian orogenesis (see van
nic rocks (420 –417 Ma, V. McNicoll, pers. Staal 1994; Valverde-Vaquero et al. 2006). The plu-
comm. 2006). tonic rocks of the proximal subgroup in Newfound-
Silurian magmatism emplaced west of the DBL land, which were emplaced immediately east of
in Newfoundland, NW New Brunswick, Maine the DBL, are comprised of 27% mafic rocks; the
and adjacent Gaspé has been discussed previously rest are felsic. The mafic rocks exhibit mainly
and is illustrated in Figures 13, 14 and 15. The calc-alkaline arc and minor arc tholeiite compo-
equivalent Late Silurian magmatic event in central sitions (Fig. 17a, b), whereas the felsic plutonic
New Brunswick (Fig. 15a– c) is mainly exposed in rocks plot exclusively within the VAG field. The
the Miramichi highlands. This magmatic phase time equivalent mafic rocks of the distal subgroup
includes the 418 + 1 Ma Mt. Elizabeth and in part have similar compositions, but also include
417 + 1 Ma North Pole Stream plutonic suites a significant proportion of rocks that exhibit conti-
(Bevier & Whalen 1990) and volcanic rocks such nental tholeiite and E-MORB characteristics,
as those of the c. 423 + 3 Ma Benjamin and suggesting a significant contribution from the litho-
Bryan Point formations (Walker et al. 1993). spheric mantle (Fig. 17a, b). Felsic plutonic rocks
These rocks generally postdate structures formed from the distal subgroup range to more Rb- and
during the Salinic orogeny, but always predate Y þ Nb-enriched compositions, such that they
development of Acadian structures in the areas straddle the boundaries between the fields for
they occur, that is, they formed NW of the migrating VAG, WPG and the syn-COLG (Fig. 17c). The
Acadian deformation front (Bradley & Tucker latter characteristics suggest that these granites
2002). Mafic rocks, which comprise c. 21% of this received a greater contribution from evolved
magmatic suite, exhibit calc-alkalic to continental high-Rb crust than the granites emplaced in
proximity to the DBL in Newfoundland (Figs 13c, supported by the generally positive (juvenile)
14c & 17), which is consistent with less tectonic 1Nd(T) values of the proximal plutons and negative
reworking of the Gander margin crust towards the (old crust contaminated) eNd (T) signatures of the
east during Ordovician-Silurian tectonism. Proto- distal plutons (Kerr et al. 1995).
liths of the plutonic rocks of the proximal subgroup In southern New Brunswick, the Acadian-related
probably involved depleted-mantle-derived Ordovi- plutonic rocks comprise an Upper Silurian suite
cian arc and associated sedimentary rocks originally dominated by intermediate to granitic rocks with
formed near the active leading edge of Ganderia minor associated gabbro (McLaughlin et al. 2003).
(Penobscot & Popelogan– Victoria arcs and their They also include the c. 423 Ma volcanic rocks
basement) prior to opening and closing of the of the Mascarene back-arc (Van Wagoner et al.
Tetagouche–Exploits back-arc basin and accretion 2002). Mafic compositions included in this mag-
of Ganderia to Laurentia (van Staal et al. 1998; matic belt, like their equivalents in Newfoundland
Zagorevski et al. 2007a). This interpretation is east of the DBL (Fig. 17), exhibit a diverse spectrum
Fig. 16. Devonian magmatism in Gaspé (Quebec) and New Brunswick. Diagrams (a, b, c, d) and abbreviations as
in Figure 7; samples have been subdivided according to age, as indicated in symbol legends, see text for discussion.
See Appendix 1 for data sources.
Fig. 17. Silurian plutons emplaced east of the Dog Bay Line (DBL) into rocks of the Gander margin in eastern
Newfoundland. Subdivision into proximal and distal is based on proximity of the plutons with respect to the DBL
(see text). Diagrams (a, b, c) and abbreviations as in Figure 7. See Appendix 1 for data sources.
of arc- to non-arc like compositions (Fig. 15a, b), data cluster consists of samples from the c. 423 Ma
but represent a higher proportion (c. 44%) of ana- Utopia and c. 422 Ma Welsford phases of the
lyzed samples. Felsic compositions (Fig. 15c) lack S. George batholith.
the Rb-enrichment observed in Newfoundland Overall, the early Acadian-related igneous rocks
(Fig. 17c) and include two distinct sample clusters, compositionally resemble late Acadian, Early
one within the VAG field and the other in the Devonian (416 –395 Ma) plutons formed further to
WPG field, though both also fall mainly within the the NW in New Brunswick (Fig. 16) and in the equi-
overlapping syn-COLG field. The Y þ Nb-enriched valent area in Newfoundland (Figs 13, 14 & 15).
The regional distribution of Early Devonian felsic rocks (Fig. 16d). Contemporaneous plutonic
Acadian magmatism suggests the existence of a rocks in eastern Newfoundland (Fig. 18) also exhi-
time-transgressive belt-like pattern (see compi- bits a similar paucity (,3% or 26 of 1182
lations of Kerr 1997; Whalen et al. 1996a; Tucker samples) of mafic rocks. The continuity of crustal
et al. 2001; Hibbard et al. 2006), which seems to shortening and lack of evidence for regional exten-
track the progressive migration of the Acadian sion during this period, combined with the predomi-
deformation front, at least locally (Bradley & nance of felsic over mafic compositions and the
Tucker 2002). This relationship needs more diverse signatures of granites (VAG, WPG and
testing, but if correct, suggests that Early Devonian syn-COLG) included, leads us to relate the
Acadian magmatism was associated with protracted Middle–Late Devonian (395 –375 Ma) magmatism
crustal shortening and related in some way to the to breakoff of the Acadian slab, which seems to have
progressive shallowing of the subducting slab, taken place immediately after the 400–395 Ma
which presumably started to happen shortly after docking of Meguma (see below), suggesting a
the arrival of buoyant Avalonia at the trench at causative relationship, rather than delamination or
c. 421 Ma. Migration of the deformation and mag- subduction of a plume, as was proposed earlier by
matic front towards the west suggest that Avalonia Murphy et al. (1999). Furthermore, the parallelism
and its attached oceanic lithosphere of the Acadian of this magmatism with the Avalonia– Laurentia
seaway (Acadian slab) remained being thrust plate boundary from Newfoundland to Massachus-
beneath composite Laurentia until at least 400 Ma, setts (Hibbard et al. 2006) would require a fortuitous
if not longer. The Acadian magmatic pulse is thus linearity of the plume head approximately parallel
associated with intra-crustal shortening and hence, to the Laurentian margin in the model of Murphy
unlikely related in any fashion to regional extension et al. (1999). However, comparison of the
and/or delamination of the lithospheric mantle (cf. Middle–Late Devonian magmatism with the better
Dostal et al. 1989, 1993). This conclusion is consis- substantiated Silurian slab-breakoff-related magma-
tent with the relatively high ratio of intrusive with tism of western Newfoundland (Fig. 13) indicates
respect to extrusive rocks in the Acadian magmatic the Devonian magmatism was dominated by
suite. Magmas intruding in thickened crust during melting of more evolved (higher LILE) crustal
contraction tend to pond and evolve at depth, prob- materials and that mafic magmas, which likely sup-
ably because they have difficulty ascending to plied the thermal flux for the large-scale crustal
crustal levels above the brittle-ductile transition. melting, probably ponded at depth and were
Critical to understanding the duration and genesis unable to move to high crustal levels, because the
of Early Devonian Acadian magmatism is the orogen was still under compression.
boundary condition imposed by the time of separ- If the postulated timing of Acadian slab-breakoff
ation and sinking of the oceanic slab that was orig- is correct, it implies that the voluminous phase of
inally attached to Avalonia (¼ breakoff of Acadian 415–395 Ma syn-convergence Acadian magmatism
slab). If this process happened, it would signal the was related to shallow or flat-slab subduction and a
end of the Acadian-related convergence. Breakoff progressively widening zone of crustal shortening in
may have been facilitated by eclogititation, and the hinterland of the orogen. In New Brunswick,
hence steepening of the leading oceanic segment where both better age constraints and more REE
attached to the trailing overriding, shallow subduct- data are available, mafic rocks formed during this
ing buoyant segment (including Avalonia) of the period exhibit calc-alkaline arc, continental tholeiite
Acadian slab (Haschke et al. 2002). The area and non-arc signatures (Fig. 16a, b). Contempora-
affected by the early Acadian magmatic phase was neous felsic rocks mostly show VAG signatures
re-intruded during Middle–Upper Devonian mag- with many samples plotting where VAG and post-
matic flare-up represented mainly by plutons, COLG granites overlap (Fig. 16c, d). Relative to
which range in age between 395 and 375 Ma in New- immediately preceding and subsequent felsic mag-
foundland, locally stitching the suture zone between matism, the paucity of WPG compositions is
Laurentia’s leading edge (southeastern Ganderia) notable, indicating little input from the lithospheric
and Avalonia. This magmatism includes S-type mantle. Another unusual feature of this felsic mag-
and I-type granites with the latter ranging to matism is that a high proportion of samples
HFSE-enriched A-type (or WPG-type) compo- exhibit elevated (La/Yb)CN values (Fig. 19),
sitions (Fig. 18c). Consanguineous magmatism which are indicative of formation under deep
also took place near Laurentia’s leading edge (south- garnet stable P –T conditions, as was suggested for
eastern Ganderia) in New Brunswick and Maine Archaean TTG magmas (cf. Martin 1986). High
(Fig. 17d), although no evidence for stitching the (La/Yb)CN ‘adakitic’ magmas have been document
Acadian suture with Avalonia. by Kay et al. (2005) as a characteristic of central
Middle– Late Devonian magmatic rocks in Chilean-Argentinian Andean flat subduction and
southern New Brunswick consists exclusively of interpreted as resulting from high-pressure
Fig. 18. Devonian plutons, emplaced into Ganderian rocks east of the Red Indian Line (RIL) in eastern Newfoundland.
Diagrams (a, b, c) and abbreviations as in Figure 7. See Appendix 1 for data sources.
metamorphism and partial melting of subducted towards the back-arc region suggests that an
fore-arc material as it entered the asthenospheric Andean flat-slab-like setting is an attractive ana-
mantle wedge beneath the arc. logue for the Early Devonian Acadian magmatism.
We discussed earlier that the fore-arc to the The principal difference is that flat subduction in
Acadian coastal arc was largely tectonically the Andes solely involves oceanic lithosphere,
removed and postulated that it was subducted whereas the Acadian flat-slab also comprises pro-
beneath the leading edge of Laurentia (represented gressive underthrusting of Avalonia beneath
at this stage by Ganderia). This feature together Laurentia. Shallowing and dehydration of the sub-
with the apparent migration of the structural front ducted Acadian slab and progressive thinning and
Fig. 19. Felsic Devonian magmatism from Gaspé (Quebec) and New Brunswick, subdivided into 2 age groups,
with flat subduction of the Acadian slab being postulated for the 416–395 Ma period, plotted on a chondrite normalized
La/Yb vs Yb diagram with fields for Archaean TTG (tonalite-trondhjemite-granodiorite), post-Archaean granitoids and
MORB from Martin (1986).
cooling of the overlying asthenospheric wedge, without such asthenospheric pockets. In general,
would over time have inhibited typical arc magma- Silurian back-arc and Salinic slab breakoff-related
tism and forced progressive retreat of the potential magmatism, which lasted until the onset of the flat-
mantle melt zone towards the rear of the now slab phase (see above) may have played an impor-
extinct coastal arc. In New Brunswick, this may tant role in the subsequent Acadian orogenesis,
help explain why during this period there is an because it may also have been an important source
apparent trailing off of magmatism, prior to the for the crustal heat necessary for the low to
Middle to late Devonian ‘flare-up’, which we medium pressure-high temperature metamorphism
related to breakoff of the Acadian slab. However, that typifies the back-arc region from the onset of
areas with thin lithosphere formed during the Acadian deformation and preferentially localized
immediately preceding Silurian extension in the deformation by thermal and/or fluid-induced weak-
back-arc region of the coastal arc and the collapsed ening (Hyndman et al. 2005).
Salinic collision zone, beneath which the Salinic
slab had broken off (van Staal & de Roo 1995;
Whalen et al. 2006) may have trapped pockets or Neoacadian
refuges of asthenosphere above the subsequent
Acadian flat-slab stage (Fig. 10b). Dehydration of The Neoacadian orogeny refers to late Early
the Acadian flat-slab and fluid-fluxing of the asthe- Devonian –Early Carboniferous (400–350 Ma) de-
nosphere in such trapped pockets may have trig- formation, metamorphism and magmatism spatially
gered their partial melting and thus may have been and genetically linked to docking of the Meguma
responsible for producing the syn-Acadian upper terrane to Laurentia (Hicks et al. 1999; van Staal
plate, arc-like magmatism during this phase. We 2005, 2007). The leading edge of the latter was rep-
introduced these asthenospheric pockets to explain resented at this stage by Acadian-accreted Avalonia.
the mantle-derived component of the syn-Acadian The name Neoacadian was originally introduced
magmatism in a flat-slab tectonic framework, by Robinson et al. (1998), to cover strong Late
which allows progressive migration of the mag- Devonian –Early Carboniferous deformation and
matic front into the orogen’s hinterland. Shallow metamorphism in southern New England, which
subduction of the cold Acadian slab would have was distinct from Acadian orogenesis in this part
rapidly displaced the fertile mantle wedge and of the northern Appalachians. Neoacadian orogen-
cooled the overlying mantle, and hence would esis, as defined by van Staal (2005, 2007), overlaps
have inhibited formation of any mantle-derived in time, but not in space with the tail end of Acadian
arc to non-arc magmatism (van Hunen et al. 2002) deformation and magmatism in Laurentia’s
hinterland near the Appalachian front in New trench of the NW-dipping subduction zone
England (see above). Such overlap is absent in New- (Fig. 10c). Such a process explains Meguma’s
foundland, because Newfoundland seems to have accretion and transfer to the upper (Laurentian)
escaped most, if not all, effects of Meguma’s accre- plate and the continuation of NW-directed thrusting
tion, which took place further to the south (van Staal along the Cobequid– Chedabucto fault system into
2005, 2007). However, a possible dynamic linkage the Carboniferous (Waldron et al. 1989; Pe-Piper
between the Neoacadian and prolonged continu- & Piper 2002). A temporal link between the break-
ation of the Acadian into the Middle Devonian off of the Avalonian slab (see above) and Meguma’s
in New England cannot be dismissed at present. docking and large-scale crustal wedging also
The name Neoacadian is therefore appropriate, suggest a genetic relationship.
because it permits a possible tectonic link with the Meguma was subjected to extensive Middle–
Acadian. Such a linkage is indeed supported by Late Devonian magmatism, which was dominated
the similarities between the tectonic processes by peraluminous (S-type) felsic granitoid rocks
responsible for both the Acadian and Neoacadian and included only rare (c. 2%) mafic rocks
orogenies (Murphy et al. 1999; van Staal 2007), (Fig. 20). Mafic rocks exhibit calc-alkaline to conti-
discussed further below. nental tholeiite compositions whereas the volumi-
The provenance and tectonic setting of Meguma nous felsic plutonic rocks span the upper portion
during the Early Palaeozoic is contentious and of the VAG field and the lower to middle portions
in need of more research. On the basis of avail- of the syn-COLG field. Compared to almost all of
able data, we favour it as having been a single our other spatial-temporal groups, this felsic pluton-
Gondwanan-derived microcontinent, but alternative ism exhibits a total lack of high Y þ Nb (WPG type)
settings such as a promontory on Gondwana itself or compositions, suggesting no involvement of
forming part of Avalonia to start with (Murphy et al. enriched lithospheric mantle. This magmatism was
2004) cannot be overruled at present. Its Palaeozoic interpreted to have occurred after transfer of
stratigraphy is markedly different from that of Ava- Meguma to Laurentia and after breakoff of the
lonia (Fig. 3), although they also share some simi- Rheic slab that was present beneath Meguma
larities, such as a large Middle Ordovician hiatus (Fig. 10c) (van Staal 2007; Moran et al. 2007).
and a palaeontological linkage by at least the Late However, comparison of its geochemical signatures
Silurian (Bouyx et al. 1997). and the relative high proportion of felsic plutonic
The dominant structural style of Meguma is rocks with other magmatic events attributed to slab-
characterized by a regional set of NW-trending shal- breakoff (e.g. Figs 13 & 14) suggest a slightly differ-
lowly plunging, upright folds and associated shear ent petrogenesis. Rather the syn- to post-collisional
zones. The folds are progressively reoriented to a S-type granitoid magmatism may be a product of
more easterly trend on approaching the dextral partial melting of deeply buried fertile Meguma
Cobequid-Chedabucto fault (CCF in Fig. 1), which sediments following crustal thickening as a result
represent the boundary with Avalonia on land. The of Avalonia’s wedging, possibly facilitated by
combination of transcurrent shearing and upright thermal input from mantle-derived mafic magmas
folding is generally interpreted to represent a that ponded at depth. If the latter is correct, mafic
dextral transpressive deformation regime formed magma emplacement is probably related to astheno-
in response to dextral oblique convergence between spheric upwelling following Rheic slab breakoff
Meguma and Avalonia. How this convergence was (Fig. 10c) or alternatively, delamination of Megu-
accommodated at depth is not well constrained. mas’s lithospheric mantle.
Geophysical evidence suggests that a wedge of Ava-
lonian crust was thrust beneath Meguma, whereas
upper mantle reflectors suggest a NW-dipping sub- Summary of orogenic events
duction zone was present beneath Nova Scotia
(Keen et al. 1991). Murphy et al. (1999) postulated Four major phases of orogenesis related to accretion
that the dip of the NW-dipping subduction zone was of ribbon-shaped crustal blocks to the Laurentian
very shallow (flat-slab) due to interaction with a margin have been recognized in the Canadian
rising plume. Regardless of the cause of the postu- segment of the northern Appalachians: the
lated flat-slab, it explains the scarcity or absence Taconic, Salinic, Acadian and Neoacadian oroge-
of late Early Devonian or younger arc magmatism nies. The Taconic is a composite of three, kinemati-
in Laurentia’s leading edge, which at this stage is cally different, but temporally and spatially closely
represented by Acadian accreted Avalonia. The related events (Taconic 1, 2 & 3), all involving
lithospheric configuration of Meguma partly under- peri-Laurentian rocks. They have been grouped
lain by Avalonian crust suggests wedging of the together under the Taconic banner, because they
downgoing Rheic plate by the leading (Avalonian) may not be separable elsewhere in the Appalachians
edge of Laurentia when Meguma entered the where critical age relationships have been obscured
Fig. 20. Late Devonian (377–368 Ma) plutonism in Meguma, Nova Scotia. Diagrams (a, b, c) and abbreviations as in
Figure 7. See Appendix 1 for data sources.
by superimposed orogenesis (van Staal et al. 2007). structures and metamorphism generated during
Taconic 1 represents a short-lived, latest Cambrian this event is cryptic and largely based on cross-
obduction (500 –493 Ma) of an infant oceanic arc cutting relationships with younger, Taconic
onto the peri-Laurentian Dashwoods microconti- 2-related arc magmatic rocks. Taconic 2 was princi-
nent (Fig. 4a) (van Staal et al. 2007) and has only pally due to dextral oblique collision between Dash-
been recognized with some degree of certainty in woods and its Notre Dame arc suprastructure with
Newfoundland. It mainly produced a regionally the downgoing Laurentian margin (Figs 4b & 8).
extensive mélange in Newfoundland (Fig. 6a). Collision had started by at least 475 Ma (Waldron
Evidence for the nature and extent of the ductile & van Staal 2001), but locally possibly already at
480 Ma (Knight et al. 1991) and A –subduction had the SE with early thrust-related structures progress-
ended before 450 Ma in most parts of the Canadian ively being steepened, in part due to refolding into
Appalachians. Penetrative regional metamorphism upright structures, during terminal collision.
appears to have been short-lived (469– 457 Ma) The Acadian orogeny started shortly after or
and has a temporal and spatial link to a widespread during the waning stages of the Salinic during the
tonalite flare-up in upper plate rocks in Newfound- Late Silurian at c. 421 Ma due to subduction of Ava-
land (Fig. 5a). These voluminous arc-type tonalites lonia beneath Laurentia following closure of the
and associated non-arc-like mafic rocks (Fig. 7), narrow intervening Acadian seaway (Fig. 10b).
are thought to be related to slab-breakoff (van The Acadian lasted at least until c. 400 Ma in New-
Staal et al. 2007). Associated reverse shear zones foundland and into the Middle –Late Devonian
were predominantly synthetic with the east-directed (c. 380 Ma) in Quebec. The dominant orthogonal
subduction of the Laurentian margin, although tectonic transport was to the west, antithetic with
antithetic structures were also generated along respect to a westerly-dipping, progressively shal-
Dashwoods’ eastern margin (Fig. 7b). The latter lowing, flat-slab-like subduction zone. The sense
structures accommodated underthrusting of supra- of obliquity was dextral. Structures and magmatism
subduction zone oceanic elements of the AAT (Figs 16 & 18) progressively became younger to the
(Fig. 5a) (Lissenberg & van Staal 2006). west (Bradley et al. 2000). The Acadian structural
Structures and metamorphism generated during setting is analogous to the present-day Andes at
Taconic 3 are localized in volcanic and sedimentary the latitude of central Chile and Argentina where
rocks that occur in a narrow belt adjacent to the the Pacific plate subducts very shallowly (flat-slab)
RIL. They formed as a result of collision between to the east, producing high (La/Yb)N ‘garnet-
the peri-Gondwanan Popelogan– Victoria arc built signature’ magmatism (Fig. 20). Syn-collision Early
upon the leading edge of Ganderia and the Lauren- Devonian magmatism also shows similarities to that
tian Red Indian Lake arc, which at this time formed associated with the Andean flat-slab segment and
the leading edge of the Laurentian plate, between progressively diminished over time until the postu-
460 and 450 Ma. The Red Indian Lake arc had lated breakoff of the Acadian slab at 395–390 Ma.
earlier re-accreted to Laurentia due to closure of The late Early –Late Devonian/Early Carbon-
the Lloyds River back-arc basin and incorporated iferous (395 –350 Ma) Neoacadian orogeny, with
into the AAT during Taconic 2 (Zagorevski et al. its voluminous, spatially restricted syn-collisional,
2006). Arc-arc collision was shortlived and lasted S-type granitoid magmatism (Fig. 18), was due to
,10 Ma (van Staal et al. 1998; Zagorevski et al. final docking of Meguma to Laurentia and stepping
2007b, c). Tectonic transport was dominantly to back of the subduction zone into the Rheic Ocean.
the SE and principally affected the upper plate The plate tectonic setting was interpreted to be
rocks (AAT). Accretion of the Popelogan– Victoria similar to the Acadian, that is, a flat-slab (Murphy
arc thus terminated Taconic orogenesis. Docking of et al. 1999).
the Popelogan –Victoria arc also implies closure of
the main Iapetan oceanic tract, leaving only the
Tetagouche–Exploits back-arc basin basins and Conclusions
the oceanic seaway between Ganderia and Avalonia
open (van Staal 2005). With the exception of localized deformation associ-
The Salinic orogeny was due to closure of the ated with major ductile-brittle fault zones, the
Tetagouche–Exploits back-arc basin (van Staal Canadian Appalachians escaped most of the intense
1994). Closure was initiated at c. 450 Ma, immedi- orogenesis associated with Alleghanian collision
ately after the Late Ordovician Taconic 3 arc– arc seen further to the south in the Appalachians.
collision, which forced the west-facing subduction Hence, they provide a window to study the distri-
zone that dipped beneath Laurentia to step back bution and nature of orogenic events that took
behind the accreted Popelogan –Victoria arc into place prior to continent– continent collision and
the backarc basin, which was floored by oceanic assembly of Pangaea. Careful dating and delineating
lithosphere. Back-arc closure produced both arc- the regional extent of structures and their associated
like mafic and felsic plutonism (Fig. 13). The Gander metamorphism and magmatism in the Canadian and
margin was accreted to Laurentia during the late immediately adjacent New England Appalachians
Llandovery– Wenlock (433 –423 Ma) during sinis- revealed the following important conclusions: (1)
tral oblique convergence. The collision appears to collisional orogenesis was episodic and comprises
have taken place slightly earlier in Newfoundland four different phases. Each phase was generally
(Whalen et al. 2006) than New Brunswick (van short-lived and linked to oblique accretion of peri-
Staal et al. 2003). Silurian syn-collision magmatism Laurentian and peri-Gondwanan crustal ribbons to
(433 –423 Ma) (Figs 13 & 14) was due to slab the Laurentian margin, which was progressively
breakoff. Tectonic transport was predominantly to expanding eastwards over time; and (2) the locus
of initial collision progressively migrated to the harden the rocks and prohibit development of long-
east, which after the Middle Ordovician was consist- lived movement zones (Lin et al. 1998). Relative
ently the direction of the orogen’s foreland (Figs 1 & hardening of the deforming material is probably
2). The Late Ordovician and younger foreland- also the main reason why collisional belts seem to
directed migration of the locus of collision-related widen over time, because new material at the
deformation and orogenesis was concurrent with margins of the belts becomes easier to deform
progressively stepping back of a west-dipping sub- than their interiors (Means 1995). An example is
duction zone behind each of the accreted blocks. provided by the Taconic 2 collision, where upper
Deviations and complications from this overall plate arc volcanic rocks and sediments were buried
simple pattern were introduced by local reactivation and heated when they were incorporated into a pro-
of older structures, widening of the A –subduction gressively widening A –subduction zone during the
channel and hinterland migration of structures and final stages of collision (Fig. 4b). Here, the upper
associated magmatic rocks during the Acadian and plate rocks were thermally weakened by a region-
possibly also during the Neoacadian. The Acadian ally very voluminous, but short-lived (5– 7 Ma)
hinterland migration, which lasted .20 Ma, is a tonalite bloom (Fig. 5b) (van Staal et al. 2007).
major deviation to the overall observed pattern. It The significance of the oblique component
is interpreted to be due to establishment of a flat-slab accommodated by the shear zones is a contentious
(Fig. 10b) with strong coupling between the upper- issue in the Appalachian –Caledonian orogen.
and downgoing lower plates, analogous to the Some workers (van Staal et al. 1998; Brem et al.
flat-slab region in the southern Andes. The resultant 2007) have speculated that some of the block or
Acadian deformation produced reverse and trans- terrane bounding faults may have accommodated
current faults, but most characteristically is large orogen-parallel movements during conver-
the penetrative steeply inclined or upright folding gence. However, with few exceptions (e.g. Malo
present throughout the northern Appalachians. & Kirkwood 1995; Zagorevski et al. 2006; Brem
Deformation during each accretionary/colli- et al. 2007; Brem 2007), evidence for major hori-
sional event appears to start in the arc and/or arc- zontal translations is commonly ambiguous and
trench gap and subsequently widens over time. there are few examples of structural duplications
Structures in the arc-trench gap and collision zone due to strike-slip (e.g. Lin et al. 2007). Nevertheless,
commonly show a dominant transport direction. a consistent sense of horizontal shear in a set of
The latter can be deduced locally from the sense coeval shear zones is generally thought to reflect
of overturning of folds, shear sense indicators in the overall obliquity of convergence between the
shear zones, and the geometry and stratigraphy of opposing blocks. If correct, the horizontal com-
fault-bounded panels (e.g. ‘old over young’). The ponent of convergence between Laurentia and the
reverse fault zones generally, but not invariably, various blocks has changed polarity from dextral
were synthetic with the polarity of subduction to sinistral several times between the Ordovician
deduced from the position of other tectonic elements and Late Devonian. For example, most Early Devo-
(e.g. position of arc, back-arc and foreland basin) nian orogen-parallel faults accommodated some
during the early stages of collision. Even during component of dextral movements, which likely rep-
the Taconic 3 arc-arc collision, which resulted resents an orogen wide partitioning of the oblique
from closure of the Iapetus main oceanic tract, and component between Laurtentia and Avalonia. Com-
involved two subduction zones with opposite plications arise where accreting crustal ribbons got
polarity (Fig. 5b) (van Staal et al. 1998), one sense caught between two different subduction zones,
of overthrusting dominates, which was synthetic which were active at the same time. For example,
with the sense of subduction of the downgoing arc the Dashwoods microcontinent was bounded by
(Zagorevski et al. 2007c). To a first-order approxi- two coeval subduction zones with opposite polarity
mation, our observations are consistent with the and obliquity of convergence (van Staal et al. 2007).
generally held belief that the lithospheric shear Careful dating of the bounding shear zones revealed
couple generated during the obliquely sliding and a dextral sense of shear along its contact with the
sinking of one plate beneath another during A – Laurentian margin (Brem et al. 2007), while a sinis-
subduction controls the overall deformation tral-oblique reverse sense of shear operated coevally
pattern during at least the early stages of collision, along its contact with the AAT on its eastern margin
which is a time when there is probably relatively (Lissenberg et al. 2005a; Lissenberg & van Staal
strong coupling between the downgoing and over- 2006; Zagorevski et al. 2006, 2007b). Dashwoods
riding plate. A non-coaxial deformation (simple thus was moving towards the south with respect to
shear dominated), in general, is a more efficient both. In general, careful structural analysis com-
deformation mechanism than a bulk coaxial defor- bined with age-dating indicate that strike-parallel
mation (pure shear), because in the latter case pen- translations were important in the straight, steeply
etrative fabric development will progressively dipping fault zones, such as the Long Range –Cabot
fault system in Newfoundland (Brem 2007; Brem and spatial criteria. The temporal subdivisions are Early
et al. 2007), However, most major fault zones are Ordovician (489–475 Ma) and Middle Ordovician
markedly curviplanar. For example, the Lloyds (470– 455 Ma). Silurian (,445 and .416 Ma) and Devo-
River– Lobster Cove fault system that separates nian (,416 and .360 Ma). In Newfoundland, the spatial
Dashwoods from the AAT (Lissenberg & van subdivisions (see Figs 1 & 2) are: (a) western Newfound-
Staal 2006; Zagorevski et al. 2007b) locally bends land [area west of the Red Indian Line (RIL)]; (b) west-
through angles of .908, which indicates it is unli- central Newfoundland [area east of the RIL but west of
kely that it could have accommodated major strike- the Dog Bay Line (DBL)]; (c) eastern Newfoundland
slip translations along its whole length. Indeed shear east of the DBL. Newfoundland data from west of the
sense indicators and old over young relationship RIL includes both coeval plutonic and volcanic rocks
suggest it was mainly a major SE-directed reverse and was compiled from Whalen et al. (2006), van Staal
fault instead. Nevertheless, some steepened and et al. (2007) and Whalen (unpublished data). Data from
relatively straight segments of this fault were sub- east of the RIL is exclusively from plutonic rocks and
sequently reused as strike-parallel movement was compiled from Kerr et al. (1994) and Kerr (1997).
zones (e.g. Lafrance 1989). Such reactivation of The very large database of Kerr et al. (1994) was filtered
fault zones complicates interpretation of their on the presence of the appropriate trace elements for
overall kinematic movement picture, because the chosen plots (see below) and on whether trace element
younger deformation tends to destroy most, if not concentrations for these elements were near detection
all evidence for the earlier movements. limits (i.e. an attempt was made to exclude unreliable or
Slab-breakoff-type magmatism was repetitive imprecise data).
(Whalen et al. 2006; van Staal et al. 2007) and In the southern Canadian Appalachians, the spatial
appears to have accompanied every accretion subdivisions are: (a) Quebec–NW New Brunswick, which
event involving a microcontinent, with the possible consists of the area east of the Baie Verte–Brompton Line
exception of Meguma. Rather than simply viewing in Gaspé to immediately west of the Fredericton trough in
them as arc batholiths formed as a result of subduc- New Brunswick; (b) southern New Brunswick (Whalen et
tion, such replication of slab-breakoff magmatism al. 1996b), which consist of the area east of the Fredericton
represents a new way of looking at voluminous, trough and west of the Caledonia fault; and (c) Nova
short-lived magmatic flare-ups of syn-orogenic Scotia’s Meguma zone, which consist of the area east of
granitoid rocks within mountain belts. The latter the Cobequid –Chedabucto fault. The Gaspé and New
model can account for the geochemical character- Brunswick dataset, which includes both Silurian to Devo-
istics of the observed syn-orogenic Appalachian nian plutonic and volcanic rock analyses, was compiled
plutonism, whereas the arc model cannot. from Bédard (1986), Dostal et al. (1989, 1993), McLeod
(1990), Whalen (1993a, b), van Wagoner et al. (2001,
This manuscript is a product of the 2005–2010 TGI-3 2002), McLaughlin et al. (2003), Yang et al. (2003) and
Appalachian program and we appreciate the continuous Wilson et al. (2005). The Meguma zone dataset consists
financial and logistical support given to us by its exclusively of plutonic rocks and is mainly from the
manager, S. Hanmer. In addition, we would like to large (1153 samples) compilation by Tate & Merrett
thank our Canadian and New England Appalachian
(1994), which was filtered for samples that included Rb,
colleagues for discussion and supplying us with many
ideas over the years through fieldtrips and intellectual Y and Nb data, giving 948 samples. As this dataset only
osmosis. Of these, we particularly like to mention included 2 samples that had ,60% silica, an additional
S. Barr, S. Castonguay, S. Colman-Sadd, J. Dewey, 14 mafic samples were compiled from Currie et al.
L. Fyffe, J. Hibbard, P. Karabinos, A. Kerr, D. Lavoie, (1998) and Tate & Clarke (1995).
S. Lin, B. Murphy, B. O’Brien, D. Rankin, D. Reusch, It has become well established that mafic (silica
P. Robinson, D. Stewart, A. Tremblay, D. West, ,60 wt.%) rock trace element signatures reflect their tec-
C. White, R. Wilson, H. Williams and R. Wintsch. tonic settings of formation. For this reason, a large array of
S. Barr, B. Clarke, M. McLeod, D. Lenz and R. Wilson mafic rock tectonomagmatic discrimination diagrams
supplied geochemical datasets in digital format. Brendan
employing a wide range of trace elements have been devel-
Murphy invited us to submit this manuscript and showed
immense patience waiting for the final version. Construc- oped for this purpose over the last 30 years. However, a
tive and thoughtful reviews by S. Barr and P. Ryan very restricted number of trace elements (e.g. only La
improved the manuscript. Geological Survey of Canada and Ce of the REE, and no Ta or Hf) were available in
contribution 20080031. the Kerr et al. (1994) Newfoundland compilation, in
contrast to the more extensive trace element array avail-
able for most New Brunswick samples. This restricted
Appendix 1 which tectonomagmatic classification diagrams could be
employed, because we wished to utilize the same plots
for both ‘transect’ areas. Based in part on past experience
In this study, geochemical data compiled from various (e.g. Whalen et al. 2006) and also for the sake of brevity,
sources has been split or grouped based on both temporal only two mafic rock plots are employed herein: (a) Wood
(1980) Th–Zr –Nb diagram; and (b) Cabanis & Lecolle Scotia. Canadian Journal of Earth Sciences, 28,
(1989) La –Y– Nb plot. Previous studies (e.g. Whalen 1769– 1779.
1993a; Whalen et al. 1998, 2001) have demonstrated that B ARR , S. M. & W HITE , C. E. 1996. Contrasts in Late
mafic end-members of plutonic suites provide both valu- Precambrian-Early Paleozoic tectonothermal history
able insights into a suite’s tectonic setting and an ‘inde- between Avalon Composite terrane sensu stricto and
other possible Peri-Gondwanan terranes in southern
pendent’ cross-check on setting indicated by felsic New Brunswick and Cape Breton. In: N ANCE , R. D.
end-member trace element signatures. As well, knowledge & T HOMPSON , M. D. (eds) Avalonian and Related
of the proportions of mafic to felsic compositions that com- Peri-Gondwanan Terranes of the Circum-North
prise a magmatic association represents key information Atlantic. Geological Society of America, Special
for constraining tectonic setting and helps establish the Paper, 304, 95–108.
role played by directly mantle-derived magmas. For exam- B ARR , S. M., W HITE , C. E. & M ILLER , B. V. 2002.
ple, magmatic suites well-established to have formed in The Kingston Terrane, southern New Brunswick,
magmatic arc settings, almost invariably include a signifi- Canada: evidence for an Early Silurian volcanic arc.
cant proportion of mafic (basaltic plus andesitic) rocks Geological Society of America Bulletin, 114, 964– 982.
B ARR , S. M., W HITE , C. E. & M ILLER , B. V. 2003. Age
(Whalen et al. 1997b). For this reason, both numbers of
and geochemistry of Late Neoproterozoic and Early
samples involved and the percentage they represent of the Cambrian igneous rocks in southern New Brunswick:
total are shown on geochemical diagrams. similarities and contrasts. Atlantic Geology, 39,
Felsic (silica .60 wt.%) rocks were plotted on both the 55–73.
Pearce (1996) Rb– Nb þ Y diagram and the Pearce et al. B ARR , S. M., R AESIDE , R. P. & W HITE , C. E. 1998.
(1984) Nb– Y diagram. As the latter was found to not con- Geological correlations between Cape Breton Island
tribute much additional information, it is not included and Newfoundland, northern Appalachian orogen.
herein. Previous experience (e.g. Whalen et al. 2006) has Canadian Journal of Earth Science, 35, 1252–1270.
shown that these granitoid tectonomagmatic discrimi- B ÉDARD , J. H. 1986. Pre-Acadian magmatic suites of the
southeastern Gaspe Peninsula. Geological Society of
nation diagrams can be very useful for subdividing grani-
America Bulletin, 97, 1177–1191.
toid rocks into groups derived from different sources B EVIER , M. L. & W HALEN , J. B. 1990. Tectonic signifi-
and/or under contrasting tectonic conditions but that cance of Silurian magmatism in the Canadian Appala-
these geochemical signatures need to be integrated with chians. Geology, 18, 411–414.
geological and other parameters to arrive at an optimum B IRD , J. M. & D EWEY , J. F. 1970. Lithosphere plate-
interpretation of granitoid tectonic conditions of for- continental margin tectonics and the evolution of the
mation. Our interpretations of what these diagrams indi- Appalachian Orogen. Geological Society of America
cate about contributing granitoid source materials in this Bulletin, 81, 1031–1060.
study follow fairly closely those made by Pearce (1996) B LACKWOOD , R. F. 1985. Geology of the Facheux Bay
area (11P/9), Newfoundland. Mineral Development
in a review of sources and settings of granitic rocks.
Division, Department of Mines and Energy, Govern-
Some relevant points concerning the Rb–Y þ Nb ment of Newfoundland and Labrador, Report 85– 4.
diagram are: (a) VAG reflect derivation from depleted B OSTOCK , H. H. 1988. Geology and petrochemistry of the
mantle sources modified by a subducted component, that Ordovician volcano-plutonic Robert’s Arm Group,
is, low Rb and Y þ Nb, but concentrations of both increase Notre Dame Bay, Newfoundland. Geological Survey
with arc maturity within the VAG field; (b) intraplate gran- of Canada, Bulletin, 369, 1– 82.
ites (WPG) show evidence of input from enriched mantle B OUCOT , A. J. 1962. Appalachian Siluro-Devonian. In:
(lithosphere and/or athenosphere) sources (i.e. high C OE , K. (ed.) Some Aspects of the Variscan Fold
Y þ Nb); (c) syn-collision granites (syn-COLG) either rep- Belt. Manchester University Press, UK, 155 –177.
B OUYX , E., B LAISE , J. ET AL . 1997. Biostratigraphie et
resent pure crustal melts or strongly crustally contaminated
paléobiogeographie du Siluro-Dévonien de la Zone
mantle melts and thus exhibit very high Rb; and (d) post- de Meguma (Nouvelle-Écosse, Canadian). Canadian
collisional granites reflect derivation from diverse tectoni- Journal of Earth Science, 34, 1295–1309.
cally juxtaposed crustal and mantle sources, including B RADLEY , D. C. 1983. Tectonics of the Acadian Orogeny
enriched-mantle materials. in New England and adjacent Canada. Journal of
Geology, 91, 381 –400.
B RADLEY , D. C. & T UCKER , R. D. 2002. Emsian synoro-
References genic paleogeography of the Maine Appalachians.
Journal of Geology, 110, 483–492.
A YDIN , N. S. 1995. Petrology of a composite mafic-felsic B RADLEY , D. C., T UCKER , R. D., L UX , D., H ARRIS , A. G.
rocks of the Fogo Island batholith: a window to mafic & M C G REGOR , D. C. 2000. Migration of the Acadian
magma chamber processes and the role of the mantle orogen and foreland basin across the northern Appala-
in the petrogenesis of granitoid rocks. Unpublished chians of Maine and adjacent areas. U.S. Geological
PhD thesis, Memorial university of Newfoundland, Survey, Professional Paper, 1624, 49.
St. John’s, Newfoundland. B REM , A. G. 2007. The late Neoproterozoic to Early
B ARR , S. M. & J AMIESON , R. A. 1991. Tectonic setting Paleozoic evolution of the Long Range Mountains,
and regional correlation of Ordovician-Silurian rocks southwestern Newfoundland. Unpublished PhD
of the Aspy terrane, Cape Breton Island, Nova thesis, University of Waterloo, Canada.
B REM , A. G., L IN , S., VAN S TAAL , C. R., D AVIS , D. D. & of Canada, Geology of Canada, 6, 631–680 (also
M C N ICOLL , V. C. 2007. The Middle Ordovician to Geological Society of America, the Geology of North
Early Silurian voyage of the Dashwoods micro- America, F-1).
continent, west Newfoundland; based on new U/Pb C URRIE , K. L. 2003. Emplacement of the Fogo Island
and 40Ar/39Ar geochronological, and kinematic batholith. Atlantic Geology, 39, 79–96.
constraints. American Journal of Science, 307, C URRIE , K. L., W HALEN , J. B., D AVIS , W. J., L ONG-
311–338. STAFFE , F. J. & C OUSENS , B. L. 1998. Geochemical
B URDEN , E. T., Q UINN , L., N OWLAN , G. S. & evolution of peraluminous plutons in southern Nova
B AILEY -N ILL , L. A. 2002. Palynology and micropa- Scotia, Canada – a pegmatite-poor suite. Lithos, 44,
leontology of the Clam Bank Formation (Lower Devo- 117– 140.
nian) of western Newfoundland. Palynology, 26, D AVID , J. & G ARIÉPY , C. 1990. Early Silurian orogenic
185–215. andesites from the central Quebec Appalachians.
B URGESS , J. L., B ROWN , M., D ALLMEYER , R. D. & VAN Canadian Journal Earth Sciences, 27, 632 –643.
S TAAL , C. R. 1995. Microstructure, metamorphism, D EAN , P. L. & S TRONG , D. F. 1978. Folded thrust faults
thermochronology and P-T-t deformation history of in Notre Dame Bay, central Newfoundland. American
the Port aux Basques gneisses, south-west Newfound- Journal of Science, 277, 97– 108.
land, Canadian Journal of Metamorphic Geology, 13, D EWEY , J. F. 2002. Transtension in arcs and orogens.
751–776. International Geology Reviews, 44, 402– 439.
C ABANIS , B. & L ECOLLE , M. 1989. Le diagramme La/ D EWEY , J. F. & S HACKLETON , R. M. 1984. A model
10– Y/15–Nb/8: un outil pour la discrimination for the evolution of the Grampian tract in the early
des series volcaniques et la mise en evidence des pro- Caledonides and Appalachians. Nature, 312(5990),
cessus de mélange et/ou de contamination crustal. 115– 121.
Comptus Rendus de L’Acadamie des Sciences, Serie D ICKINSON , W. R. & S EELY , D. R. 1979. Structure and
2, 309, 2023–2029. stratigraphy of fore-arc regions. American Association
C ASTONGUAY , S., R UFFET , G. & T REMBLAY , A. 2007. of Petroleum Geologists Bulletin, 63, 2– 31.
Dating polyphase deformation across low-grade D ICKSON , L., M C N ICOLL , V. J., N OWLAN , G. S. &
metamorphic belts: an example based on 40Ar/39Ar D UNNING , G. R. 2007. The Indian Islands Group and
muscovite age constraints from the southern Quebec its relationships to adjacent units: recent data. In:
Appalachians, Canada. Geological Society of Current Research 2007. Newfoundland and Labrador
America Bulletin, 119, 978 –992. Department of Natural Resources, Geological Survey
C ASTONGUAY , S., R UFFET , G., T REMBLAY , A. & Report, 07– 1, 1– 9.
F ERAUD , G. 2001. Tectonometamorphic evolution of D IMITROV , I., M C C UTCHEON , S. R. & W ILLIAMS , P. F.
the southern Quebec Appalachians: 40Ar/39Ar evi- 2004. Stratigraphic and structural observations in
dence for Middle Ordovician crustal thickening and Silurian rocks between Pointe Rochette and the South-
Silurian-Early Devonian exhumation of the internal east Upsalquitch River, northern New Brunswick: A
Humber zone. Geological Society of America Bulletin, progress report. In: M ARTIN , G. L. (ed.) Geological
113, 144– 160. Investigations in New Brunswick for 2003. New Bruns-
C ASTONGUAY , S., W ATTERS , S. & R AVENELLE , J. F. wick Department of Natural Resources; Minerals,
2003. Preliminary report on the structural geology of Policy and Planning Division: Mineral Resource
the Clarence Stream-Moores Mills area, southwestern Report, 2004– 4, 41– 74.
New Brunswick: implications for gold exploration. D’L EMOS , R. S., S CHOFIELD , D. I., H OLDSWORTH , R. E.
Geological Survey of Canada, Current Research, & K ING , T. R. 1997. Deep crustal and local rheological
2003-D2, 1– 10. controls on the siting and reactivation of fault and
C AWOOD , P. A., D UNNING , G. A., L UX , D. & VAN G OOL , shear zones, northeastern Newfoundland. Journal of
J. A. M. 1994. Timing of peak metamorphism and the Geological Society, London, 154, 117– 121.
deformation along the Appalachian margin of Lau- D OIG , R., N ANCE , R. D., M URPHY , J. B. & C ASSEDAY ,
rentia in Newfoundland: Silurian, not Ordovician. R. P. 1990. Evidence for Silurian sinistral accretion
Geology, 22, 399– 402. of Avalon composite terrane in Canada. Journal of
C HEW , D. M., D ALY , J. S., P AGE , L. M. & K ENNEDY , the Geological Society, London, 147, 927– 930.
M. J. 2003. Grampian orogenesis and the development D OSTAL , J., L AURENT , R. & K EPPIE , J. D. 1993. Late
of blueschist-facies metamorphism in western Ireland. Silurian– Early Devonian rifting during dextral trans-
Journal of the Geological Society, London, 160, pression in the southern Gaspé Peninsula (Quebec):
911–924. petrogenesis of volcanic rocks. Canadian Journal of
C OLMAN -S ADD , S. D. 1980. Geology of south-central Earth Sciences, 30, 2283– 2294.
Newfoundland and evolution of the eastern margin of D OSTAL , J., W ILSON , R. A. & K EPPIE , J. D. 1989. Geo-
Iapetus. American Journal of Science, 280, 991–1017. chemistry of Siluro-Devonian Tobique volcanic belt
C OLMAN -S ADD , S. P., D UNNING , G. R. & D EC , T. 1992. in northern and central New Brunswick (Canada):
Dunnage-Gander relationships and Ordovician tectonic implications. Canadian Journal of Earth
orogeny in central Newfoundland: a sediment prove- Sciences, 26, 1282–1296.
nance and U/Pb study. American Journal of Science, D UBÉ , B., D UNNING , G. R., L AUZIERE , K. & R ODDICK ,
292, 317– 355. J. C. 1996. New insights into the Appalachian Orogen
C URRIE , K. L. 1995. Plutonic rocks; Chapter 8. In: from geology and geochronology along the Cape
W ILLIAMS , H. (ed.) The Appalachian/Caledonian Ray fault zone, southwest Newfoundland. Geological
Orogen: Canada and Greenland. Geological Survey Society of America Bulletin, 108, 101–116.
D UNNING , G., O’B RIEN , S. J., C OLMAN -S ADD , S. P., H ASCHKE , M. R., S CHEUBER , E., G UNTHER , A. &
B LACKWOOD , R., D ICKSON , W. L., O’N EILL , P. P. R EUTER , K. J. 2002. Evolutionary cycles during the
& K ROGH , T. E. 1990. Silurian orogeny in the New- Andean orogeny: repeated slab breakoff and flat sub-
foundland Appalachians. Journal of Geology, 98, duction. Terra Nova, 14, 49– 55.
895– 913. H EPBURN , J. C. 2008. Mid-Paleozoic arc accretion on the
E LLIOTT , C. G., B ARNES , C. R. & W ILLIAMS , P. F. 1989. eastern side of the Appalachian orogen, eastern Massa-
Southwest New World Island stratigraphy: new fossil chusetts and adjacent areas. Abstracts with Programs.
data, new implications for the history of the Central Geological Society of America, 43th Annual
Mobile Belt, Newfoundland. Canadian Journal of Meeting, NE-section, 75.
Earth Sciences, 26, 2062–2074. H IBBARD , J. P. 1994. Kinematics of Acadian deformation
E LLIOTT , C. G., D UNNING , G. R. & W ILLIAMS , P. F. in the Northern and Newfoundland Appalachians.
1991. New constraints on the timing of deformation Journal of Geology, 102, 215–228.
in eastern Notre Dame Bay, Newfoundland, from H IBBARD , J. P. 2000. Docking Carolina: Mid Paleozoic
U/Pb zircon ages of felsic intrusions. Geological accretion in the southern Appalachians. Geology, 28,
Society of America Bulletin, 103, 125–135. 127–130.
F OX , D. & VAN B ERKEL , J. T. 1988. Mafic-ultramafic H IBBARD , J. P., VAN S TAAL , C. R. & R ANKIN , D. W.
occurrences in metasedimentary rocks of southwestern 2007. A comparative analysis of pre-Silurian building
Newfoundland. In: Current Research, Part B. Geologi- blocks of the Northern and Southern Appalachians.
cal Survey of Canada, Paper 88-1B, 41–48. American Journal of Science, 307, 23–45.
F ORTEY , R. A. & C OCKS , L. R. 2003. Palaeontological H IBBARD , J. P., VAN S TAAL , C. R., R ANKIN , D. &
evidence bearing on global Ordovician-Silurian conti- W ILLIAMS , H. 2006. Geology, Lithotectonic Map of
nental reconstructions. Earth Science Reviews, 61, the Appalachian Orogen, Canada– United States of
245– 307. America. Geological Survey of Canada, Map 2096A,
F RIEDRICH , A. M., H ODGES , K. V., B OWRING , S. A. & scale 1:1 500 000.
M ARTIN , M. W. 1999. Geochronological constraints H ICKS , R. J., J AMIESON , R. A. & R EYNOLDS , P. 1999.
on the magmatic, metamorphic and thermal evolution Detrital and metamorphic 40Ar/39Ar ages from musco-
of the Connemara Caledonides, western Ireland. vite and whole-rock samples, Meguma Supergroup,
Journal of the Geological Society, London, 156, southern Nova Scotia. Canadian Journal of Earth
1217–1230. Sciences, 36, 23–32.
F YFFE , L. R. & B EVIER , M. L. 1992. A U –Pb date on the H OLDSWORTH , R. E. 1994. Structural evolution of the
Mohannes pluton of southwestern New Brunswick. Gander-Avalon terrane boundary: a reactivated trans-
Atlantic Geology, 28, 198. pression zone in the NE Newfoundland Appalachians.
F YFFE , L. R., M CLEOD , M. J. & R UITENBERG , A. A. Journal of the Geological Society, London, 151,
1991. A geotraverse across the St. Croix-Avalon 629–646.
terrane boundary, southern New Brunswick; In: H UOT , F., H ÉBERT , R. & T URCOTTE , B. 2002. A multi-
L UDMAN , A. (ed.) Geology of the Coastal Lithotec- stage magmatic history for the genesis of the Orford
tonic Block and Neighbouring Terranes, Eastern ophiolite (Quebec, Canada): a study of the Mont
Maine and Southern New Brunswick. New England Chagnon Massif. Canadian Journal of Earth Sciences,
Intercollegiate Geological Conference, NEIGC 1991, 39, 1201–1217.
Field Trip A-2, 13– 54. H UTCHINSON , D. R., K LITGORD , K. D., L EE , M. W. &
F YFFE , L. R., P ICKERILL , R. K. & S TRINGER , P. 1999. T REHU , A. M. 1988. U.S. Geological Survey deep
Stratigraphy, sedimentology and structure of the seismic reflection profile across the Gulf of Maine.
Oak Bay and Waweig formations, Mascarene basins: Geological Society of America Bulletin, 100, 172– 184.
implications for the Paleotectonic evolution of H YNDMAN , R. D., C URRIE , C. A. & M AZZOTTI , S. P.
southwestern New Brunswick. Atlantic Geology, 35, 2005. Subduction zone backarcs, mobile belts, and
59–84. orogenic heat. GSA Today, 15, 4– 10.
G ERBI , C., J OHNSON , S. E. & A LEINIKOFF , J. N. 2006a. J AMIESON , R. A. 1988. Metamorphic P-T-t data from
Origin and orogenic role of the chain Lakes massif, western Newfoundland and Cape Breton-implications
Maine and Quebec. Canadian Journal of Earth for Taconian and Acadian tectonics. Geological
Sciences, 43, 339–366. Society of Canada, Program with Abstracts, 13, A60.
G ERBI , C., J OHNSON , S. E., A LEINIKOFF , J. N., B EDARD , J OHNSON , R. J. E. & VAN DER V OO , R. 1986. Paleomag-
J. H., D UNNING , G. & F ANNING , C. M. 2006b. Early netism of the late Precambrian Fourchu Group, Cape
Paleozoic development of the Maine-Quebec Bound- Breton Island, Nova Scotia. Canadian Journal of
ary Mountains region. Canadian Journal of Earth Earth Sciences, 23, 1673– 1685.
Sciences, 43, 367–389. K ARABINOS , P., S AMSON , S. D., H EPBURN , J. C. &
H ALL , L. & VAN S TAAL , C. R. 1999. Geology of Southern S TOLL , H. M. 1998. Taconian orogeny in the New
End of Long Range Mountains (Dashwoods Subzone), England Appalachians: collision between Laurentia
Newfoundland. Geological Survey of Canada, Open and the Shelburne Falls arc. Geology, 26, 215– 218.
File 3727, scale 1:50 000. K AY , S. M. & A BRUZZI , J. M. 1996. Magmatic evidence
H AMILTON , M. A. & M URPHY , J. B. 2004. Tectonic for Neogene lithospheric evolution of the central
significance of a Llanvirn age for the Dunn Point vol- Andean ‘flat-slab’ between 308S and 328S. Tectono-
canic rocks, Avalon terrane, Nova Scotia, Canada: physics, 259, 15–28.
implications for the evolution of the Iapetus and K AY , S. M., G ODOY , E. & K URTZ , A. 2005. Episodic arc
Rheic Oceans. Tectonophysics, 379, 199– 209. migration, crustal thickening, subduction erosion, and
magmatism in the south-central Andes. Geological Breton Island, Canadian Appalachians. American
Society of America Bulletin, 117, 67–88. Journal of Science, 307, 371–398.
K EEN , C. E., K AY , W. A., K EPPIE , D., M ARILLIER , F., L IN , S., J IANG , D. & W ILLIAMS , P. F. 1998. Transpres-
P E -P IPER , G. & W ALDRON , J. W. F. 1991. Deep sion (or transtension) zones of triclinic symmetry:
seismic reflection data from the Bay of Fundy and natural example and theoretical modeling, In: H OLDS-
Gulf of Maine: tectonic implications for the northern WORTH , R. E., S TACHAN , R. A. & D EWEY , J. F. (eds)
Appalachians. Canadian Journal of Earth Sciences, Continental transpressional and transtensional tec-
28, 1096– 1111. tonics. Geological Society, London, Special Publi-
K EPPIE , J. D. & K ROGH , T. E. 2000. 440Ma igneous cations, 135, 41– 57.
activity in the Meguma terrane, Nova Scotia, L ISS , M. J., VAN DER P LUIJM , B. A. & VAN DER V OO , R.
Canada: part of the Appalachian overstep sequence? 1994. Avalonian proximity of the Ordovician Mirami-
American Journal of Science, 300, 528–538. chi terrane, northern New Brunswick, northern Appa-
K ERR , A. 1997. Space-time composition relationships lachians: Paleomagnetic evidence for rifting and
among Appalachian-cycle plutonic suites in New- back-arc basin formation at the southern margin of
foundland. In: S INHA , K., W HALEN , J. B. & Iapetus. Tectonophysics, 227, 17– 30.
H OGAN , J. P. (eds) The Nature of Magmatism in the L ISSENBERG , C. J. & VAN S TAAL , C. R. 2006. Feedback
Appalachian Orogen. Geological Society of America, between deformation and magmatism in the Lloyd’s
Memoir, 191, 193– 220. River Fault zone, Central Newfoundland: An example
K ERR , A., H AYES , J. P., C OLMAN -S ADD , S. P., of episodic fault reactivation in an accretionary setting.
D ICKSON , W. L. & B UTLER , A. J. 1994. An integrated Tectonics, 25, TC4004.
lithogeochemical database for the granitoid plutonic L ISSENBERG , C. J., Z AGOREVSKI , A., M C N ICOLL , V. J.,
series of Newfoundland. Newfoundland Department VAN S TAAL , C. R. & W HALEN , J. B. 2005a. Assembly
of Mines and Energy, Geological Survey Branch, of the Annieopsquotch Accretionary Tract, southwest
Open File Report NFLD/2377, 1 –44. Newfoundland: Age- and geodynamic constraints
K ERR , A., J ENNER , G. J. & F RYER , B. J. 1995. Sm–Nd from syn-kinematic intrusions. Journal of Geology,
isotopic geochemistry of Precambrian to Paleozoic 113, 553–570.
granitoid suites and the deep-crustal structure of the L ISSENBERG , C. J., VAN S TAAL , C. R., B ÉDARD , J. H. &
southeast margin of the Newfoundland Appalachians. Z AGOREVSKI , A. 2005b. Geochemical constraints on
Canadian Journal of Earth Sciences, 32, 224– 245. the origin of the Annieopsquotch ophiolite belt, south-
K ING , M. S. & B ARR , S. M. 2004. Magnetic and gravity west Newfoundland. Geological Society of America
models across terrane boundaries in southern New Bulletin, 117, 1413– 1426.
Brunswick, Canada. Canadian Journal of Earth L UDMAN , A., H OPECK , J. T. & B ROCK , P. C. 1993.
Sciences, 41, 1027– 1047. Nature of the Acadian Orogeny in eastern Maine. In:
K NIGHT , I., J AMES , N. P. & L ANE , T. E. 1991. The Ordo- R OY , D. C. & S KEHAN , J. W. (eds) The Acadian
vician St. George Unconformity, northern Appala- Orogeny: Recent studies in New England, Maritime
chians: the relationship of plate convergence at the Canada, and the Autochthonous Foreland. Geological
St. Lawrence promontory to the Sauk/Tippecanoe Society of America, Special Paper, 275, 67–84.
sequence boundary. Geological Society of America M AC D ONALD , L. A., B ARR , S. M., W HITE , C. E. &
Bulletin, 103, 1200–1225. K ETCHUM , J. W. F. 2002. Petrology, age, and tectonic
K ROGH , T. E. & K EPPIE , J. D. 1990. Age of detrital zircon setting of the White Rock Formation, Meguma terrane,
and titanite in the Meguma Group, southern Nova Nova Scotia: Evidence for Silurian continental rifting.
Scotia, Canada: clues to the origin of the Meguma Canadian Journal of Earth Sciences, 39, 259–277.
Terrane. Tectonophysics, 177, 307– 323. M AC N IOCAILL , C. 2000. A new Silurian palaeolatitude
L AFRANCE , B. 1989. Structural evolution of a transpres- for eastern Avalonia and evidence for crustal rotations
sion zone in north central Newfoundland. Journal of in the Avalonian margin of southwestern Ireland.
Structural Geology, 11, 705– 716. Geophysical Journal International, 141, 661–671.
L AFRANCE , B. & W ILLIAMS , P. F. 1992. Silurian defor- M ALO , M. & K IRKWOOD , D. 1995. Faulting and progress-
mation in eastern Notre Dame Bay, Newfoundland. ive strain history of the Gaspé Peninsula in post-
Canadian Journal of Earth Sciences, 29, 1899–1914. Taconian time: A review. In: H IBBARD , J. P., VAN
L AIRD , J., T RZCIENSKI , W. E., J R . & B OTHNER , W. A. S TAAL , C. R. & C AWOOD , P. A. (eds) Current Per-
1993. High pressure Taconian and subsequent poly- spectives in the Appalachian –Caledonian Orogen.
metamorphism of southern Quebec and northern Geological Association of Canada, Special Paper, 41,
Vermont. Department of Geology and Geography, 267– 282.
University of Massachusetts, Contribution, 67–2, M ARTIN , H. 1986. Effect of steeper Archean geothermal
1–32. gradient on geochemistry of subduction-zone mag-
L ANDING , E. 1996. Avalon: Insular continent by the latest mas. Geology, 14, 753–756.
Precambrian. In: N ANCE , R. D. & T HOMPSON , M. D. M C L AUGHLIN , R. J., B ARR , S. M., H ILL , M. D., T HOMP-
(eds) Avalonian and Related Peri-Gondwanan SON , M. D., R AMEZANI , J. & R EYNOLDS , P. H. 2003.
Terranes of the Circum-North Atlantic. Geological The Moosehorn plutonic suite, southeastern Maine
Society of America, Special Paper, 304, 29– 63. and southwestern New Brunswick: Age, petrochemis-
L IN , S., D AVIS , D. D., B ARR , S. M., C HEN , Y., VAN try, and tectonic setting. Atlantic Geology, 39,
S TAAL , C. R. & C ONSTANTIN , M. 2007. U– Pb geo- 123– 146.
chronological constraints on the geological evolution M C L EOD , M. J. 1990. Geology, geochemistry, and related
and regional correlation of the Aspy terrane, Cape mineral deposits of the Saint George batholith. New
Brunswick Department of Natural Resources and of the Avalon composite terrane. Journal of Geology,
Energy, Mineral Resource Report, 5, 1 –169. 101, 375 –388.
M C L EOD , M. J., J OHNSON , S. C. & K ROGH , T. E. 2003. O’B RIEN , B. 2003. Geology of the central Notre Dame
Archived U– Pb zircon dates from southern New Bay region (parts of NTS areas 2E/3,6,11), northeast-
Brunswick. Atlantic Geology, 39, 209– 225. ern Newfoundland. Government of Newfoundland and
M C L EOD , M. J., P ICKERILL , R. K. & L UX , R. D. 2001. Labrador Department of Mines and Energy, Report,
Mafic intrusions on Campobello Island: Implications 03–03, 1 –147.
for New Brunswick-Maine correlations. Atlantic O’B RIEN , B., O’B RIEN , S. & D UNNING , G. R. 1991.
Geology, 37, 17–40. Silurian cover, Late Precambrian-Early Ordovician
M C N AMARA , A. K., M AC N IOCAILL , C., VAN DER basement, and the chronology of Silurian orogenesis
P LUIJM , B. A. & VAN DER V OO , R. 2001. West in the Hermitage Flexure (Newfoundland Appala-
African proximity of the Avalon terrane in the latest chians). American Journal of Science, 291, 760– 799.
Precambrian. Geological Society America Bulletin, O’B RIEN , S. J., O’B RIEN , B. H., D UNNING , G. R. &
113, 1161– 1170. T UCKER , R. D. 1996. Late Neoproterozoic Avalonian
M C N ICOLL , V., VAN S TAAL , C. R. & W ALDRON , J. W. F. and related peri-Gondwanan rocks of the Newfound-
2001. Accretionary history of the Northern Appala- land Appalachians. In: N ANCE , R. D. & T HOMPSON ,
chians: SHRIMP study of Ordovician-Silurian syntec- M. D. (eds) Avalonian and Related Peri-Gondwanan
tonic sediments in the Canadian Appalachians. Terranes of the Circum-North Atlantic. Geological
Abstract Volume, GAC/MAC Annual meeting, Society of America, Special Paper, 304, 9–28.
St. John’s Nfld., 1– 100. O’B RIEN , B., S WINDEN , H. S., D UNNING , G. R.,
M EANS , W. D. 1995. Shear zones and rock history. Tecto- W ILLIAMS , S. H. & O’B RIEN , F. 1997. A peri-
nophysics, 247, 157– 160. Gondwanan arc-back arc complex in Iapetus: Early-
M ILLER , B. V. & F YFFE , L. R. 2002. Geochronology of Mid Ordovician evolution of the Exploits Group,
the Letete and Waweig formations, Mascarene Group, Newfoundland. American Journal of Science, 297,
southwestern New Brunswick. Atlantic Geology, 220–272.
38, 29– 36. P ARK , A. F. & W HITEHEAD , J. 2003. Structural transect
M ILLER , B. V., F ETTER , A. H. & S TEWART , K. G. 2006. through Silurian turbidites of the fredericton Belt
Plutonism in three orogenic pulses, eastern Blue Ridge southwest of Fredericton, New Brunswick: The role
province, southern Appalachians. Geological Society of the Fredericton Fault in late Iapetus convergence.
of America Bulletin, 118, 171–184. Atlantic Geology, 39, 227– 237.
M OENCH , R. H. & A LEINIKOFF , J. N. 2003. Stratigraphy, P ARK , A. F., W ILLIAMS , P. F., R ALSER , S. & L EGER , A.
geochronology and accretionary terrane setting of two 1994. Geometry and kinematics of a major crustal
Bronson Hill arc sequences, northern New England. shear zone segment in the Appalachians of southern
Physics and Chemistry of the Earth, 28, 113– 160. New Brunswick. Canadian Journal of Earth Sciences,
M ORAN , P. C., B ARR , S. M., W HITE , C. E. & H AMILTON , 31, 1523–1535.
M. A. 2007. Petrology, age and tectonic setting of P EARCE , J. A. 1996. Sources and settings of granitic rocks.
the Seal Island Pluton, offshore southwestern Episodes, 19, 120– 125.
Nova Scotia. Canadian Journal of Earth Sciences, P EARCE , J. A., H ARRIS , N. B. W. & T INDLE , A. G. 1984.
44, 1– 12. Trace element discrimination diagrams for the tectonic
M URPHY , J. B., N ANCE , R. D. & K EPPIE , J. D. 2002. Dis- interpretation of granitic rocks. Journal of Petrology,
cussion and reply: West African proximity of the 25, 956– 983.
Avalon terrane in the latest Precambrian. Geological P EHRSSON , S., VAN S TAAL , C. R., H ERD , R. K. &
Society America Bulletin, 114, 1049–11052. M C N ICOLL , V. 2003. The Cormacks Lake complex,
M URPHY , J. B., K EPPIE , J. D., D OSTAL , J. & C OUSINS , B. Dashwoods Subzone: a window into the deeper
l. 1996. Repeated Late Neoproterozoic-Silurian lower levels of the Notre Dame Arc. In: P EREIRA , C. P. G.,
crustal melting beneath the Antigonish Highlands, W ALSH , D. G. & K EAN , B. F. (eds) Current Research
Nova Scotia: Nd-isotopic evidence and tectonic 2002. Newfoundland Department of Mines and
interpretations. In: N ANCE , R. D. & T HOMPSON , Energy, Geological Survey, Report, 2002–1,
M. D. (eds) Avalonian and Related Peri-Gondwanan 115–125.
Terranes of the Circum-North Atlantic. Geological P E -P IPER , G. & J ANSA , L. F. 1999. Pre-mesozoic base-
Society of America, Special Paper, 304, 109– 120. ment rocks offshore Nova Scotia, Canada: New con-
M URPHY , J. B., VAN S TAAL , C. R. & K EPPIE , J. D. 1999. straints on the accretion history of the Meguma
Middle to Late Paleozoic Acadian Orogeny in the terrane. Geological Society America Bulletin, 111,
northern Appalachians: A Laramide-style plume- 1773– 1791.
modified orogeny. Geology, 27, 653–656. P E -P IPER , G. & P IPER , D. J. W. 2002. A synopsis of the
M URPHY , J. B., F ERNANDEZ -S UAREZ , J., K EPPIE , J. D. geology of the Cobequid Highlands, Nova Scotia.
& J EFFRIES , T. 2004. Contiguous rather than discrete Atlantic Geology, 38, 145– 160.
Paleozoic histories for the Avalon and Meguma ter- P IASECKI , M. A. J. 1988. A major ductile shear zone in
ranes based on detrital zircon data. Geology, 32, the Bay d’Espoir area, Gander terrane, southeastern
585– 588. Newfoundland. Newfoundland Department of Mines
N ANCE , R. D. & D ALLMEYER , R. D. 1993. 40Ar/39Ar and Energy, Mineral Development Division, Report,
amphibole ages from the Kingston Complex, New 88–1, 135– 144.
Brunswick: Evidence for Silurian-Devonian tecto- P INCIVY , A., M ALO , M., R UFFET , G., T REMBLAY , A. &
nothermal activity and implications for the accretion S ACKS , P. E. 2003. Regional metamorphism of the
Appalachian Humber zone of Gaspe Peninsula: Newfoundland? Precambrian Research, 148,

40
Ar/39Ar evidence for crustal thickening during the 320– 341.
Taconian orogeny. Canadian Journal of Earth S ACKS , P. E., M ALO , M., T RZCIENSKI , W. E., P INCIVY ,
Sciences, 40, 301– 315. A. & G OSSELIN , P. 2004. Taconian and Acadian trans-
P OLLOCK , J. C., W ILTON , D. H. C., VAN S TAAL , C. R. & pression between the internal Humber Zone and the
M ORRISSEY , K. D. 2007. U– Pb detrital zircon geo- Gaspe belt in the Gaspe Peninsula: tectonic history of
chronological constraints on the early Silurian col- the Shickshock Sud fault zone. Canadian Journal of
lision of Ganderia and Laurentia along the Dog Bay Earth Sciences, 41, 635–653.
Line: The terminal Iapetan suture in the Newfoundland S AMSON , S. D., B ARR , S. M. & W HITE , C. E. 2000. Nd
Appalachians. American Journal of Science, 307, isotopic characteristics of terranes within the Avalon
399–433. Zone, southern New Brunswick. Canadian Journal of
P OTTER , J., L ONGSTAFFE , F. J. & B ARR , S. M. 2008. Earth Sciences, 37, 1039– 1052.
Regional 18O depletion of Neoproterozoic igneous S CHENK , P. E. 1997. Sequence stratigraphy and
rocks from Avalonia, Cape Breton Island and southern provenance on Gondwana’s margin: The Meguma
New Brunswick, Canada. Geological Society of zone (Cambrian to Devonian) of Nova Scotia,
America Bulletin, 120, 347 –367. Canada. Geological Society America Bulletin, 109,
P RICE , J. R., B ARR , S. M., R AESIDE , R. P. & R EYNOLDS , 395– 409.
P. H. 1999. Petrology, tectonic setting, and 40Ar/39Ar S CHOFIELD , D. I. & D’L EMOS , R. S. 2000. Granite petro-
(hornblende) dating of the Late Ordovician–Early genesis in the Gander Zone, NE Newfoundland:
Silurian Belle Cote Road orthogneiss, western Cape mixing of melts from multiple sources and the role of
Breton Highlands, Nova Scotia. Atlantic Geology, 35, lithospheric delamination. Canadian Journal of Earth
1–17. Sciences, 37, 535–547.
P RIGMORE , J. K., B ULLER , A. J. & W OODCOCK , N. H. S CHULTZ , K. J., S TEWART , D. B., T UCKER , R. D.,
1997. Rifting during separation of eastern Avalonia P OLLOCK , J. & A YUSO , R. A. 2008. The Ellsworth
from Gondwana: evidence from subsidence analysis. terrane coastal Maine: Geochronology, geochemistry,
Geology, 25, 203– 206. and Nd–Pb isotopic composition – Implications for
Q UINN , L., B ASHFORTH , A. R., B URDEN , E. T., the rifting of Ganderia. Geological Society of America
G ILLESPIE , H., S PRINGER , R. K. & W ILLIAMS , Bulletin, 120, 1134– 1158.
S. H. 2004. The Red Island Road Formation: Early S TOCKMAL , G. S. & W ALDRON , J. W. F. 1993. Structural
Devonian terrestrial fill in the Anticosti foreland and tectonic evolution of the Humber Zone, western
basin, western Newfoundland. Canadian Journal of Newfoundland: Implications of cross sections
Earth Sciences, 41, 587 –602. through the Appalachian structural front, Port au Port
R ATCLIFFE , N. M., H AMES , W. E. & S TANLEY , R. S. Peninsula. Tectonics, 12, 1056–1075.
1998. Interpretation of ages of arc magmatism and col- S WINDEN , H. S., J ENNER , G. A. & S ZYBINSKI , Z. A.
lisional tectonics in the Taconian Orogen of western 1997. Magmatic and tectonic evolution of the
New England. American Journal of Earth Sciences, Cambrian-Ordovician Laurentian margin of Impetus:
298, 791– 797. Geochemical and isotopic constraints from the Notre
R ATCLIFFE , N. M., H ARRIS , A. G. & W ALSH , G. J. Dame subzone, Newfoundland. In: S INHA , K.,
1999. Tectonic and regional metamorphic impli- W HALEN , J. B. & H OGAN , J. P. (eds) The Nature of
cations of the discovery of Middle Ordovician cono- Magmatism in the Appalachian Orogen. Geological
donts in cover rocks east of Green Mountain massif, Society of America, Memoir, 191, 367 –395.
Vermont. Canadian Journal of Earth Sciences, 36, S ZYBINSKI , Z. A. 1995. Paleotectonic and structural
371–382. setting of thye western Notre Dame Bay area, New-
R ICHTER , D. A. & R OY , D. C. 1974. Sub-greenschist foundland Appalachians. Unpublished PhD thesis,
metamorphic assemblages in northern Maine. Cana- Memorial University of Newfoundland, St. John’s,
dian Mineralogist, 12, 469–474. Newfoundland.
R OBINSON , P., T UCKER , R. D., B RADLEY , D., B ERRY , T ATE , M. C. & C LARKE , D. B. 1995. Petrogenesis and
H. N. & O SBERG , P. H. 1998. Paleozoic orogens in regional tectonic significance of Late Devonian mafic
New England, USA. Geologiska Fó´reningens Stock- intrusions in the Meguma Zone, Nova Scotia. Cana-
holm Forhandlingar, 120, 119 –148. dian Journal of Earth Sciences, 32, 1883– 1898.
R OGERS , N. & VAN S TAAL , C. R. 2003. Volcanology T ATE , M. C. & M ERRETT , D. 1994. Compilation of major
and tectonic setting of the northern Bathurst Mining oxide, trace element and rare-earth element analyses
Camp: Part II – Mafic volcanic constraints on for Late Devonian peraluminous granitoids in the
back-arc opening. Economic Geology, Monograph, Meguma Zone, Nova Scotia. Nova Scotia Department
11, 181–202. of Natural Resources, Open File Report, 94– 14, 71.
R OGERS , N., VAN S TAAL , C. R. & V ALVERDE - T HURLOW , J. G., S PENCER , C. P., B OERNER , D. E.,
V AQUERO , P. 2005. Geology, Noel Paul’s Brook, R EED , L. E. & W RIGHT , J. A. 1992. Geophysical
Newfoundland (NTS 12-A/9). Geological Survey of interpretation of a high resolution reflection seismic
Canada Open File OF4547, scale 1:50 000, 1 sheet. survey at the Buchans mine, Newfoundland. Canadian
R OGERS , N., VAN S TAAL , C. R., M C N ICOLL , V., Journal of Earth Sciences, 29, 2022– 2037.
P OLLOCK , J., Z AGOREVSKI , A. & W HALEN , J. T REMBLAY , A. & P INET , N. 1994. Distribution and
2006. Neoproterozoic and Cambrian arc magmatism characteristics of Taconian and Acadian deformation,
along the eastern margin of the Victoria Lake Super- southern Quebec Appalachians. Geological Society of
group: A remnant of Ganderian basement in central America Bulletin, 106, 1172– 1181.
T UCKER , R. D., O SBERG , P. H. & B ERRY , H. N. 2001. Orogen. Geological Association of Canada, Special
The geology of a part of Acadia and the nature of the Paper, 41, 367–389.
Acadian orogeny across central and eastern Maine. VAN S TAAL , C. R. & H ATCHER , R. D., J R . 2009. Global
American Journal of Science, 301, 205 –260. setting of Ordovician orogenesis. In: F INNEY , S. C. &
V ALVERDE -V AQUERO , P., D UNNING , G. & VAN S TAAL , B ERRY , W. B. N. (eds) Global Ordovician Earth
C. R. 2000. The Margaree orthogneiss (Port aux System. Geological Society of America, Special
Basques Complex, SW Newfoundland) evolution of a Paper (in press).
peri-Gondwanan, Mid Arenig– Early Llanvirn, mafic- VAN S TAAL , C. R., R AVENHURST , C., W INCHESTER ,
felsic igneous complex. Canadian Journal of Earth J. A., R ODDICK , J. C. & L ANGTON , J. P. 1990. Post-
Sciences, 37, 1691– 1710. Taconic blueschist suture in the northern Appalachians
V ALVERDE -V AQUERO , P. & VAN S TAAL , C. R. 2001. of northern New Brunswick, Canada. Geology, 18,
Relationships between the Dunnage–Gander zones 1073– 1077.
in the Victoria Lake– Peter Strides Pond area. In: VAN S TAAL , C. R., W INCHESTER , J. A. & B EDARD , J. H.
P EREIRA , C. P. G., W ALSH , D. G. & K EAN , B. F. 1991. Geochemical variations in Ordovician volcanic
(eds) Current Research 2001. Newfoundland Depart- rocks of the northern Miramichi Highlands and their
ment of Mines and Energy, Geological Survey tectonic significance. Canadian Journal of Earth
Report, 2000– 1, 1– 9. Sciences, 28, 1031–1049.
V ALVERDE -V AQUERO , P., VAN S TAAL , C. R., VAN S TAAL , C. R., W INCHESTER , J. A., B ROWN , M. &
M C N ICOLL , V. & D UNNING , G. 2006. Middle Ordo- B URGESS , J. L. 1992. A reconnaissance geotra-
vician magmatism and metamorphism along the verse through southwestern Newfoundland. Current
Gander margin in Central Newfoundland. Journal of Research, Part D, Geological Survey of Canada,
the Geological Society, London, 163, 347 –362. Paper 92-1D, 133–144.
V AN DER P LUIJM , B. A. 1986. Geology of eastern New VAN S TAAL , C. R., D UNNING , G., V ALVERDE , P.,
World Island, Newfoundland: An accretionary B URGESS , J. & B ROWN , M. 1994. Arenig and
terrane in the northeastern Appalachians. Geological younger evolution of the Gander margin: A compari-
Society of America Bulletin, 97, 932–945. son of the New Brunswick and Newfoundland seg-
V AN DER V ELDEN , A. J., VAN S TAAL , C. R. & C OOK , ments. Geological Association of Canada, NUNA
F. A. 2004. Crustal structure, fossil subduction and conference: New perspectives in the Appalachian–
the tectonic evolution of the Newfoundland Appala- Caledonian Orogen, Program and Abstracts. Atlantic
chians: Evidence from a reprocessed seismic reflection Geology, 30, 178 –179.
survey. Geological Society America Bulletin, 116, VAN S TAAL , C. R., S ULLIVAN , R. W. & W HALEN , J. B.
1485–1498. 1996. Provenance and tectonic history of the Gander
V AN DER V OO , R. & J OHNSON , R. J. E. 1985. Paleomag- Margin in the Caledonian/Appalachian Orogen: Impli-
netism of the Dunn Point Formation (Nova Scotia): cations for the origin and assembly of Avalonia. In:
High paleolatitudes for the Avalon terrane in the N ANCE , R. D. & T HOMPSON , M. D. (eds) Avalonian
Late Ordovician. Geophysical Research Letters, 12, and Related Peri-Gondwanan Terranes of the Circum-
337– 340. North Atlantic. Geological Society of America, Special
V AN H UNEN , J., VAN DER B ERG , A. P. & V LAAR , N. J. Paper, 304, 347– 367.
2002. On the role of subducting oceanic plateaus in VAN S TAAL , C. R., D EWEY , J. F., M AC N IOCAILL , C. &
the development of shallow flat subduction. Tectono- M C K ERROW , S. 1998. The Cambrian-Silurian tec-
physics, 352(3–4), 317– 333. tonic evolution of the northern Appalachians:
VAN S TAAL , C. R. 1994. The Brunswick subduction History of a complex, southwest Pacific-type segment
complex in the Canadian Appalachians: Record of of Iapetus. In: B LUNDELL , D. J. & S COTT , A. C.
the Late Ordovician to Late Silurian collision (eds) Lyell: the Past is the Key to the Present. Geologi-
between Laurentia and the Gander margin of Avalon. cal Society, London, Special Publications, 143,
Tectonics, 13, 946– 962. 199–242.
VAN S TAAL , C. R. 2005. The Northern Appalachians. VAN S TAAL , C. R., R OGERS , N. & T AYLOR , B. E. 2001.
In: S ELLEY , R. C., C OCKS , L. R. & P LIMER , I. R. Formation of low temperature mylonites and phyllo-
(eds) Encyclopedia of Geology. Elsevier, Oxford, 4, nites by alkali-metasomatic weakening of felsic volca-
81–91. nic rocks during a progressive, subduction-related
VAN S TAAL , C. R. 2007. Pre-Carboniferous tectonic evol- deformation. Journal of Structural Geology, 23,
ution and metallogeny of the Canadian Appalachians, 903–921.
In: G OODFELLOW , W. D. (ed.) Mineral Deposits of VAN S TAAL , C. R., W ILSON , R. A. ET AL . 2003. Geology
Canada: A Synthesis of Major Deposit-types, District and tectonic history of the Bathurst Supergroup and
Metallogeny, the Evolution of Geological Provinces, its relationships to coeval rocks in southwestern
and Exploration Methods. Geological Association of New Brunswick and adjacent Maine – a synthesis.
Canada, Mineral Deposit Division, Special Publi- In: G OODFELLOW , W. D., M C C UTCHEON , S. R. &
cation, 5, 793–818. P ETER , J. M. (eds) Massive Sulfide Deposits of the
VAN S TAAL , C. R. & DE R OO , J. A. 1995. Mid-Paleozoic Bathurst Mining Camp, New Brunswick, and Northern
tectonic evolution of the Appalachian Central Mobile Maine. Economic Geology, Monograph, 11, 37– 60.
Belt in northern New Brunswick, Canada: Collision, VAN S TAAL , C. R., M C N ICOLL , V., V ALVERDE -
extensional collapse, and dextral transpression. In: V AQUERO , P., B ARR , S., F YFFE , L. R. & R EUSCH ,
H IBBARD , J., VAN S TAAL , C. R. & C AWOOD , P. (eds) D. N. 2004. Ganderia, Avalonia, and the Salinic and
Current Perspectives in the Appalachian-Caledonian Acadian orogenies. Abstracts with Programs,
Geological Society of America, 39th Annual Meeting, W ALDRON , J. W. F., W HITE , C. E., B ARR , S. M.,
NE-section, 128– 129. S IMONETTI , A. & H EAMAN , L. M. 2009. Provenance
VAN S TAAL , C. R., V ALVERDE -V AQUERO , P., Z AGOR- of the meguma terrane, Nova Scotia: rifted margin of
EVSKI , A., B OUTSMA , S., P EHRSSON , S., VAN early Paleozoic Gondwana. Canadian Journal of
N OORDEN , M. & M C N ICOLL , V. 2005. Geology, Earth Sciences, 46, 1– 8.
King George IV Lake, Newfoundland (NTS 12-A/04). W ALKER , J. A., G OWER , S. & M C C UTCHEON , S. R.
Geological Survey of Canada, Open File 1665, scale 1993. Antinouri Lake-Nicolas Denys project, Glouce-
1:50 000, 1 sheet. ster and Restigouche counties, New Brunswick. In:
VAN S TAAL , C. R., W HALEN , J. B. ET AL . 2007. The Notre A BBOTT , S. A. (ed.) Current research 1993. New
Dame arc and the Taconic Orogeny in Newfoundland. Brunswick Department of Natural Resources and
In: H ATCHER , R. D., J R ., C ARLSON , M. P., Energy, Minerals and Energy Division, Information
M C B RIDE , J. H. & M ARTÍNEZ C ATALÁN , J. R. circular 93-1, 58– 70.
(eds) 4-D Framework of Continental Crust. Geological W EST , D. P., L UDMAN , A. & L UX , D. R. 1992. Silurian
Society of America, Memoir, 200, 511– 552. age for the Pocomoonshine Gabbro-Diorite, south-
VAN S TAAL , C. R., C URRIE , K. L., R OWBOTHAM , G., eastern Maine and its regional tectonic implications.
G OODFELLOW , W. & R OGERS , N. 2008. Pressure– American Journal of Earth Sciences, 292, 253–273.
temperature paths and exhumation of Late W EST , D. P., B EAL , H. M. & G ROVER , T. 2003. Silurian
Ordovician– Early Silurian blueschists and associated deformation and metamorphism of Ordovician arc
metamorphic nappes of the Salinic Brunswick subduc- rocks of the Casco Bay Group, south-central Maine.
tion complex, northern Appalachians. Geological Canadian Journal of Earth Sciences, 40, 887–905.
Society of America Bulletin, 120, 1455–1477. W HALEN , J. B. 1990. Geology of a northern portion of the
VAN S TAAL , C. R., C ASTONGUAY , S., M C N ICOLL , V., Central Plutonic Belt, New Brunswick. Geological
B REM , A., H IBBARD , J., S KULSKI , T. & J OYCE , N. Survey of Canada, Map 1751A, scale 1:100 000.
2009. Taconic arc– continent collision confirmed in W HALEN , J. B. 1993a. Geology, petrography and geo-
the Newfoundland Appalachians. Geological Society chemistry of Appalachian granites in New Brunswick
of America, 44th Annual Meeting, NE– section, p. 4. and Gaspesie, Quebec. Geological Survey of Canada,
V AN W AGONER , N. A., L EYBOURNE , M. I., D ADD , Bulletin, 436, 124.
K. A. & H USKINS , M. L. A. 2001. The Silurian(?) Pas- W HALEN , J. B. 1993b. Petrographic, geochemical,
samaquoddy Bay mafic dyke swarm, New Brunswick: and isotopic data on granites in New Brunswick and
Petrogenesis and tectonic implications. Canadian Gaspesie, Quebec. In digital format and sample
Journal of Earth Sciences, 38, 1565–1578. location maps. Geological Survey of Canada, Open
V AN W AGONER , N. A., L EYBOURNE , M. I., D ADD , File 2570.
K. A., B ALDWIN , D. K. & M C N EIL , W. 2002. Late W HALEN , J. B., F YFFE , L. R., L ONGSTAFFE , F. &
Silurian bimodal volcanism of southwestern New J ENNER , G. A. 1996a. The position and nature of the
Brunswick, Canada: Products of continental Gander–Avalon boundary, southern New Brunswick,
extension. Geological Society of America Bulletin, based on geochemical and isotopic data from granitoid
114, 400– 418. rocks. Canadian Journal of Earth Sciences, 33,
W ALDRON , J. W. F. & VAN S TAAL , C. R. 2001. Taconic 129– 139.
orogeny and the accretion of the Dashwoods block: A W HALEN , J. B., J ENNER , G. A., L ONGSTAFFE , F. J. &
peri-Laurentian microcontinent in the Iapetus Ocean. H EGNER , E. 1996b. Nature and evolution of the
Geology, 29, 811– 814. eastern margin of Iapetus: Geochemical and isotopic
W ALDRON , J. W. F., P IPER , D. J. W. & P E -P IPER , G. constraints from Siluro-Devonian granitoid plutons in
1989. Deformation of the Cape Chignecto Pluton, the New Brunswick Appalachians. Canadian Journal
Cobequid Highlands, Nova Scotia: Thrusting at the of Earth Sciences, 33, 140– 155.
Meguma-Avalon boundary. Atlantic Geology, 25, W HALEN , J. B., J ENNER , G. A., L ONGSTAFFE , F. J.,
51– 62. G ARIÉPY , C. & F RYER , B. J. 1997a. Implications of
W ALDRON , J. W. F., A NDERSON , S. D., C AWOOD , P. A., granitoid geochemical and isotopic (Nd, O, Pb) data
G OODWIN , L. B., H ALL , J., J AMIESON , R. A., from the Cambrian-Ordovician Notre Dame arc for
P ALMER , S. E., S TOCKMAL , G. S. & W ILLIAMS , the evolution of the Central Mobile belt, Newfound-
P. F. 1998. Evolution of the Appalachian Laurentian land Appalachians. In: S INHA , K., W HALEN , J. B. &
margin; Lithoprobe results in western Newfoundland. H OGAN , J. P. (eds) The Nature of Magmatism in the
Lithoprobe East transect–Le transect est du projet Appalachian Orogen. Geological Society of America,
Lithoprobe: Ottawa, ON, Canada. Canadian Journal Memoir, 191, 367–395.
of Earth Sciences, 35, 1271– 1287. W HALEN , J. B., VAN S TAAL , C. R., L ONGSTAFFE , F. J.,
W ALDRON , J. W. F., M URPHY , J. B., M ELCHIN , M. J. & G ARIEPY , C. & J ENNER , G. A. 1997b. Insights
D AVIS , G. 1996. Silurian tectonics of Western into tectonostratigraphic zone identification in
Avalonia: Strain-corrected subsidence history of the southwestern Newfoundland based on isotopic (Nd,
Arisaig Group, Nova Scotia. Journal of Geology, O, Pb) and geochemical data. Atlantic Geology, 33,
104, 677– 694. 231– 241.
W ALDRON , J. W. F., H ENRY , A. D., B RADLEY , J. C. & W HALEN , J. B., S YME , E. C. & S TERN , R. A. 1998. Geo-
P ALMER , S. E. 2003. Development of a folded thrust chemical and Nd isotopic evolution of Paleoprotero-
stack: Humber Arm Allochthon, Bay of Islands, New- zoic arc-type magmatism in the Flin Flon Belt,
foundland Appalachians. Canadian Journal of Earth Trans-Hudson Orogen, Canada. Canadian Journal of
Sciences, 40, 237– 253. Earth Sciences, 35, 227–250.
W HALEN , J. B., M C N ICOLL , V. J., VAN S TAAL , C. R., W ILSON , R. A., K AMO , S. & B URDEN , E. T. 2005.
L ISSENBERG , C. J., L ONGSTAFFE , F. J., J ENNER , G. Geology of the Val d’Amour Formation: revisiting
A. & VAN B REEMEN , O. 2006. Spatial, temporal and the type area of the Dalhousie Group, northern New
geochemical characteristics of Silurian collision-zone Brunswick. In: M ARTIN , G. L. (ed.) Geological Inves-
magmatism: An example of a rapidly evolving mag- tigations in New Brunswick for 2004. New Brunswick
matic system related to slab break-off. Lithos, 89, Department of Natural Resources; Minerals, Policy
377– 404. and Planning Division, Mineral Resource Report,
W HITE , C. E., B ARR , S. M., B EVIER , M. L. & K AMO , S. 2005-1, 167– 212.
1994. A revised interpretation of the Cambrian and W ILSON , R. A., VAN S TAAL , C. R. & K AMO , S. 2008.
Ordovician rocks in the Bourinot belt of central Cape Lower Silurian subduction-related volcanic rocks in
Breton Island, Nova Scotia. Atlantic Geology, 30, the Chaleurs Group, northern New Brunswick,
123– 142. Canada. Canadian Journal of Earth Sciences, 45,
W HITE , C. E., B ARR , S. M., R EYNOLDS , P. H., G RACE , E. 981–998.
& M C M ULLIN , D. W. A. 2006. The Pocologan meta- W OOD , D. A. 1980. The application of a Th-Hf-Ta
morphic suite: High-pressure metamorphism in a Silur- diagram to problems of tectonomagmatic classification
ian fore-arc complex, Kingston terrane, southern New and to establishing the nature of crustal contamination
Brunswick. Canadian Mineralogist, 44, 905–927. of basaltic lavas of the British Tertiary volcanic
W ILLIAMS , H. 1979. Appalachian Orogen in Canada. province. Earth and Planetary Science Letters, 50,
Canadian Journal of Earth Sciences, 16, 792– 807. 11–30.
W ILLIAMS , H. & S T . J ULIEN , P. 1982. The Baie Verte Y ANG , X. M., L ENTZ , D. R. & M C C UTCHEON , S. R.
Brompton Line: Early Paleozoic continent ocean inter- 2003. Petrochemical evolution of subvolcanic
face in the Canadian Appalachians. In: S T . J ULIEN , P. granitoid intrusions within the Late Devonian Mount
& B ELAND , J. (eds) Major Structural Zones and Faults Pleasant Caldera, southwestern New Brunswick,
of the Northern Appalachians. Geological Association Canada: Comparison of Au versus Sn-W-Mo-
of Canada, Special Paper, 24, 177–208. polymetallic mineralization systems. Atlantic
W ILLIAMS , H., D ALLMEYER , R. D. & W ANLESS , R. K. Geology, 39, 97– 121.
1976. Geochronology of the Twillingate Granite and Z AGOREVSKI , A., R OGERS , N., VAN S TAAL , C. R.,
Herring Neck Group, Notre Dame Bay, Newfound- M C N ICOLL , V., L ISSENBERG , C. J. & V ALVERDE -
land. Canadian Journal of Earth Sciences, 13, V AQUERO , P. 2006. Lower to Middle Ordovician
1591–1601. evolution of peri-Laurentian arc and back-arc
W ILLIAMS , H., C OLMAN -S ADD , S. P. & S WINDEN , H. S. complexes in the Iapetus: constraints from the
1988. Tectonicstratigraphic subdivisions of central Annieopsquotch Accretionary Tract, Central New-
Newfoundland. Current Research, Part B. Geological foundland. Geological Society of America Bulletin,
Survey of Canada, Paper, 88-1B, 91–98. 118, 324 –342.
W ILLIAMS , H., C URRIE , K. L. & P IASECKI , M. A. J. Z AGOREVSKI , A., VAN S TAAL , C. R., M C N ICOLL , V. C.
1993. The Dog Bay Line: A major Silurian tectonic & R OGERS , N. 2007a. Upper Cambrian to Upper
boundary in northeast Newfoundland. Canadian Ordovician peri-Gondwanan island arc activity in the
Journal of Earth Sciences, 30, 2481–2494. Victoria Lake Supergroup, Central Newfoundland:
W ILLIAMS , H., L AFRANCE , B., D EAN , P. L., W ILLIAMS , Tectonic development of the Ganderian margin.
P. F., P ICKERING , K. T. & VAN DER P LUIJM , B. A. American Journal of Science, 307, 339–370.
1995. Badger Belt. In: W ILLIAMS , H. (ed.) The Appa- Z AGOREVSKI , A., VAN S TAAL , C. R. & M C N ICOLL , V.
lachian/Caledonian Orogen: Canada and Greenland. 2007b. Distinct Taconic, Salinic and Acadian defor-
Geological Survey of Canada, Geology of Canada, 6, mation along the Iapetus suture zone, Newfoundland
403– 413 (also Geological Society of America, the Appalachians. Canadian Journal of Earth Sciences,
Geology of North America, F-1). 44, 1567–1585.
W ILLIAMS , P. F., E LLIOTT , C. G. & L AFRANCE , B. D. Z AGOREVSKI , A., VAN S TAAL , C. R., M C N ICOLL , V.,
1988. Structural geology and melanges of eastern R OGERS , N. & V ALVERDE -V AQUERO , P. 2007c. Tec-
Notre Dame Bay, Newfoundland. Geological Associ- tonic architecture of an arc-arc collision zone, New-
ation of Canada, Annual Meeting, St. John’s, New- foundland Appalachians. In: C LIFT , P. & D RAUT , A.
foundland Field Trip Guide Book, Trip B2. (eds) Formation and Applications of the Sedimentary
W ILSON , R. 2003. Geochemistry and petrogenesis of Record in Arc-Collision Zones. Geological Society of
Ordovician arc-related mafic volcanic rocks in the America, Special Paper, 436, 309–333.
Popelogan Inlier, northern New Brunswick. Canadian Z AGOREVSKI , A., R OGERS , N., VAN S TAAL , C. R.,
Journal of Earth Sciences, 40, 1171–1189. M C C LENNAGHAN , S. & H ASLAM , R. 2007d. Tecto-
W ILSON , R., B URDEN , E. T., B ERTRAND , R., A SSELIN , nostratigraphic relationships in the Buchans area: A
E. & M CRACKEN , A. D. 2004. Stratigraphy and composite of Ordovician and Silurian terranes. In:
tectono-sedimentary evolution of the Late Ordovician P EREIRA , C. P. G., W ALSH , D. G. & K EAN , B. F.
to Middle Devonian Gaspe belt in northern New Bruns- (eds) Current Research 2007. Newfoundland Depart-
wick: evidence from the Restigouche area. Canadian ment of Mines and Energy, Geological Survey
Journal of Earth Sciences, 41, 527– 551. Report, 07– 1, 103–116.
From Rodinia to Pangaea: ophiolites from NW Iberia as witness
for a long-lived continental margin
SONIA SÁNCHEZ MARTÍNEZ1, RICARDO ARENAS1*, JAVIER FERNÁNDEZ-SUÁREZ1
& TERESA E. JEFFRIES2
1
Departamento de Petrologı́a y Geoquı́mica e Instituto de Geologı́a Económica (CSIC),
Universidad Complutense, 28040 Madrid, Spain
2
Department of Mineralogy, The Natural History Museum, London SW7 5BD, UK
*Corresponding author (e-mail: arenas@geo.ucm.es)
Abstract: The ophiolites preserved in the Variscan suture of NW Iberia (Galicia) show a broad
variability in lithology, geochemistry and chronology. This wide variety rules out the simplest
plate tectonic scenario in which these ophiolites would have been exclusively related to the
oceanic domain closed during the final Pangaea assembly, that is the Rheic Ocean. The ophiolitic
units from Galicia also provide important data about the palaeogeography immediately preceding
the opening of this ocean, and some information about pre-Gondwanan supercontinent cycles.
Six different ophiolites can be distinguished in the allochthonous complexes of Galicia: the
Purrido, Somozas, Bazar, Vila de Cruces, Moeche and Careón units. The Purrido Ophiolite is con-
stituted by metagabbroic amphibolites with igneous protoliths dated at 1159 + 39 Ma (Mesopro-
terozoic), and geochemical affinities typical of island-arc tholeiites. These mafic rocks can be
interpreted as one of the scarce members of the pre-Rodinian ophiolites, and they were probably
generated in a back-arc setting in the periphery of the West African Craton. The Somozas Ophio-
litic Mélange consists of a mixing of submarine volcanic rocks (pillow-lavas, submarine breccias,
pillow-breccias, hyaloclastites), diabases, gabbros, microgabbros, diorites and granitoids, sur-
rounded by a matrix of serpentinites or, less frequently, phyllites. Two granitic samples from
this mélange yield U–Pb ages ranging between c. 527 and 503 Ma (Cambrian), which together
with the characteristic arc signatures obtained in all the studied igneous rocks suggest that this
ophiolite was generated in a peri-Gondwanan volcanic arc. The Bazar Ophiolite is formed by
different tectonic slices with high temperature amphibolites, granulites, metagabbros and ultrama-
fic rocks. The amphibolites are the most abundant rock type and show typical N– MORB compo-
sitions with igneous protoliths dated at 498 + 2 Ma (Cambrian). The high-temperature
metamorphism affecting some parts of the unit has been dated at c. 480 Ma (lower Ordovician),
and it is considered to be related to the development of an oceanic accretionary complex under
the volcanic arc represented by the upper units of the allochthonous complexes of Galicia. Consid-
ering the most common palaeogeographic reconstructions for the Cambrian period, it is suggested
that the oceanic lithosphere represented by the Bazar Ophiolite was formed into the peri-Gondwa-
nan oceanic domain prior to the rifting of the Avalonian microcontinent, that is the Iapetus–Torn-
quist Ocean. According to current data about the Vila de Cruces Unit, it can be interpreted as a
composite terrane, whose lithologies have U– Pb ages ranging from 1176– 497 Ma, but constituted
by metaigneous rocks with arc signatures. This dataset has been interpreted in relation to the devel-
opment of a back-arc basin around the Cambrian –Ordovician limit, involving a Mesoproterozoic
basement and the reactivation of a former suture. The opening of this back-arc basin can also be
identified as the birth of the Rheic Ocean, and probably it would also include the lithological suc-
cession belonging to the Moeche Unit, although its basic rocks exhibit compositions with more
oceanic character. Finally, the Careón Ophiolite includes remnants of an oceanic lithosphere gen-
erated in a supra-subduction zone setting at 395 + 2 Ma (middle Devonian). This ophiolite was
formed in a contractive Rheic Ocean, shortly preceding the closure of this ocean. This is the
only ophiolite in Galicia that can be related to mature stages of the Rheic Ocean, although as it
is commonly observed in other regions the N–MORB crust is not preserved. This common
oceanic crust has disappeared during subduction, probably in an intra-oceanic setting and during
the generation of the igneous section preserved in the Careón Ophiolite.
The axial zone of the Variscan belt includes differ- (Fig. 1). These terranes include several ophiolitic
ent exotic terranes extending from the western units of different age with a complex tectonothermal
Iberian Peninsula through the French Armorican evolution involving high-pressure events of pre-
and Central Massifs to the Bohemian Massif Variscan and Variscan ages. Considering that the
318 S. SÁNCHEZ MARTÍNEZ ET AL.
Fig. 1. Sketch showing the relationships between the Appalachian, Caledonian and Variscan belts at the end of the
Variscan convergence. The striped area represents the approximate extent of Avalonia. LBM, London–Brabant Massif;
STA, Silesian Terrane Assemblage. The approximate location of the NW Iberia map of Figure 2 is also shown. From
Martı́nez Catalán et al. (2007a).
FROM RODINIA TO PANGAEA: OPHIOLITES FROM NW IBERIA 319
Variscan orogenic belt was generated as a result of two ensembles: the lower and upper ophiolitic
the collision of Laurussia and Gondwana during units (Arenas et al. 2007b) with the Somozas Ophio-
the final assembly of Pangaea, the ophiolites pre- litic Mélange being considered as a structurally
served along the Variscan suture zone should be different case (Fig. 2). Other ophiolitic units exist
related to the oceanic domain that was closed in the in the Trás-os-Montes region, in northern Portugal,
course of the collision, that is the Rheic Ocean in the allochthonous complexes of Morais and Brag-
(Matte 2001; Stampfli & Borel 2002; Winchester ança. These ophiolites have not been included in this
et al. 2002; Murphy & Nance 2003). This interpret- review, but recent descriptions of these mafic
ation, although consistent with the overall plate ensembles and new U –Pb data are reported in Pin
tectonic scenario, is simplistic in light of recent et al. (2006; and references therein). The mafic
data on the NW Iberia ophiolites. These ophiolites units involved in the Variscan suture in NW Iberia
show a remarkable geological and geochemical appear between two allochthonous terranes of con-
diversity and offer an outstanding case study to trasting geological features, the so-called basal and
explore not only the nature of the evolved Rheic upper units (Fig. 2).
Ocean (Sánchez Martı́nez et al. 2007a) but also The upper units are constituted by metasedimen-
the initial stages of its development (Arenas et al. tary rocks intruded by igneous rocks with a wide
2007a), some features of the peri-Gondwanan pre- compositional range (gabbros to granitoids). The
Rheic oceanic realms and even palaeogeographic thickness of these units may reach tens of kilo-
scenarios related to the assembly and dispersal of metres. The upper units have been interpreted as
the Rodinia supercontinent (Sánchez Martı́nez remnants of a magmatic arc whose main activity
et al. 2006). took place at the Cambrian– Ordovician boundary
This work is aimed at providing a synthetic view (c. 500 Ma). The signature of detrital zircons in
of the geological, geochronological and geochem- greywackes from turbidite beds located towards
ical features of the NW Iberian ophiolites and the top of these units suggests that the arc was gen-
their interpreted palaeogeographic and palaeotec- erated in the periphery of the northern margin of
tonic implications. It is focused on the ophiolitic west Gondwana (Fernández Suárez et al. 2003).
units from Galicia, NW Spain, which are included This arc would have rifted from the Gondwanan
in the allochthonous complexes of Cabo Ortegal margin during the initial stages of the opening of the
and Órdenes (Figs 1 & 2). The striking diversity Rheic Ocean and drifted northwards as an island arc
of the ophiolites from NW Iberia is an excellent that eventually collided with Laurussia (Gómez
reminder of the complexity of the terranes with Barreiro et al. 2007). The tectonothermal evolution
oceanic affinity involved in orogenic sutures. At of the upper units is complex and of a polymeta-
the same time, such diversity offers a natural morphic nature (Fernández-Suárez et al. 2002).
example to study processes that occurred at the The arc records a metamorphic episode of inter-
palaeocontinental margins and oceanic realms mediate pressure accompanied by intense defor-
involved in the Variscan collision. If we assume mation that took place between 498 and 480 Ma
that the features of the Variscan suture preserved (Abati et al. 1999, 2007). This event is interpreted
in NW Iberia are not unique at global scale, it is as generated as a consequence of the accretionary
logical to consider that the complexity of oceanic dynamics of the arc. A later HP– HT event is
units preserved in such zones of different orogenic recorded in the lower part of the upper units and
belts may be greater than it is commonly assumed has been dated at 420–400 Ma (Gómez Barreiro
and reported. This may spur further studies that et al. 2006; Fernández-Suárez et al. 2007) and it
should lead to a better and more complete under- has been interpreted as a result of the collision
standing of ancient oceanic realms. (and understacking) of the arc with/under Laurus-
sia, broadly coeval with the collision of Avalonia.
The upper units also record Variscan tectonothermal
The Variscan suture of NW Iberia events related to the final stages of Pangaea
assembly.
The Variscan suture in NW Iberia is preserved in The basal units consist of a rock ensemble of
several allochthonous complexes of synformal continental affinity constituted by metasedimentary
structure and is highlighted by an ensemble of rocks, orthogneisses and mafic rocks that have been
mafic rock units of ophiolitic affinity (Arenas interpreted to represent a section of the most exter-
et al. 1986). In Galicia (NW Spain), six ophiolitic nal part of the Gondwanan margin (Martı́nez
units have been described so far, these are named Catalán et al. 1996). The granitic orthogneisses
the Purrido, Somozas (ophiolitic mélange), Bazar, are abundant, mostly with early Ordovician ages
Vila de Cruces, Moeche and Careón (Fig. 2). and the igneous protoliths show both calc-alkaline
Based on their structural position, five of these six (older intrusions) and peralkaline (younger intru-
ophiolitic units have been traditionally grouped in sions) geochemical affinity. The presence of these
Fig. 2. Schematic geological map and cross-section of the allochthonous complexes of Galicia, showing the
distribution and structural relationships of the most important terranes. Six different ophiolitic units can be distinguished
in the Cabo Ortegal and Órdenes complexes, the Purrido, Somozas, Moeche, Vila de Cruces, Careón and
Bazar ophiolites.
granitoids and coeval mafic rocks are evidence (816 + 15 Ma and 428 + 5 Ma). The latter over-
for magmatic activity at the Gondwanan margin laps with the LA–ICP–MS U –Pb age obtained in
after the separation of the volcanic arc represented two rutile crystals extracted from the overlying
by the upper units. The tectonothermal evolution paragneisses (412 + 19 Ma and 428 + 11 Ma;
of the basal units is characterized by the ubiquitous Sánchez Martı́nez 2009). This age is interpreted to
presence of high-pressure and low- to intermediate- represent a metamorphic event that affected both
temperature metamorphism reaching the (glauco- the amphibolites and the paragneisses.
phane bearing) eclogite and blueschist facies The common amphibolites of the Purrido Ophio-
(Arenas et al. 1995, 1997). The age of this meta- lite have geochemical features that fit those of
morphic episode is dated at c. 370 Ma (Rodrı́guez island-arc tholeiites, as shown in the Th –Hf– Ta
et al. 2003). This event is considered to be generated diagram of Figure 3b (Wood 1980). This interpret-
by the northward directed subduction of the external ation is consistent with normalized abundance
margin of Gondwana under Laurussia, heralding the patterns of immobile trace elements (Pearce 1996)
Variscan collision s.s. The basal units record the ear- (Fig. 3c). Although this pattern is flat and close to
liest Variscan deformation event recognized in the unity (Th and Ce concentrations are similar to
basement of western Europe, this event being those of N –MORB, whereas Nb, Zr, Ti and Y are
coeval with the closure of the Rheic Ocean. This variably depleted), there is a larger negative Nb
ocean possible attained its maximum width at the anomaly suggesting a subduction-related origin
time of the collision of the upper units with Laurus- component (Pearce 1996).
sia and from that point it contracted until it was
closed with the subduction of the Gondwanan Somozas Ophiolitic Mélange
margin under Laurussia.
In the eastern part of the Cabo Ortegal Complex, at
the thrust front of the allochthonous complexes of
The ophiolitic units of NW Iberia Galicia, there is a tectonic mélange of considerable
Purrido Ophiolite size (Fig. 2). In detail, the mélange is formed by
complex tectonic imbrications involving an ophio-
The Purrido ophiolitic unit, c. 300 m thick, is litic unit and slices of the basal units and the
exposed in the western part of the Cabo Ortegal Parauthocton. The mélange contains metric to hec-
Complex, under the lower part of the upper units tometric lens-shaped tectonic blocks and slices con-
affected by high-P and high-T metamorphism stituted by ophiolites, metasediments and high-T
(Fig. 2). It is a monotonous mafic sequence consti- felsic orthogneisses and amphibolites wrapped
tuted by common and garnet-bearing amphibolites around by a low-T, highly sheared matrix of serpen-
showing an intense planolinear fabric and mineral tinites or phyllites. The upper main slice of the
associations typical of medium-pressure amphibo- mélange reaches 500 m in thickness and represents
lite facies. No relics of pre-metamorphic lithologies a typical ophiolitic mélange with minor metasedi-
have been identified so far but their textural aspect mentary rocks and few high-T tectonic blocks.
(often resembling that of flasergabbros) suggests a The ophiolitic lithologies of the Somozas Mélange
possible gabbroic protolith. The massive amphibo- are submarine metavolcanic rocks (pillow-lavas,
lite unit is bound at its top by garnet – staurolite – submarine breccias, pillow-breccias and hyaloclas-
kyanite medium pressure paragneisses that crop tites), diabases, gabbros, micro-gabbros, diorites,
out within the Carreiro shear zone (Fig. 3a). Detrital granitoids and highly serpentinized spinel-bearing
zircon age data indicate a maximum sedimentation ultramafic rocks (Fig. 4a). The serpentinites are
age of c. 470 Ma for the sedimentary precursor of the most abundant rock type in the ophiolitic
the paragneisses (Sánchez Martı́nez 2009). The mélange. The igneous rocks are always affected
nature of the contact between amphibolites and by a hydrothermal metamorphism that did not oblit-
paragneisses is uncertain although these gneisses erate the igneous textures but caused generalized
are different from those within the overlying units greenschist to amphibolite facies recrystallization.
affected by high-P and high-T metamorphism and, A sample taken from a hectometric tectonic
like the amphibolites themselves, cannot be con- block of granitoid was dated using U– Pb–LA –
sidered part of the upper units. ICP– MS geochronology (Arenas et al. 2007c)
U– Pb dating of zircons from these amphibolites giving an age of 527 + 5 Ma (Cambrian). Another
by LA –ICP–MS (Sánchez Martı́nez et al. 2006) sample of granite from a different hectometric
yielded a poorly constrained Mesoproterozoic crys- tectonic block that also includes gabbros and
tallization age of 1159 + 39 Ma and older zircons diorites yielded an age of 503 + 5 Ma age (Cam-
whose ages range from 1265– 1658 Ma that have brian) using U– Pb (SHRIMP–RG). A sample of
been interpreted as xenocrysts. The amphibolite conglomerates taken from a large tectonic block
sample studied also contained younger zircons included in serpentinites yielded several age
Fig. 3. Lithological constitution and geochemistry of the Purrido Ophiolite. (a) Characteristic section of the ophiolite.
(b) Th– Hf–Ta diagram (Wood 1980) with the projection of the most representative amphibolites. (c) Normal mid-
ocean-ridge basalt (N–MORB)-normalized trace-element pattern of the common amphibolites (average composition);
selected elements and normalizing values according to the criteria of Pearce (1996).
groups suggesting that sedimentation of this rock The ophiolite involved in the Somozas Mélange
occurred in the periphery of the West-African was the first to be studied with some detail in the
Craton (LA–ICP– MS; 59 concordant or sub- NW Iberian Massif. Arenas (1988) described its dis-
concordant detrital zircons). This conglomerate tribution, lithological constitution and geochemical
exhibits an age group with a large number of characteristics, suggesting that it could be con-
zircons (19 zircons) with ages between 630 and sidered as a typical mid-ocean ridge ophiolite (harz-
497 Ma, probably representing Pan-African events burgite ophiolite type, HOT). However, some of the
and also the activity in the volcanic arc where the studied ophiolitic lithologies showed geochemical
ophiolite was generated. The maximum age of sedi- features that are not entirely compatible with this
mentation of this conglomerate as inferred from the origin. New geochronological and geochemical data
two youngest concordant zircons is 465 + 5 Ma. of the igneous and sedimentary lithologies involved
This age can be considered as a reference age for in the Somozas Mélange allow this interpretation to
the end of the magmatic activity in the volcanic-arc be refined. According to their chemical character-
located in the periphery of Gondwana. istics two different groups of igneous rocks can be
Fig. 4. Lithological constitution and geochemistry of the Somozas Ophiolitic Mélange (a) Typical section of the
ophiolitic mélange. (b) Th– Hf–Ta diagram (Wood 1980) with the projection of the most representative lithologies of
the ophiolite; they include basalts, basaltic andesites, diabases and two main types of gabbros and diorites. (c) Normal
mid-ocean-ridge basalt (N– MORB) normalized trace-element patterns (average composition); selected elements
and normalizing values after Pearce (1996).
distinguished. A first group is formed by gabbros, between these igneous rocks is difficult since they
diorites, granitoids and basalts –basaltic andesites occur as large tectonic blocks in a mélange.
with calc-alkaline affinities and close similarity with However, the common basaltic rocks have never
igneous suites generated in volcanic arcs (Fig. 4b). been found intruded by calc-alkaline dykes or pluto-
The second group is constituted by gabbros, diabase nic bodies. Moreover, a key outcrop on the coast
dikes and common basaltic rocks with chemical shows calc-alkaline submarine volcanic rocks
compositions typical of island –arc tholeiites, and intruded by a network of basaltic dykes belonging
therefore also interpreted as generated in supra- to the second compositional group. Therefore, it is
subduction zone settings (Fig. 4b). The general possible to interpret that the island –arc tholeiites
patterns exhibited by selected trace elements nor- are younger than the calc-alkaline igneous rocks,
malized to the N –MORB composition are consist- and they were probably generated after a mature
ent with this interpretation (Fig. 4c). The volcanic arc located in the periphery of Gondwana,
recognition of the original intrusive relationships possibly at its most external margin.
Bazar Ophiolite and contain mineral associations indicative of low-

to intermediate-pressure conditions (plagioclase þ
The Bazar Ophiolite is located in the westernmost clinopyroxene þ orthopyroxene þ hornblende þ
exposures of the Ordenes Complex. It is constituted ilmenite + garnet + olivine). A different lithologi-
by several imbricate slices that contain gabbroic cal section exists at the base of the Carballo–Bazar
rocks, with minor ultramafic rocks at the base of slice; it is constituted by relatively well preserved
the unit (Dı́az Garcı́a 1990). The main tectonic gabbros, leucogabbros, tonalites and ultramafic
slice (the Carballo–Bazar slice) reaches a thickness rocks (Fig. 5a). This section also shows a lower-T
of 4000 m and is composed of amphibolites and metamorphism belonging to the low-T part of the
flasergabbros displaying a high-T foliation that amphibolite facies, and it is therefore interpreted
evolved from an initial granulite-facies stage. as a different slice.
Metre-scale boudin of well-preserved mafic granu- U –Pb –LA–ICP– MS dating has been per-
lites are found within the metagabbros (Fig. 5a). formed on a sample of amphibolite from the
These boudin are wrapped by the high-T foliation Carballo –Bazar slice (Sánchez Martı́nez 2009). A
Fig. 5. Lithological constitution and geochemistry of the Bazar Ophiolite (a) Typical section of the ophiolite; it shows
three main different slices although more cryptic imbrications probably do exist into the Carballo-Bazar slice. (b) Th–
Hf–Ta diagram (Wood 1980) with the projection of the most representative lithologies of the ophiolite; they include
amphibolites, mafic granulites and low-T metagabbros. (c) Normal mid-ocean-ridge basalt (N–MORB) normalized
trace-element patterns (average composition); selected elements and normalizing values after Pearce (1996).
population of 36 zircons gave ages mostly concor- consistent with a metabasaltic origin for the greens-
dant that allowed the definition of two clusters at chists. There are scarce intercalations of mylonitic
498 + 2 Ma (Cambrian) and 483 + 2 Ma (lower orthogneisses of tonalitic composition and metre-
Ordovician). These ages are interpreted as the crys- to decametre-scale lenses of medium- to fine-
tallization age of the mafic protolith and the age of grained metagabbros. The lower tectonic slice is
the high-T metamorphism, respectively. Granulites of ultramafic nature and is in contact with the rest
preserved as boudin also contain zircon, although of the unit by an extensional detachment (Fig. 6a).
it is less abundant than in the studied amphibolite The unit shows variable metamorphic features,
sample. Seven zircons separated from a granulite from low-grade rocks in the Vila de Cruces slice
boudin in the Castriz slice (Fig. 5a) were analysed. to medium grade in the Sampayo slice. In addition,
Five of these zircons yielded concordant to sub- a high-P metamorphic gradient has been identified
concordant ages between 458+8 and 489 + 6 Ma in the upper slice, where the schists (although very
and are considered to be related to the granulite retrograded) contain mineral associations with
metamorphic event. Two zircons yielded Neopro- garnet and without biotite.
terozoic ages around 620– 625 Ma. Although more Geochronological data suggest a complex scen-
data are needed to clarify the meaning of these ario that may involve a composite terrane rather
older ages, they are an indication that the protolith than a single ophiolitic unit. A tonalitic orthogneiss
ages of the granulite boudin may be significantly yielded a U –Pb age of 497 + 4 Ma (Arenas et al.
older than those of the host mafic rocks (Sánchez 2007a). Two gabbro lenses yielded overlapping
Martı́nez 2009). albeit imprecise U –Pb ages of 1176 + 85 and
The chemical composition of the common 1168 þ 14/250 Ma (Sánchez Martı́nez 2009).
amphibolites in the Bazar Ophiolite indicates an Unfortunately, the greenschists (volumetrically
oceanic affinity with N –MORB (i.e. basalts gener- dominant lithology of the unit), have not been
ated in typical mid-ocean ridges or in broad and dated yet, nor the maximum sedimentation age of
evolved back-arc basins; Fig. 5b, c), whereas the the clastic metasedimentary intercalations (zircons
mafic granulites are transitional between MOR and were not found in several studied samples of micas-
WP basalts with normalized trace element patterns chists). However, Sm–Nd data on the greenschists
comparable to those of T–MORB generated in are compatible with a Palaeozoic age for these
plume–ridge interactions (Pearce 1996). Metagab- rocks (Sánchez Martı́nez 2009). Even though it is
bros associated with ultramafic rocks from the evident that further isotopic and radiometric data
lower slice show a clearly different composition, are needed to clarify the origin and meaning of
characteristic of island –arc tholeiites (Fig. 5b, c). this unit, at this stage, it is possible to assume
Although, as expressed above, further chronological with confidence that Vila de Cruces likely rep-
and geochemical work is needed, the data obtained resents a composite unit in which at least a Mesopro-
so far suggest that the Bazar Ophiolite may represent terozoic element (coeval with that of the Purrido
a composite terrane with a complex evolution. The ophiolite, see above) and a Palaeozoic mostly
chronological evidence is consistent with the geo- mafic element are present. The regional foliation
chemical diversity within the mafic units. of the schists, acquired during the Variscan regional
deformation has been dated at 363–367 Ma
Vila de Cruces Ophiolite (40Ar/39Ar in two phyllite samples Dallmeyer
et al. 1997).
The Vila de Cruces Ophiolitic unit is located to the The geochemical features of the main lithologies
S– SE of the Ordenes Complex, lying under of the Vila de Cruces Unit are shown in Figure 6b, c.
the Careón Ophiolite and above the basal units of The mafic rocks, both the greenschists and the
the allochthonous complexes (Fig. 2). Together with Mesoproterozoic metagabbros, have compositions
the Moeche Ophiolite, it has been included in the compatible with those of island arc tholeiites, fol-
group of the lower ophiolites (Fig. 2; Arenas et al. lowing the discriminant diagram of Wood (1980)
2007b). This ophiolite is c. 4000 m in thickness (Fig. 6b). This signature is also highlighted by
and is a complex structure characterized by the their marked Nb-negative anomaly (Fig. 6c). The
existence of several tectonic slices, including the supra-subduction origin of these mafic rocks is
Sampayo slice with internal imbrications (Fig. 6a). also consistent with the volcanic arc signature
The unit is also characterised by the existence of of the tonalitic orthogneisses (Fig. 6b; Sánchez
two generations of often large recumbent folds Martı́nez 2009).
(Martı́nez Catalán et al. 2002; Arenas et al. 2007a).
The Vila de Cruces Ophiolite is dominated by Moeche Ophiolite
mylonitized greenschists, with intercalations of
metapelites and metacherts. These alternations of This unit has traditionally been correlated with
mafic rocks and metasedimentary rocks are the Vila de Cruces Unit, given the similar
Fig. 6. Lithological constitution and geochemistry of the Vila de Cruces Ophiolite (a) Typical section of the ophiolite,
with three main tectonic slices, which also show internal imbrication and recumbent folding. (b) Th– Hf–Ta diagram
(Wood 1980) with the projection of the most representative lithologies of the ophiolite; including common and
heterogeneous greenschists, metagabbros and orthogneisses. (c) Normal mid-ocean-ridge basalt (N –MORB)
normalized trace-element patterns (average composition) of the mafic lithologies; selected elements and normalizing
values after Pearce (1996).
lithological constitution of both units (Arenas et al. Sánchez Martı́nez (2009) has obtained U –Pb
2007b), although Vila the Cruces is considerably detrital zircon ages from micaschists that indicate
thicker than Moeche (Fig. 7a). The Moeche Unit maximum sedimentation ages between c. 469 and
is constituted by greenschists of possible metabasal- 455 Ma (Middle–Upper Ordovician). Older detrital
tic origin with abundant intercalations of phyllites zircon age clusters are compatible with a prove-
and micaschists, some metagabbro bodies and nance in the West African Craton and the Cadomian
some serpentinite intercalations towards the upper magmatic arc rocks, that is the original sediments
part of the unit. Its internal structure is poorly were deposited in the periphery of northern Gond-
known, and therefore the possible existence of wana. Although no radiometric dates are available
tectonic imbrications and/or recumbent folds for the metaigneous rocks of the unit, preliminary
cannot be ruled out (by analogy with the Vila de Sm– Nd data obtained on the mafic schists
Cruces Unit). (Sánchez Martı́nez 2009) are compatible with a
Fig. 7. Lithological constitution and geochemistry of the Moeche Ophiolite (a) Characteristic section of the ophiolite.
(b) Th– Hf–Ta diagram (Wood 1980) with the projection of the greenschists of this unit. (c) Normal mid-ocean-ridge
basalt (N– MORB) normalized trace-element patterns (average composition) of the greenschists; selected elements and
normalizing values after Pearce (1996).
Palaeozoic age. The regional foliation has been Careón Ophiolite

dated at 364 Ma (Dallmeyer et al. 1997, coeval
with that of the Vila de Cruces Unit). The Careón Ophiolite is exposed in the SE part of the
The main compositional features of the greens- Ordenes Complex and exhibits a varied lithological
chists of this unit are shown in the diagrams of sequence. It is made up of three main tectonic
Figure 7b, c. They have a composition transitional slices, of which the intermediate (Careón) slice is
between N –MORB and island-arc tholeiites. In the the thickest and contains the most varied lithological
Hf –Th– Ta diagram of Wood (1980) they plot in ensemble with ultramafic rocks at the base and a
the area straddling both compositional fields sequence of isotropic gabbros towards the top
(Fig. 7b). They show a slight enrichment in trace ele- (Fig. 8a). Both the ultramafic rocks and the gabbros
ments with respect to N –MORB (Fig. 7c), with a are intruded by a swarm of diabase dykes and pegma-
slight Nb–negative anomaly suggesting a supra- toid gabbros. This ultramafic – mafic rocks assem-
subduction environment. Therefore, the mafic sch- blage with no record of volcanic or sedimentary
ists of the Moeche Unit are compositionally different rocks is remarkably different from that of typical
from their putative analogues in the Vila de Cruces mid-ocean ridge ophiolites. Dı́az Garcı́a et al.
Unit, given that they show an attenuated arc signature (1999) interpreted the Careón Unit as an ophiolite
and a drift towards N –MORB compositions. generated in a supra-subduction zone (SSZ) setting.
Fig. 8. Lithological constitution and geochemistry of the Careón Ophiolite. (a) Characteristic section of the ophiolite.
(b) Th–Hf –Ta diagram (Wood 1980) with the projection of two types of metagabbros of the Careón slice, the diabase
dykes and the amphibolites of the Orosa slice. (c) Normal mid-ocean-ridge basalt (N– MORB) normalized
trace-element patterns (average composition); selected elements and normalizing values after Pearce (1996).
At the contact zones between the different tectonic yielded a crystallization age of 395 + 2 Ma
slices, there is evidence of thermal effects, even the (middle Devonian; Dı́az Garcı́a et al. 1999), while
development of metamorphic soles with high-T another sample of leucogabbro from the same
recrystallization. This is particularly well developed slice yielded an identical age of 395 + 3 Ma (Pin
and preserved at the contact between the Careón et al. 2002). These ages indicate that this ophiolite
and Vilouriz slices, where a 2 m-thick level with records the final stages of the Rheic Ocean when it
corundum crystals up to 2 cm was developed. was well into its contractional phase (e.g. Sánchez
These features of the contacts are interpreted as an Martı́nez et al. 2007a). On the other hand
40
indication that the tectonic slices were stacked Ar/39Ar dating of the amphiboles defining the
when the whole section of oceanic lithosphere was amphibolite facies foliation yielded an age of
still very hot, that is shortly after its generation. 376.8 + 0.4 Ma (Upper Devonian; Dallmeyer et al.
U –Pb dating of zircons from a sample of leuco- 1997) that is interpreted as indicative of the chronol-
gabbro from the upper part of the Careón slice ogy of the accretion of the ophiolite slices to the
southern margin of Laurussia. It must be noted that the equator and latitude 408S (Fig. 9a; Kröner &
the accretion ages of 363 –367 Ma obtained for the Cordani 2003). Data provided here are consistent
Vila de Cruces and Moeche units fit the structural with the idea that the Purrido Ophiolite could have
position of the latter under the Careón Unit, indicat- been generated in relation to an active arc system
ing a later (younger) accretion. in the eastern margin of Amazonia-West Africa
Figure 8b, c show the geochemical features of (Fig. 9a).
the Careón Ophiolite. All the lithologies have com- An enigmatic aspect in the evolution of the
positions typical of island-arc tholeiites. The meta- Purrido Ophiolite is related to its preservation, as
gabbros show pronounced Nb-negative anomalies, its mafic character and probable generation in an
and slightly more transitional compositions towards oceanic setting should have made its survival diffi-
N –MORB in the diabase dykes and the amphibolites cult. However, part of the ophiolite has been pre-
from the Orosa slice. In the case of the Careón slice served to the present day, and it is involved in the
(where the oceanic lithosphere section is more com- Variscan suture of NW Iberia without apparent con-
plete) the transition from island-arc compositions in nection with other Mesoproterozoic rocks of conti-
the gabbros towards N –MORB in the diabases is nental nature. In principle there are two main
interpreted as the geochemical signature of the litho- hypotheses that may account for the preservation
spheric extension above a subduction zone. of the ophiolite: a fast obduction over the continen-
tal margin of the West African Craton, or remaining
stable far away from the subduction zone in some
Discussion: origin of the ophiolites kind of marginal basin located between a volcanic
arc system and the West African Craton margin
of NW Iberia itself. It is to be hoped that future investigations
A Mesoproterozoic ophiolite in NW Iberia focused on the study of the structural history of
the Purrido Ophiolite will help to clarify this issue.
According to the data presented in previous works, It can be expected that the structural history of the
the gabbroic protoliths of the Purrido Ophiolite Purrido amphibolites is not as simple as it may
have yielded a crystallization age of 1159 + 39 Ma. seem, and their evolution could have been polyde-
Taking into account that this is a mafic unit built formational. These amphibolites show a rather
up by 300 m of massive amphibolites, this assem- simple amphibolite-facies plano-linear fabric that
blage can be interpreted as a pre-Rodinian ophiolite. is inconsistent with the complete absence of
This group of ophiolites older than 1100 –1000 Ma, igneous relics within the unit. There is no other
the estimated assembly age for the Rodinia super- metaplutonic unit in Galicia with the same thick-
continent (Dalziel 1991; Hoffman 1991), is rela- ness, where igneous bodies with primary textures
tively small, as only around 35 cases have been and/or mineralogy have not been preserved. This
reported so far (Moores 2002). Considering the geois recorded in units affected even by granulite-facies
chemical affinity of its mafic rocks, the Purrido high-temperature metamorphism, and it is to be
Ophiolite can be assigned to the supra-subduction expected to be more common when rocks only
zone type, suggesting that it was generated in relation reach medium-grade metamorphic conditions, as is
to the activity of a volcanic arc, probably in a the case with the Purrido Ophiolite. Thus, the appar-
back-arc setting or a marginal basin. This tectonic ent simplicity of the deformational history of the
setting is also compatible with the presence of inher- Purrido Ophiolite may not be real, and we cannot
ited zircons with ages ranging between 1658 and rule out its obduction over the leading margin of a
1265 Ma. Thus, this unit would represent a fragment continent in motion soon after its generation in a
of the plutonic section of a SSZ ophiolite. marginal basin. On the other hand, we can neither
The Purrido Ophiolite does not show evidence of discount the possibility that it was preserved
having been affected by Mesoproterozoic high- roughly intact within the back-arc. The alternative
temperature metamorphic events. It is not related hypothesis, of the obduction of the ophiolite
to other rocks with continental-affinity of the same during the continental dispersal that followed the
age either. Therefore it is unlikely that the Purrido Rodinia fragmentation, after a long stability period
Unit was involved in the Grenville orogenic belt in a marginal basin (Fig. 9b), cannot be completely
(developed between 1300 and 1000 Ma), and it can ruled out before having more information about the
be thus interpreted as a section of peri-arc oceanic tectonothermal history of the unit. It can be pointed
lithosphere located away from the realm of this out that the fragmentation of Rodinia started at
mobile belt. Considering the currently accepted 750 Ma (e.g. Torsvik 2003) and that the dated
palaeogeographic reconstructions of the continents Purrido amphibolite sample had a single zircon
immediately before the assembly of Rodinia, it that yielded a concordant age of 816 + 15 Ma.
seems that at c. 1100 Ma Amazonia and West Africa The presence of Palaeo-Proterozoic or even
defined a single continental domain located between Archaean continental basement in the Variscan
Fig. 9. (a) Distribution of the main continental blocks at 1100 Ma, during the stages previous to the development of the
Grenvillian orogen and the resulting assembly of the Rodinia supercontinent (Kröner & Cordani 2003; based in
Pisarevski et al. 2003). The frame indicates the most probable region considered for the generation of the mafic
rocks constituting the Purrido Ophiolite. Abbreviations: Am, Amazonia; Au, Australia; Ba, Baltica; C, Congo; Gr,
Greenland; In, India; Ka, Kalahari; La, Laurentia; Ma, Mawson craton; W, West Africa; P, Pampean terrane; R, Rockall;
RP, Rı́o de la Plata; SC, South China; SF, São Francisco; Si, Siberia; T, Tarim. (b) Probable location of the Purrido
Ophiolite (framed region) in a reconstruction of the Rodinia supercontinent prior to its fragmentation (c. 750 Ma).
Rodinia reconstruction by Torsvik (2003).
Belt of western Europe was described a long time Palaeozoic, possibly without being affected by sig-
ago. This is the case of the Icart gneisses (located nificant deformation.
on the Channel Islands and Brittany; Samson & The Purrido Ophiolite is in contact with detrital
D’Lemos 1998) and the submerged granulites of metasedimentary rocks from the Carreiro shear
the Bay of Biscay (Guerrot et al. 1989), which in zone, which have a maximum sedimentation age
both cases are interpreted to be lithological assem- of 470 + 7 Ma (Lower–Middle Ordovician bound-
blages related to the West African Craton. Consider- ary) and were likely deposited in the periphery of
ing the information provided in this work, it is the West African Craton as suggested by detrital
clear that the basement of the Variscan Belt also zircon U –Pb age clusters (Sánchez Martı́nez
includes – at least – Mesoproterozoic mafic sec- 2009). It is not possible to determine if these sedi-
tions with oceanic affinity, as the Purrido Ophiolite, ments were initially deposited over the mafic
which probably obducted over a continental margin, rocks, although the geological link existing
or remained stable next to it during the Rodinia between both lithological assemblages and their
assembly. We hypothesize that during the stages similar tectonothermal evolution could be inter-
immediately preceding the fragmentation and dis- preted in this way. If this interpretation is correct,
persal of Rodinia, the Purrido Ophiolite was located and taking into account the data provided by the
next to the eastern margin of the West African other ophiolitic units of Galicia that will be dis-
Craton, in a position equivalent to that indicated in cussed below (see also Scotese 2001; Stampfli &
Figure 9b (Sánchez Martı́nez et al. 2006). In other Borel 2002), the expected scenario at the early
words, we favour the simplest possible interpret- Palaeozoic in the northern Gondwanan margin
ation and we situate this Mesoproterozoic terrane would be determined by the development of mar-
in a position analogous to that currently occupied ginal basins behind an active volcanic arc. This vol-
by its remnants. canic arc has been traditionally situated at the
Reviewing the palaeographic reconstruction for inception of its development at the external
the times following the Rodinia fragmentation, margin of Gondwana (see Arenas et al. 2007a, b),
some of the most outstanding events are connected although an origin from an intra-oceanic subduction
with the amalgamation of continents to constitute zone in the proximity of Gondwana cannot be ruled
the present central and southern Africa. These out. The drift of the arc and its retreat from Gond-
different collisions are Neoproterozoic in age and wana, together with the rifting of other microconti-
are usually grouped under the term Pan-African nents from the main continent, as was the case of
orogeny. Dalziel (1997) suggested the existence of Avalonia (Stampfli & Borel 2002; Murphy et al.
a supercontinent of ephemeral life during the late 2006), led to the development of a new major
Neoproterozoic, that he named Pannotia (see also oceanic domain, the Rheic Ocean, that determined
Murphy & Nance 2004). During its assembly, the to a large extent the Palaeozoic paleogeography of
eastern margin of Amazonia –West Africa remained the domain presently occupied by western Europe
essentially stable and without significant orogenic and eastern North America. This situation corre-
activity, although it seems clear that the develop- sponding to the period between the Cambro–
ment of an important magmatic arc, the so-called Ordovician boundary and the middle Ordovician is
Cadomian–Avalonian Arc, took place along its shown in Figure 10c, d.
northern margin (Figs 9 & 10a). Pannotia broke up Figure 11 represents an interpretive model for
into four large continents, Laurentia, Baltica, the development of a marginal basin system that
Siberia and Gondwana at 550 Ma ago, but this par- occurred at the early stages of the opening of the
tition happened without the margin of Amazonia – Rheic Ocean. This sketch does not reflect a specific
West Africa being significantly affected moment, but a period of time characterized by the
(Fig. 10b). In other words, the data available seem development of the volcanic arc and the beginning
to indicate that this large continental margin of its rift from Gondwana, ranging from c. 500–
behaved coherently and was rather stable, at least 480 Ma. One of the most uncertain aspects of this
over a period between 1100 Ma (Fig. 9a) and model is the width of the marginal basin, which is
550 –520 Ma (Fig. 10b). A little modified large con- at present impossible to constrain. However, it is
tinental margin, the Amazonia –West African possible that it was not too wide at this age
margin, may have participated in the amalgamation (Fig. 10), especially taking into account that the
and fragmentation of two supercontinents (Rodinia Mesoproterozoic terrane was probably covered by
and Pannotia), to finally constitute most of the sediments at c. 470 Ma, just postdating the
northern margin of Gondwana at c. 550 Ma. Thus moment represented in this diagram, and whose
there are reasons to surmise that, on the basis of probable source area was the West African Craton
the apparent stability of the Amazonia –West or sediments derived from it. Accordingly, it is con-
African margin, the Purrido Ophiolite could have ceivable that the back-arc was covered by
remained next to this margin until the early sediments.
Fig. 10. Schematic evolution of continents and oceanic domains between the late Proterozoic (c. 600 Ma) and the
Middle Ordovician (c. 465 Ma) (a) Reconstruction of the Pannotia supercontinent; the frame shows the most probable
region for the location of the NW Iberia Proterozoic terranes. (b) Fragmentation of Pannotia in four main continents
(Gondwana, Laurentia, Baltica and Siberia) during the early Cambrian. (c) Reconstruction of the continental blocks
during the stages preceding the opening of the Rheic Ocean (c. 490 Ma). It shows the beginning of the generation of the
marginal basin system that gave rise to this ocean. The Rheic Ocean opened in the Northern margin of Gondwana
as a consequence of the separation and drift of Avalonia and other peri-Gondwanan terranes. The section of the
arc-system probably originating the upper units of the allochthonous complexes of Galicia appears coloured in red.
(d) The Rheic Ocean domain prior to the closure of the Iapetus and Tornquist oceans and the final assembly of Laurussia.
Modified and adapted after Arenas et al. (2007a), Gómez Barreiro (2007) and Martı́nez Catalán et al. (2007a); based on
the palaeogeographic reconstruction by Winchester et al. (2002).
Fig. 11. Diagram illustrating the general model suggested in this paper for the beginning of the opening of the
Rheic Ocean. The figure shows the position and most probable meaning of four of the ophiolitic units of Galicia, the
Purrido, Bazar, Vila de Cruces and Moeche units. Explanation in the text.
In Figure 11, we have considered that the Rheic the north of the mid-Rheic ridge). However, data
Ocean opened in a region where the Mesoprotero- relating to the tectonothermal evolution of the
zoic terrane represented by the Purrido Ophiolite Purrido Ophiolite and the metasedimentary rocks
was already integrated, and that this opening split involved in the Carreiro shear zone seem to indicate
this terrane, so that one part of it was separated that the second interpretation is more plausible, and
from Gondwana becoming part of the basement of therefore it is depicted in Figure 11. The sample of
the drifting volcanic arc, and another part remained metasedimentary rock from the Carreiro shear
attached to the continental margin, thus retaining its zone contains rutile dated by LA-ICP-MS, which
original position. The clues for this tentative yielded U–Pb ages of 412 +19 and 428 + 11 Ma;
interpretation lie, on the one hand, in the structural moreover, the sample of Purrido amphibolite
position occupied within the Variscan suture by also dated by LA-ICP-MS contains a zircon of
the units with Mesoproterozoic remnants, and on c. 428 + 5 Ma. Regardless of the existence of
the other hand, in the tectonothermal evolution other younger zircons in both lithologies (Variscan
recorded by the terrane itself, arguments that will reworking), these U – Pb data, though scarce,
be addressed below. suggest that both ensembles underwent a meta-
In principle, the most straightforward way to morphic event during the Middle– Late Silurian
know the concrete position of a terrane of the oro- that happened when the Rheic Ocean was still at
genic wedge in relation to the suture, that is to the the height of its spreading and probably at the
Rheic Ocean, is to determine its structural relation- time when its width was maximum. This meta-
ship with the ophiolites derived from the true and morphic event clearly links the tectonothermal evol-
characteristic Rheic domain, assuming that the ution of the Purrido Ophiolite and metasedimentary
polarity of the accretion in the orogenic wedge is rock located on top, with that of the upper units of
known. The only ophiolite in Galicia that can be the allochthonous complexes of the NW Iberian
assigned to an evolved Rheic Ocean is that rep- Massif. More precisely, the ages are analogous to
resented by the Careón Unit, since the other ophio- those obtained in the basal slice of the upper units,
litic units were developed in early stages of the affected by high-P and high-T metamorphism, that
history of this ocean or they are unrelated to it. have been interpreted in connection with the accre-
Taking into account that the Purrido and Careón tion of an arc to the southern margin of Laurussia
units are not in contact with each other, it is imposs- during the stages preceding the onset of the contrac-
ible to know if the Purrido Unit remained attached to tion of the Rheic Ocean (Gómez-Barreiro et al.
Gondwana during the opening of the Rheic Ocean or 2006; Sánchez Martı́nez et al. 2007a; Arenas et al.
if it drifted towards Laurussia (i.e. it was situated to 2007b; Fernández-Suárez et al. 2007). Accordingly,
the data presented in this paper seem to indicate that the Gondwanan edge, the birth date of this ocean
a part of the Mesoproterozoic terrane, represented can be constrained to a relatively narrow period.
by the Purrido Unit, should be part of the basement This is the time period ranging from the peak
of the volcanic arc that split up from Gondwana at activity in a mature volcanic arc (c. 527–503 Ma)
the Cambrian –Ordovician boundary. Together with to the development of a basin which includes detri-
this arc, it drifted to the north, until its final accretion tus derived from rocks representing the final mag-
to the southern margin of Laurussia in the Silurian, matic activity in the arc (c. 465 + 5 Ma; Arenas
the time when the Rheic Ocean probably started et al. 2007c).
to contract.
Remnants of the Cambrian peri-arc ocean?
Remnants of a Cambrian mature The results obtained in the Bazar Ophiolite have
peri-Gondwanan volcanic arc been quite surprising because of their complexity
(Sánchez Martı́nez 2009; Sánchez Martı́nez et al.
In the Somozas Ophiolitic Mélange, the ophiolitic 2007b). When it comes to analysing the significance
lithologies appear imbricated with continental of the new data, it has to be considered that the
rocks probably related to the palaeomargin of Gond- Bazar Unit has traditionally been correlated with
wana (basal units of the allochthonous complexes the Careón Unit, which was generated during
and Schistose Domain). These ophiolites are those middle Devonian times (Arenas et al. 2007b) and
occupying the lowest structural position in NW considered a typical representative of the litho-
Iberia (Arenas et al. 2007b), as they appear below sphere of the Rheic Ocean. The new U– Pb geochro-
other ophiolitic units developed in the southern nological data are preliminary in some of the
domain of the Rheic Ocean. Based on its geological lithologies belonging to the Bazar Unit. However,
context, the original paleogeographic position of the the new U –Pb data for the mafic rocks seem to indi-
ophiolites involved in the Somozas Mélange was cate that the Bazar Unit represents a composite
likely located in the oceanic domain attached to mafic unit, constituted by materials of Neoprotero-
the Gondwana margin or even in the more external zoic (c. 624 Ma) to late Cambrian (c. 498 Ma) age.
margin of this supercontinent, depending on the This diversity of ages is also consistent with the
meaning of this ophiolite. According to the compo- compositional variability between the different
sitional features of the rocks that constitute this lithological types (from N –MORB to island-arc
ophiolitic mélange, we suggest an origin in a peri- tholeiites generated at supra-subduction zone set-
Gondwanan volcanic arc. As discussed in the pre- tings), as well as with their different metamorphic
vious section, this arc had a long period of activity evolution.
and reached a mature stage, which in the Somozas The most common rock type in the Bazar Unit is
mélange is manifest in the presence of granitoid high-temperature amphibolite, which contains a
rocks. Based on the U –Pb zircon ages obtained large amount of zircon with U –Pb ages grouped
from these granitoids within the mélange, the arc into two populations. The first, with a mean age of
was probably active during most of the Cambrian 498 + 2 Ma (Upper Cambrian), is interpreted to
period (c. 530– 503 Ma; Arenas et al. 2007c). reflect the crystallization age of the gabbroic proto-
The metaigneous lithologies of the Somozas liths, while the second, with a mean age of
Ophiolitic Mélange belong to two compositional 483 + 2 Ma (Lower Ordovician), has been inter-
groups linked by their relationship with a volcanic preted to date the granulitic metamorphism under-
arc. The older group has calc-alkaline compositions gone by the unit. If we consider that the Bazar
whereas the younger igneous rocks display compo- amphibolites have a chemical composition typical
sitional features of island-arc tholeiites. The mutual of N– MORB, and that hence they were probably
relationship between both groups can be established generated in a mid-ocean ridge, it is necessary to
by the presence of a network of diabases belonging consider a scenario where the common oceanic
to the second group intruding the calc-alkaline sub- lithosphere was accreted and underwent high-
marine volcanics. Furthermore, the evolution from a temperature metamorphism. Considering the
calc-alkaline magmatism typical of mature arcs to a nature of metamorphism, the most probable tectonic
younger one which shows affinities with island-arc setting would be an active and maybe relatively
tholeiites, can be related to the opening of an juvenile volcanic arc, below which the accretion
intra-arc basin. The progressive widening of this of the oceanic lithosphere that fed the magmatism
basin could explain the rifting of the external part of this arc (see e.g. the work of Peacock 1990)
of the arc, and its subsequent drift leaving behind took place. Considering the palaeogeographic
a new oceanic domain, the Rheic Ocean. According models accepted for the peri-Gondwanan realm
to the materials involved in the Somozas Mélange, during Cambrian and Ordovician times, as well as
which were originally located in the periphery of the data obtained in the other units of the
allochthonous complexes of the NW Iberian Massif, lenticular inclusions of metric to decametric size
it is likely that the Bazar amphibolites derive from within the strongly deformed greenschists that rep-
the oceanic lithosphere exterior to the peri- resent the predominant lithology in this unit. Unfor-
Gondwanan arc system (Fig. 11). This lithosphere tunately, the age of these greenschists is not well
was accreted under this arc system at a subduction constrained, hindering their interpretation.
zone dipping towards the arc and Gondwana, and However, available Sm–Nd data indicate that at
that would be the cause generating the arcs them- least part of these greenschists could be Palaeozoic.
selves, the peri-Gondwanan marginal basin system Meanwhile, their geochemical characteristics are
and the supra-subduction geochemical signatures homogeneous and indicative of an origin in a supra-
identified in the Palaeozoic metabasites of Vila de subduction zone setting. The Vila de Cruces Unit
Cruces and Moeche units. In other words, the also contains orthogneisses dated at 497 + 4 Ma
Bazar amphibolites probably represent the crust of (Cambrian) with typical volcanic arc geochemical
the Iapetus-Tornquist Ocean, being consumed and features, as well as intercalations of metasedimen-
reducing its width while it was giving ground to tary rocks (in which no zircon was found), but con-
the new ocean floor that was generated at the spread- sidering their lithological resemblance and
ing Rheic domain (Fig. 11). The structural position analogous tectonothermal evolution, it is reasonable
of the Bazar amphibolites in NW Iberian Massif to link them with those present in the Moeche Unit
reflects the described dynamic scenario, since they (that will be discussed below), that yield Palaeozoic
are located below the arc-derived terrane (the maximum sedimentation ages and were deposited
upper units) and on the side that should be facing around the periphery of the West African Craton.
the exterior ocean (the western part of these units). In short, the Vila de Cruces Unit represents a
The meaning of the Neoproterozoic rocks is composite terrane in which a Mesoproterozoic
more uncertain. However, if confirmed with mafic basement similar to that forming the Purrido
further data, their presence would add evidence for Ophiolite was involved, and new Palaeozoic
the protracted and complex evolution of the Proter- crust possibly developed around the Cambrian –
ozoic margins of Amazonia –West Africa, as Ordovician boundary, as the result of oceanic mag-
already highlighted by previous studies (e.g. matism associated to a supra-subduction zone. The
Sánchez Martı́nez et al. 2006). Mesoproterozoic mafic basement should be inten-
sely intruded by magmas generated in the supra-
subduction domain during the development of this
Two different elements of an evolving new Palaeozoic crust, until being reduced to a few
back-arc basin remnants within the Palaeozoic rocks. These rem-
nants would be represented by the metagabbros
The model favoured here considers that the Vila dated at c. 1170 Ma. Therefore, the Vila de Cruces
de Cruces Unit remained attached to the margin Unit contains the part of the Mesoproterozoic base-
of Gondwana until the formation of Pangaea ment that remained attached to Gondwana when the
(Fig. 11). This interpretation is based on the struc- Rheic Ocean opened (Fig. 11). It has to be high-
tural position occupied by the unit within the Varis- lighted that according to the data provided in this
can orogenic wedge, where it appears below the work it seems clear that this ocean opened in the
ophiolites of the Rheic Ocean represented by the proximity of an ancient suture, and possibly along
Careón Unit (Fig. 2). Moreover, evidence for it. This possibility was already proposed by
the metamorphic event developed at 410– 430 Ma Murphy et al. (2006), albeit these authors developed
has not been found in this unit, which further sup- this idea to suggest that Avalonia was initially
ports the idea that it did not accrete to the southern accreted to the margin of Amazonia–West African
margin of Laurussia, as the Purrido Unit, and that Craton, generating the suture from which the sub-
its Palaeozoic tectonothermal evolution is exclu- sequent separation and drifting of Avalonia from
sively Variscan and related to the main stages of Gondwana, opening the Rheic Ocean (Fig. 10),
the development of Pangaea. took place. Also, according to our data, it is possible
As discussed in previous sections, the Vila de that the opening of this ocean took place in a region
Cruces Unit shows a great complexity, since it is from the northern margin of Gondwana character-
constituted by lithologies very different in age. ized by the existence of more than one Proterozoic
Thus, it represents a composite terrane formed suture. It is pertinent to indicate here that different
during a long period of time. This unit contains rem- data suggest the possible existence of a Mesoproter-
nants of a Mesoproterozoic terrane which is similar ozoic basement in the lower crust of NW Iberia.
to (or the same as) that represented by the Purrido Mesoproterozoic ages are common in detrital
Ophiolite: two metagabbros of the Vila de Cruces zircons and micas from the core of the Ibero–
Unit have been dated at c. 1170 Ma (Sánchez Armorican arc (Gutiérrez-Alonso et al. 2005, and
Martı́nez 2009). These metagabbros appear as references therein). Moreover, most of the
Ordovician volcanic rocks from NW Iberia yield present. Taking into account their geochemical fea-
Mesoproterozoic Nd model ages (Murphy et al. tures, the greenschists of this unit appear as basaltic
2008). The nature and extent of such a basement types similar to N –MORB but with compositional
are unknown, but it is conceivable that some corre- features indicating that their generation took place
lation with the Mesoproterozoic rocks of the Vila de close to a subduction zone. As a result, we interpret
Cruces Unit may exist. In this way, this unit can that the origin of the Moeche Unit should be located
include the thinned Mesoproterozoic basement at a more outboard position within the peri-
located in the transitional crust of the more external Gondwanan marginal basin system (Fig. 11). In
margin of Gondwana (Fig. 11). This basement, as other words, this unit can be considered as transi-
part of the more complex Vila de Cruces Unit, tional between the Vila de Cruces Unit and the
would have been thrust over the continental typical lithosphere of the Rheic Ocean. It must be
margin (basal units of the NW Iberia allochthonous clarified that this kind of oceanic lithosphere that
complexes) during the early stages of the Variscan would be characteristic of the Rheic realm has
collision. neither been found in the allochthonous complexes
The Moeche Unit is made up of strongly of Galicia nor is it likely to have been generated
deformed greenschists alternating with phyllite during the Middle Ordovician. In any case, it
levels. It thus bears a close lithological similarity should be constituted by typical N –MORB basalts
to the Vila de Cruces Unit, with which it is usually lacking alternating levels of detrital rocks.
correlated (Arenas et al. 2007b). The correlation
between these units is also supported by an analo-
gous tectonothermal evolution, with identical A Middle Devonian supra-subduction
40
Ar/39Ar ages for the main regional foliation ophiolite from a contractive Rheic Ocean
(Dallmeyer et al. 1997) and the absence of the meta-
morphic event dated around 410– 430 Ma in the The age of the Careón Ophiolite has been known
Purrido Ophiolite. However, there are no Palaeozoic since a U –Pb dating study by TIMS revealed that
orthogneisses in the Moeche Unit like those forming the metagabbros of this unit were generated
part of the Vila de Cruces Unit, nor metagabbros of 395 + 2 Ma ago, that is, during the Middle Devo-
Mesoproterozoic age like those appearing in that nian (Dı́az Garcı́a et al. 1999). If the Careón Ophio-
unit. Nevertheless, taking into account the limited lite is the sole section of lithosphere of the Rheic
extent of the outcrops of these metagabbros and Ocean preserved in the Variscan suture of NW
our limited sampling, we cannot rule out that not- Iberian Massif, except for the infant stages of the
yet-found remnants of this Mesoproterozoic mafic generation of the peri-Gondwanan marginal basins,
terrane are present in the Moeche Unit. it is too young to be representative of the typical
As in the case of the Vila de Cruces Unit, there is lithosphere of this ocean. Moreover, the lithological
an uncertainty about the age of the mylonitic greens- constitution of the Careón Ophiolite cannot be con-
chists, that represent the main lithological type of sidered characteristic of the most common oceanic
this terrane, to interpret the meaning of the Moeche lithospheres. This led Dı́az Garcı́a et al. (1999) to
Unit. These mafic greenschists appear interbedded interpret this ophiolite as generated in a supra-
with levels of phyllite whose detrital zircons indicate subduction zone setting. This interpretation was
maximum sedimentation ages ranging between 470 also favoured by Pin et al. (2002) who studied the
and 455 Ma (Middle–Upper Ordovician), and a pro- chemical composition of the most representative
venance from source areas located in the proximity metabasites of the Careón Unit. This idea is also con-
of the West African Craton. In other words, these sistent with the new data based on the study and geo-
phyllites represent detrital sediments probably chemical interpretation of a group of metabasite
deposited during middle Ordovician times. Regard- samples (Sánchez Martı́nez et al. 2007a).
ing the mafic greenschists, their textural features According to the palaeogeographic models
indicate that they probably derive from basic volca- accepted for the Middle Ordovician, the width of
nic rocks, and their distribution in the field alternat- the Rheic Ocean should increase quickly as Avalo-
ing with levels of Ordovician sediments suggests nia and other related peri-Gondwanan terranes
similar ages. The Sm –Nd isotopic composition of (as that represented by the upper units of the
these greenschists is also more compatible with a allochthonous complexes of Galicia) drifted to the
Palaeozoic age, although our data are very scarce north away from Gondwana (Fig. 10c, d). This
and should be completed in the future. widening of the Rheic Ocean was taking place
Based on the above, the data presented in this while the Iapetus and Tornquist oceans were under-
paper allow us to interpret the Moeche Unit as an going a fast contraction. In general, it is accepted
ensemble of rocks mostly of Palaeozoic age, that the Rheic Ocean started to close during
although it cannot be completely ruled out that rem- Ludlow times (c. 420 Ma), right after the accretion
nants of the Mesoproterozoic mafic terrane may be of Avalonia to Laurussia and the resulting closure
of the Iapetus Ocean (Stampfli & Borel 2002). At other equivalent allochthonous terranes in the
the same time, it seems that another peri- French Massif Central and in the Bohemian
Gondwanan terrane derived from a volcanic arc Massif (see Sánchez Martı́nez et al. 2007a). It is
was accreted to the southern margin of Baltica– characterized by the presence of a high-pressure
Avalonia (Gómez-Barreiro et al. 2007). That high-temperature metamorphic event dated by U –
terrane constitutes the upper units of the allochtho- Pb and 40Ar/39Ar at 400–425 Ma (Gómez-Barreiro
nous complexes of NW Iberian Massif, but also et al. 2006; Fernández-Suárez et al. 2007), probably
Fig. 12. Palaeographic reconstruction of the Rheic Ocean realm at the Silurian– Devonian boundary (modified after
Sánchez Martı́nez et al. 2007a). It shows the generation of new oceanic lithosphere associated with intra-oceanic
subduction directed to the NE. It also shows the situation in the southern margin of Laurussia of the accreted arc from
which the upper units of the allochthonous complexes of NW Iberia are considered to derive. The general position of the
main continents is a modification from the original palaeogeographical reconstruction by Stampfli & Borel (2002).
recording its accretion to the southern margin of upper units of the allochthonous complexes of NW
Laurussia. As discussed in previous sections, the Iberia, is there evidence for large late Silurian or
detection of this metamorphic event in the Purrido Devonian volcanic arcs developed from subduction
Ophiolite and in the paragneisses of the Carreiro to the north. The same general absence of volcanic
shear zone allowed us to interpret that both litho- arcs of this age is typical for the terranes located
logical ensembles moved together, being part of in the northern margin of Gondwana, which were
this terrane, therefore including remnants of a finally involved in the Variscan orogeny in Europe
mafic Mesoproterozoic terrane. (Fig. 12). However, our data require significant gen-
It is a matter of controversy whether a single (the eration of oceanic lithosphere in the early to middle
Rheic Ocean; Linnemann et al. 2004; Murphy et al. Devonian, while the Rheic Ocean was evidently
2006) or multiple oceanic domains (Franke 2000; contracting.
Winchester et al. 2002) existed to the south of Ava- Considering the lithological section of the
lonia in the Silurian and Devonian. It seems clear Careón Ophiolite, its suprasubduction zone geo-
that the southern continental margin originally bor- chemical affinity, and the general characteristics of
dering the Rheic Ocean is presently represented by the European Variscan belt, we suggest that the
the Saxo– Thuringia and Ossa –Morena zones of Rheic Ocean was closed mainly by intraoceanic
the European Variscan belt. Important arc-related subduction directed to the north (Fig. 12; Sánchez
magmatism c. 360 –335 Ma has been studied in Martı́nez et al. 2007a). This subduction was prob-
the Mid-German Crystalline Rise (Saxo – Thuringia ably located near the northern margin of the
Zone), where it has been attributed to subduction ocean, and its development involved consumption
toward the south (Altherr et al. 1999). The same of old and cold N –MORB-type oceanic lithosphere
age and tectonic setting have been suggested for and the generation of limited volumes of new
the Late Devonian –Dinantian volcanism described oceanic lithosphere of suprasubduction zone type
in the French Massif Central (Pin & Paquette (Fig. 12). This interpretation is compatible with
2002). This subduction directed to the south and the rarity or virtual absence of common MOR–
the associated magmatism are younger than the type ophiolites, like those associated with divergent
first deformation and coeval high-P metamorphism tectonic settings (Boudier & Nicolas 1985) in the
affecting the most external margin of Gondwana European Variscan belt (see Sánchez Martı́nez
(dated at 365 –370 Ma in NW Iberia; Rodrı́guez et al. 2007a; Arenas et al. 2007a). The model also
et al. 2003). Therefore, they mainly occurred after explains the scarcity of older (pre-Silurian) ophio-
the closure of the Rheic Ocean and have been inter- lites that could be related to early stages of the evol-
preted in relation to the opening and later closure of ution of the Rheic Ocean. In this respect, it is worth
a foredeep basin (Martı́nez Catalán et al. 1997). The mentioning that the old ophiolitic units preserved in
opening and closure of this basin may explain the Galicia, that is those of Ordovician or older age, rep-
double vergence of the European Variscan Belt resent an outstanding example on a continental
(Matte 1991). The south-facing part of the belt scale, and allow the history of the oceanic domains
shows the oldest tectonothermal evolution and pre- involved in the evolution of the western European
serves information about the closure of the Rheic basement to be traced.
Ocean, which would have been coeval with north-
directed subduction (Matte 1991; Martı́nez This research has been funded by projects CGL2004-
Catalán et al. 1997, 2007b). On the other hand, the 04306-CO2-02/BTE and CGL2007-65338-CO2-01/BTE
north-facing part of the belt is younger, and its of the Spanish Agency Dirección General de Investigación
(Ministerio de Educación y Ciencia). The authors thank
development was probably preceded by subduction
G. G. Alonso and F. Pereira for insightful reviews.
toward the south and probably also by an important
extensional event. The general absence of large
Silurian– Devonian volcanic arcs associated with References
the closure of the Rheic Ocean is also a character- A BATI , J., D UNNING , G. R., A RENAS , R., D ÍAZ G ARCÍA ,
istic of the European margin of Laurussia – F., G ONZÁLEZ C UADRA , P., M ARTÍNEZ C ATALÁN ,
Avalon. In this way, the Devonian volcanic rocks J. R. & A NDONAEGUI , P. 1999. Early Ordovician
described in southern Avalonia in England have orogenic event in Galicia (NW Spain): evidence from
been interpreted as originating in an extensional U–Pb ages in the uppermost unit of the Órdenes
setting (Floyd 1982), as has been voluminous Complex. Earth and Planetary Science Letters, 165,
213–228.
latest Devonian –Visean volcanism described in
A BATI , J., C ASTIÑEIRAS , P., A RENAS , R., F ERNÁNDEZ -
the continuation of Avalonia in the South Portu- S UÁREZ , J., G ÓMEZ -B ARREIRO , J. & W OODEN , J.
guese zone of SW Iberia. So, it can be concluded 2007. Using SHRIMP zircon dating to unravel tecto-
that neither in the southern margin of Avalonia, nothermal events in arc environments. The early
nor in the related peri-Gondwanan terranes accreted Palaeozoic arc of NW Iberia revisited. Terra Nova,
to the south of Laurussia, as that represented in the 19, 1 –8.
A LTHERR , R., H ENES -K LAIBER , U., H EGNES , E. & D ALZIEL , I. W. D. 1997. Neoproterozoic-Paleozoic
S ATIR , M. 1999. Plutonism in the Variscan Odenwald geography and tectonics: review, hypothesis, environ-
(Germany): from subduction to collision. International mental speculation. Geological Society of America
Journal of Earth Sciences, 88, 422 –443. Bulletin, 109, 16–42.
A RENAS , R. 1988. Evolución petrológica y geoquı́mica de D ÍAZ G ARCÍA , F. 1990. La geologı́a del sector occidental
la unidad alóctona inferior del complejo metamórfico del Complejo de Órdenes (Cordillera Hercı́nica, NW
básico-ultrabásico de Cabo Ortegal (Unidad de de España). Nova Terra, 3, 1– 230.
Moeche) y del Silúrico paraautóctono, Cadena Hercı́- D ÍAZ G ARCÍA , F., A RENAS , R., M ARTÍNEZ C ATALÁN ,
nica Ibérica (NW de España). Corpus Geologicum J. R., G ONZÁLEZ DEL T ÁNAGO , J. & D UNNING , G.
Gallaeciae, 4, 1– 543. 1999. Tectonic evolution of the Careón ophiolite
A RENAS , R., G IL I BARGUCHI , J. I. ET AL . 1986. Tectonos- (Northwest Spain): a remnant of oceanic lithosphere
tratigraphic units in the complexes with mafic and in the Variscan belt. Journal of Geology, 107,
related rocks of the NW of the Iberian Massif. Hercy- 587– 605.
nica, II, 87– 110. F ERNÁNDEZ -S UÁREZ , J., C ORFU , F. ET AL . 2002. U–Pb
A RENAS , R., R UBIO P ASCUAL , F. J., D ÍAZ G ARCÍA , F. & evidence for a polyorogenic evolution of the HP– HT
M ARTÍNEZ C ATALÁN , J. R. 1995. High-pressure units of the NW Iberian Massif. Contributions to
micro-inclusions and development of an inverted Mineralogy and Petrology, 143, 236– 253.
metamorphic gradient in the Santiago Schists F ERNÁNDEZ -S UÁREZ , J., D ÍAZ G ARCÍA , F., J EFFRIES ,
(Órdenes Complex, NW Iberian Massif, Spain): evi- T. E., A RENAS , R. & A BATI , J. 2003. Constraints on
dence of subduction and syn-collisional decompres- the provenance of the uppermost allochthonous
sion. Journal of Metamorphic Geology, 13, 141–164. terrane of the NW Iberian Massif: inferences from det-
A RENAS , R., A BATI , J., M ARTÍNEZ C ATALÁN , J. R., rital zircon U–Pb ages. Terra Nova, 15, 138–144.
D ÍAZ G ARCÍA , F. & R UBIO P ASCUAL , F. 1997. F ERNÁNDEZ -S UÁREZ , J., A RENAS , R., A BATI , J.,
P– T evolution of eclogites from the Agualada Unit M ARTÍNEZ -C ATALÁN , J. R., W HITEHOUSE , M. J. &
(Órdenes Complex, NW Iberian Massif, Spain): J EFFRIES , T. E. 2007. U–Pb chronometry of polyme-
Implications for crustal subduction. Lithos, 40, tamorphic high-pressure granulites: an example from
221–242. the allochthonous terranes of the NW Iberian Variscan
A RENAS , R., M ARTÍNEZ C ATALÁN , J. R., S ÁNCHEZ belt. In: H ATCHER , R. D., J R ., C ARLSON , M. P.,
M ARTÍNEZ , S., F ERNÁNDEZ -S UÁREZ , J., A NDONAE- M C B RIDE , J. H. & M ARTÍNEZ C ATALÁN , J. R.
GUI , P. & P EARCE , J. A. 2007a. The Vila de Cruces (eds) 4-D Framework of Continental Crust. Geological
Ophiolite: a remnant of the early Rheic Ocean in the Society of America Memoirs, 200, 469–488.
Variscan suture of Galicia (NW Iberian Massif). F LOYD , P. A. 1982. Chemical variation in Hercynian
Journal of Geology, 115, 129–148. basalts relative to plate tectonics. Journal of the Geo-
A RENAS , R., M ARTÍNEZ C ATALÁN , J. R. ET AL . 2007b. logical Society, London, 139, 507 –520.
Paleozoic ophiolites in the Variscan suture of Galicia F RANKE , W. 2000. The mid-European segment of the Var-
(northwest Spain): distribution, characteristics and iscides: tectonostratigraphic units, terrane boundaries
meaning. In: H ATCHER , R. D., J R ., C ARLSON , M. P., and plate tectonic evolution. In: F RANKE , W.,
M C B RIDE , J. H. & M ARTÍNEZ C ATALÁN , J. R. H AAK , V., O NCKEN , O. & T ANNER , D. (eds) Oro-
(eds) 4-D Framework of Continental Crust. Geological genic Processes: Quantification and Modelling in the
Society of America Memoirs, 200, 425– 444. Variscan Belt. Geological Society, London, Special
A RENAS , R., S ÁNCHEZ M ARTÍNEZ , S., C ASTIÑEIRAS , P., Publications, 179, 35–61.
F ERNÁNDEZ -S UÁREZ , J. & J EFFRIES , T. 2007c. G ÓMEZ -B ARREIRO , J., W IJBRANS , J. R. ET AL . 2006.
40
Geochemistry and geochronology of the ophiolite Ar/39Ar laserprobe dating of mylonitic fabrics in a
involved in the Somozas Mélange: new insights polyorogenic terrane of NW Iberia. Journal of the Geo-
on the birth of the Rheic Ocean. In: A RENAS , R., logical Society, London, 163, 61– 73.
M ARTÍNEZ C ATALÁN , J. R., A BATI , J. & S ÁNCHEZ G ÓMEZ -B ARREIRO , J., M ARTÍNEZ C ATALÁN , J. R. ET AL .
M ARTÍNEZ , S. (eds) The International Geoscience 2007. Tectonic evolution of the upper allochthon
Programme IGCP 497, ‘The Rheic Ocean: Its Origin, of the Órdenes Complex (northwestern Iberian
Evolution and Correlatives’, Galicia Meeting 2007: Massif): structural constraints to a polyorogenic peri-
Field Trip Guide & Conference Abstracts Volume, Gondwanan terrane. In: L INNEMANN , U., N ANCE ,
151–153. Publicaciones del Instituto Geológico y R. D., K RAFT , P. & Z ULAUF , G. (eds) The Evolution
Minero de España. of the Rheic Ocean: From Avalonian–Cadomian Active
B OUDIER , F. & N ICOLAS , A. 1985. Harzburgite and lher- Margin to Alleghenian –Variscan collision. Geological
zolite subtypes in ophiolitic and oceanic environments. Society of America Special Papers, 423, 315–332.
Earth and Planetary Science Letters, 76, 84–92. G UERROT , C., P EUCAT , J. J., C APDEVILA , R. & D OSSO ,
D ALLMEYER , R. D., M ARTÍNEZ C ATALÁN , J. R. ET AL . L. 1989. Archean protoliths within Early Proterozoic
1997. Diachronous Variscan tectonothermal activity granulitic crust of the west European Hercynian
in the NW Iberian Massif: evidence from 40Ar/39Ar belt: possible relics of the west African craton.
dating of regional fabrics. Tectonophysics, 277, Geology, 17, 241–244.
307–337. G UTIÉRREZ -A LONSO , G., F ERNÁNDEZ -S UÁREZ , J.,
D ALZIEL , I. W. D. 1991. Pacific margins of Laurentia and C OLLINS , A. S., A BAD , I. & N IETO , F. 2005. Amazo-
East Antarctica-Australia as a conjugate rift pair: evi- nian Mesoproterozoic basement in the core of the
dence and implications for an Eocambrian superconti- Ibero-Armorican arc: 40Ar/39Ar detrital mica ages
nent. Geology, 19, 598– 601. complement the zircońs tale. Geology, 33, 637– 640.
H OFFMAN , P. F. 1991. Did the breakout of Laurentia turn M URPHY , J. B., G UTIÉRREZ -A LONSO , G. ET AL . 2006.
Gondwana inside-out? Science, 252, 1409–1412. Origin of the Rheic Ocean: rifting along a Neoprotero-
K RÖNER , A. & C ORDANI , U. 2003. African, southern zoic suture? Geology, 34, 325–328.
Indian and South American cratons were not part of M URPHY , J. B., G UTIÉRREZ -A LONSO , G., F ERNÁNDEZ -
the Rodinia supercontinent: evidence from field S UÁREZ , J. & B RAID , J. A. 2008. Probing crustal and
relationships and geochronology. Tectonophysics, mantle lithosphere origin through Ordovician volcanic
375, 325– 352. rocks along the Iberian passive margin of Gondwana.
L INNEMANN , U., M C N AUGHTON , N. J., R OMER , R. L., Tectonophysics, 461, 166– 180.
G EHMLICH , M., D ROST , K. & T ONK , C. 2004. West P EACOCK , S. M. 1990. Numerical-simulation of meta-
African provenance for Saxo-Thuringia (Bohemian morphic pressure-temperature-time paths and fluid
Massif): Did Armorica ever leave pre-Pangean production in subducting slabs. Tectonics, 9,
Gondwana? – U/Pb-SHRIMP zircon evidence and 1197– 1211.
the Nd-isotopic record. International Journal of P EARCE , J. A. 1996. A users guide to basalt discrimination
Earth Sciences, 93, 683– 705. diagrams. In: W YMAN , D. A. (ed.) Trace Element Geo-
M ARTÍNEZ C ATALÁN , J. R., A RENAS , R., D ÍAZ G ARCÍA , chemistry of Volcanic Rocks: Application for Massive
F., R UBIO P ASCUAL , F. J., A BATI , J. & M ARQUÍNEZ , Sulphide Exploration. Short Course Notes, Geological
J. 1996. Variscan exhumation of a subducted Paleozoic Association of Canada, 12, 79–113.
continental margin: the basal units of the Órdenes P IN , C. & P AQUETTE , J. L. 2002. Le magmatisme basique
Complex, Galicia, NW Spain. Tectonics, 15, 106–121. calcoalcalin d’âge dévono-dinantien du nord du Massif
M ARTÍNEZ C ATALÁN , J. R., A RENAS , R., D ÍAZ G ARCÍA , Central, témoin d’une marge active hercynienen: argu-
F. & A BATI , J. 1997. Variscan accretionary complex ments géochimiques et isotopiques Sr/Nd. Geodina-
of northwest Iberia: terrane correlation and succession mica Acta, 15, 63– 77.
of tectonothermal events. Geology, 25, 1103– 1106. P IN , C., P AQUETT , J. L., S ANTOS Z ALDUEGUI , J. F. &
M ARTÍNEZ C ATALÁN , J. R., D ÍAZ G ARCÍA , F. ET AL . G IL I BARGUCHI , J. I. 2002. Early Devonian supra-
2002. Thrust and detachment systems in the Órdenes subduction zone ophiolite related to incipient colli-
Complex (northwestern Spain): implications for the sional processes in the Western Variscan Belt: the
Variscan-Appalachian Dynamics. In: M ARTÍNEZ Sierra de Careón unit, Órdenes Complex, Galicia. In:
C ATALÁN , J. R., H ATCHER , R. D., A RENAS , R., M ARTÍNEZ C ATALÁN , J. R., H ATCHER , R. D.,
D ÍAZ G ARCÍA , F. (eds) Variscan-Appalachians A RENAS , R. & D ÍAZ G ARCÍA , F. (eds)
Dynamics: the Building of the Late Paleozoic Base- Variscan-Appalachians Dynamics: the Building of
ment. Geological Society of America Special Papers, the Late Paleozoic Basement. Geological Society of
364, 163– 182. America Special Papers, 364, 54–71.
M ARTÍNEZ C ATALÁN , J. R., A RENAS , R. ET AL . 2007a. P IN , C., P AQUETTE , J. L., Á BALOS , B., S ANTOS , F. J. &
The rootless Variscan suture of NW Iberia (Galicia, I BARGUCHI , J. I. 2006. Composite origin of an early
Spain). Field trip guide. In: A RENAS , R., M ARTÍNEZ Variscan transported suture: ophiolitic units of the
C ATALÁN , J. R., A BATI , J. & S ÁNCHEZ M ARTÍNEZ , Morais Nappe Complex (north Portugal). Tectonics,
S. (eds) The International Geoscience Programme 25, 1 –19.
IGCP 497, ‘The Rheic Ocean: Its Origin, Evolution P ISAREVSKY , S. A., W INGATE , M. T. D., P OWELL , C.
and Correlatives’, Galicia Meeting 2007: field MC A., J OHNSON , S. P. & E VANS , D. A. D. 2003.
Trip Guide & Conference Abstracts Volume, 3 –117. Models of Rodinia assembly and fragmentation. In:
Publicaciones del Instituto Geológico y Minero de Y OSHIDA , M., W INDLEY , B. F. & D ASGUPTA , S.
España. (eds) Proterozoic East Gondwana: Supercontinent
M ARTÍNEZ C ATALÁN , J. R., A RENAS , R. ET AL . 2007b. Assembly and Break-up. Geological Society, London,
Space and time in the tectonic evolution of the north- Special Publications, 206, 35–55.
western Iberian Massif. Implications for the compre- R ODRÍGUEZ , J., C OSCA , M. A., G IL I BARGUCHI , J. I. &
hension of the Variscan belt. In: H ATCHER , D ALLMEYER , R. D. 2003. Strain partitioning and pres-
R. D., J R ., C ARLSON , M. P., M C B RIDE , J. H. & ervation of 40Ar/39Ar ages during Variscan exhuma-
M ARTÍNEZ C ATALÁN , J. R. (eds) 4-D framework of tion of a subducted crust (Malpica-Tui complex, NW
Continental Crust. Geological Society of America Spain). Lithos, 70, 111–139.
Memoirs, 200, 403– 423. S AMSON , S. D. & D’L EMOS , R. 1998. U– Pb geochronol-
M ATTE , P. 1991. Accretionary history and crustal evol- ogy and Sm-Nd isotopic composition of Proterozoic
ution of the Variscan belt in Western Europe. Tectono- gneisses, Channel Island, UK. Journal of the Geologi-
physics, 196, 309– 337. cal Society, London, 155, 609– 618.
M ATTE , P. 2001. The Variscan collage and orogeny (480– S ÁNCHEZ M ARTÍNEZ , S. 2009. Geoquı́mica y Geocrono-
290 Ma) and the tectonic definition of the Armorica logı́a de las ofiolitas de Galicia. Nova Terra, 37,
microplate: a review. Terra Nova, 13, 122– 128. 1– 351.
M OORES , E. M. 2002. Pre-1 Ga (pre-Rodinian) ophiolites: S ÁNCHEZ M ARTÍNEZ , S., J EFFRIES , T., A RENAS , R.,
their tectonic and environmental implications. Geo- F ERNÁNDEZ -S UÁREZ , J. & G ARCÍA -S ÁNCHEZ , R.
logical Society of America Bulletin, 114, 80–95. 2006. A pre-Rodinian ophiolite involved in the Varis-
M URPHY , J. B. & N ANCE , R. D. 2003. Do supercontinents can suture of Galicia (Cabo Ortegal Complex, NW
introvert or extrovert?: Sm–Nd isotopic evidence. Spain). Journal of the Geological Society, London,
Geology, 31, 873– 876. 163, 737 –740.
M URPHY , J. B. & N ANCE , R. D. 2004. How do supercon- S ÁNCHEZ M ARTÍNEZ , S., A RENAS , R., D ÍAZ G ARCÍA ,
tinents assemble? American Scientist, 92, 324–333. F., M ARTÍNEZ C ATALÁN , J. R., G ÓMEZ -B ARREIRO ,
J. & P EARCE , J. A. 2007a. Careón Ophiolite, NW dynamic plate boundaries and restored synthetic
Spain: suprasubduction zone setting for the youngest oceanic isochrons. Earth and Planetary Science
Rheic Ocean floor. Geology, 35, 53– 56. Letters, 196, 17– 33.
S ÁNCHEZ M ARTÍNEZ , S., A RENAS , R., F ERNÁNDEZ - T ORSVIK , T. H. 2003. The Rodinia jigsaw puzzle. Science,
S UÁREZ , J., J EFFRIES , T., G ARCÍA -S ÁNCHEZ , R. & 300, 1379– 1381.
A BATI , J. 2007b. The Bazar Ophiolite (NW Iberia): a W INCHESTER , J. A., P HARAOH , T. C. & V ERNIERS , J.
remnant of the Iapetus-Tornquist Ocean in the Varis- 2002. Paleozoic amalgamation of Central Europe: an
can suture?. In: A RENAS , R., M ARTÍNEZ C ATALÁN , introduction & synthesis of new results from recent
J. R., A BATI , J. & S ÁNCHEZ M ARTÍNEZ , S. (eds) geological and geophysical investigations. In:
The International Geoscience Programme IGCP 497, W INCHESTER , J. A., P HARAOH , T. C. & V ERNIERS ,
‘The Rheic Ocean: its Origin, Evolution and Correla- J. (eds) Palaeozoic Amalgamation of Central Europe.
tives’, Galicia Meeting 2007: Field Trip Guide & Con- Geological Society, London, Special Publications,
ference Abstracts Volume, 150– 151. Publicaciones del 201, 1– 18.
Instituto Geológico y Minero de España. W OOD , D. A. 1980. The application of a Th–Hf– Ta
S COTESE , C. R. 2001. Atlas of Earth History: PALEOMAP diagram to problems of tectomagmatic classification
Project. Arlington, Texas, 1– 52. and to establishing the nature of crustal contamination
S TAMPFLI , G. M. & B OREL , G. D. 2002. A plate tectonic of basaltic lavas of the British Tertiary Volcanic Pro-
model for the Paleozoic and Mesozoic constrained by vince. Earth and Planetary Science Letters, 50, 11–30.
Rheic Ocean mafic complexes: overview and synthesis
J. BRENDAN MURPHY1*, GABRIEL GUTIÉRREZ-ALONSO2, R. DAMIAN NANCE3,
JAVIER FERNÁNDEZ-SUÁREZ4, J. DUNCAN KEPPIE5, CECILIO QUESADA6,
JAROSLAV DOSTAL7 & JAMES A. BRAID1
1
Department of Earth Sciences, St. Francis Xavier University, Antigonish, Nova Scotia,
B2G 2W5 Canada
2
Departamento de Geologı́a, Universidad de Salamanca, 33708 Salamanca, Spain
3
Department of Geological Sciences, Ohio University, Athens, Ohio 45701, USA
4
Departamento de Petrologı́a y Geoquı́mica, Universidad Complutense,
28040, Madrid, Spain
5
Departamento de Geologı́a Regional, Instituto de Geologia, Universidad Nacional
Autonoma de Mexico, 04510 Mexico D.F. México
6
IGME, c/Rı́os Rosas 23, 28003, Madrid, Spain
7
Department of Geology, St. Mary’s University, Halifax, Nova Scotia, B3H 3C3 Canada
*Corresponding author (e-mail: bmurphy@stfx.ca)
Abstract: The Rheic Ocean formed during the Late Cambrian –Early Ordovician when
peri-Gondwanan terranes (e.g. Avalonia) drifted from the northern margin of Gondwana, and
was consumed during the collision between Laurussia and Gondwana and the amalgamation of
Pangaea. Several mafic complexes, from the Acatlán Complex in Mexico to the Bohemian
Massif in eastern Europe, have been interpreted to represent vestiges of the Rheic Ocean. Most
of these complexes are either Late Cambrian –Early Ordovician or Late Palaeozoic in age. Late
Cambrian –Early Ordovician complexes are predominantly rift-related continental tholeiites,
derived from an enriched c. 1.0 Ga subcontinental lithospheric mantle, and are associated with
crustally-derived felsic volcanic rocks. These complexes are widespread and virtually coeval
along the length of the Gondwanan margin. They reflect magmatism that accompanied the early
stages of rifting and the formation of the Rheic Ocean, and they remained along the Gondwanan
margin to form part of a passive margin succession as Avalonia and other peri-Gondwanan terranes
drifted northward. True ophiolitic complexes of this age are rare, a notable exception occurring in
NW Iberia where they display ensimatic arc geochemical affinities. These complexes were thrust
over, or extruded into, the Gondwanan margin during the Late Devonian–Carboniferous collision
between Gondwana and Laurussia (Variscan orogeny). The Late Palaeozoic mafic complexes
(Devonian and Carboniferous) preserve many of the lithotectonic and/or chemical characteristics
of ophiolites. They are characterized by derivation from an anomalous mantle which displays time-
integrated depletion in Nd relative to Sm. Devonian ophiolites pre-date closure of the Rheic Ocean.
Although their tectonic setting is controversial, there is a consensus that most of them reflect narrow
tracts of oceanic crust that originated along the Laurussian margin, but were thrust over Gondwana
during Variscan orogenesis. The relationship of the Carboniferous ophiolites to the Rheic Ocean
sensu stricto is unclear, but some of them apparently formed in a strike-slip regimes within a
collisional setting directly related to the final stages of the closure of the Rheic Ocean.
Introduction Baltica –Gondwana became known as the Iapetus

Ocean (Harland & Geyer 1972) and represents the
Since the publication of Tuzo Wilson’s provo- first application of modern plate tectonic principles
cative (1966) paper ‘Did the Atlantic close and to the pre-Mesozoic world. Since that time, volumes
then re-open?’ the Palaeozoic evolution of the of research have been published on vestiges of the
Appalachian– Caledonide–Variscan orogen has Iapetus Ocean preserved as ophiolitic complexes,
been key to understanding the development of primarily along the Laurentian margin.
Pangaea. The Palaeozoic ‘Proto-Atlantic’ ocean he More recent palaeogeographic reconstructions,
envisaged between the rocks with ‘Pacific fauna’ primarily based on faunal and palaeomagnetic
of Laurentia and those with ‘Atlantic fauna’ of data, have seen the orthogonal opening and closing
model of Wilson (1966) superceded by more actua-

listic models (Figs 1 & 2) involving terrane transfer
from Gondwana to Baltica and Laurentia prior to the
terminal collision between Gondwana and Laurus-
Fig. 1. Palaeozoic reconstructions showing the localities where the sequences documenting the evolution of northern flank of the Gondwanan margin (southern flank of
sia (Keppie 1985; McKerrow & Scotese 1990;
the Rheic Ocean) are preserved shown in a Pangaean reconstruction. Ac–Oax, Acatlán– Oaxaquia; Fl, Florida; C, Carolinia; A, Avalonia; O-M, Ossa Morena; NW-I,
Cocks & Fortey 1988; Cocks & Torsvik 2002;
Murphy et al. 2006a; Gómez-Barreiro et al. 2007).
These studies show: (i) that the ophiolitic complexes
were formed and obducted in the early stages (by the
Early Ordovician) of development of the Iapetus
Ocean (e.g. van Staal et al. 1998); (ii) that the ophio-
lites are generally supra-subduction zone (back-arc
and fore-arc) bodies (Jenner & Swinden 1993;
MacLachlan & Dunning 1998; Bédard et al. 1998;
Bédard & Stevenson 1999; Pin et al. 2006;
Sánchez Martı́nez et al. 2007); and (iii) that
closure of the Iapetus Ocean had occurred by the
mid-Silurian and so preceded the formation of
Pangaea by .100 Ma (e.g. Williams 1979; Keppie
1985, 1993). The northern realm of the Iapetus
Ocean was closed by the Early– mid-Silurian col-
lision between Baltica and Laurentia to form Laur-
ussia (Scandian orogeny: Roberts & Gee 1985;
Roberts & Stephens 2000; Winchester et al. 2002;
Brueckner & Van Roermund 2004). In the central
and southern realms of the Iapetus Ocean,
however, closure is attributed to the accretion of
smaller terranes (such as Avalonia, Ganderia and
Carolinia) that had rifted away from their former
positions along the margin of Gondwana (e.g.
Murphy et al. 1995, 2005b; Hibbard 2000). In
fact, we now know that much of the Appalachian-
Caledonide orogen underlain by rocks containing
the ‘Atlantic fauna’ of Wilson (1966) belongs to
NW Iberia; Arm, Armorica; MC, Massif Central; BM, Bohemian massif.
Avalonia, a terrane that separated from Gondwana

in the Late Cambrian –Early Ordovician, giving rise
to the Rheic Ocean between Avalonia and Gond-
wana. In Ordovician times, Avalonia consequently
separated a closing Iapetus Ocean to the north from
an opening Rheic Ocean to the south (Fig. 2). As a
result, it was not the demise of the Iapetus Ocean
that gave rise to Pangaea, but the closure of its
successor, the Rheic Ocean.
Despite its obvious importance to the assembly
of Pangaea, studies of the Rheic Ocean have lagged
behind those of the Iapetus Ocean. A variety of
mafic complexes, interpreted by most authors as
ophiolites, are potentially the remnants of this
ocean and record important tectonothermal events
during its evolution. However, a modern synthesis
of these complexes and their tectonic significance
is lacking. In this paper, we first provide the regional
context in which Rheic Ocean ophiolites and other
mafic complexes can be evaluated. We then provide
an overview of these complexes, assess whether
they represent ophiolites, and interpret their signifi-
cance in terms of understanding Pangaean geology.
Although more data are available for some mafic
RHEIC OCEAN MAFIC COMPLEXES: OVERVIEW AND SYNTHESIS 345
Fig. 2. Palaeozoic reconstructions (modified from Stampfli et al. 2002) showing the location of Avalonian-type terranes
along the Gondwanan margin at 495 Ma, their separation from Gondwana and northward drift of c. 2000 km by
465 Ma, and their accretion to Baltica by 440 Ma, followed by Laurentia, leading to the closure of the Iapetus Ocean.
Closure of the Rheic ocean was accomplished by northwesterly-directed subduction beneath the Laurussian margin
(e.g. 395 Ma) leading to the amalgamation of Pangaea (e.g. 350 Ma).
complexes than others, a first-order similarity exists Winchester et al. 2002; van Staal et al. 2008),
within and between them that allows their regional and to collision between Laurentia and Baltica to
tectonic significance to be evaluated. form Laurussia (events assigned to the Scandian
orogeny, Roberts & Gee 1985).
Subduction initiation within the Iapetus Ocean
Evolution of the Rheic Ocean: global was broadly coeval with the c. 500 Ma formation
context of the Rheic Ocean. The Rheic Ocean originated
when ribbon terranes (collectively called peri-
The tectonic evolution of the Palaeozoic Era is Gondwanan) drifted from the Gondwanan margin.
dominated by early Palaeozoic ocean develop- By 460 Ma, for example, one of these terranes
ment and continental dispersal, followed by con- (Avalonia) had drifted 2000 km north of the Gond-
vergence in the mid- to Late Palaeozoic, which wanan margin (Hamilton & Murphy 2004). As these
culminated in continental collisions that led to terranes formed during the inception of the Rheic
the Late Carboniferous –Early Permian amalga- Ocean, their locations broadly coincided with the
mation of the supercontinent Pangaea (Fig. 2). boundary between the contracting Iapetus Ocean
The Appalachian –Caledonide–Variscan orogen of to the north and the expanding Rheic Ocean to the
North America and Europe formed within the int- south (Fig. 2, Cocks & Torsvik 2002; Stampfli &
erior of Pangaea as the result of a series of orogenic Borel 2002; Murphy et al. 2006a). By the end of
pulses related to the closure of two key oceans, the the Ordovician, most studies suggest that these peri-
Iapetus Ocean by mid-Silurian, and the Rheic Gondwanan terranes had collided with Baltica;
Ocean in the Carboniferous. Subduction of the however, the timing of their accretion to Laurentia
Iapetus Ocean initiated along its northern flank in is controversial. Some authors favour accretion at
the Late Cambrian and led to arc –continent col- various times between the Early Silurian and Late
lisions and ophiolite obduction (events assigned to Devonian (e.g. van Staal et al. 1998; Hibbard et al.
the Taconic orogeny, Williams 1979; Keppie 2002; van Staal et al. 2008), some favour a major
1985; Cawood et al. 1996; van Staal et al. 1998; Silurian accretionary event, the Salinic orogeny
(e.g. Keppie et al. 2003b; Murphy et al. 2002b), of Gondwana that acted as an indenter during
whereas other favour accretion in the Devonian, Late Palaeozoic collisional orogenesis (Quesada
the Acadian orogeny (Hatcher 2002). Irrespective of et al. 1991; Quesada & Dallmeyer 1994). In both
these important details, there is a general consensus interpretations, the Palaeozoic rocks are genetically
that: (i) from the Late Silurian –Late Carbonife- related to one another and to the evolution of the
rous, the Rheic Ocean separated Gondwana from Rheic Ocean. However, the contrasting views
Laurussia; (ii) the closure of this ocean led to the affect the interpretation of the palaeogeography of
assembly of Pangaea through the collision of these the northern Gondwanan margin during Variscan
two continental landmasses; and (iii) this collision orogenesis and the tectonic setting associated with
gave rise to the Variscan orogeny in Europe and emplacement of these mafic complexes.
the Alleghanian orogeny in North America. It is
also widely accepted that the Rheic Ocean did not
close completely during Variscan orogenesis, but Acatlán Complex, Mexico
remained open to the east, where it is known as
the Paleotethys Ocean (e.g. Stampfli et al. 2002). General Geology
Mafic complexes that potentially represent
vestiges of Rheic Ocean geology occur from The Acatlán Complex forms the Mixteca terrane of
the Acatlán Complex in Mexico (e.g. Nance et al. southern Mexico (Fig. 3a) and is dominated by eclo-
2006, 2007; Keppie et al. 2008a) to the Bohemian gitic sedimentary and meta-igneous rocks (Piaxtla
Massif in eastern Europe (e.g. Oliver et al. 1993; Suite) and tectonic slices of phyllites, psammites
Zelazniewicz et al. 1998) to the NE (Fig. 1; current and basalt. The complex is faulted against Mesopro-
co-ordinates). The Acatlán Complex is the largest terozoic (c. 1 Ga) granulite facies gneisses of the
inlier of Palaeozoic rocks in Mexico, outcropping Oaxacan Complex, which forms the basement of
over an area equivalent in size to Massachusetts southern Oaxaquia (Ortega-Gutiérrez et al. 1995,
(Fig. 3). The origin of the Acatlán Complex is 1999), a terrane that was likely attached to the north-
controversial. According to some authors, its mafic ern portion of the Amazonian craton throughout
assemblages represent vestiges of the Iapetus Ocean the Palaeozoic (Keppie & Ramos 1999; Keppie
(e.g. Yàñez et al. 1991; Ortega-Gutiérrez et al. 2004). Oaxaquia extends northwards along the
1999; Talavera-Mendoza et al. 2005; Vega-Granillo backbone of Mexico to the suture zone with
et al. 2007). However, recent age data suggest Laurentia near the US –Mexican border (Fig. 3a).
that much of its tectonothermal evolution post- The Oaxacan Complex is unconformably overlain
dates closure of the Iapetus Ocean, and that its by latest Cambrian–earliest Ordovician platformal
history is more compatible with an evolution strata (Tiñu Formation) that contain Gondwanan
within the Rheic Ocean (Murphy et al. 2006b; fauna (Landing et al. 2007). These, in turn, are over-
Nance et al. 2006, 2007; Middleton et al. 2007; lain by Carboniferous –Permian carbonates and
Keppie et al. 2008a). clastic rocks (Centeno-Garcia & Keppie 1999).
Rheic Ocean mafic complexes are not exposed The boundary between the Mixteca and Oaxaquia
in northern South America or in North America. terranes is a Permian dextral shear zone (Fig. 3b;
Geophysical data from the southeastern United Elı́as-Herrera & Ortega-Gutiérrez 2002).
States, however, has led to the identification of The evolution of the Acatlán Complex spans the
deep seismic reflection and magnetic anomalies Ordovician– Jurassic with tectonothermal activity
(known as the Brunswick and East Coast reflecting: (i) the development of an Ordovician
Anomalies), which are interpreted as mafic bodies rift-passive margin on an Oaxacan-aged basement,
that reflect the subsurface expression of the Rheic which corresponds with the southern flank of the
Ocean suture in the southern Appalachian orogen Rheic Ocean (Keppie 2004; Miller et al. 2007;
(McBride & Nelson 1988, 1991). Keppie et al. 2008a; Morales-Gámez et al. 2008;
In Europe, mafic complexes are found within the Ramos-Arias et al. 2008); (ii) subduction-related,
Variscan orogen (Fig. 4), which is a c. 1000 km- eclogite facies metamorphism and exhumation
wide curvilinear belt with several fault-bounded tec- in the Late Devonian –Mississippian (Middleton
tonostratigraphic zones that can be traced the length et al. 2007) during and after the amalgamation of
of the orogen (e.g. Franke 1989). Mafic complexes Pangaea; (iii) subduction along the paleo-Pacific
occur predominantly in terranes adjacent to the margin following the assembly of Pangaea (Nance
northern (e.g. southern Iberia, Lizard Complex) et al. 2007; Keppie et al. 2008a); and (iv) Jurassic
and southern (NW Iberia, Massif Central) margins plume-related activity coeval with Pangaea
of the Rheic Ocean. Curvature of the tectono- breakup and the opening of the Gulf of Mexico
stratigraphic zones is interpreted either as an oro- (Keppie et al. 2004a; Nance et al. 2006).
cline (Weil et al. 2001; Gutiérrez-Alonso et al. The Acatlán Complex contains remnants of
2003) or as the result of a pre-orogenic promontory oceanic lithosphere preserved as high-pressure
Fig. 3. Simplified geological map of the Acatlán Complex, southern Mexico (modified from Ortega-Gutiérrez et al.
1999; Keppie et al. 2004b, 2008a; Nance et al. 2006).
rocks (serpentinites, blueschists and eclogites) that et al. 2004a, b, 2006, 2007, 2008a; Murphy et al.
have been emplaced into low grade psammites and 2006b; Nance et al. 2006, 2007; Middleton et al.
pelites containing mid-ocean ridge (MORB) and 2007; Miller et al. 2007; Ramos-Arias et al. 2008;
ocean island (OIB) pillow basalts interbedded with Morales-Gámez et al. 2008).
cherts. These rocks have been interpreted as vestiges The Carboniferous protolith ages of some of
of the Iapetus Ocean (Ortega-Gutiérrez et al. 1999; these mafic-ultramafic rocks clearly post-date the
Talavera-Mendoza et al. 2005; Vega-Granillo et al. Iapetus Ocean. However, the occurrence of Mid-
2007), but other studies suggest that they are more continent (USA) fauna in the Mississippian rocks
likely to have originated in the Rheic Ocean overlying the c. 1.0 Ga Oaxacan Complex indicates
(Keppie & Ramos 1999; Keppie 2004; Keppie that Pangaea had already amalgamated by the
Fig. 4. Tectonostratigraphic zonation of the Variscan orogen in Europe (after Franke 2000; Martı́nez Catalán
et al. 2007).
beginning of the Carboniferous (Navarro-Santillán 2004). Analysis of chromites within the serpentinite
et al. 2002). The Acatlán Complex lies c. 1000 km suggests that the protoliths were formed in a supra-
to the south of the suture between Oaxaquia– subduction zone setting. Unfortunately, no protolith
Amazonia and Laurentia (Keppie 2004), so or metamorphic age data are available. However,
whereas some of these mafic–ultramafic rocks along strike to the north, in the San Francisco de
may have originated within the Rheic Ocean, their Ası́s area, retrograde eclogite facies rocks have
metamorphism and exhumation probably occurred yielded a concordant U –Pb TIMS zircon age of
on the western margin of Pangaea within the Pan- 346 + 2 Ma, 350–330 Ma U –Pb SHRIMP ages
thalassa Ocean. for decompression migmatites, and a c. 350 Ma
40
The oldest dated igneous rocks in the Acatlán Ar/39Ar muscovite plateau age (Middleton et al.
Complex are c. 440 –480 Ma granitoid rocks and 2007) that are inferred to date peak metamorphism
amphibolites (Keppie et al. 2004a, 2008b; Talavera- and rapid exhumation. Here, the protoliths include
Mendoza et al. 2005; Miller et al. 2007). Mafic units rift-passive margin sedimentary rocks intruded
within the Acatlán Complex occur in several fault by a bimodal igneous assemblage dated at 470–
slices and may preserve oceanic vestiges of different 420 Ma (U –Pb SHRIMP ages for megacrystic
ages and tectonic settings (Vega-Granillo et al. granitoids, Murphy et al. 2006b) and 442 + 2 Ma
2007). Hence, the complex has been subdivided (U –Pb SHRIMP ages for zircon cores in amphi-
into lithodemes (Keppie et al. 2006, 2008a; bolite dikes: Elı́as-Herrera et al. 2007). Further
Middleton et al. 2007; Ramos-Arias et al. 2008; north, in the Patlanoaya area, retrogressed eclogite
Morales-Gámez et al. 2008; Grodzicki et al. 2008) gave an U –Pb zircon age of 353 + 1 Ma, and blues-
and tectonic synthesis of these units is a long-term chists have yielded c. 342 Ma 40Ar/39Ar ages for
goal of current research. Here, we summarize both glaucophane and phengite (Elı́as-Herrera
recent advances in our understanding of some of et al. 2007). Similarly, Vega-Granillo et al. (2007)
these complexes in order to illustrate the variability. reported 40Ar/39Ar plateau ages for phengite of
c. 347–333 Ma from several other high-pressure
Ordovician – Early Silurian complexes suites. Thus, the subduction-related metamorphism
is Mississippian in age. The rift-passive margin
The largest ultramafic body, the Tehuitzingo serpen- tectonic setting for the protoliths of most of the
tinite, occurs in the middle of the Acatlán Complex high-pressure rocks is almost identical to that of
(Fig. 3), and contains small eclogitic lenses indi- the latest Cambrian–earliest Ordovician Tiñu
cating high-pressure metamorphism probably Formation, which lies unconformably upon the
associated with a subduction zone (Proenza et al. c. 1 Ga Oaxacan Complex (Murphy et al. 2005b).
These relationships suggest that the leading edge of the mafic rocks are interpreted to be differentiated
the continental margin was first subducted and then continental tholeiites derived from an enriched
obducted over the inner continental margin (Nance mantle source. Zircon cores have yielded an age of
et al. 2006, 2007). In this model, the high-pressure 442 + 2 Ma (Elı́as-Herrera et al. 2004), which is
blueschist –eclogitic rocks mark an oceanic suture. interpreted as the time of intrusion.
However, the geological record of the Acatlán The megacrystic granites have crustal signatures
Complex on either side of the high-pressure rocks and 1.5–1.8 Ga TDM ages that are very similar to
is remarkably similar. This has led Keppie et al. those of the Oaxacan Complex (Patchett & Ruiz
(2008a) to conclude that high pressure rocks rep- 1987; Ruiz et al. 1988; Weber & Köhler 1999)
resent material that was removed by subduction and to Ordovician granitoid plutons in the northern
erosion from the upper plate, carried down the Acatlán Complex (Talavera-Mendoza et al. 2005;
subduction channel and then extruded into the Miller et al. 2007). Collectively, these data
upper plate. suggest that the Ası́s granitic magma was derived
The Ası́s Lithodeme of the Piaxtla Suite (Fig. 3) by crustal anatexis of the Oaxaquia basement,
is composed mainly of medium-to high-grade which is inferred to underlie the Acatlán Complex
metapsammitic and metapelitic rocks, and thin (Keppie 2004; Keppie et al. 2008a). Although the
bands of amphibolite (interpreted as dykes, available geochronological data are not sufficient
Murphy et al. 2006b; Middleton et al. 2007) that to resolve the relative timing of basaltic and grani-
are intruded by megacrystic granite, the margins toid magmatism in the Ası́s Lithodeme, regional
of which are mylonitized. Geochemical and Sm– considerations are consistent with the interpretation
Nd and U –Pb isotopic data for the metasedimentary that the amphibolites and granitoids represent part
rocks closely match those of the Tiñu Formation of a bimodal suite.
overlying the neighbouring Oaxacan Complex. The continental affinity of some of the rocks
Both contain a significant mafic component (sugg- in the Ası́s Lithodeme (Murphy et al. 2006b;
esting proximal sources, e.g. Nesbitt & Young Middleton et al. 2007) suggests that these rocks
1996), Mesoproterozoic TDM ages (1.5– 1.83 Ga) are not ophiolites. Instead, the Ası́s Lithodeme
(Murphy et al. 2006a, b), and c. 990 –1200 Ma igneous rocks are interpreted as a bimodal suite
detrital zircons (Gillis et al. 2005; Middleton et al. formed during continental rifting and crustal exten-
2007). Since the Tiñu Formation data are interpreted sion associated with the opening of the Rheic Ocean
to reflect derivation mainly from the underlying (Murphy et al. 2006b; Keppie et al. 2008a, b).
Oaxacan Complex, the protoliths of the Ası́s Litho- Recent geochronological data shows that these
deme metasedimentary rocks have been attributed rocks underwent eclogite facies metamorphism at
to a similar rift-passive margin setting (Keppie c. 346 Ma, followed by amphibolite facies meta-
et al. 2001, 2003a; Solari et al. 2003; Ortega- morphism and migmatization associated with
Obregón et al. 2003; Cameron et al. 2004; rapid uplift in the Visean (Middle Mississippian,
Murphy et al. 2006b). Although poorly constrained, Elı́as-Herrera et al. 2004; Middleton et al. 2007;
the c. 700 –470 Ma age range for deposition of Vega-Granillo et al. 2007). These data are attributed
the Ası́s Lithodeme, together with its geochemical, to either subduction and obduction of the leading
isotopic and detrital zircon population charac- edge of Gondwana during closure of the Rheic
teristics, are consistent with the correlation of its Ocean and the amalgamation of Pangaea (Nance
protoliths with the Tiñu Formation, and conse- et al. 2007), or to subduction erosion followed by
quently an origin along the Gondwanan margin of high-pressure metamorphism and extrusion into
the Rheic Ocean. the upper plate (Keppie et al. 2008a).
The amphibolites display geochemical trends
typical of differentiated continental tholeiites Cosoltopec Formation
(Murphy et al. 2006b). They contain high FeOt/
MgO and TiO2, slight enrichment in LREE and The Cosoltepec Formation (Fig. 3b) forms a major
trace element patterns typical of derivation from part of the Acatlán Complex and consists of unfos-
an enriched mantle source (Murphy et al. 2006b) siliferous phyllites and psammites with tectonic
with no obvious evidence of crustal contamination slices of oceanic basalt and intercalated red chert
(such as negative Nb, Ta or Zr anomalies). Sm– (Keppie et al. 2007). The formation, as currently
Nd isotopic analyses yield 1Nd(t) ranging from defined, is a composite unit (Keppie et al. 2008b;
þ2.8 to þ 4.6 (t ¼ 475 Ma) and TDM from 0.75– Morales-Gámez et al. 2008; Ramos-Arias et al.
1.27 Ga (Murphy et al. 2006b). The 1Nd(t) values 2008). Some metasedimentary rocks assigned to
are considerably lower than values expected for the formation are intruded by mid-Upper Ordovician
juvenile magmas from a depleted mantle source granitoids (Miller et al. 2007; Keppie et al. 2008b),
and are interpreted to reflect derivation from the whereas metasedimentary rocks from the type
sub-continental lithospheric mantle. Taken together, area have yielded detrital zircons as young as
c. 410 Ma (Talavera-Mendoza et al. 2005). In con- Formation and its mafic lenses derived from the
trast to the Ası́s lithodeme, an oceanic affinity has overriding plate and the Ası́s lithodeme eclogitic
been proposed for this formation, based primarily rocks forming a subducted part of the upper plate.
upon the occurrence of pillow basalts in the for- The occurrence of Mid-continent (USA) fauna
mation (e.g. Ramı́rez Espinoza 2001). Recently in Mississippian rocks resting above the Tiñu
published geochemical data indicate that the Formation (Navarro-Santillán et al. 2002) suggests
basalts are predominantly MORB and continental that Pangaea had amalgamated by this time, imply-
tholeiites with flat or depleted LREE patterns, and ing that these events took place on the western
basalts and andesites with OIB affinities, which margin of Pangaea.
have distinctly fractionated LREE-enriched patterns
(Keppie et al. 2007, 2008b). Current age data indi-
cate deposition during the Ordovician and Carbon- Iberia
iferous (Keppie et al. 2008b; Morales-Gámez General geology
et al. 2008; Grodzicki et al. 2008). In places, these
units are inferred to be unconformably overlain by Palaeozoic rocks of Iberia are divided into tectono-
either uppermost Devonian (Fammenian) strata stratigraphic zones (e.g. Lotze 1945; Julivert et al.
(Vachard & Flores de Dios 2002) or the Permian 1972; Farias et al. 1987; Quesada 1991; Fig. 4)
Tecomate Formation (Keppie et al. 2008a). based on their Lower Palaeozoic sedimentary differ-
The widespread tectonic interleaving with the ences, which are interpreted to reflect their relative
clastic metasediments was interpreted by Keppie proximity to the Gondwanan margin. Boundaries
et al. (2007) to reflect either deposition of the sedi- between these zones are major Variscan thrusts
mentary rocks directly upon the ocean-floor basalts and reverse faults that were in some cases reacti-
or extrusion of the basalts in areas removed from vated extensionally in the aftermath of the Variscan
continentally-derived sediments. However, some orogeny. The Cantabrian Zone (CZ) preserves a
mafic dykes intrude the clastic rocks (Keppie et al. coastal environment, whereas the West Asturian –
2008b; Ramos-Arias et al. 2008; Morales-Gámez Leonese (WALZ), Central Iberian (CIZ), Schistose
et al. 2008). Deformation and greenschist to sub- Domain (also known as the Galicia-Tras – os-
greenschist facies metamorphism commenced Montes Domain, SGTM) and Ossa–Morena (OMZ)
immediately prior to deposition of the late Famme- zones preserve the more outboard tectonostratigra-
nian strata, but continued into the Mississippian phy (Julivert et al. 1972; Quesada 1991; Ribeiro
(Ramos-Arias et al. 2008). et al. 1990; Pérez Estaún et al. 1990; Quesada
In the western portion of the Acatlán Complex, et al. 1991; Aramburu et al. 2002; Martı́nez
basalts occur in an unnamed sequence that was for- Catalán et al. 1997, 1999; Marcos et al. 2004;
merly assigned to the Cosoltepec Formation near Gutiérrez-Marco et al. 1999; Robardet 2002, 2003;
Olinalá (Fig. 3b). Here, the rocks consist of two tec- Robardet & Gutiérrez-Marco 2004). The SGTM is
tonically interleaved units (Grodzinski et al. 2008). considered to be parauthochthonous (Farias et al.
The Canoas unit consists of interbedded psammites 1987; Ribeiro et al. 1990) and structurally overlies
and pelites, the youngest concordant detrital zircon the authochthonous CIZ with which it shares
from which yields an age of 462 + 15 Ma (Middle igneous and stratigraphic affinities and is interpreted
Ordovician). The Coatlaco unit consists of inter- to consist of the most internal parts of the Gondwa-
bedded quartzite and pillow basalts. An average of nan passive margin. Structurally overlying the
the 8 youngest detrital zircons in the Coatlaco quart- Gondwana northern passive margin as klippen,
zite yields an age of 357 + 8 Ma (c. Devono– Variscan allochthonous complexes in NW Iberia
Carboniferous boundary), which is interpreted as include ophiolitic mafic complexes whose correla-
the present best estimate for the depositional age tives extend into the Massif Central of France. In
of these basalts. The basalts are tholeiitic with southern Iberia, ophiolitic complexes (e.g. the
within-plate and MORB-like affinities that together Beja– Acebuches ophiolite (BAO), Aracena massif)
with the interbedded cherts suggest an oceanic tec- occur as discontinuous, dismembered bodies along
tonic setting. The basaltic magmatism was coeval the boundary between the Pulo do Lobo Zone
with rapid exhumation of eclogitic rocks of the (PDLZ) and the OMZ (e.g. Silva et al. 1990). The
Piaxtla Suite, and Grodinski et al. (2008) conse- PDLZ is also characterized by an accretionary
quently suggest a link between their eruption and prism of Middle–Late Devonian age (Robardet
the exhumation of the Piaxtla Suite. 2003; Crowley et al. 2000) and together with the
Tectonic juxtaposition of the low-grade clastic BAO is thought to represent a suture reflecting the
rocks with the eclogite-bearing rocks of the Ası́s closure of the Rheic Ocean which spatially juxta-
Lithodeme suggests that the Late Devonian – posed the Iberian para-autochthon (OMZ) and the
Mississippian deformation was related to exhuma- South Portuguese Zone (SPZ), a suspect exotic
tion following subduction, with the Cosoltepec terrane. Although the oldest exposed rocks in the
SPZ are Devonian in age, several studies suggest the rocks preserve either an Eifelian –Givetian uncon-
terrane is underlain by Avalonian basement, which formity or a hiatus with some volcanic rocks
was attached to Laurussia at the time of Variscan present (Puschmann 1967; Gutiérrez-Alonso et al.
orogenesis (Oliveira & Quesada 1998; de la Rosa 2008). The Upper Devonian sedimentary sequences
et al. 2002). Although highly controversial, some extend from the CZ (coastal environment) to the
authors also propose the presence of another Varis- OMZ (outermost shelf) and record a significant
can suture in the intensely deformed Unidad Central diachronous increase in subsidence towards the
that bounds the CIZ and the OMZ, locating a narrow coast, and herald the loading of the margin during
ocean subsidiary to the Rheic, between both the progressive collision between Gondwana and
zones (see Simancas et al. 2002, and references Laurussia (Dallmeyer et al. 1997; Martı́nez
therein). Catalán et al. 2007).
Much of the passive margin sequence was sub-
Palaeozoic stratigraphy along the Northern ducted and exhumed during Variscan orogenesis
Gondwana Margin of Iberia and occurs as high- to low-pressure (eclogite,
amphibolite and greenschist facies) units in alloch-
Palaeozoic mafic complexes in Iberia occur in thonous complexes that are preserved as klippen in
different tectonostratigraphic zones that record NW Iberia (Martı́nez Catalán et al. 1997; Arenas
different aspects of Rheic Ocean evolution. et al. 2007a, b).
Mafic complexes are predominantly either Late
Cambrian –Early Ordovician or Devonian in age; Palaeozoic igneous activity along the
although some Mesoproterozoic mafic rocks have northern Gondwanan margin
been recently discovered (Sánchez Martı́nez et al.
2006). The CZ includes several Lower Palaeozoic volcanic
Early Cambrian rocks are dominated by silici- events, the most common of which are Lower
clastic rocks that were deposited on top of Ediacaran Ordovician in age (Loeschke & Zeidler 1982;
strata. In most locations, the contact between Early Heinz et al. 1985; Gallastegui et al. 1992; Barrero
Cambrian and Ediacaran successions is an uncon- & Corretgé 2002; Gutiérrez-Alonso et al. 2003).
formity. The Early and Middle Cambrian succes- The most voluminous volcanics of this age,
sions comprise the base of a passive margin however, occur in the northern CIZ, including the
sequence that continued throughout most of the ‘Ollo de Sapo’ belt, which are dominated by felsic
Palaeozoic. The Late Cambrian– Middle Ordovi- volcanics interpreted as rift-related intra-crustal
cian stratigraphy is preserved in the CIZ, WALZ melts (Valverde-Vaquero & Dunning 2000; Castro
and OMZ and is dominated by a rift-to-drift tran- et al. 1999, 2003; Dı́ez Montes 2006; Bea et al.
sition, most notably by the Lower Ordovician later- 2006; Zeck et al. 2007). Late Ordovician mafic vol-
ally extensive, mostly detrital sequence, known as canic activity has a continental, within-plate, alkalic
the Armorican Quartzite, overlying Llanvirn black signature. Sm–Nd isotopic data yield 1Nd values
shales (e.g. Young 1990; Gutiérrez-Alonso et al. ranging from þ1.0 to þ1.1, which are well below
2007) and accompanying igneous suites. Although the values for contemporary depleted mantle and
heterogeneously deformed by Variscan oroge- are interpreted to reflect derivation from an old
nesis, NW Iberia preserves the most complete (c. 0.9– 1.1 Ga) sub-continental mantle lithosphere,
Late Cambrian– Devonian passive margin sequence, whereas the coeval felsic rocks were generated by
which sits on top of Mesoproterozoic (c. 1.1– partial melting of Mesoproterozoic–Palaeoprotero-
1.4 Ga) and Late Neoproterozoic (c. 750– 550 Ma) zoic crust (Murphy et al. 2008).
basement rocks (Fernández-Suárez et al. 2000, Five allochthonous units described in NW
2002a, b; Gutiérrez-Alonso et al. 2003, 2005). Iberia (Ortegal, Ordenes, Malpica –Tui, Morais
The stratigraphy records the rift-to-drift transition and Bragança, Fig. 5) share a common organization
represented by the Late Cambrian –Ordovician that indicate their collective origin as a single
Armorican quartzite and overlying black shales. allochthonous slice obducted on top of the Gond-
Silurian and Devonian rocks outcrop extensively wana margin. The five allochthonous units represent
in all zones except the WALZ, predominantly in the product of the subsequent dismemberment of
the core of some synclinal structures in the CIZ the obducted slice and are preserved in the cores
and OMZ (Arenas et al. 2007a, b; Martı́nez Catalán of late Variscan synforms. The SGTM in this slice
et al. 1997, 1999, 2007). Silurian strata include a consists of a c. 475 Ma thick, interbedded silici-
wide variety of lithologies that are mostly pelagic, clastic volcanic sequence (Valverde-Vaquero et al.
including black shales (Robardet & Gutiérrez- 2005), and represents the outermost portion of the
Marco 2004). Devonian strata have nearshore Gondwanan passive margin (Farias et al. 1987;
and offshore intercalations and, in the CZ, include Martı́nez-Catalán et al. 1996; Marcos et al. 2002;
abundant reef deposits. Early–Middle Devonian Pereira et al. 2006). Above this unit, and present
in all the complexes, the Basal Unit represents Two ophiolitic units of Early Ordovician and
the basement of the Gondwanan margin and is com- Devonian age structurally overlie the Basal Unit
posed of metasediments of unknown age intruded (Martı́nez Catalán et al. 1997). The Lower Ophiolite
by a 490– 460 Ma bimodal igneous suite with alka- is variously known as the Moeche (Ortegal
line affinities (Pin et al. 1992; Montero 1993; Santos Complex), Vila de Cruces (Ordenes Complex, where
Zalduegui et al. 1995; Abati et al. 2003) that are it has been dated at 497 + 4 Ma, Arenas et al.
interpreted to be related to the rifting episode that 2007a) or the Izeda-Remondes (Morais Complex)
caused the opening of the Rheic Ocean (Ribeiro & and also occurs in the Bragança Complex. These
Floor 1987; Pin et al. 1992). Sm/Nd isotopic complexes preserve a complete ophiolitic suite that
values indicate that the mafic rocks were derived was metamorphosed under greenschist conditions.
from partial melting of a c. 1.0 Ga sub-continental Available geochemical data indicate that the ophio-
lithospheric mantle (Murphy et al. 2008). The lites of the Ortegal and the Ordenes complexes have
Basal Unit is thought to represent crust that was a supra-subduction zone affinity (Sánchez Martı́nez
rifted from Gondwana during the formation of the et al. 2007; Arenas et al. 2007a), whereas the ophio-
Rheic Ocean, and returned to Gondwana during lites in the Morais Complex have a MORB signature
the collision between Gondwana and Laurussia. (Pin et al. 2006). Sm –Nd isotopic values range
The Basal Unit underwent HP conditions around between þ8.0 and þ9.0, suggesting a MORB
370 Ma (e.g. Arenas et al. 1995; Gil Ibarguchi source reservoir that was depleted in Nd relative to
1995; Santos Zalduegui et al. 1995; Martı́nez Sm on a time-integrated basis (Pin et al. 2006).
Catalán et al. 1996; Rubio Pascual et al. 2002; High-pressure metamorphism has been recognized
Rodrı́guez et al. 2003; Rodrı́guez Aller 2005), an only in the Vila de Cruces unit (Arenas et al.
age that is considered as the collision age between 2007a). The main deformation event has been
Gondwana and Laurussia in this sector of the dated at c. 365 Ma (Dallmeyer & Gil Ibarguchi
Variscan belt. 1990; Dallmeyer et al. 1997).
Fig. 5. Geologic map of NW Iberia (after Farias et al. 1987; Martı́nez Catalán et al. 1997, 2007) emphasizing the
allochthonous complexes that contain mafic complexes.
The Upper Ophiolite includes the Careón and (Arenas et al. 1986; Andonaegui et al. 2002; Santos
Bazar units (Ordenes Complex), the Morais– Zalduegui et al. 2002; Abati et al. 2003; Castiñeiras
Talhinas Unit (Morais Complex), the Soeira Unit 2005), possibly adjacent to a continental margin
(Bragança Complex) and Purrido Unit (Ortegal (Peucat et al. 1990). They may reflect either residues
Complex). The Careón Unit has been dated at of anatexis (Drury 1980), crystal fractionates from
c. 395 Ma (Dı́az-Garcı́a et al. 1999; Pin et al. 2002) melts derived from a primitive mantle source
and the Morais– Talhinas Unit between c. 396 and (Gravestock 1992) or crystallization of a stratiform
405 Ma (Pin et al. 2006). Geochemical data indi- gabbroic complex at the base of the continental crust
cate that these ophiolites were formed in a supra- (Galán & Marcos 1997). The IP unit is composed
subduction setting (Pin et al. 2006; Sánchez- of terrigenous Ordovician sediments, younger
Martinez et al. 2007). Sm –Nd isotopic data yield than 480 Ma (Fernández-Suárez et al. 2003) and
1Nd values that vary from þ7.5 to þ8.9 and are c. 500 Ma (Abati et al. 1999) calcalkaline igneous
higher than the contemporary depleted mantle, indi- rocks.
cating derivation from an anomalously depleted The HT –LP protoliths are c. 490 Ma (conven-
source (Pin et al. 2006). Amphibolite facies meta- tional and SHRIMP U– Pb on zircon; Peucat et al.
morphism in the Careón unit yielded an age of 1990; Ordóñez-Casado et al. 2001; Fernández-
c. 375 Ma (Dallmeyer et al. 1997), but older ages of Suárez et al. 2007; and a Sm–Nd isochron, Santos
c. 385 Ma were found in Ortegal (Peucat et al. 1990) Zalduegui et al. 2002), and a similar age has been
and Bragança and Morais (Dallmeyer et al. 1991). proposed for the magmatic activity recorded in
Until recently, the aforementioned units were the IP units. This age is overprinted by a strong
correlated with the Purrido Unit (Vogel 1967; subduction-related HP metamorphism before
Azcárraga 2000) located in the Ortegal Complex. c. 400 Ma (Schäfer et al. 1993; Santos Zalduegui
However, recent U –Pb (zircon) data (Sánchez et al. 1996; Ordóñez-Casado et al. 2001; Fernández-
Martı́nez et al. 2006) suggest a 1.1 Ga age for at Suárez et al. 2002a, b, 2007; Roger & Matte 2005;
least part of this unit. Like the other ophiolites, Gómez Barreiro et al. 2006) and the subsequent
this unit underwent amphibolite facies metamorph- exhumation-related HT event at c. 390 Ma
ism at c. 390 Ma (Peucat et al. 1990). However, (Dallmeyer et al. 1991, 1997; Valverde Vaquero
the Purrido ophiolite may be a composite unit. The & Fernández 1996; Gómez Barreiro et al. 2006;
1Nd calculated for a crystallization age of 1.1 Ga is Fernández-Suárez et al. 2007). The ultramafic
þ10.1, which is well above the depleted mantle lithologies present in the HP– HT unit have been
value for that age (DePaolo 1988), suggesting that variously interpreted as the vestiges of sub-
the protolith age of the sample is much younger ducted, imbricated oceanic slabs (Gil Ibarguchi
than 1.1 Ga (Murphy & Gutiérrez-Alonso 2008). et al. 1990), residual harzburgites related to ophi-
The high 1Nd, together with the negative Nb–Ta olitic complexes (Santos Zalduegui et al. 1995),
anomalies suggest that the juvenile mantle source exhumed subcontinental mantle (Peucat et al.
was contaminated by a subduction component that 1990; Ábalos et al. 1996), subducted back-arc ophio-
was itself derived from juvenile crust (Murphy & lites (Peucat et al. 1990) or the crust –mantle inter-
Gutiérrez-Alonso 2008). Such contamination is face beneath a magmatic arc (Moreno et al. 2001).
characteristic of the incipient stages of oceanic arc Mafic to intermediate samples from the Upper
development (e.g. Stern 2004). The high 1Nd Unit have arc-related major and trace element
values are also characteristic of several Devonian patterns, and are characterized by relatively low
and Carboniferous ophiolites in the Variscan orogen, 1Nd (21.2 to þ2.0, t ¼ 395 Ma) and high TDM
including the Morais Complex (Pin et al. 2006), values which suggest derivation from the sub-
the Lizard Complex (Davies et al. 1984), Aracena continental lithospheric mantle (SCLM) (Murphy
(Castro et al. 1996) and Massif Central (Pin & & Gutiérrez-Alonso 2008). Ultramafic HP– HT
Paquette 2002). These complexes have 1Nd values rocks include eclogite and peridotite and Sm–Nd iso-
that are all characterized by 1Nd values that are topic systematics suggest contrasting sources. The
equivalent to or slightly higher than the isotopic eclogite has high 1Nd (þ9.9, t ¼ 395), suggesting
composition of the contemporary depleted mantle derivation from a juvenile mantle source. The perido-
(DePaolo 1981, 1988) suggesting derivation from tite has much lower 1Nd and 147Sm/144Nd, values that
a mantle source that was depleted in Nd relative to are similar to the SCLM source.
Sm on a time-integrated basis. Both HP and IP units probably represent part
Structurally overlying the ophiolites are highly of the arc that separated the closing Iapetus ocean
metamorphosed continental rocks, divided into a from the spreading Rheic ocean and was accreted
high pressure-high temperature (HP –HT) unit and to, and partially subducted beneath, the Laurussian
an intermediate pressure (IP) unit. The protoliths margin in early Devonian times (Fernández-Suárez
and geochemical affinities of the HP –HT unit are et al. 2007; Gómez Barreiro et al. 2007) reaching
broadly consistent with an ensialic arc environment a pressure of at least 1.8 GPa (Gil Ibarguchi et al.
1990; Mendı́a Aranguren 2000). The emplacement implications, the timing of deposition and defor-
of the allochthonous complexes, including the mation of units within the PDLZ remain poorly
ophiolitic mafic units, occurred after the 375 – constrained. Limited dating of spores in the upper
365 Ma collision of the Gondwana margin with flysch sequences yielded Late Devonian –early
the trench (Santos Zalduegui et al. 1995; Rodrı́guez Carboniferous ages (Eden 1991; Oliveira &
et al. 2003). Quesada 1998), suggesting an Early– Middle Devo-
nian age for the basal formation of the accretionary
SW Iberia complex.
Mafic complexes are preserved along the north-
In southwestern Iberia, mafic complexes occur ern margin of the PDLZ in the southern part of the
along the contact between the OMZ and the SPZ Aracena Metamorphic Belt (Munhá et al. 1986;
as a sequence of dismembered amphibolites (Beja Quesada & Dallmeyer 1994; Castro et al. 1996).
Acebuches Ophiolite, Munhá et al. 1986). The Amphibolites (BAO) in the Aracena metamorphic
occurrence of these potential oceanic remnants belt have MORB-like geochemical characteristics
together with the bounding PDLZ oceanic domain and 1Nd values (t ¼ 350 Ma) close to or above
(proposed accretionary prism; Eden 1991) to the the depleted mantle curve at þ7.9 to þ9.2. They
south, suggests the presence of a potential Variscan are thought to have originated in an ocean ridge
suture zone linking the OMZ (Gondwanan para- setting (Castro et al. 1996) located adjacent to
autochthon) to the SPZ (Laurussia, Avalonia?). the Gondwanan margin (Quesada et al. 2006).
The OMZ consists of a Neoproterozoic conti- The Beja mafic complex has also been interpreted
nental arc basement that was accreted to the CIZ as a potential ophiolite (e.g. Andrade et al. 1976)
(Gondwana) during the Cadomian orogeny. This or a remnant of an ensimatic arc. U– Pb (zircon)
basement is overlain by Cambrian –Ordovician rift- data yield a 350 Ma protolith age (Pin & Paquette
related bimodal volcanic and sedimentary rocks 1997; Pin et al. 2008), and cooling below 500 8C
(e.g. Quesada 1990, 1998). Trace element data at 340 Ma (Ar/Ar hornblende, Dallmeyer &
suggest that the mafic rocks were derived from an Tucker 1993; Pin et al. 2008) reflects exhumation
asthenospheric source variably enriched in LILE relative to the colliding SPZ during Variscan oro-
and LREE similar to that of basalts found in genesis. 1Nd values for the gabbros (t ¼ 350 Ma)
modern ocean island and continental to oceanic show a wide range from þ4.0 to 26.1 (Pin &
rift settings (Sanchez-Garcia et al. 2003). Felsic Paquette 1997). The higher values are from mafic
magmas are attributed to crustal anatexis. The cumulates, whereas the lower values, which corre-
magmatism is attributed to a mantle plume that is spond to higher Sr initial values, are attributed to
genetically related to the opening of the Rheic crustal contamination. 1Nd values from metabasalts
Ocean (Sánchez-Garcı́a et al. 2003). (þ8 to þ9), are higher than that of the model
The SPZ consists of pelites and quartzites depleted mantle curve of DePaolo (1981). These
of upper Devonian age, bimodal volcanics of isotope data, together with trace element and REE
Tournaisian –lower Visean age (forming the analyses, are consistent with a MORB environment
‘Iberian Pyrite Belt’ with its associated important (Pin & Paquette 1997). Recent Shrimp U/Pb
massive sulphide ore-bodies) and flysch deposits (zircon) age dates from the MORB amphibolites
of upper Visean and Namurian – Westphalian age. of the BAO and Beja Gabbro suggest a crystalliza-
During the Variscan Orogeny it has been suggested tion age for the mafic protoliths of 332 + 3 –
that the SPZ underwent a transition from a passive 340 + 4 Ma (Azor et al. 2008).
margin shelf-type environment to a syn-orogenic To date, most researchers contend that the SPZ is
flysch type setting (Silva 1990; Quesada et al. underlain by Avalonian basement. As Avalonia had
1991). This strong sedimentary and tectonic polarity accreted to Laurussia by the mid-Silurian (Quesada
is dominated by SW vergent folding, thrust dis- 1990), the BAO and PDLZ potentially preserve a
placement and a general decrease of deformation Rheic Ocean suture and record the accretionary pro-
intensity in the same direction, suggesting that this cesses that reflect the amalgamation of Pangaea. A
syntectonic sedimentation propagated toward the plethora of tectonic models have been ascribed to
SW (Silva 1990; Quesada et al. 1991) during the accommodate the current spatial juxtaposition of
Variscan Orogeny. the SPZ and OMZ based on this Avalonian affinity
The PDLZ, which occurs along the northern (Quesada & Dallmeyer 1994; Castro et al. 1996;
margin of the SPZ, has been interpreted as an Onézime et al. 2003). However, the recent discov-
accretionary complex (Eden 1991). The PDLZ is ery of relatively young amphibolite protolith ages
tectonically overlain by the potential ophiolitic (Azor 2008) suggests the BAO is not genetically
units (BAO), suggesting a change from a passive linked to other oceanic units around the Rheno-
to an active margin along the OMZ during the Hercynian belt (i.e. Lizard). Although this does
mid-late Palaeozoic. Despite these plate scale not preclude the possibility that the SPZ is underlain
by Avalonian basement, it does suggest the need for ophiolitic complexes of the Pulo de Lobo shear
a re-evaluation of the genesis and tectonothermal zone of Iberia and the Rhenohercynian zone of
history of the BAO oceanic units. northern Germany (Figs 4 & 7).
Geophysical evidence suggests that the Lizard
Lizard Complex, U.K. Complex is a thin-skinned, gently southward-
dipping structural slice ,1 km thick (e.g. Doody
The Lizard Complex of SW England (Fig. 6) has & Brooks, 1986; BIRPS & ECORS 1986). The
long been interpreted as an Early Devonian ophio- complex predominantly consists of peridotite,
lite (e.g. Bromley 1979; Barnes & Andrews 1984; amphibolite and gabbro with local sheeted dykes,
Davies 1984; Floyd 1984; Gibbons & Thompson subordinate granite and metasediment. The lowest
1991; Cook et al. 2000, 2002; Sandeman et al. structural unit, known as the Basal Unit, contains
1997, 2000; Nutman et al. 2001) obducted onto hornblende schists that are interpreted to represent
the southern continental margin of Avalonia. pillow lavas, dykes and interflow sediments. The
According to the tectonostratigraphic zonation of Basal Unit structurally overlies the early Ordovician
the Variscan orogen (Franke 1989), the Lizard Man O’War Gneiss, interpreted as part of the
Complex is potentially correlative with the Avalonian continental margin.
Fig. 6. Simplified geological map of southwestern Britain modified from British Geological Survey.
Fig. 7. Tectonostratigraphy of the Bohemian Massif and adjoining areas in the Mid-European Variscan orogen
(after Linnemann et al. 2004). Inset: Simplified map showing the large suture zones between Bohemia, Avalonia
and Baltica.
Although the age of the ophiolite is widely later stages of Rheic Ocean evolution. Recently,
believed to be 397 + 2 Ma (U –Pb, zircon; Clark Cook et al. (2000) proposed that the Lizard perido-
et al. 1998), the age of obduction is controversial. tites formed part of a non-volcanic rifted margin
A gneiss, thought to represent basement melted that was exhumed from depths of 15.7–7.5 kbar
during obduction of the ophiolite, has yielded an along extensional shear zones before intrusion
age of 376 Ma (Sandeman et al. 2000). However, of the gabbro and the sheeted dykes. If so, the peri-
on the basis of SHRIMP data, Nutman et al. dotite would not be a cogenetic member of the
(2001) argued that obduction occurred between ophiolitic complex.
400 and 380 Ma, that is soon after crystallization The ophiolite preserves pre-obduction fabrics
of the oceanic crust. This was followed by a later that record high-temperature deformation within
episode of thrusting that post-dates deposition of an oceanic environment (Gibbons & Thompson
underlying Late Devonian clastic rocks. 1991; Roberts et al. 1993). For example, the perido-
The Lizard Complex has MORB-like geochem- tite contains steep mylonitic fabrics and stretching
istry (e.g. Cooke et al. 2000) and 1Nd values lineations that are cut by the gabbros and sheeted
(t ¼ 397 Ma) for the mafic dykes and plagiogranite dykes, which are interpreted to have formed
ranging from þ9.0 to þ11.8 (Davies 1984) which during thinning of the upper mantle prior to the for-
are higher than the model depleted mantle of mation of oceanic crust. The Lizard Complex is
equivalent (i.e. c. 397 Ma) age (DePaolo 1981, widely believed to represent a vestige of a narrow
1988). The high 1Nd values implies derivation (c. 400 km) Red Sea-type oceanic tract, which
from a highly depleted mantle at c. 397 Ma which, may (Linnemann et al. 2004) or may not (Tait
together with regional constraints, suggest an et al. 2000) have been attached to Gondwana at
oceanic lithospheric source generated during the the time. The width of the putative ocean tract is
apparently too small to be resolved palaeomagneti- The relationship between the Saxo–Thuringian
cally. Kinematic data along intra-oceanic shear and Moldanubian zones is unclear because the
zones together with and the obliquity of the NW– contact is obscured by Late Mesozoic and Early
SE orientation of the sheeted dykes to the E–W Cenozoic strata. Matte (2001) interpreted these
faults in Armorica is interpreted to reflect dextral zones as a coherent microplate (Armorica),
transtension along an intra-continental strike –slip whereas Franke (2000) proposed that an oceanic
zone between Avalonia and Armorica (Badham tract existed between them. Palaeomagnetic data
1992; Cook et al. 2000, 2002). If so, the Lizard have been used to support (Tait et al. 2000) and
Complex may not be a vestige of the Rheic Ocean refute (McKerrow et al. 2000) the existence of
proper, but may instead represent a narrow such an oceanic tract. Amphibolites are abundant
oceanic tract, similar to the modern Gulf of Califor- in the Moldanubian zone of the eastern Bohemian
nia, formed adjacent to the northern margin of the Massif, and are traditionally divided into three
Rheic Ocean as suggested by Davies (1984). units: the Rehberg ophiolite, the Buschandlwand
amphibolite and the Raabs Group. The Rehberg
Eastern Europe ophiolite consists of metamorphosed (amphibolite
facies) ultramafic to mafic plutons, a gabbro/dyke
The Bohemian Massif (Fig. 7) is divided into four complex, basaltic, andesitic and rhyolitic volcanics,
tectonostratigraphic zones; from north to south, and siliciclastic sediments. Geochemical data are
these are the Rhenohercynian zone, the Mid- consistent with a supra-subduction zone environ-
German Crystalline Zone, and the Saxothuringian ment (Höck et al. 1997). These traits are similar to
and Moldanubian zones (e.g. Franke 2000). The the Letovice ophiolite in Moravia, which contains
Rhenohercynian zone is correlated with East Avalo- both MORB and tholeiitic island arc basalts and
nia (Figs 4 & 7), and therefore was part of Laurussia is interpreted to have formed in a narrow basin
in the Late Palaeozoic (Fig. 2). The major crustal floored by oceanic crust between the Moldanubian
units south of the Mid-German Crystalline Zone block to the west and the Brunovistulian–Moravian
belong to the Bohemian Massif, which is interpreted block to the east (Höck et al. 1997).
40
to be the northern margin of the peri-Gondwanan Ar/39Ar data from the Letovice ophiolite yield
Armorican Microplate that was tectonically dis- ages of 328 + 7 and 332 + 5 Ma, which are inter-
membered during Late Devonian –Carboniferous preted to date the age of metamorphism (MacIntyre
orogenesis (Matte 2001). The Mid-German Crystal- et al. 1992). They are consistent with cooling ages
line Zone is therefore interpreted to preserve part of documented by Dallmeyer et al. (1992) from the
the suture zone of the Rheic Ocean (Fig. 2). Moldanubian and Moravian zones, respectively
Although potential source rocks are not exposed, (Höck et al. 1997). The protolith ages of the
detrital chrome-spinel in Frasnian to Visean grey- ophiolites are unknown. However, in the Sudetes
wackes in the Harz Mountains of the Rheno- Mountains to the NE, ophiolites of the Central
Hercynian zone is interpreted to have been derived Sudetic terrane (Cymerman et al. 1997) occur
from pre-Givetian Alpine-type peridotites in the between the Saxo–Thuringian and Mondanubian
German Mid-Crystalline Rise (Ganssloser 1999). zones and have yielded c. 420 Ma crystallization
Saxo–Thuringia is considered to reflect the ages (Śle˛ża ophiolite, U –Pb zircon, Oliver et al.
northern margin of the Armorican plate and is 1993) and c. 350 Ma Sm–Nd whole rock ages
underlain by Late Neoproterozoic arc and back-arc (Pin et al. 1988) interpreted to reflect the age
complexes, and 540 Ma post-tectonic granitoids of metamorphism (Cymerman et al. 1997). Else-
developed during the Cadomian Orogeny where within the Central Sudetic terrane, the
(Linnemann et al. 2000; Linnemann & Romer Klodzko Metamorphic Complex (KMC) contains
2002). These rocks are unconformably overlain by Neoproterozoic –Devonian metasedimentary and
a Lower –Middle Cambrian platformal succession, meta-igneous rocks exposed in a sequence of thrusts.
followed disconformably in the Upper Cambrian This complex is interpreted to preserve remnants of
and Lower Ordovician by a thick (c. 3 km) trans- the northern Gondwanan margin that was tele-
gressive sequence thought to reflect the drift of scoped during the Variscan orogeny (Cymerman
Avalonia from the Gondwanan margin (Linnemann et al. 1997). Kryza et al. (2003) identified two domi-
et al. 2004). A West African affinity for Saxo– nant Palaeozoic meta-igneous units within the KMC:
Thuringia is suggested by Cambro–Ordovician (1) a Cambro– Ordovician bimodal, rift-related
fauna and the presence of latest Ordovician suite similar to several suites developed along the
glaciomarine diamictites of the Saharan glaciation northern periphery of Gondwana (most notably in
(Linnemann et al. 2004). The Lower Silurian– the Ossa Morena Zone of Iberia; Sánchez-Garcı́a
Lower Carboniferous stratigraphy is dominated by et al. 2003); and (2) a Devonian, predominantly
deep-water clastic and carbonate sediments depos- mafic unit in the northern part of the complex,
ited along the southern flank of the Rheic Ocean. which has a 1Nd value of þ6.8 (assuming a youngest
possible crystallization age of 400 Ma), typical of is interpreted to reflect incipient rifting of a volcanic
magmas derived from a depleted mantle source. arc along a continental margin associated with
According to Franke (2000), the oceanic rocks southward-dipping subduction (Pin & Paquette
between the Saxo– Thuringian and Moldanubian 2002). In the western massif, c. 370–380 Ma arc
zones represent an oceanic tract of unknown plutons (Pin & Paquette 2002) are interpreted to
width. Alternatively, they have been interpreted as reflect either northward- (Shaw et al. 1993) or
nappes transported southward from the Rheic southward-directed subduction (Faure et al. 1997).
suture zone (Kroner & Hahn 2003). In the South Vosges, a series of thrusts expose
Late Devonian tholeiitic basalts along with siliceous
Massif Central shales, radiolarian cherts and rare limestone, a
sequence that is thought to reflect deposition in an
The Massif Central is a portion of the Gondwanan ocean basin (Schneider et al. 1990).
margin that underwent polyphase orogenic activity
between the Late Silurian and Late Carboniferous. Discussion
It is dominated by allochthonous complexes includ-
ing ophiolites and high-pressure metamorphic units. Although the study of mafic complexes within the
Early orogenic activity produced coesite-bearing Rheic Ocean has lagged behind that of their
eclogitic rocks formed at a depth of c. 90 km counterparts in the Iapetus Ocean, the available
between 420 and 400 Ma (Lardeaux et al. 2001). data are sufficient to identify some first-order
Unroofing to a depth of c. 30 km by 380 –360 Ma characteristics: (1) those of Late Cambrian– Early
was accompanied by crustal extension and gravita- Ordovician age are related to the initial rifting of
tional collapse (Burg et al. 1994; Faure 1995). the Rheic Ocean along the northern Gondwanan
Middle–Late Devonian nappes contain a gneissic margin; (2) those of Devonian and Carboniferous
series consisting of Late Neoproterozoic –Early age are ophiolitic and represent fragments of
Palaeozoic metasediments, intruded by Early Ordo- Rheic or Rheic-related ocean floor, probably gener-
vician plutons (e.g. Ledru et al. 1994). Early ated along the Avalonian margin of Laurussia; and
Ordovician bimodal igneous rocks are interpreted (3) those of Carboniferous age appear to be oceanic
by Pin & Marini (1993) to be a local representative in origin, but their relationship to the Rheic Ocean,
of a regional rifting event. The mafic magmas have a sensu stricto, is unclear.
plume-related chemistry and the felsic magmas are Late Cambrian –Early Ordovician mafic com-
attributed to crustal melting. plexes occur from Mexico to Bohemia, indicating
Voluminous Devonian –Carboniferous magma- that rifting was coeval along much of the northern
tism occurs throughout the Massif Central. In the Gondwanan margin. In most of these complexes,
NE, the massif is dominated by mid- to Late the basalts are differentiated continental tholeiites
Devonian calc-alkaline to tholeiitic volcanic and and are part of a bimodal suite. Where isotopic
plutonic rocks (Thiéblemont & Cabanis 1986). data are available (Mexico, NW Iberia) the basalts
Low-grade Late Devonian –Early Carboniferous appear to have been derived from either the sub-
(c. 366 –358 Ma: U – Pb zircon, Pin & Paquette continental lithospheric mantle (SCLM) that was
1997) bimodal basalts, rhyolites and genetically metasomatically altered at c. 1.0 Ga, or from
related trondhjemite intrusions occur in the juvenile sources.
allochthonous Brévenne Series, which was tectoni- Basalts derived from SCLM are typically part of
cally juxtaposed against high-grade rocks to the a bimodal complex, with varying proportions of
north and south by Carboniferous dextral strike-slip felsic magmas that were derived by crustal
motion (Costa et al. 1993). The metabasalts are melting. Passive margin sedimentation accompa-
characterized by an enrichment in incompatible nied this rift-related magmatism, as documented,
elements (e.g. Th and LREE), depletion in high from west to east, by siliciclastic rocks at the base
field strength elements, and positive 1Ndt values of the Tiñu Formation in Mexico and by the Armor-
(from þ5 to þ8 at t ¼ 360 Ma). Taken together, ican Quartzite in mainland Europe. The genetic
these data are consistent with a MORB mantle relationship of the basalts with crustally derived
source that was contaminated by subduction- melts and their association with passive margin
derived fluids. The metarhyolites are enriched in sediments indicates that they are not ophiolites.
LILE, have high 1Ndt values (from þ4.7 to þ6.8) Instead, these complexes developed along the
and are interpreted to have been derived from the passive margin of northern Gondwana and reflect
basalts by fractional crystallization (Pin & Paquette the rift and drift of Avalonia and related peri-
1997). The trondhjemites, on the other hand, have Gondwanan terranes during the formation of the
much lower 1NDt values (from 21.0 to þ2.2) and Rheic Ocean. Passive margin sedimentation along
their genesis is thought to involve an important this margin appears to have continued virtually
contribution of continental crust. The magmatism uninterrupted until the Devonian.
Late Cambrian –Early Ordovician ophiolites are Modanubian Zone of Bohemia have chemistries
rare. The best documented examples occur in NW consistent with the settings of these Late Palaeozoic
Iberia and are interpreted to be remnants of the ophiolites, although whether these ophiolites
Rheic Ocean (or subsidiary oceans) that were formed before or during collisional processes that
obducted during the Variscan Orogeny. These accompanied Rheic Ocean closure is unclear.
ophiolite bodies typically have an ensimatic island Devonian and Carboniferous mafic complexes
arc tholeiitic chemistry and are interpreted to have appear to be isolated, individually enigmatic bodies,
been generated in a back-arc environment during although some common traits are shared. Interpret-
the first stages of the opening of the Rheic Ocean ations of these complexes depend to some degree
(Arenas et al. 2007a). In Mexico, mafic complexes on resolution of the controversy about the existence
between c. 480 and 400 Ma are not well documen- of a separate Armorican microplate. If this micro-
ted. However, unit formerly assigned to the Cosol- plate existed, then some of the mafic complexes
topec Formation clearly contain Ordovician and may reflect the formation of that plate, or the evol-
Late Devonian –Carboniferous volcanic rocks with ution of the ocean (known as the Palaeotethys
MORB and OIB chemical affinities and oceanic Ocean), which is thought to have formed when the
sediments that are candidates for ophiolites. Armorican plate separated from Gondwana.
The Devonian (e.g. Lizard, southern Britain) Deformation and metamorphism of Rheic Ocean
and Carboniferous (Aracena, Beja-Acebuches; SW mafic complexes occurred in the Carboniferous
Iberia) mafic complexes are ophiolites with geo- associated with the collision of Gondwana with
chemical affinities and juvenile Sm– Nd isotopic Laurasia. Late Cambrian –Early Ordovician sequ-
characteristics that are typical of MORB compo- ences lack regional-scale tectonothermal activity
sitions (e.g. Castro et al. 1996; Davies 1984; Cook until the Carboniferous deformation and meta-
et al. 2000). Unlike the Ordovician ophiolites, morphism. From the Acatlán Complex of Mexico
there is nothing in the geochemistry or isotopic sys- to the Bohemian Massif, metamorphism up to eclo-
tematics of these complexes that suggests an gite facies followed by partial melting collectively
arc-related environment. Indeed these ophiolites reflect subduction and/or subduction erosion fol-
typically have higher 1Nd than the theoretical lowed by rapid exhumation of the leading edge
contemporary depleted mantle value, indicating of the Gondwanan margin. These relationships
derivation from a similar, but isotopically unusual suggest that this sequence was on the lower plate
mantle, with time-integrated depletion of Nd relative during continental collision (e.g. Martı́nez Catalán
to Sm. A mantle composition as unusually depleted et al. 2007; Middleton et al. 2007), or the upper
as this could form if a significant fraction of basalt plate removed by subduction erosion (Keppie
was removed from it during an earlier (unspecified) et al. 2008a).
tectonothermal episode. However, the relationship The classic exposures in NW Iberia indicate that
of the Devonian and Carboniferous ophiolites to the Gondwanan passive margin succession was
Rheic ocean evolution may be different. The Devo- overthrust by nappes, which include ophiolite com-
nian Lizard Complex, for example, is widely plexes derived from the remnants of the Rheic
believed to represent a narrow oceanic tract formed Ocean and possibly from Laurussia (Martı́nez-
by dextral strike-slip motion along the southern Catalán et al. 1996, 2007). The nappes also
margin of Avalonia. Its Devonian age indicates contain ophiolitic complexes that are interpreted to
that its origin pre-dates the closure of the Rheic be remnants of the Rheic Ocean that were obducted
Ocean. The relationship of the Carboniferous during the Variscan orogeny.
Aracena and Beja ophiolites to the Rheic ocean In conclusion, the expression of the Rheic
sensu stricto is enigmatic, because they potentially Ocean suture zone can be traced from the Acatlán
post-date the collision between the SPZ and the Complex in Mexico, through Iberia and the Mid-
OMZ, which represents closure of the Rheic Ocean German Crystalline Rise, to the northern margin
in SW Iberia (Quesada 1990; Azor et al. 2008). of the Bohemian massif. Mafic complexes docu-
These ophiolites may have formed in a sinistral ment important aspects of the evolution of this
strike slip environment associated with oblique ocean from its origin to the genesis of ophiolites
subduction and collision between these zones emplaced during its closure. Although much work
(Quesada 1990) that, according to Pin et al. (2008), remains to be done, the available data demonstrate
may also have involved slab breakoff, or a post- that a more precise knowledge of the evolution of
collisional transtensional event (Azor et al. 2008). these mafic complexes will help to constrain the
Unfortunately, Sm –Nd isotopic data from other geodynamic setting that attended the formation of
mafic complexes, required to test either the regional Pangaea.
extent of this unusual mantle, or the processes that
gave rise to it, are lacking. However, several ophio- JBM acknowledges the continuing support of NSERC.
lite complexes, such as those of NW Iberia and the (Canada) through Discovery and Research Capacity
Development grants. G.G.-A. funding comes from Spanish (Northwest Iberian Massif). Journal of Geology, 115,
Education and Science Ministry Project Grant 129–148.
CGL2006-00902 (ODRE.) and the Mobility Program A RENAS , R., M ARTÍNEZ C ATALÁN , J. R. ET AL . 2007b.
Grant PR2007-0475 and the hospitality of St. Francis Paleozoic ophiolites in the Variscan suture of Galicia
Xavier University. RDN acknowledges funding from (northwest Spain). In: H ATCHER , R. D., J R .,
NSF (EAR-0308105) and an Ohio University Baker C ARLSON , M. P., M C B RIDE , J. H. & M ARTÍNEZ
Award. JDK would like to acknowledge a Papiit grant C ATALÁN , J. R. (eds) 4-D Framework of Continental
IN100108 and a CONACyT grant (CB –2005–1: 24894). Crust. Geological Society, London, Memoirs, 200,
We are grateful to D. Thorkelson and A. Hynes for 403–423.
constructive reviews. A ZCÁRRAGA , J. 2000. Evolución tectónica y metamórfica
de los mantos inferiores de grado alto y alta presión
del Complejo de Cabo Ortegal. Serie Nova Terra, 17,
References 1– 306.
A ZOR , A., R UBATTO , D., S IMANCAS , J. F., G ONZÁLEZ
Á BALOS , B., A ZCÁRRAGA , J., G IL I BARGUCHI , J. I., L ODEIRO , F., M ARTÍNEZ P OYATOS , D., M ARTÍN
M ENDIA , M. & S ANTOS Z ALDUEGUI , J. F. 1996. P ARRA , L. M. & M ATAS , J. 2008. Rheic Ocean ophio-
Flow stress, strain rate and effective viscosity evalu- litic remnants in southern Iberia questioned by
ation in a high-pressure metamorphic nappe (Cabo SHRIMP U– Pb zircon ages on the Beja-Acebuches
Ortegal, Spain). Journal of Metamorphic Geology, amphibolites. Tectonics, 27, TC5006, doi: 10.1029/
14, 227–248. 2008TC002306, 1– 11.
A BATI , J., D UNNING , G. R., A RENAS , R., D ÍAZ G ARCÍA , B ARNES , R. P. & A NDREWS , J. R. 1984. Hot or cold
F., G ONZÁLEZ C UADRA , P., M ARTÍNEZ C ATALÁN , emplacement of the Lizard Complex? Journal of the
J. R. & A NDONAEGUI , P. 1999. Early Ordovician Geological Society, London, 141, 37– 39.
orogenic event in Galicia (NW Spain): evidence from B ARRERO , M. & C ORRETGÉ , L. G. 2002. La diferenciación
U– Pb ages in the uppermost unit of the Ordenes petrográfica y geoquı́mica de las rocas ı́gneas del
Complex. Earth and Planetary Science Letters, 165, Arroyo Farandón (Asturias). Geogaceta, 32, 135–138.
213– 228. B ADHAM , J. P. N. 1982. Strike-slip orogens– an expla-
A BATI , J., A RENAS , R., M ARTÍNEZ C ATALÁN , J. R. & nation for the Hercynides. Journal of the Geological
D ÍAZ G ARCÍA , F. 2003. Anticlockwise P– T Patch of Society, London, 139, 493– 504.
Granulites from the Monte Castelo Gabbro (Órdenes B EA , F., M ONTERO , P., T ALAVERA , C. & Z INGER , T.
Complex, NW Spain). Journal of Petrology, 44, 2006. A revised Ordovician age for the oldest magma-
305– 327. tism of Central Iberia: U– Pb ion-microprobe and
A NDONAEGUI , P., G ONZÁLEZ DEL T ANAGO , J., LA-ICPMS dating of the Miranda do Douro ortho-
A RENAS , R., A BATI , J., M ARTÍNEZ C ATALÁN , J. R., gneiss. Geologica Acta, 4, 395– 401.
P EINADO , M. & D ÍAZ G ARCÍA , F. 2002. Tectonic B ÉDARD , J. H. & S TEVENSON , R. 1999. The Caldwell
setting of the Monte Castelo gabbro (Ordenes Group lavas of southern Quebec: MORB-like tholeiites
Complex, northwestern Iberian Massif): evidence for associated with the opening of the Iapetus Ocean.
an arc-related terrane in the hangingwall to the Variscan Canadian Journal of Earth Sciences, 36, 999–1019.
suture. In: M ARTÍNEZ C ATALÁN , J. R., H ATCHER , B ÉDARD , J. H., L AUZIÈRE , K., T REMBLAY , A. &
R. D., J R ., A RENAS , R. & D ÍAZ G ARCÍA , F. (eds) S ANGSTER , A. 1998. Evidence for forearc seafloor-
Variscan-Appalachian Dynamics: The Building of the spreading from the Betts Cove ophiolite, New-
Late Paleozoic Basement. Geological Society of foundland: oceanic crust of boninitic affinity.
America, Special Paper, 364, 37– 56. Tectonophysics, 284, 233– 245.
A RAMBURU , C., T RUYOLS , J. ET AL . 2002. El Paleozoico BIRPS & ECORS 1986. Deep seismic reflection profiling
Inferior de la Zona Cantábrica. In: R ABANO , I., between England, France and Ireland. Journal of the
G UTIÉRREZ -M ARCO , J. C. & S AAVEDRA , J. (eds) Geological Society, London, 143, 45– 52.
El Paleozoico Inferior de Ibero-América. Universidad B ROMLEY , A. V. 1979. Ophiolitic origin of the Lizard
de Extremadura, 397–421. Complex. Cambourne School of Mines Journal, 79,
A RENAS , R., G IL I BARGUCHI , J. I. ET AL . 1986. Tectonos- 25–38.
tratigraphic units in the complexes with mafic and B RUECKNER , H. K. & V AN R OERMUND , H. L. M.
related rocks of the NW of the Iberian Massif. 2004. Dunk tectonics: a multiple subduction/eduction
Hercynica, 2, 87– 110. model for the evolution of the Scandinavian Caledo-
A RENAS , R., R UBIO P ASCUAL , F. J., D ÍAZ G ARCÍA , F. & nides. Tectonics, 23(TC2004), 1– 20, doi: 10.1029/
M ARTÍNEZ C ATALÁN , J. R. 1995. High-pressure 2003TC001502.
micro-inclusions and development of an inverted B URG , J. P., V AN D EN D RIESSCHE , J. & B RUN , J. P.
metamorphic gradient in the Santiago Schists (Ordenes 1994. Syn- to post-thickening extension in the
Complex, NW Iberian Massif, Spain): evidence of sub- Variscan Belt of Western Europe: mode and structural
duction and syn-collisional decompression. Journal of consequences. Géologie de la France, 3, 33– 51.
Metamorphic Geology, 13, 141 –164. C AMERON , K. L., L ÓPEZ , R., O RTEGA -G UTIÉRREZ , F.,
A RENAS , R., M ARTÍNEZ C ATALÁN , J. R., S ÁNCHEZ S OLARI , L. A., K EPPIE , J. D. & S CHULZE , C. 2004.
M ARTÍNEZ , S., F ERNÁNDEZ -S UÁREZ , J., A NDO- U–Pb constraints and Pb isotopic compositions of
NAEGUI , P., P EARCE , J. A. & C ORFU , F. 2007a. leached feldspars: constraints on the origin and
The Vila de Cruces Ophiolite: a Remnant of the evolution of Grenvillian rocks from eastern and
Early Rheic Ocean in the Variscan Suture of Galicia southern Mexico. In: T OLLO , R. P., C ORRIVEAU , L.,
M C L ELLAND , J. & B ARTHOLOMEW , M. J. (eds) Pro- fragmentation of the Armorican Terrane Assemblage.
terozoic Tectonic Evolution of the Grenville Orogen in Terra Nova, 12, 171 –180.
North America. Geological Society of America C YMERMAN , Z., P IASECKI , M. A. J. & S ESTON , R. 1997.
Memoir, 197, 755– 770. Terranes and terrane boundaries in the Sudetes.
C ASTIÑEIRAS , P. 2005. Origen y evolución tectonotermal Geological Magazine, 134, 717– 725.
de las unidades de O Pino y Cariño (Complejos alócto- D ALLMEYER , R. D. & G IL I BARGUCHI , J. I. 1990. Age
nos de Galicia). Laboratorio Xeolóxico de Laxe, Serie of amphibolitic metamorphism in the ophiolitic unit
Nova Terra, 28, 1– 279. of the Morais allochthon (Portugal): implications for
C ASTRO , A., F ERNÁNDEZ , C., DE LA R OSA , J. D., early Hercynian orogenesis in the Iberian Massif.
M ORENO V ENTAS , I. & R OGERS , G. 1996. Signifi- Journal of the Geological Society, London, 147,
cance of MORB-derived amphibolites from the 873– 878.
Aracena metamorphic belt, southwest Spain. Journal D ALLMEYER , R. D. & T UCKER , R. D. 1993. U– Pb zircon
of Petrology, 37, 235–260. age for the Lagoa augen gneiss, Morais complex, Por-
C ASTRO , A., P ATIÑO D OUCE , A. E., C ORRETGÉ , L. G., tugal: tectonic implications. Journal of the Geological
DE LA R OSA , J. D., E L -B IAD , M. & E L -H MIDI , H. Society, London, 150, 405– 410.
1999. Origin of peraluminous granites and granodior- D ALLMEYER , R. D., R IBEIRO , A. & M ARQUES , F. 1991.
ites, Iberian massif, Spain: an experimental test of Polyphase Variscan emplacement of exotic terranes
granite petrogénesis. Contributions to Mineralogy (Morais and Bragança Massifs) onto Iberian succes-
and Petrology, 135, 255– 276. sions: evidence from 40Ar/39Ar mineral ages. Lithos,
C ASTRO , A., C ORRETGÉ , L. G., DE LA R OSA , J. D., 27, 133 –144.
F ERNÁNDEZ , C., L ÓPEZ , S., G ARCÍA -M ORENO , O. D ALLMEYER , R. D., N EUBAUER , F. & H ÖCK , V. 1992.
& C HACÓN , H. 2003. The Appinite-Migmatite Chronology of late Paleozoic tectonothermal activity
Complex of Sanabria, NW.Iberian Massif, Spain. in the southeastern Bohemian massif, Austria (Mol-
Journal of Petrology, 44, 1309–1344. danubian and Moravo-Silesian zones) – 40Ar/39Ar
C AWOOD , P. A., VAN G OOL , J. A. M. & D UNNING , G. R. mineral age controls. Tectonophysics, 210, 135–153.
1996. Geological development of eastern Humber and D ALLMEYER , R. D. & M ARTÍNEZ C ATALÁN , J. R. 1997.
western Dunnage zone: corner Brook–Glover Island Diachronous Variscan tectonothermal activity in the
region. Canadian Journal of Earth Sciences, 33, NW Iberian Massif: evidence from 40Ar/39Ar dating
182–198. of regional fabrics. Tectonophysics, 227, 307– 337.
C ENTENO -G ARCÍA , E. & K EPPIE , J. D. 1999. Latest D AVIES , G. R. 1984. Isotopic evolution of the Lizard
Paleozoic-early Mesozoic structures in the central Complex. Journal of the Geological Society, London,
Oaxaca terrane of southern Mexico: deformation 141, 3– 14.
near a triple junction. Tectonophysics, 301, 231– 242. D E LA R OSA , J. D., J ENNER , G. A. & C ASTRO , A. 2002.
C LARK , A. H., S COTT , D. J., S ANDEMAN , H. A. & A study of inherited zircons in granitoid rocks from the
B ROMLEY , A. V. 1998. Siegenian generation of the South Portuguese and Ossa-Morena Zones, Iberian
Lizard ophiolite: U– Pb zircon age data for plagiogra- Massif: support for the exotic origin of the South
nite, Porthkerris, Cornwall. Journal of the Geological Portuguese Zone. Tectonophysics, 352, 245– 256.
Society, London, 155, 595–598. D E P AOLO , D. J. 1981. Neodymium isotopes in the
C OCKS , L. R. M. & F ORTEY , R. A. 1988. Lower Colorado Front Range and crust-mantle evolution in
Paleozoic facies and faunas around Gondwana. In: the Proterozoic. Nature, 29, 193– 196.
A UDLEY -C HARLES , M. G. & H ALLAM , A. (eds) D E P AOLO , D. J. 1988. Neodymium Isotope Geochemistry:
Gondwana and Tethys. Geological Society, London, An Introduction. Springer Verlag, New York.
Special Publications, 37, 183–200. D ÍAZ G ARCÍA , F., A RENAS , R., M ARTÍNEZ C ATALÁN ,
C OCKS , L. R. M. & T ORSVIK , T. H. 2002. Earth geogra- J. R., G ONZÁLEZ DEL T ÁNAGO , J. & D UNNING ,
phy from 500 to 400 million years ago: a faunal and G. R. 1999. Tectonic evolution of the Careón ophiolite
palaeomagnetic review. Journal of the Geological (Northwest Spain): a remnant of oceanic lithosphere
Society, London, 159, 631–644. in the Variscan belt. Journal of Geology, 107,
C OOK , C. A., H OLDSWORTH , R. E., S TYLES , M. T. & 587– 605.
P EARCE , J. A. 2000. Pre-emplacement structural D ÍEZ M ONTES , A. 2006. La geologı́a del Dominio ‘Ollo
history recorded by mantle peridotites: an example de Sapo’ en las comarcas de Sanabria y Terra do
from the Lizard Complex, SW England. Journal of Bolo. PhD thesis, University of Salamanca, Spain.
the Geological Society, London, 157, 1049– 1064. D OODY , J. J. & B ROOKS , M. 1986. Seismic refraction
C OOK , C. A., H OLDSWORTH , R. E. & S TYLES , M. T. investigation of the structural setting of the Lizard
2002. The emplacement of peridotites and associated and Start complexes, SW England. Journal of the
oceanic rocks from the Lizard Complex, southwest Geological Society, London, 143, 135–140.
England. Geological Magazine, 139, 2745. D RURY , S. A. 1980. The geochemistry of high-pressure
C OSTA , S., M ALUSKI , H. & L ARDEAUX , J. M. 1993. gneisses from Cabo Ortegal (NW Spain): residues of
40
Ar– 39Ar chronology of Variscan tectono- deep anatexis. Geol Mijbouw, 59, 61– 64.
metamorphic events in an exhumed crustal nappe: E DEN , C. 1991. Tectonostratigraphic analysis of the
the Monts du Lyonnais complex (Massif Central, northern extend of the oceanic exotic terrane, north-
France). Chemical Geology, 105, 339– 359. western Huelva Province, Spain. PhD thesis, Earth
C ROWLEY , Q. G., F LOYD , P. A., W INCHESTER , J. A., Sciences, Southampton.
F RANKE , W. & H OLLAND , J. G. 2000. Early Palaeo- E LÍAS -H ERRERA , M. & O RTEGA -G UTIÉRREZ , F. 2002.
zoic rift-related magmatism in Variscan Europe: Caltepec fault zone: an Early Permian dextral
transpressional boundary between the Proterozoic F LOYD , P. A. 1984. Geochemical characteristics and
Oaxacan and Paleozoic Acatlán complexes, southern comparison of the basic rocks of the Lizard Complex
Mexico, and regional implications. Tectonics, 21(3), and the basaltic lavas within the Hercynian troughs
4 –1 to 4 –18, doi: 10.1029/200TC001278. of SW England. Journal of the Geological Society,
E LÍAS -H ERRERA , M., O RTEGA -G UTIÉRREZ , F., London, 141, 61– 70.
S ÁNCHEZ -Z AVALA , J. L., R EYES -S ALAS , A. M., F RANKE , W. 1989. Variscan plate-tectonics in Central-
M ACÍAS -R OMO , C. & I RIONDO , A. 2004. New Europe – current ideas and open questions. Tectono-
geochronological and stratigraphic data related to physics, 169, 221–228.
the Paleozoic evolution of the high-pressure F RANKE , W. 2000. The mid-European segment of the Var-
Piaxtla Grop, Acatlán Complex, southern Mexico. IV iscides: tectonometamorphic units, terrane boundaries
reunion nacional de ciencias de la tierra, Libro de and plate tectonic evolution. Geological Society,
Resumenes, 1– 150. London, Special Publications, 179, 35– 61.
E LÍAS -H ERRERA , M., M ACÍAS -R OMO , C., O RTEGA - G ALÁN , G. & M ARCOS , A. 1997. Geochemical evolution
G UTIÉRREZ , F., S ÁNCHEZ -Z AVALA , J. L., of high-pressure mafic granulites from the Bacariza
I RIONDO , A. & O RTEGA -R IVERA , A. 2007. Conflict- formation (Cabo Ortegal Complex, NW Spain): an
ing Stratigraphic and Geochronologic Data From the example of a heterogeneous lower crust. Geologische
Acatlán Complex: ‘Ordovician’ Granites Intrude Rundschau, 86, 539 –555.
Metamorphic and Sedimentary Rocks of Devonian- G ALLASTEGUI , G., M ARTÍN P ARRA , L. M., DE P ABLO
Permian age. Eos Transactions of the American Geo- M ACIÁ , J. G. & R ODRÍGUEZ F ERNÁNDEZ , L. R.
physical Union, 88(23), Joint Assembly Supplement, 1987. Las metavulcanitas del Dominio Esquistoso de
Abstract T41A-12. Galicia tras os Montes: petrografı́a, geoquı́mica y
F ARIAS , P., G ALLASTEGUI , G. ET AL . 1987. Aportaciones ambiente geotectónico (Galicia, NO de España). Cua-
al conocimiento de la litoestratigrafı́a y estructura dernos do Laboratorio Xeoloxico de Laxe, Coruña, 15,
de Galicia Central. Faculdade de Ciências do Porto. 127–139.
Memorias do Museu e Laboratório Mineralógico e G ANSSLOSER , M. 1999. Detrital chromian spinels in
Geológico, 1, 411–431. Rhenohercynian greywackes and sandstones (Givetian-
F AURE , M. 1995. Late orogenic Carboniferous extensions Visean, Variscides, Germany) as indicators of ultrama-
in the Variscan French Massif Central. Tectonics, 14, fic source rocks. Geological Magazine, 136, 437– 451.
132– 153. G IBBONS , W. & T HOMPSON , L. 1991. Ophiolitic mylo-
F AURE , M., L ELOIX , C. & R OIG , J. Y. 1997. L’évolution nites in the Lizard complex: ductile extension in the
polycyclique de la chaı̂ne hercynienne. Bulletin de la lower crust. Geology, 19, 1009–1012.
Societé.Géologique de France, 168, 695– 705. G ILLIS , R. J., G EHRELS , G. E., R UIZ , J. & F LORES DE
F ERNÁNDEZ -S UÁREZ , J., G UTIÉRREZ A LONSO , G., D IOS G ONZÁLEZ , A. 2005, Detrital zircon provenance
J ENNER , G. A. & T UBRETT , M. N. 2000. New ideas of Cambrian-Ordovician and Carboniferous strata of
on the Proterozoic– Early Paleozoic evolution of NW the Oaxaca terrane, southern Mexico. Sedimentary
Iberia: insights from U– Pb detrital zircon ages. Pre- Geology, 182, 87–100.
cambrian Research, 102, 185– 206, doi: 10.1016/ G IL I BARGUCHI , J. I. 1995. Petrology of jadeite metagra-
S0301-9268(00)00065-6. nite and associated orthogneiss from the Malpica-Tuy
F ERNÁNDEZ -S UÁREZ , J., G UTIÉRREZ A LONSO , G., Allochton (Northwest Spain). European Journal of
C OX , R. & J ENNER , G. A. 2002a. Assembly of the Mineralogy, 7, 403– 415.
Armorica microplate: a strike-slip terrane delivery? G IL I BARGUCHI , J. I. & D ALLMEYER , R. D. 1991. Hercy-
Evidence from U– Pb ages of detrital zircons. nian blueschist metamorphism in North Portugal:
Journal of Geology, 110, 619– 626, doi: 10.1086/ tectonothermal implications. Journal of Metamorphic
341760. Geology, 9, 539– 549.
F ERNÁNDEZ -S UÁREZ , J., C ORFU , F. ET AL . 2002b. U–Pb G ÓMEZ B ARREIRO , J., W IJBRANS , J. R., C ASTINEIRAS ,
evidence for a polyorogenic evolution of the HP– HT P., M ARTÍNEZ C ATÁLAN , J. R., A RENAS , R., D ÍAZ
units of the NW Iberian Massif. Contributions to G ARCÍA , F. & A BATI , J. 2006. Ar-40/Ar-39 laserp-
Mineralogy and Petrology, 143, 236– 253. robe dating of mylonitic fabrics in a polyorogenic
F ERNÁNDEZ -S UÁREZ , J., D ÍAZ -G ARCIÁ , F., J EFFRIES , terrane of NW Iberia. Journal of the Geological
T. E., A RENAS , R. & A BATI , J. 2003. Constraints on Society, London, 163, 61–73.
the peovenance of the uppermost allochthonous G ÓMEZ -B ARREIRO , J., M ARTÍNEZ C ATALÁN , J. R.,
terrane of the NW Iberian Massif. Inferences A RENAS , R., C ASTIÑEIRAS , P., A BATI , J., D ÍAZ
from detrital zircon U –Pb age. Terra Nova, 15, G ARCÍA , F. & W IJBRANS , J. R. 2007. Tectonic
138– 144. evolution of the upper allochthon of the Ordenes
F ERNÁNDEZ -S UÁREZ , J., A RENAS , R., A BATI , J., Complex (northwestern Iberian Massif): structural
M ARTÍNEZ C ATALÁN , J. R., W HITEHOUSE , M. J. & constraints to a polyorogenic peri-Gondwanan
J EFFRIES , T. 2007. U– Pb chronometry of poly- terrane. In: L INNEMANN , U., N ANCE , R. D., K RAFT ,
metamorphic high-pressure granulites: an example P. & Z ULAUF , G. (eds) The Evolution of the Rheic
from the allochthonous terranes of the NW Ocean: From Avalonian– Cadomian Active Margin
Iberian Variscan belt. In: H ATCHER , R. D., J R ., to Alleghenian– Variscan Collision. Geological
C ARLSON , M. P., M C B RIDE , J. H. & M ARTÍNEZ Society of America Special Paper, 423, 315– 332.
C ATALÁN , J. R. (eds) Four-D evolution of continental G RAVESTOCK , P. J. 1992. The chemical causes of
crust. Geological Society of America Memoir, 200, uppermost mantle heterogenities. PhD thesis, Open
469– 488. University.
G RODZICKI , K. R., N ANCE , R. D., K EPPIE , J. D., H EINZ , W., L OESCHKE , J. & V AVRA , G. 1985. Phreato-
D OSTAL , J. & M URPHY , J. B. 2008. Structural, geo- magmatic volcanism during the Ordovician of the
chemical and geochronological analysis of metasedi- Cantabrian Mountains. Geologische Rundschau, 74,
mentary and metavolcanic rocks of the Coatlaco area, 623– 639.
Acatlán Complex, southern Mexico. Tectonophysics, H IBBARD , J. 2000. Docking Carolina: Mid-Paleozoic
461, 311– 323. accretion in the southern Appalachians. Geology, 28,
G UTIÉRREZ -A LONSO , G. & F ERNÁNDEZ -S UÁREZ , J. 127– 130.
1996. Geologı́a y Geoquı́mica del granitoide pre- H IBBARD , J. P., S TODDARD , E. F., S TODDARD , E. F.,
Varisco de Puente de Selce (Antiforme del Narcea, S ECOR , D. T. & D ENNIS , A. J. 2002. The Carolina
Asturias). Revista de la Sociedad Geológica de Zone: overview of Neoproterozoic to Early Paleozoic
España, 9, 227– 239. peri-Gondwanan terranes along the eastern flank of
G UTIÉRREZ -M ARCO , J. C., A RAMBURU , C. ET AL . 1999. the southern Appalachians. Earth Science Review,
Revisión bioestratigráfica de las pizarras del Ordovı́- 57, 299 –339.
cico Medio en el noroeste de España zonas Cantábrica, H ÖCK , V., M ONTAG , O. & L EICHMANN , J. 1997. Ophio-
Asturoccidental-leonesa y Centroibérica septentrional. lite remnants at the eastern margin on the Bohemian
Acta Geologica Hispanica, 34, 3 –87. Massif, and their bearing on the tectonic evolution.
G UTIÉRREZ -A LONSO , G., F ERNÁNDEZ -S UÁREZ , J., Mineralogy and Petrology, 60, 267–287.
J EFFRIES , T. E., J ENNER , J. E., T UBRETT , M. N., J ENNER , G. A. & S WINDEN , H. S. 1993. The Pipestone
C OX , R. & J ACKSON , S. E. 2003. Terrane accretion Pond Complex, central Newfoundland: complex mag-
and dispersal in the northern Gondwana margin: matism in an eastern Dunnage ophiolite. Canadian
an Early Paleozoic analogue of a long lived active Journal of Earth Sciences, 30, 434– 448.
margin. Tectonophysics, 65, 221– 232, doi: 10.1016/ J ULIVERT , M., M ARCOS , A. & T RUYOLS , J. 1972.
S0040-1951(03)00023-4. L’evolutión paléogéographique du nord-ouest de
G UTIÉRREZ -A LONSO , G., F ERNÁNDEZ -S UÁREZ , J., l’Espagne pendant l’Ordovicien-Silurien. Bulletin de
C OLLINS , A. S., A BAD , I. & N IETO , F. 2005. Amazo- la Societé Géologique et Minéralogique de Bretagne,
nian Mesoproterozoic basement in the core of 4, 1– 7.
the Ibero-Armorican Arc: 40Ar/39Ar detrital mica K EPPIE , J. D. 1985. The Appalachian Collage. In: G EE ,
ages complement the zircon’s tale. Geology, 33, D. G. & S TURT , B. (eds) The Caledonide Orogen,
637–640. Scandinavia, and Related Areas. J. Wiley & Sons,
G UTIÉRREZ -A LONSO , G., F ERNÁNDEZ -S UÁREZ , J., New York, 1217–1226.
G UTIÉRREZ -M ARCO , J. C., C ORFU , F., M URPHY , J. K EPPIE , J. D. 1993. Synthesis of Paleozoic deformational
B. & S UÁREZ , M. 2007. U–Pb depositional age for events and terrane accretion in the Canadian Appala-
the upper Barrios Formation (Armorican Quartzite chians. Geologische Rundschau, 82, 381– 431.
facies) in the Cantabrian zone of Iberia: Implications K EPPIE , J. D. 2004. Terranes of Mexico revisited: A 1.3
for stratigraphic correlation and paleogeography. In: billion year odyssey. International Geology Review,
L INNEMANN , U., N ANCE , R. D., K RAFT , P. & 46, 765 –794.
Z ULAUF , G. (eds) The Evolution of the Rheic Ocean: K EPPIE , J. D. & R AMOS , V. S. 1999. Odyssey of terranes
From Avalonian– Cadomian Active Margin to in the Iapetus and Rheic Oceans during the Paleozoic.
Alleghenian– Variscan Collision. Geological Society In: K EPPIE , J. D. & R AMOS , V. A. (eds) Laurentia-
of America Special Paper, 423, 287– 296, doi: Gondwana Connections Before Pangea. Geological
10.1130/2007.2423(13). Society of America, Special Paper, 336, 267–276.
G UTIÉRREZ -A LONSO , G., M URPHY , J. B., F ERNÁNDEZ - K EPPIE , J. D., D OSTAL , J., O RTEGA -G UTIÉRREZ , F. &
S UÁREZ , J. & H AMILTON , M. 2008. A Devonian L ÓPEZ , R. 2001. A Grenvillian arc on the margin of
age for the El Castillo Volcanic rocks (Salamanca, Amazonia: evidence from the southern Oaxacan
Central Iberian Zone): implications for rifting along Complex, southern Mexico. Precambrian Research,
the northern Gondwana margin and the evolution of 112, 165–181.
the Rheic ocean. Tectonophysics, 461, 157–165. K EPPIE , J. D., D OSTAL , J., C AMERON , K. L., S OLARI ,
H AMILTON , M. A. & M URPHY , J. B. 2004. Tectonic sig- L. A., O RTEGA -G UTIÉRREZ , F. & L ÓPEZ , R. 2003a.
nificance of a Llanvirn age for the Dunn Point volcanic Geochronology and geochemistry of Grenvillian
rocks, Avalon terrane, Nova Scotia, Canada: impli- igneous suites in the northern Oaxacan Complex,
cations for the evolution of the Iapetus and Rheic southern Mexico: tectonic implications. Precambrian
oceans. Tectonophysics, 379, 199– 209. Research, 120, 365– 389.
H ARLAND , W. B. & G AYER , R. A. 1972, The Arctic K EPPIE , J. D., N ANCE , R. D., M URPHY , J. B. & D OSTAL ,
Caledonides and earlier oceans. Geological Magazine, J. 2003b. Tethyan, Mediterranean, and Pacific ana-
109, 289– 314. logues for the Neoproterozoic–Paleozoic birth and
H ATCHER , R. D. 2002. Alleghenian (Appalachian) development of peri-Gondwanan terranes and their
orogeny, a product of zipper tectonics: rotational, transfer to Laurentia and Laurussia. Tectonophysics,
trans pressive continent-continent collision and 365, 195–220.
closing of ancient oceans along irregular margins. K EPPIE , J. D., S ANDBERG , C. A., M ILLER , B. V.,
In: M ARTÍNEZ C ATALÁN , J. R., H ATCHER , S ÁNCHEZ -Z AVALA , J. L., N ANCE , R. D. & P OOLE ,
R. D., J R ., A RENAS , R. & D ÍAZ G ARCÍA , F. (eds) F. G. 2004a. Implications of latest Pennsylvanian
Variscan– Appalachian Dynamics: The Building of to Middle Permian paleontological and U–Pb
the Late Paleozoic Basement. Geological Society of SHRIMP data from the Tecomate Formation to
America, Special Paper, 364, 199–208. re-dating tectonothermal events in the Acatlán
Complex, southern Mexico. International Geology and Nd– Sr–Pb isotopic characterisation of marginal
Review, 46, 745– 753. basins with constraints to geotectonic setting and
K EPPIE , J. D., N ANCE , R. D. ET AL . 2004b. Mid-Jurassic provenance. Tectonophysics, 352, 33– 64.
tectonothermal event superposed on a Paleozoic geo- L INNEMANN , U., G EHMLICH , M. ET AL . 2000. Cadomian
logical record in the Acatlán Complex of southern subduction to Early Palaeozoic rifting: the evolution of
Mexico: hotspot activity during the breakup of Saxo-Thuringia at the margin of Gondwana in the light
Pangea. Gondwana Research, 7, 239– 260. of single zircon geochronology and basin development
K EPPIE , J. D., N ANCE , R. D., F ERNÁNDEZ -S UÁREZ , J., (Central European Variscides, Germany). In: F RANKE ,
S TOREY , C. D., J EFFRIES , T. E. & M URPHY , J. B. W., H AAK , V., O NCKEN , O. & T ANNER , D. (eds) Oro-
2006. Detrital zircon data from the eastern Mixteca genic Processes: Quantification and Modelling in the
terrane, southern Mexico: evidence for an Ordovician- Variscan Belt. Geological Society, London, Special
Mississippian continental rise and a Permo-Triassic Publications, 179, 131–153.
clastic wedge adjacent to Oaxaquia. International L INNEMANN , U., M C N AUGHTON , N. J., R OMER , R. L.,
Geology Review, 48, 97– 111. G EHMLICH , M., D ROST , K. & T ONK , C. 2004. West
K EPPIE , J. D., D OSTAL , J. & E LIÁS -H ERRERA , M. 2007. African provenance for Saxo-Thuringia (Bohemian
Ordovician-Devonian oceanic basalts in the Cosolte- Massif): did Armorica ever leave pre-Pangean Gond-
pec Formation, Acatlán Complex, southern Mexico: wana? – U/Pb-SHRIMP zircon evidence and the
vestiges of the Rheic Ocean? In: L INNEMANN , U., Nd-isotopic record. International Journal of Earth
N ANCE , R. D., K RAFT , P. & Z ULAUF , G. (eds) The Sciences, 93, 683– 705.
Evolution of the Rheic Ocean: From Avalonian– L OESCHKE , W. H. J. & Z EIDLER , N. 1982. Early Paleo-
Cadomian active margin to Alleghenian –Variscan zoic sills in the Cantabrian Mountains (Spain) and
collision. Geological Society of America, Special their geotectonic environment. Jahrbuch für Geologie
Paper, 423, 477–488. und Paläontologie, Abhandlungen, 7, 419–439.
K EPPIE , J. D., D OSTAL , J., M URPHY , J. B. & N ANCE , L OTZE , F. 1945. Zur Gliderung der Varisziden der
R. D. 2008a. Synthesis and tectonic interpretation of Iberischen Meseta. Geotektonische Forschungen, 6,
the westernmost Paleozoic Variscan orogen in 78–92.
southern Mexico: from rifted Rheic margin to active M AC I NTYRE , R. M., B OWES , D. R., H AMIDULLAH , S. &
Pacific margin. Tectonophysics, doi: http://dx.doi. O NSTOTT , T. C. 1992. K –Ar and Ar –Ar isotopic
org/10.1016/j.tecto.2008.01.012. study on amphiboles from meta-ophiolite complexes,
K EPPIE , J. D., D OSTAL , J. ET AL . 2008b. Ordovician– Eastern Bohemian Massif. Proceedings, 1st Inter-
earliest Silurian rift tholeiites in the Acatlán Com- national Conference Bohemian Massif, Prague 1988,
plex, southern Mexico: evidence of rifting on the 195–199.
southern margin of the Rheic Ocean. Tectonophysics, M AC L ACHLAN , K. & D UNNING , G. 1998. U–Pb ages and
461, 130– 156; doi: 10.1016/j.tecto.2008.01.010. tectono-magmatic evolution of Middle Ordovician vol-
K RONER , U. & H AHN , T. 2003. Sedimentation, canic rocks of the Wild Bight Group, Newfoundland
Deformation und Metamorphose im Saxothuringikum Appalachians. Canadian Journal of Earth Sciences,
während der variszsichen Orogenese: Die komplexe 35, 998– 1017.
Entwicklung von Nord-Gondwana während konti- M ARCOS , A., F ARIAS , P., G ALÁN , G., F ERNÁNDEZ , F. J.
nentaler Subduktion und schiefer Kollision. In: & L LANA -F ÚNEZ , S. 2002. The tectonic framework
L INNEMANN , U. (ed.) Das Saxothuringikum. Staatliche of the Cabo Ortegal Complex: a slab of lower crust
Naturhistorische Sammlungen Dresden, 133– 146. exhumed in the Variscan Orogen (NW Iberian Penin-
K RYZA , R., M AZUR , S. & P IN , C. 2003. Subduction- sula). In: M ARTÍNEZ C ATALÁN , J. R., H ATCHER ,
and non-subduction-related igneous rocks in the R. D., J R ., A RENAS , R. & D ÍAZ G ARCÍA , F. (eds)
Central-European Variscides: geochemical and Nd Variscan– Appalachian Dynamics: The Building of
isotope evidence for a composite origin of the Kodzko- the Late Paleozoic Basement. Geological Society of
Metamorphic Complex, Polish Sudetes. Geodinamica America, Special Paper, 364, 143–162.
Acta, 16, 39– 57. M ARCOS , A., M ARTÍNEZ C ATALÁN , J. R., G UTIÉRREZ -
L ANDING , E., K EPPIE , J. D. & W ESTROP , S. 2007. Term- M ARCO , J. C. & P ÉREZ -E STAÚN , A. 2004. Estratigra-
inal Cambrian and lowest Ordovician of Mexican West fı́a y paleogeografı́a. In: V ERA , J. A. (ed.) Geologı́a de
Gondwana: biotas and sequence stratigraphy of the España. S.G.E.-I.G.M.E., Madrid, 49–52.
Tiñu Formation. Geological Magazine, 144, 900–936. M ARTÍNEZ C ATALÁN , J. R., H ACAR R ODRÍGUEZ , M. P.,
L ARDEAUX , J. M., L EDRU , P., D ANIEL , I. & D UCHENE , V ILLAR A LONSO , M. P., P ÉREZ -E STAÚN , A. &
S. 2001. The Variscan French Massif Central – a G ONZÁLEZ L ODEIRO , F. 1992. Paleozoic extensional
new addition to the ultra-high pressure metamorphic tectonics in the limit between the West Asturias-
‘club’: exhumation processes and geodynamic conse- Leonese and the Central Iberian Zones of the Bariscan
quences. Tectonophysics, 332, 143 –167. Fold-Belt in NW Spain. Geologische Rundschau, 81,
L EDRU , P., A UTRAN , A. & S ANTALLIER , D. 1994. 545–560.
Lithostratigraphy of Variscan terranes in the French M ARTÍNEZ C ATALÁN , J. R., A RENAS , R., D ÍAZ G ARCÍA ,
Massif Central. A basic for paleogeographical recon- F., R UBIO P ASCUAL , F. J., A BATI , J. & M ARQUÍNEZ ,
struction. In: K EPPIE , J. D. (ed.) Pre-Mesozoic J. 1996. Variscan exhumation of a subducted Paleozoic
geology in France and related areas. Springer continental margin: the basal units of the Órdenes
Verlag, Berlin, 276–288. Complex, Galicia, NW Spain. Tectonics, 15, 106– 121.
L INNEMANN , U. & R OMER , R. L. 2002. The Cadomian M ARTÍNEZ C ATALÁN , J. R., A RENAS , R., D ÍAZ G ARCÍA ,
Orogeny in Saxo-Thuringia, Germany: geochemical F. & A BATI , J. 1997. Variscan accretionary complex
of northwest Iberia: terrane correlation and M ORENO , T., G IBBONS , W., P RICHARD , H. & L UNAR , R.
succession of techtonothermal events. Geology, 25, 2001. Platiniferous chromitite and the tectonic setting
1103–1106. of ultramafic rocks in Cabo Ortegal (north-west
M ARTÍNEZ C ATALÁN , J. R., A RENAS , R., D ÍAZ G ARCÍA , Spain). Journal of the Geological Society, London,
F. & A BATI , J. 1999. Allochthonous units in the Var- 158, 601–614.
iscan belt of NW Iberia. Terranes and accretionary M ILLER , B. V., D OSTAL , J., K EPPIE , J. D., N ANCE , R. D.,
history. In: S INHA , A. K. (ed.) Basement Tectonics. O RTEGA -R IVERA , A. & L EE , J. W. K. 2007. Ordovi-
Kluwer Academic Publishers, 13, 65– 84. cian calc-alkaline granitoids in the Acatlán Complex,
M ARTÍNEZ C ATALÁN , J. R., A RENAS , R. ET AL . 2007. southern Mexico: geochemical and geochronological
Space and time in the tectonic evolution of the data and implications for tectonic of the Gondwanan
northwestern Iberian Massif. Implications for the margin of the Rheic Ocean. In: L INNEMANN , U.,
comprehension of the Variscan belt. In: H ATCHER , N ANCE , R. D., Z ULAF , G. & K RAFT , P. (eds) The
R. D., J R ., C ARLSON , M. P., M C B RIDE , J. H. & Evolution of the Rheic Ocean: From Avalonian–
M ARTÍNEZ C ATALÁN , J. R. (eds) 4-D framework of Cadomian Active Margin to Alleghenian– Variscan
Continental Crust. Geological Society, London, Collision. Geological Society of America, Special
Memoirs, 200, 1 –435. Paper, 423, 465– 475.
M ATTE , P. 2001. The Variscan collage and orogeny M UNHÁ , J., O LIVEIRA , J. T., R IBEIRO , A., O LIVEIRA , V.,
(480–290 Ma) and the tectonic definition of the Q UESADA , C. & K ERRICH , R. 1986. Beja-Acebuches
Armorica microplate: a review. Terra Nova, 13, ophiolite: characterization and geodynamic signifi-
122–128. cance. Maleo, 2(2), 13, 30.
M C B RIDE , J. & N ELSON , D. 1991. Deep seismic reflec- M URPHY , J. B. & G UTIÉRREZ -A LONSO , G. 2008.
tion constraints on Palaeozoic crustal structure and The origin of the Variscan NW Iberian allochthonous
definition of the Moho in the buried southern Appala- complexes in NW Iberia: Sm/Nd isotopic constraints.
chian orogen. Continental lithosphere: deep Seismic Canadian Journal of Earth Sciences, 45, 651–668.
Reflections. Geodynamics, 22, 1– 20. M URPHY , J. B., R ICE , R. J., S TOKES , T. R. & K EPPIE , D.
M C B RIDE , J. & N ELSON , D. 1988. Intergration of F. 1995. The St. Mary’s Basin, central mainland
COCORP deep reflection and magnetic anomaly Nova Scotia: late Paleozoic basin formation and defor-
analysis in the southeastern US: implications for the mation along the Avalon-Meguma Terrane boundary,
origin of the Brunswick and East Coast magnetic Canadian Appalachians. In: H IBBARD , J., V AN
anomalies. Geological Society of America Bulletin, S TAAL , C. & C AWOOD , P. (eds) New perspectives in
100, 436– 455. the Caledonian –Appalachian orogen. Geological
M C K ERROW , W. S. & S COTESE , C. R. (eds) 1990. Association of Canada, Special Paper, 41, 409– 420.
Palaeozoic palaeogeography and biogeography. Geo- M URPHY , J. B., K EPPIE , J. D., B RAID , J. & N ANCE , R. D.
logical Society of London, London, Memoirs, 12. 2005b. Geochemistry of the Tremadocian Tiñu
M C K ERROW , W. S., M AC N IOCAIL , C., A HLBERG , P. E., Formation (southern Mexico): provenance in the
C LAYTON , G., C LEAL , C. J. & E AGAR , M. C. 2000. underlying 1 Ga Oaxacan Complex on the southern
The Late Palaeozoic relations between Gondwana margin of the Rheic Ocean. International Geology
and Laurussia. Geological Society, London, Special Review, 47, 887–900.
Publications, 179, 9 –20. M URPHY , J. B., G UTIÉRREZ -A LONSO , G. ET AL . 2006a.
M IDDLETON , M., K EPPIE , J. D., M URPHY , J. B., M ILLER , Origin of the Rheic ocean: rifting along a Neoprotero-
B. V., N ANCE , R. D., O RTEGA -R IVERA , A. & L EE , zoic Suture? Geology, 34, 325 –328.
J. K. W. 2007. P– T– t constraints on exumation M URPHY , J. B., K EPPIE , J. D., N ANCE , R. D., M ILLER , B.
following subduction in the Rheic Ocean from eclogi- V., M IDDLETON , M., F ERNÁNDEZ -S UÁREZ , J. &
tic rocks in the Acatlán Complex of southern Mexico. J EFFRIES , T. E. 2006b. Geochemistry and U–Pb pro-
In: L INNEMANN , U., N ANCE , R. D., K RAFT , P. & tolith ages of eclogitic rocks of the Ası́s Lithodeme,
Z ULAUF , G. (eds) The Evolution of the Rheic Piaxtla Suite, Acatlán Complex, southern Mexico:
Ocean: From Avalonian–Cadomian Active Margin to tectonothermal activity along the southern margin of
Alleghenian– Variscan Collision. Geological Society the Rheic Ocean. Journal of the Geological Society,
of America, Special Paper, 423, 489– 509. London, 163, 683–695.
M ONTERO , M. P. 1993. Geoquı́mica y petrogénesis del N ANCE , R. D., M ILLER , B. V., K EPPIE , J. D., M URPHY ,
complejo peralcalino de la Sierra del Galiñeiro J. B. & D OSTAL , J. 2006. The Acatlán Complex,
(Pontevedra, España). PhD thesis, University of southern Mexico: record of Pangea assembly to
Oviedo, Spain. breakup. Geology, 34, 857–860.
M ORALES -G ÁMEZ , M., K EPPIE , J. D. & N ORMAN , M. N ANCE , R. D., M ILLER , B. V., K EPPIE , J. D., M URPHY ,
2008. Ordovician-Silurian rift-passive margin on the J. B. & D OSTAL , J. 2007. Vestige of the Rheic
Mexican margin of the Rheic Ocean overlain by Ocean in North America: the Acatlán Complex of
Permian periarc rocks: evidence from the Acatlán southern México. In: L INNEMANN , U., N ANCE ,
Complex, southern Mexico. In: P EREIRA , M. F., R. D., Z ULAF , G. & K RAFT , P. (eds) The Evolution
B OZKURT , E., Q UESADA , C. & S TRACHAN , R. (eds) of the Rheic Ocean: From Avalonian–Cadomian
The Foundations and Birth of the Rheic Ocean: active margin to Alleghenian –Variscan collision.
Avalonian –Cadomian Orogenic Processes and Geological Society of America, Special Paper, 423,
Early Palaeozoic Rifting at the North-Gondwana 437– 452.
Margin. Tectonophysics, 461, 291–310; http://dx. N AVARRO -S ANTILLÁN , D., S OUR -T OVAR , F. &
doi.org/10.1016/j.tecto.2008.01.014 C ENTENO -G ARCÍA , E. 2002. Lower Mississippian
(Osagean) brachiopods from the Santiago Formation, high-pressure nappe (northwestern Spain). Tectono-
Oaxaca, Mexico: stratigraphic and tectonic impli- physics, 177, 263–292.
cations. Journal of South American Earth Sciences, P EREIRA , M. F., C HICHORRO , M., L INNEMANN , U.,
15, 327–336. E GUILUZ , L. & S ILVA , J. B. 2006. Inherited arc signa-
N ESBITT , H. W. & Y OUNG , G. M. 1996. Petrogenesis of ture in the Ediacaran and Early Cambrian basins of the
sediments in the absence of chemical weathering: Ossa-Morena Zone (Iberian Massif, Portugal): paleo-
effects of abrasion and sorting on bulk composition geographic link with West-Central European and
and mineralogy. Sedimentology, 43, 341– 358. North African Cadomian correlatives. Precambrian
N UTMAN , A. P., G REEN , D. H., C OOK , C. A., S TYLES , Research, 144, 297–315.
M. T. & H OLDSWORTH , R. E. 2001. SHRIMP U–Pb P IN , C. & M ARINI , F. 1993. Early Ordovician continental
zircon dating of the exhumation of the Lizard Perido- break-up in Variscan Europe: Nd– Sr isotope and trace
tite and its emplacement over crustal rocks: constraints element evidence from bimodal igneous associations
for tectonic models. Journal of the Geological Society, of the Southern Massif Central, France. Lithos, 29,
London, 158, 809– 820. 177–196.
O LIVEIRA , J. T. & Q UESADA , C. 1998. A comparison of P IN , C. & P AQUETTE , J.-L. 1997. A mantle-derived
stratigraphy, structure, and palaeogeography of the bimodal suite in the Hercynian belt: Nd isotope and
South Portuguese zone and south-west England, Euro- trace element evidence for a subduction-related rift
pean Variscides. Geoscience in South-west England, 9, origin of the Late Devonian Brevenne metavolcanics,
141– 150. Massif Central (France). Contributions to Mineralogy
O LÍVER , G., C ORFU , F. & K ROGH , T. 1993. U– Pb ages & Petrology, 129, 222– 238.
from southwest Poland: evidence for a Caledonian P IN , C. & P AQUETTE , J.-L. 2002. Sr–Nd isotope and trace
suture zone between Baltica and Gondwana. Journal element evidence for a Late Devonian active margin in
of the Geological Society, London, 150, 355–369. northern Massif Central (France). Geodynamica Acta,
O NÉZIME , J., C HARVET , J., F AURE , M., B OURDIER , J. L. 15, 63–77.
& C HAUVET , A. 2003. A new geodynamic inter- P IN , C., M AJEROWICZ , A. & W OJCIECHOWSKA , I. 1988.
pretation for the South Portuguese Zone (SW Upper Palaeozoic oceanic crust in the Polish Sudetes:
Iberia) and the Iberian Pyrite Belt genesis. Tectonics, Nd– Sr isotope and trace element evidence. Lithos,
22, 4. 21, 195– 209.
O RDÓÑEZ -C ASADO , B., G EBAUER , D., S CHÄFER , H.-J., P IN , C., O RTEGA C UESTA , L. A. & G IL I BARGUCHI , J. I.
G IL I BARGUCHI , J. I. & P EUCAT , J. J. 2001. A single 1992. Mantle derived, early Paleozoic A-type meta-
Devonian subduction event for the HP/HT meta- granitoids from the NW Iberian massif: Nd isotope
morphism of the Cabo Ortegal Complex within the and trace-element constraints. Bulletin de la Societe
Iberian Massif. Tectonophysics, 332, 359– 385. Geologique de la France, 163, 483– 494.
O RTEGA -G UTIÉRREZ , F., R UIZ , J. E. & C ENTENO - P IN , C., P AQUETTE , J. L., S ANTOS Z ALDUEGUI , J. F. &
G ARCÍA , E. 1995. Oaxaquia, a Proterozoic microcon- G IL I BARGUCHI , J. I. 2002. Devonian supra-
tinent accreted to North America during the late Paleo- subduction zone ophiolite related to incipient colli-
zoic. Geology, 23, 1127–1130. sional processes in the Western Variscan Belt: the
O RTEGA -G UTIÉRREZ , F., E LÍAS -H ERRERA , M., Sierra de Careón Unit, Órdenes Complex, Galicia. In:
R EYES -S ALAS , M., M ACÍAS -R OMO , C. & L ÓPEZ , M ARTÍNEZ C ATALÁN , J. R., H ATCHER , R. D.,
R. 1999. Late Ordovician-Early Silurian continental A RENAS , R. & D ÍAZ G ARCÍA , F. (eds) Variscan-
collision orogeny in southern Mexico and its bearing Appalachian Dynamics: The Building of the Late
on Gondwana-Laurentia connections. Geology, 27, Paleozoic Basement. Geological Society of America,
719– 722. Special Paper, 364, 57– 71.
O RTEGA -O BREGÓN , C. & K EPPIE , J. D. ET AL . 2003. P IN , C., P AQUETTE , J. L., Á BALOS , B., S ANTOS
Geochronology and geochemistry of the 917 Ma, Z ALDUEGUI , J. F. & G IL I BARGUCHI , J. I. 2006.
calc-alkaline Etla granitoid pluton (Oaxaca, southern Composite origin of an early Variscan transported
México): evidence of post-Grenvillian subduction suture: ophiolitic units of the Morais Nappe Complex
along the northern margin of Amazonia. International (north Portugal). Tectonics, 25, TC5001, doi:
Geology Review, 45, 596– 610. 10.1029/2006TC001971.
P ATCHETT , P. J. & R UIZ , J. 1987. Nd isotopic ages of P IN , C., F ONSECA , P. E., P AQUETTE , J.-L., C ASTRO , P. &
crustal formation and metamorphism in the Precam- M ATTE , P. 2008. The c. 350 Ma Beja Igneous
brian of eastern and southern Mexico. Contributions Complex: a record of transcurrent slab break-off in
to Mineralogy and Petrology, 96, 523 –528. the Southern Iberia Variscan Belt? Tectonophysics,
P ÉREZ -E STAÚN , A., B ASTIDA , F., M ARTÍNEZ - 461, 356 –377; doi: 10.1016/j.tecto.2008.06.001.
C ATALÁN , J. R., G UTIÉRREZ -M ARCO , J. C., P ROENZA , J. A., O RTEGA -G UTIÉRREZ , F., C AMPRUBÍ ,
M ARCOS , A. & P ULGAR , J. A. 1990. West Asturias- A., T RITLLA , J., E LÍAS -H ERRERA , M. & R EYES -
Leonese Zone: Stratigraphy. In: D ALLMEYER , R. D. S ALAS , M. 2004. Paleozoic serpentinite-enclosed
& M ARTÍNEZ -G ARCÍA , E. (eds) Pre-Mesozoic chromitites from Tehuitzingo (Acatlán Complex,
Geology of Iberia. Springer Verlag, Berlin, 92–102. southern Mexico): a petrological and mineralogical
P EUCAT , J. J., B ERNARD -G RIFFITHS , J., G IL study. Journal of South America Earth Sciences, 16,
I BARGUCHI , J. I., D ALLMEYER , R. D., M ENOT , 649–666.
R. P., C ORNICHET , J. & I GLESIAS P ONCE DE L EÓN , P USCHMANN , H. 1967. Zum Problem derSchichlücken
M. 1990. Geochemical and geochronological cross im Devon der Sierra Morena (Spanien). Geologische
section of the deep Variscan crust: the Cabo Ortegal Rundschau, 56, 528 –542.
Q UESADA , C. 1990. Precambrian terranes in the R OBARDET , M. & G UTIÉRREZ -M ARCO , J. C. 2004.
Iberian Variscan foldbelt. In: S TRACHAN , R. A. & The Ordovician, Silurian and Devonian sedimentary
T AYLOR , G. K. (eds) Avalonian and Cadomian rocks of the Ossa-Morena Zone (SW Iberian
Geology of the North Atlantic. Blackie and Son, Peninsula, Spain). Journal of Iberian Geology, 30,
Glasgow, 109–133. 73–92.
Q UESADA , C. 1991. Geological constraints on the Paleo- R OBERTS , D. & G EE , D. G. 1985. An introduction to
zoic tectonic evolution of tectonostratigraphic terranes the structure of the Scandinavian Caledonides. In:
in Iberia. Tectonophysics, 185, 225–245. G EE , D. G. & S TURT , B. A. (eds) The Caledonide
Q UESADA , C. 1998. A reappraisal of the structure of the Orogen– Scandinavia and Related Areas. Wiley,
Spanish segment of the Iberian pyrite belt. Mineralium Chichester, 55– 68.
Deposita, 33, 31–44. R OBERTS , D. & S TEPHENS , M. G. 2000. Caledonian
Q UESADA , C. & D ALLMEYER , R. D. 1994. Tectonother- Orogenic Belt. In: L UNDQVIST , T. & A UTIO , S.
mal evolution of the Badajoz– Córdoba shear zone (eds) Description to the Bedrock Map of Central
(SW Iberia): characteristics and 40Ar/39Ar mineral Fennoscandia (Mid-Norden). Geological Survey of
age constraints. Tectonophysics, 231, 195– 213. Finland, Special Papers, 28, 79– 108.
Q UESADA , C., B ELLIDO , F. ET AL . 1991. Terranes within R OBERTS , S., A NDREWS , J. R., B ULL , J. M. & S ANDER-
the Iberian Massif: correlations with West African SON , D. J. 1993. Slow-spreading ridge-axis tectonics:
sequences. In: D ALLMEYER , R. D. & L ECORCHÉ , J. evidence from the Lizard complex, UK. Earth and
P. (eds) The West African Orogens and Circum– Planetary Science Letters, 116, 101–112.
Atlantic Correlations. Springer-Verlag, Berlin, R ODRÍGUEZ A LLER , J. 2005. Recristalización y defor-
267–294. mación de litologı́as supracorticales sometidas a mata-
Q UESADA , C., S ÁNCHEZ -G ARCÍA , T., B ELLIDO , F., morfismo de alta presión (Complejo de Malpica-Tuy,
L ÓPEZ -G UIJARRO , R., A RMENDÁRIZ , M. & B RAID , NO del Macizo Ibérico). Laboratorio Xeolóxico de
J. 2006. Field trip guide (Spain). Introduction: the Laxe, Serie Nova Terra, 29, 1– 572.
Ossa– Morena Zone – from Neoproterozoic arc R ODRÍGUEZ , J., C OSCA , M. A., G IL I BARGUCHI , J. I. &
trough Early Paleozoic rifting to late Paleozoic D ALLMEYER , R. D. 2003. Strain partitioning and
orogeny. In: P EREIRA , M. F. & Q UESADA , C. (eds) preservation of 40Ar/39Ar ages during Variscan
Ediacaran to Visean crustal growth processes in the exhumation of a subducted crust (Malpica-Tui
Ossa–Morena Zone (SW Iberia). Évora Meeting complex, NW Spain). Lithos, 70, 111– 139.
2006, Conference abstracts and field trip guide. Insti- R OGER , F. & M ATTE , P. 2005. Early Variscan HP
tuto Geológico Minero de España, 51–73. metamorphism in the western Iberian Allochthon-A
R AMÍREZ E SPINOZA , J. 2001. Tectono-magmatic evol- 390 Ma U– Pb age for the Braganca eclogite (NW
ution of the Paleozoic Acatlán Complex, southern Portugal). International Journal of Earth Sciences,
Mexico and its correlation with the Appalachian 94, 173 –179.
system. Unpublished PhD thesis, University of R UBIO P ASCUAL , F. J., A RENAS , R., D ÍAZ G ARCÍA , F.,
Arizona, 1 –170. M ARTÍNEZ C ATALÁN , J. R. & A BATI , J. 2002.
R AMOS -A RIAS , M. A., K EPPIE , J. D., O RTEGA -R IVERA , Eclogites and eclogite– amphibolites from the San-
A. & L EE , J. W. K. 2008. Extensional deformation on tiago Unit Ordenes Complex, NW Iberian Massif,
the western margin of Pangea, Patlanoaya area, Spain): a case study of contrasting high-pressure
Acatlán Complex, southern Mexico. Tectonophysics, metabasites in a context of crustal subduction. In:
448(1–4), 60–76. M ARTÍNEZ C ATALÁN , J. R., H ATCHER , R. D.,
R IBEIRO , M. L. & F LOOR , P. 1987. Magmatismo A RENAS , R. & D ÍAZ G ARCÍA , F. (eds) Variscan–
peralcalino no Maciço Hespérico: sua distribuçao e Appalachian Dynamics: The Building of the Late
significado geodinámico. In: B EA , F., C ARNICERO , Paleozoic Basement. Geological Society of America,
A., G ONZALO , J. C., L ÓPEZ -P LAZA , M. & Special Papers, 364, 105– 124.
R ODRÍGUEZ A LONSO , M. D. (eds), Geologı́a de los R UIZ , J., P ATCHETT , P. J. & O RTEGA -G UTIÉRREZ , F.
granitoides y rocas asociadas del Macizo Hespérico. 1988. Proterozoic and Phanerozoic basement terranes
Rueda, Madrid, 211– 221. of Mexico from Nd isotopic studies, Geological
R IBEIRO , A., P EREIRA , E. & D IAS , R. 1990. Structure Society of America Bulletin, 100, 274–281.
in the NW of the Iberian Peninsula. In: D ALLMEYER , S ÁNCHEZ M ARTÍNEZ , S., A RENAS , R., D ÍAZ G ARCÍA ,
R. D. & M ARTÍNEZ -G ARCÍA , E. (eds) Pre-Mesozoic F., M ARTÍNEZ C ATALÁN , J. R. & G ÓMEZ -
Geology of Iberia. Springer Verlag, Berlin, 221– 236. B ARREIRO , J. 2007. Careón ophiolite, NW Spain:
R OBARDET , M. 2002. Alternative approach to the suprasubduction zone setting for the youngest Rheic
Variscan Belt in Southwestern Europe: preorogenic Ocean floor. Geology, 7, 53– 56; doi: 10.1130/
paleobiogeographical constraints. In: M ARTÍNEZ - G23024A.
C ATALÁN , J. R., H ATCHER , R. D., A RENAS , R. & S ÁNCHEZ -G ARCÍA , T., B ELLIDO , F. & Q UESADA , C.
D ÍAZ G ARCÍA , F. (eds) Variscan– Appalachian 2003. Geodynamic setting and geochemical signatures
Dynamics: The Building of the Late Paleozoic Base- of Cambrian-Ordovician rift-related igneous rocks
ment. Geological Society of America, Special Paper, (Ossa-Morena Zone, SW Iberia). Tectonophysics,
364, 1 –15. 365, 233–255.
R OBARDET , M. 2003. The Armorica microplate: fact or S ÁNCHEZ -M ARTÍNEZ , S, J EFFRIES , T., A RENAS , R.,
fiction? Critical review of the concept and contradic- F ERNÁNDEZ -S UÁREZ , J. & G ARCÍA -S ÁNCHEZ , R.
tory paleobiogeographical data. Palaeogeography, 2006. A pre-Rodinian ophiolite involved in the Varis-
Palaeoclimatology & Palaeoecology, 195, 125– 148. can suture of Galicia (Cabo Ortegal Complex, NW
Spain). Journal of the Geological Society, London, Mexico: roots of an orogen. Tectonophysics, 365,
163, 737– 740. 257–282.
S ÁNCHEZ -M ARTÍNEZ , S., A RENAS , R., D ÍAZ G ARCÍA , S TAMPFLI , G. M., VON R AUMER , J. F. & B OREL , G. D.
F., M ARTÍNEZ C ATALÁN , J. R. & G ÓMEZ 2002. Paleozoic evolution of pre-Variscan terranes:
B ARREIRO , J. 2007. Careón ophiolite, NW Spain: from Gondwana to the Variscan collision. Geological
Suprasubduction zone setting for the youngest Rheic Society of America, Special Paper, 364, 263– 280.
Ocean floor. Geology, 35, 53–56. S TAMPFLI , G. M. & B OREL , G. D. 2002. A plate tectonic
S ANDEMAN , H. A., C LARK , A. H., S COTT , D. J. & model for the Paleozoic and Mesozoic constrained
M ALPAS , J. G. 2000. The Kennack Gneiss of the by dynamic plate boundaries and restored synthetic
Lizard Peninsula, Cornwall, S. W. England: comin- oceanic isochrones. Earth and Planetary Science
gling and mixing of mafic and felsic magmas during Letters, 196, 17– 33.
Givetian continental incorporation of the Lizard ophio- S TERN , R. J. 2004. Subduction initiation: spontaneous and
lite. Journal of the Geological Society, London, 157, induced. Earth and Planetary Science Letters, 226,
1227–1242. 275–292.
S ANDEMAN , H. A., C LARK , A. H., S TYLES , M. T., T AIT , J., S CHÄTZ , M., B ACHTADSE , V. & S OFFEL , H.
S COTT , D. J., M ALPAS , J. G. & F ARRAR , E. 1997. 2000. Palaeomagnetism and Palaeozoic palaeoge-
Geochemistry and U–Pb and 40Ar/39Ar geochronol- ography of Gondwana and European terranes. Geologi-
ogy of the Man of War Gneiss, Lizard Complex, cal Society, London, Special Publications, 179, 21–34.
SW England: pre-Hercynian arc-type crust with a T ALAVERA -M ENDOZA , O., R UÍZ , J., G EHRELS , G. E.,
Sudetan-Iberian connection. Journal of the Geological M EZA -F IGUEROA , D. M., V EGA -G RANILLO , R. &
Society, London, 154, 403 –411. C AMPA -U RANGA , M. F. 2005. U–Pb geochronology
S ANTOS Z ALDUEGUI , J. F., S CHÄRER , U. & G IL of the Acatlán Complex and implications for the
I BARGUCHI , J. I. 1995. Isotope constraints on the age Paleozoic paleogeography and tectonic evolution of
and origin of magmatism and metamorphism in the southern Mexico. Earth and Planetary Science
Malpica-Tuy allochthon, Galicia, NW-Spain. Chemi- Letters, 235, 682– 699.
cal Geology, 121, 91– 103. T HIÉBLEMONT , D. & C ABANIS , B. 1986. Découverte
S ANTOS Z ALDUEGUI , F. J., S CHARER , U., G IL I BARGU- d’une association de volcanites d’arc et de basaltes
CHI , J. I. & G IRARDEAU , J. 2002. Genesis of de type ‘MORB’ dans la formation paléo-volcanique
pyroxenite-rich peridotite at Cabo Ortegal (NW silurienne de la Meilleraie, Vendée, France. Comptes
Spain): geochemical and Pb –Sr –Nd isotope data. Rendus Academie de Sciences, Paris, 311(II),
Journal of Petrology, 43, 17–43. 1221– 1226.
S CHÄFER , H. J., G EBAUER , D., G IL I BARGUCHI , J. I. & V ACHARD , D. & F LORES DE D IOS , A. 2002. Discovery
P EUCAT , J. J. 1993. Ion-microprobe U–Pb zircon of latest Devonian/earliest Mississippian microfossils
dating on the HP/HT Cabo Ortegal Complex (Galicia, in San Salvador Patlanoaya (Puebla, Mexico): biogeo-
NW Spain): preliminary results. Terra Abstracts, 5, 22. graphic and geodynamic consequences. Compte Rendu
S CHNEIDER , J. L., H ASSENFORDER , B. & P AICHELER , Geoscience, 334, 1095– 1101.
J.-C. 1990. Une ou plusieurs ‘lignes des klippes’ dans V ALVERDE V AQUERO , P. & F ERNÁNDEZ , F. J. 1996.
les Vosges du Sud (France)? Nouvelles données sur Edad de enfriamiento U– Pb en rutilos del Gneiss de
la nature des ‘klippes’ et leur signification dans la Chimparra (Cabo Ortegal, NO de España). Geogaceta,
dynamique varisque. Comptes Rendus Academie de 20(2), 475– 478.
Sciences, Paris, 311(II), 1221–1226. V ALVERDE -V AQUERO , P. & D UNNING , G. R. 2000. New
S HAW , A., D OWNES , H. & T HIRLWALL , M. F. 1993. The U–Pb ages for Early Ordovician magmatism in Central
quartz-diorites of Limousin: elemental and isotopic Spain. Journal of the Geological Society, London, 157,
evidence for Devono-Carboniferous subduction in the 15–26.
Hercynian belt of the French Massif Central. Chemical V ALVERDE -V AQUERO , P., M ARCOS , A., F ARIAS , P. &
Geology, 107, 1 –18. G ALLASTEGUI , G. 2005. U– Pb dating of Ordovician
S ILVA , J. B., O LIVEIRA , J. T. & R IBEIRO , A. 1990. felsic volcanism in the Schistose Domain of the
Structural outline of the South Portuguese Zone. In: Galicia-Trás-os-Montes Zone near Cabo Ortegal
D ALLMEYER , R. D. & M ARTÍNEZ G ARCÍA , E. (eds) (NW Spain). Geologica Acta, 3(1), 27– 37.
Pre-Mesozoic Geology of Iberia. Springer Verlag, V AN S TAAL , C. R., D EWEY , J. F., M AC N IOCAILL , C. &
348– 362. M C K ERROW , W. S. 1998. The Cambrian-Silurian tec-
S IMANCAS , J. F., G ONZÁLEZ L ODEIRO , F., E XPÓSITO , I., tonic evolution of the Northern Appalachians and
A ZOR , A. & M ARTÍNEZ P OYATOS , D. 2002. Opposite British Caledonides: history of a complex, west and
subduction polarities connected by transform. faults southwest Pacific-type segment of Iapetus. In:
in the Iberian Massif and western European Variscides. B LUNDELL , D. & S COTT , A. C. (eds) Lyell: The
In: M ARTÍNEZ C ATALÁN , J. R., H ATCHER , R. D., Past is the Key to the Present. Geological Society,
A RENAS , R. & D ÍAZ G ARCÍA , F. Variscan– London, Special Publications, 143, 199– 242.
Appalachian dynamics: the building of the late V AN S TAAL , C. R., C URRIE , K. L., R OWBOTHAN , G.,
Paleozoic basement. Geological Society of America, R OGERS , N. & G OODFELLOW , W. 2008. P –T paths
Special Paper, 364, 253– 262. and exhumation of Late Ordovician-Early Silurian
S OLARI , L. A., K EPPIE , J. D., O RTEGA -G UTIÉRREZ , F., blueschists and associated metamorphic nappes of the
C AMERON , K. L., L ÓPEZ , R. & H AMES , W. E. Salinic Brunswick subduction complex, northern
2003. 990 and 1100 Ma Grenvillian tectonothermal Appalachians. Geological Society of America Bulletin,
events in the northern Oaxacan Complex, southern 120, 1455–1477; doi: 10.1130/B26324.1.
V EGA -G RANILLO , R., T ALAVERA -M ENDOZA , O., W INCHESTER , J. A. & F LOYD , P. A. 1977. Geochemical
M EZA -F IGUEROA , D., R UIZ , J., G EHRELS , G. E., discrimination of different magma series and their
L ÓPEZ -M ARTÍNEZ , M. & DE LA C RUZ -V ARGAS , differentiation products using immobile elements.
J. C. 2007. Pressure– temperature– time evolution of Chemical Geology, 20, 325–343.
Paleozoic high-pressure rocks of the Acatlán Complex Y ÁÑEZ , P., R UIZ , J., P ATCHETT , P. J., O RTEGA -
(southern Mexico): implications for the evolution of G UITÉRREZ , F. & G EHRELS , G. 1991. Isotopic
the Iapetus and Rheic Oceans. Geological Society of studies of the Acatlan complex, southern Mexico:
America, Bulletin, 119, 1249– 1264. implications for Paleozoic North America tectonics.
V OGEL , D. E. 1967. Petrology of an eclogite- and Geological Society of America Bulletin, 103, 817–828.
pyrigarnite-bearing polymetamorphic rock complex Y OUNG , T. 1990. Ordovician sedimentary facies and
at Cabo Ortegal, NW Spain. Leidse Geologische Mede- faunas of Southwest Europe: palaeogeographic and
delingen, 40, 121– 213. tectonic implications. In: M C K ERROW , W. S. &
W EBER , B. & K ÖHLER , H. 1999. Sm/Nd, Rb/Sr and S COTESE , C. R. (eds) Palaeozoic Palaeogeography
U–Pb geochronology of a Grenville terrane in and Biogeography. Geological Society, London,
southern Mexico: origin and geologic history of the Memoirs, 12, 421–430.
Guichicovi complex. Precambrian Research, 96, Z ECK , H. P., W HITEHOUSE , M. J. & U GIDOS , J. M. 2007.
245–262. 496 + 3 Ma zircon ion microprobe age for pre-
W EIL , A. B., V AN DER V OO , R. & VAN DER P LUIJM , B. Hercynian granite, Central Iberian Zone, NE Portugal
2001. New paleomagnetic data from the southern (earlier claimed 618 + 9). Geological Magazine,
Cantabria-Asturias Arc, northern Spain: implications 144, 21–31.
for true oroclinal rotation and the final amalgamation Z ELAZNIEWICZ , A., D ORR , W. & D UBINSKA , E. 1998.
of Pangea. Geology, 29, 991– 994. Lower Devonian oceanic crust from U–Pb zircon evi-
W ILLIAMS , H. 1979. Appalachian orogen in Canada. dence and Eo-Variscan event in the Sudetes. In:
Canadian Journal of Earth Sciences, 16, 792– 798. F RANKE , W. & H ERZBERG , S. (eds) SPP Orogene
W ILSON , J. T. 1966. Did the Atlantic Ocean close and then Processe 7. Kolloquium, Giessen. Terra Nostra,
re-open? Nature, 211, 676 –681. 98(2), 173–176.
The palaeomagnetically viable, long-lived and all-inclusive
Rodinia supercontinent reconstruction
DAVID A. D. EVANS
Department of Geology & Geophysics, Yale University, New Haven, CT 06520-8109, USA
(e-mail: dai.evans@yale.edu)
Abstract: Palaeomagnetic apparent polar wander (APW) paths from the world’s cratons at
1300– 700 Ma can constrain the palaeogeographic possibilities for a long-lived and all-inclusive
Rodinia supercontinent. Laurentia’s APW path is the most complete and forms the basis for super-
position by other cratons’ APW paths to identify possible durations of those cratons’ inclusion in
Rodinia, and also to generate reconstructions that are constrained both in latitude and longitude
relative to Laurentia. Baltica reconstructs adjacent to the SE margin of Greenland, in a standard
and geographically ‘upright’ position, between c. 1050 and 600 Ma. Australia reconstructs adja-
cent to the pre-Caspian margin of Baltica, geographically ‘inverted’ such that cratonic portions
of Queensland are juxtaposed with that margin via collision at c. 1100 Ma. Arctic North
America reconstructs opposite to the CONgo þ São Francisco craton at its DAmaride –Lufilian
margin (the ‘ANACONDA’ fit) throughout the interval 1235– 755 Ma according to palaeomag-
netic poles of those ages from both cratons, and the reconstruction was probably established
during the c. 1600– 1500 Ma collision. Kalahari lies adjacent to Mawsonland following collision
at c. 1200 Ma; the Albany– Fraser orogen continues along-strike to the Sinclair-Kwando-Choma-
Kaloma belt of south-central Africa. India, South China and Tarim are in proximity to Western
Australia as previously proposed; some of these connections are as old as Palaeoproterozoic
whereas others were established at c. 1000 Ma. Siberia contains a succession of mainly
sedimentary-derived palaeomagnetic poles with poor age constraints; superposition with the
Keweenawan track of the Laurentian APW path produces a position adjacent to western India
that could have persisted from Palaeoproterozoic time, along with North China according to its
even more poorly dated palaeomagnetic poles. The Amazonia, West Africa and Rio de la Plata
cratons are not well constrained by palaeomagnetic data, but they are placed in proximity to
western Laurentia. Rift successions of c. 700 Ma in the North American COrdillera and
BRAsiliano-Pharuside orogens indicate breakup of these ‘COBRA’ connections that existed for
more than one billion years, following Palaeoproterozoic accretionary assembly. The late Neopro-
terozoic transition from Rodinia to Gondwanaland involved rifting events that are recorded
on many cratons through the interval c. 800 –700 Ma and collisions from c. 650– 500 Ma. The
pattern of supercontinental transition involved large-scale dextral motion by West Africa
and Amazonia, and sinistral motion plus rotation by Kalahari, Australia, India and South China,
in a combination of introverted and extroverted styles of motion. The Rodinia model presented
here is a marked departure from standard models, which have accommodated recent discordant
palaeomagnetic data either by excluding cratons from Rodinia altogether, or by decreasing
duration of the supercontinental assembly. I propose that the revised model herein is the only poss-
ible long-lived solution to an all-encompassing Rodinia that viably accords with existing
palaeomagnetic data.
Motivation information: for Pangaea by seafloor magnetic ano-

malies with a supportive record from continent-
A full understanding of ancient orogens requires based palaeomagnetic studies (e.g. Torsvik et al.
an accurate palaeogeographic framework. By their 2008), and for times before Pangaea thus far only
nature, orogens destroy prior geologic information by the labourious continent-based method.
(via metamorphism, erosion and subduction) and Most reconstructions of the early Neoprotero-
thus challenge our efforts to reconstruct their his- zoic supercontinent Rodinia (Fig. 1), involving
tories. Reconstructing supercontinents is the great- connections between western North America and
est palaeogeographic challenge of all, combining Australia þ Antarctica (Moores 1991; Karlstrom
patchworks of this partially destroyed information et al. 1999; Burrett & Berry 2000; Wingate et al.
from a series of orogens with a complementary 2002) and eastern North America adjacent to
record of rifting and passive margin development, Amazonia (Dalziel 1991; Hoffman 1991; Sadowski
and with quantitative kinematic data. The latter & Bettencourt 1996), have in the former case
record is best constrained by palaeomagnetic been negated or superseded by subsequent
372 D. A. D. EVANS
geochronological and palaeomagnetic results Rainbird et al. 1998; Sears & Price 2003) are incom-
(Pisarevsky et al. 2003a, b), and in the latter patible with recent palaeomagnetic data (reviewed
instance suffered from minimal palaeomagnetic by Pisarevsky & Natapov 2003), as are reconstruc-
support (Tohver et al. 2002, 2006). Configurations tions that directly juxtapose Laurentia and Kalahari
that directly adjoin Laurentia and Siberia (e.g. (Hanson et al. 2004). Kalahari and the composite
Fig. 1. (a) Base maps for Rodinian cratons, reconstructed to their relative early Jurassic Pangaea configuration in
present South American co-ordinates. Gondwanaland fragments are reconstructed according to McElhinny et al.
(2003); these and other Euler parameters are listed in Table 1. North and South China, although distinct cratons in
Rodinia, are united in this Early Jurassic configuration as indicated by Yang & Besse (2001). Late Mesoproterozoic
(‘Grenvillian’) orogens are shaded grey. Truncated Mercator projection. Panels (b– d) show previous Rodinia models in
the present North American (i.e. Laurentian) reference frame. References are Dalziel (1997), Pisarevsky et al. (2003a)
and Li et al. (2008).
RODINIA SUPERCONTINENT RECONSTRUCTION 373
374 D. A. D. EVANS
Congo þ São Francisco craton have been suggested puzzle is tantalizingly solvable. Most of the geologi-
to be excluded from Rodinia altogether (Kröner & cal comparisons are purely of regional extent, con-
Cordani 2003; Cordani et al. 2003). sidering only two or three cratons (examples cited
Four general classes of options exist for how to above), generally within the context of Hoffman’s
deal with discrepant palaeomagnetic data: (1) con- (1991) global model. In these geologically-based
sider potential shortcomings in those data, regarding juxtapositions, palaeomagnetic data have been
either palaeomagnetic reliability or age constraints; used in a subsidiary capacity, or ignored altogether,
(2) add loops in the APW path to accommodate the despite the fact that palaeomagnetism is currently
outlying poles; (3) exclude cratons from member- the only strictly quantitative method available for
ship in Rodinia entirely; or (4) restrict the duration reconstructing Rodinia and earlier supercontinents.
of that membership so that it falls between the Although in a global sense Rodinia assembled
discrepant palaeomagnetic pole ages. The standard in the late Mesoproterozoic and fragmented in
Rodinia model (Hoffman 1991) and its relatively the mid-Neoproterozoic (Hoffman 1991; Dalziel
minor variations have been strained to the limits 1997; Condie 2002), thereby existing through the
of temporal and spatial constraints, so that some interval 1000– 800 Ma, there are numerous indica-
have questioned whether Rodinia even existed at tions of locally earlier assembly or later breakup.
all (Meert & Torsvik 2003); and yet there is still For example, only one side of Laurentia was
the persistent global tectonostratigraphic evidence deformed by orogeny in the late Mesoproterozoic:
for late Mesoproterozoic convergence of cratons, the Grenvillian (¼ proto-Appalachian) margin, and
followed by mid-late Neoproterozoic rifting and this belt did not evolve to a rifted passive margin
passive margin development. Given the abundance until after 600 Ma (Cawood et al. 2001). Northern
of focused studies yielding a wealth of new tectono- Laurentia experienced the c. 1600 Ma Forward
stratigraphic, geochronological and palaeomagnetic Orogeny (Maclean & Cook 2004) followed by
data during the last two decades (summarized by extensional events with associated large igneous
Pisarevsky et al. 2003a; Meert & Torsvik 2003; Li provinces at 1270 Ma (LeCheminant & Heaman
et al. 2008), the search for Rodinia may benefit 1989) and 720 Ma (Heaman et al. 1992). The
from an entirely fresh perspective. Here I introduce western margin assembled in the Palaeoproterozoic
a novel, long-lived Rodinia that is compatible with (Karlstrom et al. 2001; Ross 2002) and did not
the most reliable palaeomagnetic data from all of rift until the middle or latest Neoproterozoic
the dozen or so largest cratons during the interval (Link et al. 1993; Colpron et al. 2002). Rifting of
1300–700 Ma, with minor allowances on the ages Rodinia along these northern and western Lauren-
of a single set of poles from the São Francisco tian margins, then, split the proto-Laurentian con-
craton. I show that given these palaeomagnetic tinent through terrains that had been joined for
data, the new reconstruction is the only general about a billion years. In these instances, if we can
model of Rodinia that could have existed for this find the correct Rodinia juxtapositions, we have
length of time with all of the largest cratons included also solved part of the configuration of Nuna,
in its assembly. The model serves as a palaeo- which is Rodinia’s Palaeoproterozoic superconti-
magnetic end-member starting point for further nental predecessor (Hoffman 1996). Many other
testing and, if desired, relaxation on the assumptions examples of this type exist around the world, essen-
of longevity or inclusion of all the largest cratons. tially wherever a Neoproterozoic rifted margin does
The present analysis is thus most similar to that of not coincide with a ‘Grenvillian’ orogen (Fig. 1).
Weil et al. (1998) in seeking a unified Rodinia When we test Rodinia models with palaeomagnetic
model that conforms to the original concept of late data, therefore, we must in some cases consider
Mesoproterozoic assembly and mid-Neoproterozoic results from rocks as old as c. 1800 Ma (e.g.
dispersal, while incorporating all of the most Idnurm & Giddings 1995).
reliable palaeomagnetic data. Axisymmetry of the Earth’s time-averaged geo-
magnetic field implies that when individual palaeo-
magnetic poles from two continents are compared,
Methods their relative palaeolongitude remains uncon-
strained. This shortcoming to palaeomagnetically-
Many Rodinia reconstructions have been based based palaeogeographic reconstructions has led
primarily on comparisons of the geological records to illustrations of Rodinia and older supercontinents
among Meso –Neoproterozoic cratons. For this that show only a set of latitude-constrained options,
time interval, we can identify 13 large cratons further unconstrained by the unknown geomag-
(Fig. 1), plus many smaller terranes (e.g. Kolyma, netic polarity states of the compared palaeomag-
Barentsia, Oaxaquia, Yemen). With only a dozen netic data, a degree of freedom for nearly every
or so large pieces and an abundant well-preserved Precambrian reconstruction (Hanson et al. 2004).
Meso-Neoproterozoic rock record, the Rodinia Among these degrees of freedom in palaeolongitude
and hemispheric ambiguity, two or more cratons A quantitatively viable Rodinia may be found by
are juxtaposed in several allowable positions of investigating possible APW superpositions and
direct contact for the specific age of pole compari- determining whether the resulting juxtapositions
son. If a similar reconstruction emerges from are geologically reasonable for the time intervals
several adjacent time slices, then a long-lived under consideration. This method requires equal
direct connection between the cratons can be con- APW track lengths between coeval poles on any
sidered viable. Examples of this method, called two given cratons; thus it is conceivable that no
the ‘closest approach’ technique, are found in APW comparisons will be possible between those
Buchan et al. (2000, 2001), Meert & Stuckey blocks and that they must have been in relative
(2002) and Pesonen et al. (2003). motion throughout the interval under consideration.
A more powerful method of reconstructing Likewise, there is no guarantee that direct juxta-
ancient supercontinents relies on the coherent positions of cratons will emerge: some pole com-
motion of all component cratons as part of that parisons may result in substantial or complete
supercontinent, for the duration of their conjunction geographic overlap between two or more cratons,
within a single lithospheric plate. Throughout which are unallowable, and others may indicate
the time interval when constituent cratons are wide separations between blocks, requiring the pres-
assembled into a supercontinent, and if that assem- ence of intervening blocks (or occurrence of rapid
blage is in motion relative to the Earth’s magnetic true polar wander; see Evans (2003) to legitimize
field reference frame (due to plate tectonics or the initial hypothesis of common APW.
true polar wander, or both), then all elements of Accurate Neoproterozoic craton outlines are
the landmass will share the same palaeomagnetic important not only for correct geometric fits in
APW path. After the supercontinent disaggregates, Rodinia reconstructions, but they also indicate
the APW paths diverge (Powell et al. 1993), whether certain palaeomagnetic results from mar-
but their older segments carry a record of the ginal foldbelts apply to a craton or to its alloch-
earlier supercontinental motion. As we approach thonous terranes. Cratonic outlines, drawn in
the problem from the present, we see that each accordance with a broad range of tectonic and strati-
craton’s APW path contains segments alternating graphic studies that are too numerous to cite here,
between times of individual plate motion and mem- are generally chosen to lie within craton-marginal
bership in successive supercontinents. When the orogens at the most distal extent of recognizable
cratons are reconstructed to their correct positions stratigraphic connections to each adjacent block.
in a supercontinent, the APW paths superimpose Cratons that have split into fragments during the
atop one another (Evans & Pisarevsky 2008). breakup of Pangaea (e.g. Laurentia þ Greenland þ
Examples of this type of analysis are found in Rockall, or Kalahari þ Falkland þ Grunehogna þ
Weil et al. (1998) and Piper (2000), although both Ellsworth) must first be reassembled according to
of those studies preceded important new palaeo- seafloor-spreading data combined with geological
magnetic data that disallow some of their cratonic ‘piercing points.’ Post-Pangaean fragments are
juxtapositions. The modified Palaeopangaea recon- restored to each other according to standard recon-
struction of Piper (2007) achieves broad-brush structions (Table 1), with the exception of Kalahari:
palaeomagnetic APW concordance among several following restoration of the Falkland Islands
cratons, merely as a result of pole averaging (e.g. (Grunow et al. 1991), Grunehogna is reconstructed
Siberia), misquoted ages (e.g. Bangemall sills of to align the Natal and Maud orogenic fronts in
Australia), or rotation parameters yielding some- the manner suggested by Jacobs & Thomas
what acceptable pole matches but differing dramati- (2004), and the Ellsworth þ Haag province is then
cally from the simple cartoon depiction of the rotated to fit into the Natal embayment. The Siberian
reconstruction (e.g. Amazonia, São Francisco þ craton shows restoration of a 258 internal rotation
Plata, West Africa and Tanzania þ Kalahari) or between its northwestern and southeastern (Aldan)
even producing unacceptable geometric overlaps portions, associated with development of the Devo-
(northern Australia directly atop Kalahari, and por- nian Vilyuy aulacogen, to resolve discrepancies
tions of North China directly atop eastern India, in older palaeomagnetic data (Smethurst et al.
in the ‘primitive’ or pre-1100 Ma reconstruction). 1998; Gallet et al. 2000). Craton boundaries in Ant-
As discussed below, Laurentia has the most com- arctica are particularly uncertain, and the present
plete palaeomagnetic APW path for the interval of analysis uses conservative estimates of minimal
c. 1300 –750 Ma that is most relevant for testing areas attached to each block. Smaller blocks
Rodinia reconstructions. In this paper I use the with limited to no palaeomagnetic data, such
most reliable palaeomagnetic poles from non- as Precordillera– Cuyania, Oaxaquia, Barentsia,
Laurentian cratons to compare with the Laurentian Azania and various poorly exposed blocks in
reference APW path and thereby to constrain the South America (Dalziel 1997; Collins & Pisarevsky
possible configurations of a long-lived Rodinia. 2005; Fuck et al. 2008), are not described in
376 D. A. D. EVANS
Table 1. Pre-Mesozoic reassemblies of Rodinian cratons
Craton fragment Euler rotn. Reference
8N 8E 8CCW
Rotations to Laurentia
Greenland 67.5 241.5 213.8 Roest & Srivastava 1989
Rockall Plateau 75.3 159.6 223.5 Srivastava & Roest 1989;
Royer et al. 1992
Rotations to Kalahari
Falkland Islands 245.3 349.2 156.3 Grunow et al. 1991
Grunehogna 205.3 324.5 58.6 After Jacobs & Thomas 2004
Ellsworth-Haag 248.9 102.8 82.8 Geometric fit
Rotation to Congo
São Francisco 46.8 329.4 55.9 McElhinny et al. 2003
Rotation to West Africa
São Luis 53.0 325.0 51.0 McElhinny et al. 2003
Rotations to India
Enderby Land 204.8 016.6 93.2 McElhinny et al. 2003
Eastern Madagascar 18.8 026.3 62.1 McElhinny et al. 2003
Southern Somalia 28.9 040.9 64.8 McElhinny et al. 2003
Rotation to Australia
Terre Adélie 01.3 037.7 30.3 McElhinny et al. 2003
Reconstruction of Siberia
NW Siberia to Aldan shield 60.0 115.0 225.0 Fit pre-Devonian poles from
Smethurst et al. 1998;
Gallet et al. 2000
detail but are mentioned below where appropriate. it adequately represents the general trend with
Cratons and palaeomagnetic poles are rotated to which the most important poles from other cratons
geometric accuracy via the software created by may be compared. The sense of vorticity of the
Cogné (2003). All calculations assume a geocentric Grenville APW loop, at c. 1000 Ma, has been
axial-dipolar magnetic field, recently verified for debated (Weil et al. 1998). The present compilation
the Proterozoic using a compilation of evaporite follows Pisarevsky et al. (2003a) in selecting the
palaeolatitudes that gave subtropical values as most reliable (Q . 3 in the scheme of Van der
expected (Evans 2006) and a planetary sphere of Voo 1990) results from late Keweenawan sedimen-
constant radius. tary rocks, in stratigraphic order, which generates a
southward leg of the loop at c. 1808 longitude,
followed by the well-dated Haliburton ‘A’ pole at
Laurentia 1015+15 Ma (Warnock et al. 2000), and then by
a northward leg at ,1808E longitude. This clock-
Reliable Precambrian palaeomagnetic data are wise sense of the Grenville loop is compatible
currently so sparse that in only a few instances can with the earlier interpretation of Hyodo & Dunlop
we assemble poles into coherent APW paths. In (1993) but not that of Weil et al. (1998). Another
the Rodinia interval, only Laurentia has a well- set of reliable Laurentian poles is determined for
established path, with ages of c. 1270–1000 Ma the interval 780–720 Ma, summarized by Buchan
and tracking younger, with imprecise cooling ages et al. (2000) and Pisarevsky et al. (2003a), plus
from the Grenville Province (Fig. 2; Weil et al. more recent data from stratified successions in
1998; Pisarevsky et al. 2003a). The APW path western United States (e.g. Weil et al. 2004,
shown in Figure 2 also includes the 1750 Ma 2006). Within the intervening 200 Ma gap of no
grand mean of Irving et al. (2004) and representa- ‘key’ poles, a recent result from poorly dated
tive ‘key’ poles from c. 1450 Ma, as listed by but palaeomagnetically stable sedimentary rocks
Buchan et al. (2000). Although this is not a comin Svalbard (including a positive soft-sediment
plete set of data from the 1750–1270 Ma interval, slump fold test guaranteeing primary remanence)
Fig. 2. Apparent polar wander path for Laurentia between 1270 and 720 Ma (poles listed in Table 2), which is used as a
reference curve for superimposing ‘key’ poles from other cratons during the Rodinian interval (Buchan et al. 2001).
Note the large age gap between 1015 and 780 Ma, which is hypothesized here to include a large APW loop at 800 Ma
as observed from other cratons (see text and Figs 8 & 10) and from recent high-quality data from Svalbard (Maloof
et al. 2006).
suggests a large APW loop, hitherto unrecognized been suggested with minor variations throughout
for early Neoproterozoic Laurentia (Maloof et al. the last three decades (Patchett et al. 1978;
2006). Regarding these results, it is important to Piper 1980; Bond et al. 1984; Gower et al. 1990;
note that this loop is underpinned by data from Hoffman 1991; Dalziel 1997; Weil et al. 1998;
continuous stratigraphic sections in Svalbard, Hartz & Torsvik 2002; Pisarevsky et al. 2003a;
thus eliminating uncertainties in the reconstruction Cawood & Pisarevsky 2006). The principal vari-
of Svalbard to Laurentia, or local rotations, as ations among these reconstructions are the latitude
trivial explanations for the divergent pole positions. of juxtaposition along the Greenland margin and
Global correlations of this new, c. 800 Ma, APW the orientation of Balitca such that various
loop are discussed in various sections below with margins (e.g. Caledonide, Timanian, Uralian) are
the relevant data from other cratons. proposed to participate in the direct conjunction
Lack of data from the c. 1000– 800 Ma interval with Greenland (Buchan et al. 2000; Cawood &
of the Laurentian APW path renders many Rodinian Pisarevsky 2006). The reconstruction favoured
cratonic juxtapositions currently untested; for here is nearly identical to that proposed by
example, the various reconstructions of Australia þ Pisarevsky et al. (2003a), but with a tighter
Mawsonland against particular segments of the fit. Pisarevsky et al. (2003a) opted for a several
Laurentian Cordilleran margin (SWEAT, hundred-kilometre gap between the present-day
AUSWUS, AUSMEX) all fail palaeomagnetic margins of SE Greenland and the Norwegian
comparisons at c. 750 Ma (Wingate & Giddings Caledonides, in order to account for palinspastic
2000) and c. 1200 Ma (Pisarevsky et al. 2003b), restoration of Caledonide shortening. These same
but any one of those reconstructions could margins, however, experienced a counteracting
be salvaged if it assembled after c. 1000 Ma amount of extension during Cenozoic initiation of
and fragmented by c. 800 Ma. In this option for Atlantic Ocean opening as well as Eocambrian
dealing with the discrepant palaeomagnetic data as opening of Iapetus; the three post-Rodinian altera-
described above, only Li et al. (1995, 2002, 2008) tions to the Greenland and Baltic margins may
have developed a tectonically reasonable hypo- well have nullified one another, so a tight fit is
thesis by inserting the South China block between preferred here (Fig. 3).
Laurentia and Australia. In that model, the Palaeomagnetic poles within the 1100–850 Ma
Sibao orogen represents the suture between interval from Baltica broadly superimpose onto
the Australia þ Yangtze craton and Cathaysia þ the Laurentian Grenville APW loop when the two
Laurentia in a collision at c. 900 Ma (Li et al. 2008). cratons are restored to their proposed Rodinian
reconstruction (Fig. 3). The distribution of Baltic
poles from this interval has been described
Baltica in terms of a so-called Sveconorwegian APW
loop, with discussions on the ages of individual
The least controversial component of the revised poles and whether a complete loop is actually cir-
Rodinia reconstruction proposed herein is the place- cumscribed by the data (Walderhaug et al. 1999;
ment of Baltica adjacent to eastern Laurentia, as has Brown & McEnroe 2004; Pisarevsky & Bylund
378 D. A. D. EVANS
Fig. 3. Palaeomagnetic poles from Baltica (grey), compared with the Laurentian reference APW path. As with Figures
4– 11, dark-grey shaded cratonic areas are late Mesoproterozoic (‘Grenvillian’) orogens. (a) Poles with their 95%
confidence ellipses and ages in Ma, as listed in Table 2, plus Evans & Pisarevsky (2008) for the 1780– 1270 Ma data. (b)
Superposition of the 1780– 1270 Ma poles with the pre-Rodinian APW path from Laurentia, rotating Baltica and its
poles by (47.58N, 001.58E, þ498CCW) after Evans & Pisarevsky (2008). This pre-Rodinia configuration is essentially
the same as ‘NENA’ by Gower et al. (1990). (c) Rodinian reconstruction (1070– 615 Ma) that superimposes the
Sveconorwegian APW loop atop the Grenville APW loop of Laurentia (despite recently recognized age mismatches),
using the Euler rotation parameters given in Table 3.
2006). Here I tentatively adopt the simple expla- to all other proposed Rodinian juxtapositions
nation postulated by Pisarevsky & Bylund (2006) of Laurentia and Baltica (Cawood & Pisarevsky
that the high-latitude Sveconorwegian poles rep- 2006).
resent a single, post-900 Ma overprint affecting Palaeomagnetic poles from Baltica and Lauren-
the southernmost regions of Norway and Sweden, tia during the preceding interval c. 1750–1270 Ma
despite lack of independent supporting evidence are incompatible with the preferred c. 1100 –
for such an event (Brown & McEnroe 2004). As a 600 Ma reconstruction, and suggest instead a modi-
more complex alternative, there might be several fied fit with Baltica’s Kola– Timanian margin
oscillatory ‘Sveconorwegian’ loops in the Baltic adjacent to East Greenland (Fig. 3). This fit is essen-
APW path, which would be geodynamically tially the geologically-based Northern Europe þ
explained best by multiple episodes of true polar North America (NENA) reconstruction of Gower
wander (TPW; Evans 2003). More detailed palaeo- et al. (1990), confirmed palaeomagnetically by
magnetic and thermochronological studies of the Buchan et al. (2000) and Evans & Pisarevsky
two cratons from this time interval are needed to (2008) for the pre-Rodinian interval. Relative to
resolve these questions. Laurentia, Baltica rotated clockwise c. 708 about
The proposed reconstruction of Baltica adjacent an Euler pole near Scoresby Sund, some time
to SE Greenland at c. 1100–600 Ma, like that of between 1270 and 1050 Ma, in approximately the
Pisarevsky et al. (2003a), brings the Sveconorwe- same sense as was first proposed by Patchett et al.
gian orogen in southern Scandinavia close to the (1978) and Piper (1980). New palaeomagnetic
Grenville orogen in Labrador with minor right- results from the 1122 Ma Salla Dyke in northern
stepping offset (Gower et al. 2008). It also unites Finland are more compatible with a pre-rotation
the loci of precisely coeval 615 Ma Long Range reconstruction than a post-rotation reconstruction,
dykes in Labrador (Kamo et al. 1989; Kamo & suggesting that the rotation occurred after, or even
Gower 1994) and Egersund dykes in southernmost coincident with, dyke emplacement at c. 1120 Ma
Norway (Bingen et al. 1998). Palaeomagnetic (Salminen et al. 2009). The proposed rotation is
results from both of these dyke swarms have consistent with the broad-scale tectonic asymmetry
yielded scattered results spanning a wide range of of Baltica (orogeny in west, rifting in east) through
inclinations, rendering palaeolatitude comparisons the Mesoproterozoic interval (Bogdanova et al.
difficult (Murthy et al. 1992; Walderhaug et al. 2008). Below it will be shown how this rota-
2007); however, they are as consistent with the tion created a broad gulf along the edge of the
reconstruction introduced here as they are for that Rodinia-encircling ocean, Mirovia (McMenamin
of Pisarevsky et al. (2003a) and Li et al. (2008), & McMenamin 1990), which became an isolated
and this general class of reconstructions is superior sea following further Rodinian amalgamation.
Australia 1 Mawsonland 2002), and the Mundine Well Dykes at 755 Ma

(Wingate & Giddings 2000). The latter result is
The semi-contiguous Albany– Fraser and Musgrave supplemented by a pole from oriented borehole
belts are commonly considered as part of a late core of the Browne Formation (estimated age
Mesoproterozoic suture zone among three cons- c. 830–800 Ma), which is the only result among
tituent Australian cratons (western, northern and several reported by Pisarevsky et al. (2007) with
southern; Myers et al. 1996), or between a pre- adequate statistics on the mean direction. Other
viously united western þ northern craton (Li 2000) palaeomagnetic poles from Australia during the
and the southern, ‘Mawson Continent’ (Cawood & Meso –Neoproterozoic interval are problematic, as
Korsch 2008). The latter entity extends from the discussed by Wingate & Evans (2003): they suffer
Australian Gawler craton s.s. into Terre Adélie in from any combination of poor geochronology,
Antarctica, and possibly as far south as the lack of tilt control, and unknown timing of the
Transantarctic Mountains near the Miller Range magnetic remanence acquisition. Similarly, a more
(Goodge et al. 2001; Fitzsimons 2003; Payne recent result from the Alcurra dykes in the Mus-
et al. 2009). Here the term ‘Mawsonland’ is for- grave belt (Schmidt et al. 2006) also suffer from
mally introduced as a more succinct synonym to lack of tectonic control, either relative to the palaeo-
‘Mawson Continent.’ The Albany–Fraser belt is horizontal or in the sense of vertical-axis rotation
truncated on its western end by the late Neoprotero- of the Musgrave region. The principal conclusion
zoic Pinjarra orogen (Fitzsimons 2003) and on its of the latter study, that Australia did not assemble
eastern end by the Palaeozoic Lachlan– Thompson until after 1070 Ma, should be treated with caution
accretionary orogen (Li & Powell 2001), although until further palaeomagnetic studies of Australia’s
some local vestiges of Mesoproterozoic orogeny constituent cratons are undertaken.
or magmatism are documented in northern Queens- The great-circle angular distance between
land (Blewett & Black 1998) and around Tasmania the two key poles (32.58) is identical within error
(Berry et al. 2005; Fioretti et al. 2005). These of the angular distance between the two age-
truncations can provide important piercing points correlative interpolated positions on the Laurentian
for Rodinia reconstructions, because the Albany– APW path, and this permits the working hypothesis
Fraser–Musgrave orogen contains, along its entire that both cratons could have been part of a single
length, two episodes of tectonomagmatic activity Rodinia plate throughout the intervening time inter-
dated at c. 1320 and c. 1200–1150 Ma (White val. Under this assumption, these two poles can be
et al. 1999; Clark et al. 2000). Early consolidation superimposed on the Laurentian APW path in
of the Australia þ Mawsonland continent allows it two options, depending on choice of geomagnetic
to be considered as a single entity in post-1200 Ma polarity (Fig. 4). One option points the Albany-
Rodinia reconstructions. Fraser orogen directly into the centre of the northern
Two important (‘key’) palaeomagnetic poles are margin of Laurentia (Fig. 4a), which appears incom-
available for the c. 1100–750 Ma Rodinian interval: patible with the lack of an equivalently aged orogen.
the Bangemall basin sills at 1070 Ma (Wingate et al. Although Hoffman (1991) depicted the Racklan
Fig. 4. Superposition of ‘key’ poles from Australia, at 1070 and 755 Ma (Table 2), and the Laurentian APW path (in the
North American reference frame) according to one polarity option (a) that produces pronounced geologic mismatches
between Western Australia and northern Canada, and the alternative, allowable polarity option (b) indicating the
preferred position adjacent to the Uralian margin of Baltica. Euler parameters for the rotation of Australia and its poles to
Laurentia are (221.88N, 318.48E, þ155.38CCW) in panel (A), and (212.58N, 064.08E, þ134.58CCW) in panel (B).
380 D. A. D. EVANS
orogeny in that region as a Mesoproterozoic event, central Africa and the São Francisco craton in
subsequent work indicates a Palaeoproterozoic eastern Brazil are almost universally considered
age for that and related events (Thorkelson et al. to represent a single tectonic entity in Rodinian
2001; Maclean & Cook 2004). There is a poorly times (e.g. Brito Neves et al. 1999; Alkmim et al.
understood post-Racklan orogenic event in Yukon 2001). The Congo craton itself is transected by a
(Corn Creek orogeny; Thorkelson et al. 2005), Mesoproterozoic orogen, the Kibaran belt, which
but its precise timing and regional extent are divides the poorly-known western two-thirds of
unknown. Similarly, although Hoffman (1991) and the craton that is largely covered by the Phanerozoic
Dalziel (1997) extended the Grenville orogen north- Congo basin (Daly et al. 1992) from the relatively
ward along the margin of East Greenland, more well-exposed Tanzania and Bangweulu massifs
recent work in that area – plus the once-contiguous (and bounding Palaeoproterozoic belts) in the east
eastern Svalbard – dates the ‘Grenvillian’ tectono- (De Waele et al. 2008). The Kibaran belt has been
magmatic activity at c. 950 Ma (Watt & Thrane viewed as either an ensialic orogen (e.g. Klerkx
2001; Johansson et al. 2005), far younger than the et al. 1987) or a subduction-accretionary margin
Albany-Fraser belt and negating that potential pier- followed by c. 1080 Ma continental collision (e.g.
cing point. The reconstruction of Australia relative Kokonyangi et al. 2006). At the southeastern
to Laurentia as shown in Figure 4a is also incompa- extremity of the craton, the Irumide belt records
tible with the only viable option for the Congo þ c. 1100– 1000 Ma deformation and magmatism
São Francisco craton, as will be shown below. (Johnson et al. 2005; De Waele et al. 2008).
The second option for a continuously connected Reliable palaeomagnetic data from the aggre-
Laurentia and Australia þ Mawsonland in 1100– gate Congo þ São Francisco craton are sparse. In
750 Ma Rodinia (Fig. 4b) requires a gap between this paper, two poles from Congo are used as the
the two cratons that is neatly filled by Baltica key tie points to the Laurentian master APW
in the reconstruction presented above. In this curve: the post-Kibaran (e.g. Kabanga –Musongati)
fit, which juxtaposes the southern Urals and pre- layered mafic–ultramafic intrusions (Meert et al.
Caspian depression against northern Queensland, 1994) and the Mbozi gabbro (Meert et al. 1995).
poles from earlier ages of 1200–1140 Ma do not The former pole is constrained by Ar/Ar dating
superimpose (Fig. 4b), requiring that the postulated at c. 1235 Ma, despite crystallization ages of the
connection was not established until c. 1100 Ma. complexes as early as c. 1400–1350 Ma (Maier
Blewett & Black (1998) documented evidence of et al. 2007). The intrusions lie along the boundary
c. 1100 Ma tectonomagmatism in the Cape River between para-autochtonous Tanzania craton to the
province of northern Queensland, which could east, and an orogenic internal zone to the west
testify to the inferred collision with Baltica at that (Tack et al. 1994). Using the Kabanga –Musongati
time – almost synchronously with Baltica’s rotation pole to represent the entire Congo craton requires
as described above. Although Proterozoic geology that the subsequent c. 1080 Ma deformation was
of the pre-Caspian depression is entombed by ensialic rather than collisional. Despite the fact
c. 15 km of overlying Phanerozoic sedimentary that the palaeomagnetically studied Mbozi gabbro
cover (Volozh et al. 2003), the para-autochthonous is not directly dated, the later-stage syenites in the
Bashkirian anticlinorium of the southern Urals complex are now constrained by a 748+6 Ma
exposes the Riphean stratotype succession of zircon U –Pb age (Mbede et al. 2004), and this
the Baltic craton that can be subdivided into may serve as an approximation of the age of palaeo-
three unconformity-bounded successions. Angular magnetic remanence. In addition to these poles,
unconformity between the Middle and Upper the 795+7 Ma (Ar/Ar; Deblond et al. 2001)
Riphean successions has commonly been attributed Gagwe– Kabuye lavas have yielded a result that
to a rift event (Maslov et al. 1997) but it could appears reliable yet is widely separated from the
also be the distal expression of collisional tectonism slightly younger Mbozi pole (Meert et al. 1995).
at c. 1100 Ma between Baltica and Australia Two groups of poles from dykes in Bahia, Brazil
as proposed herein. The Beloretzk terrane, with (D’Agrella-Filho et al. 2004) are also included in
two stages of deformation bracketing eclogiti- the aggregate Congo þ São Francisco APW path.
zation, all between 1350 and 970 Ma (Glasmacher These groups of poles, with Ar/Ar ages of c. 1080
et al. 2001), could be a sliver of the proposed and 1020 Ma, suggest high-latitude positions for
collision zone. the Congo þ São Francisco craton that appeared
to negate any direct long-lived Rodinian connec-
tions with Laurentia (Weil et al. 1998; Pisarevsky
Congo 1 São Francisco et al. 2003a; Cordani et al. 2003), although colli-
sion between the two blocks at c. 1000 Ma was
Sharing many tectonic similarities since the considered possible (D’Agrella-Filho et al. 1998).
Archaean–Palaeoproterozoic, the Congo craton in As discussed below, these poles are of crucial
importance for testing the radical Rodinia revisions Congo þ São Francisco and Laurentia produces
proposed in this paper. the juxtaposition of Arctic North America with
Given the 1235 Ma Kabanga –Musongati and CONgo at its DAmaride margin (‘ANACONDA’).
c. 750 Ma Mbozi poles superimposed atop the This reconstruction places the two groups of poles
coeval Sudbury dykes and c. 750 Ma poles from from Bahia dykes (D’Agrella-Filho et al. 2004)
Laurentia (Table 2), the two polarity options for atop the c. 1100 Ma Keweenawan poles from Laur-
this long-lived reconstruction of the two blocks entia. This would imply that the c. 1080–1010 Ma
are shown in Figure 5. In the first option (Fig. 5a), Ar/Ar ages from these dykes and baked country
there is substantial overlap between the two rocks (Renne et al. 1990; D’Agrella-Filho et al.
cratons that cannot be avoided by minor adjustments 2004) are inaccurately low, a reflection of large
to the rotations within the uncertainty limits of the scatter in the raw Ar datasets and thus potentially
poles. This implies that the reconstruction, although ubiquitous Ar-loss in those rocks. Interestingly,
palaeomagnetically accurate, is not geologically palaeomagnetic polarity reversal asymmetries that
possible. Other Congo þ São Francisco poles are are documented in two studies of São Francisco
also shown in Figure 5a, to illustrate that they too dykes in Bahia, Brazil (D’Agrella-Filho et al.
fall off the Laurentian APW path in the reconstruc- 1990, 2004), precisely superimpose on reversal
tion of this first polarity option. asymmetries among Keweenawan rocks in Lauren-
The second polarity option for superimpo- tia (Halls & Pesonen 1982) in the ANACONDA
sing the 1235 Ma and c. 750 Ma poles between reconstruction; this suggests a geomagnetic origin
Table 2. Palaeomagnetic poles used in this study
Craton/Rock unit* Age (Ma)† Unrot. A958 Pole reference
8N 8E
Laurentia
Franklin –Natkusiak (FN) c. 720 08 163 4.0 Buchan et al. 2000
Kwagunt Fm (Kw) 742+ 6 18 166 7.0 Weil et al. 2004
Galeros Fm (G) .Kw 202 163 6.0 Weil et al. 2004
Tsezotene sills (Ts) 779+ 2 02 138 7.0 Park et al. 1989
Wyoming dykes (Wy) c. 784 13 131 4.0 Harlan et al. 1997
Haliburton ‘A’ (HA) 1015+ 15 233 142 6.0 Warnock et al. 2000
Chequamegon (C) c. J 212 178 5.0 McCabe & Van der Voo 1983
Jacobsville (J) ,F 209 183 4.0 Roy & Robertson 1978
Freda (F) ,N 02 179 4.0 Henry et al. 1977
Nonesuch (No) c. 1050? 08 178 4.0 Henry et al. 1977
Lake Shore traps (LS) 1087+ 2 22 181 5.0 Diehl & Haig 1994
Unkar intrusions (Ui) c. 1090 32 185 8.0 Weil et al. 2003
Portage Lake (PL) 1095+ 3 27 181 2.0 Halls & Pesonen 1982
Upper Nth Shore (uNS) 1097+ 2 32 184 5.0 Halls & Pesonen 1982
Upper Osler R (uO-r) 1105+ 2 43 195 6.0 Halls 1974
Logan sills (Lo) 1108+ 1 49 220 4.0 Buchan et al. 2000
Abitibi dykes (Ab) 1141+ 1 43 209 14.0 Ernst & Buchan 1993
Upper Bylot (uB) c. 1200? 08 204 3.0 Fahrig et al. 1981
Sudbury dykes (Sud) 1235 þ7/23 203 192 3.0 Palmer et al. 1977
Mackenzie dykes (Mac) 1267+ 2 04 190 5.0 Buchan & Halls 1990
Baltica
Hunnedalen dykes (Hun) c. 850 41 042 10.0 Walderhaug et al. 1999
Rogaland anorth (Rog) c. 900? 42 020 9.0 Brown & McEnroe 2004
Gällared Gneiss (GG) 985? 44 044 6.0 Pisarevsky & Bylund 1998
Hakefjorden (Hak) 916+ 11 205 069 4.0 Stearn & Piper 1984
Goteborg –Slussen (Got) 935+ 3 07 062 12.0 Pisarevsky & Bylund 2006
Dalarna dykes (Dal) 946+ 1 205 059 15.0 Pisarevsky & Bylund 2006
Karlshamn –Fajo (Karls) c. 950 213 067 16.0 Pisarevsky & Bylund 2006
Laanila Dolerite (Laa) c. 1045 02 032 15.0 Mertanen et al. 1996
Bamble intrus. (Bam) c. 1070 203 037 15.0 Brown & McEnroe 2004
(Continued)
382 D. A. D. EVANS
Table 2. Continued
Craton/Rock unit* Age (Ma)† Unrot. A958 Pole reference
8N 8E
Australia
Mundine Well (MW) 755+ 3 45 135 4.0 Wingate & Giddings 2000
Browne Fm (Br) c. 830? 45 142 7.0 Pisarevsky et al. 2007
Bangemall sills (Bang) 1070+ 6 34 095 8.0 Wingate et al. 2002
Lakeview dolerite c. 1140 210 131 17.5 Tanaka & Idnurm 1994
Mernda Morn mean c. 1200 248 148 15.5 this study‡
Congo
Mbozi complex (Mb) 748+ 6 46 325 9.0 Meert et al. 1995†
Gagwe lavas (Gag) 795+ 7 225 273 10.0 Meert et al. 1995
Kabanga-Musongati (KM) 1235+ 5 17 293 5.0 Meert et al. 1994
São Francisco
Bahia dykes W-dn (B-w) see text 09 281 7.0 D’Agrella-Filho et al. 2004
Bahia dykes E-up (B-e) see text 212 291 6.0 D’Agrella-Filho et al. 2004
Kalahari
Namaqua mean (Nam) c. 1000 209 150 18.0 Gose et al. 2006
Kalkpunt Fm (KP) c. 1090 57 003 7.0 Briden et al. 1979†
Umkondo mean (Um) 1110+ 3 64 039 3.5 Gose et al. 2006
India
Malani Rhyolite (MR) c. 770 75 071 8.5 Torsvik et al. 2001
Harohalli dykes (Har) 815/1190 25 078 8.5 Malone et al. 2008
Majhgawan kimb (Majh) 1074+ 14? 37 213 15.5 Gregory et al. 2006
Wajrakarur kimb (Waj) c. 1100 45 059 11.0 Miller & Hargraves 1994†
South China
Liantuo mean (Li) 748+ 12 204 341 13.0 Evans et al. 2000
Xiaofeng dykes (X) 802+ 10 214 271 11.0 Li et al. 2004
Tarim
Beiyixi volcanics (Bei) 755+ 15 218 014 4.5 Huang et al. 2005†
Aksu dykes (Ak) 807+ 12 219 308 6.5 Chen et al. 2004
Svalbard
Svanbergfjellet (Svan) c. 790? 226 047 6.0 Maloof et al. 2006
Grusdievbreen u. (Gru2) c. 800? 01 073 6.5 Maloof et al. 2006
Grusdievbreen l. (Gru1) c. 805? 220 025 11.5 Maloof et al. 2006
North China
Nanfen Fm (Nan) c. 790? 217 121 11.0 Zhang et al. 2006
Cuishuang Fm (Cui) c. 950? 241 045 11.5 Zhang et al. 2006
Siberia
Kandyk Fm (Kan) c. 990 203 177 4.0 Pavlov et al. 2002
Milkon Fm (Mil) 1025? 206 196 4.0 Pisarevsky & Natapov 2003
Malgina Fm (Mal) 1040? 225 231 3.0 Gallet et al. 2000
Amazonia
Aguapei sills (Agua) c. 980 264 279 8.5 D’Agrella-Filho et al. 2003†
Fortuna Fm (FF) 1150? 260 336 9.5 D’Agrella-Filho et al. 2008
Nova Floresta sills (NF) c. 1200 225 345 5.5 Tohver et al. 2002
*Abbreviations for rock units correspond to the poles depicted in Figures 2–10.
†
Ages are queried where highly uncertain or estimated in part by position on the APW path. Ages are cited fully in Buchan et al. (2000) or
Pisarevsky et al. (2003a), or the pole references, except where otherwise noted: Zig-Zag Dal – Midsommersø from Upton et al. (2005);
Western Channel diabase from Hamilton & Buchan (2007); Mbozi complex from Mbede et al. (2004); Kunene anorthosite from Mayer
et al. (2004) and Drüppel et al. (2007); Harohalli dykes from Malone et al. (2008); Wajrakarur kimberlites from Kumar et al. (2007);
Beiyixi volcanics from Xu et al. (2005); Kalkpunt Formation estimated from Pettersson et al. (2007).
‡
Combined calculation of Fraser dyke VGP (Pisarevsky et al. 2003b) with Ravensthorpe dykes of Giddings (1976) and one additional dyke
at Narrogin (A. V. Smirnov & D. A. D. Evans, unpublished data).
Fig. 5. Non-‘key’ palaeomagnetic poles from Congo þ São Francisco during the interval c. 1370–750 Ma (Table 2).
The 1235 and 755 Ma poles from Congo are used to generate two possible superpositions onto the Laurentian APW
path (in the North American reference frame). One polarity option (a) generates a complete overlap between the two
cratons and is thus not allowed, whereas the other (b) produces the ‘ANACONDA’ juxtaposition described in the text.
Note that the Bahia dykes poles fall on the c. 1100 Ma segment of the Laurentian path, and a testable prediction of
ANACONDA is that the existing Ar/Ar ages of 1080– 1010 Ma from these dikes are too young (see text). Euler
parameters for the rotation of Congo þ São Francisco and its poles from the present African reference frame to
Laurentia are (20.08N, 334.98E, þ162.58CCW) in panel (A), and (224.08N, 249.08E, þ128.58CCW) in panel (B).
for the asymmetries (Pesonen & Nevanlinna 1981). igneous activity at c. 1.38 Ga is recorded by the
A joint palaeomagnetic and U –Pb restudy of these Kunene complex and coeval Hart River magmatism
dykes is currently underway (Catelani et al. 2007). in Canada (Thorkelson et al. 2001) and
Other poles shown in Figure 5b are from the Midsommersø dolerite and cogenetic Zig-Zag Dal
Kunene anorthosite (Piper 1974), which with new basalt (Upton et al. 2005). Along the southeastern
precise U –Pb ages on consanguineous felsic rocks margin of Congo, potassic magmatism at 1360–
of 1380–1370 Ma (Mayer et al. 2004; Drüppel 1330 Ma (Vrana et al. 2004) correlates broadly in
et al. 2007) deserves refinement with modern age with the c. 1350 Ma Mashak volcanics in the
palaeomagnetic techniques; and the Gagwe lavas southern Urals (Maslov et al. 1997).
(Meert et al. 1995) with an updated Ar/Ar age of The ANACONDA reconstruction also identi-
795+ 7 Ma (Deblond et al. 2001). If the Gagwe fies potential long-sought counterparts to the giant
pole is primary, then the ANACONDA fit requires 1.27 Ga Mackenzie/Muskox/Coppermine large
a large APW loop shared among all elements of igneous province in Canada, with correlatives in
Rodinia at c. 800 Ma. Such a loop has not tradi- Greenland and Baltica (Ernst & Buchan 2002). In
tionally been accepted for Laurentia, as some eastern Africa, the late-Kibaran intrusive complexes
anomalous results of that age have been interpreted described above were once thought to be of the same
instead as suffering from local vertical-axis rotation age (Tack et al. 1994), but are now known to be
(e.g. Little Dal lavas; Park & Jefferson 1991). either older (Tohver et al. 2006) or younger (De
However, a large APW loop in the Laurentian Waele et al. 2008). Nonetheless, a dyke of similar
path at c. 800 Ma has now been demonstrated by tectonic setting in Burundi is dated by Ar/Ar at
high-quality palaeomagnetic results from Svalbard c. 1280– 1250 Ma (Deblond et al. 2001), and
(Maloof et al. 2006; see above) and is generally Tohver et al. (2006) raise the possibility that the
supported by data from several other cratons that c. 1380 Ma zircons in the layered intrusions are
imply large APW shifts between c. 800 and xenocrysts. Thus, more comprehensive geochronol-
c. 750 Ma (Li et al. 2004). ogy of this region is warranted. In Brazil, the
The ANACONDA reconstruction juxtaposes Niquelândia and related mafic-ultramafic com-
several intriguingly similar geological features plexes have numerous age constraints, the most
among the São Francisco, Congo and northern recent study suggests emplacement ages at
Laurentian cratons. Extensive Palaeoproterozoic– 1250+20 Ma (Pimentel et al. 2004). The latter
Mesoproterozoic orogeny in the southern Angola– intrusions lie within the late Neoproterozoic Brasilia
Congo craton (Seth et al. 2003) adjoins crust belt, adjacent to the São Francisco craton, and in the
with a similarly aged interval of deformation in present model are considered not grossly allochtho-
arctic Canada (Thorkelson et al. 2001; MacLean nous relative to that craton (Pimentel et al. 2006).
& Cook 2004) and suggests initial amalgamation Concordance of the Laurentian and Congo þ São
of ANACONDA at c. 1.6–1.5 Ga. Post-collisional Francisco APW paths younger than this age
384 D. A. D. EVANS
indicates that extension at 1270–1250 Ma failed to becoming ‘orphaned’ during mid-Neoproterozoic

separate the cratons, rather than opening a postu- Rodinia breakup.
lated ‘Poseidon’ ocean (Jackson & Iannelli 1981). The juvenile Hf and Nd signatures of 1.4 Ga
These magmatic loci could represent early stages A-type granites preserved as clasts and detrital
of the rifting that is required by palaeomagnetic zircons in Transantarctic Mountains sediments
data to have rotated Baltica away from Congo and (Goodge et al. 2008) have been used to support a
toward southern Greenland in the late Mesoprotero- connection with western Laurentia in the SWEAT
zoic, as discussed above. Baltica’s rotation, coupled juxtaposition. However, Goodge et al. (2008; their
with arrival of Australia at c. 1100 Ma as discussed fig. 3a) illustrate other regions of the world with
above, isolated a craton-sized tract of remnant ocean comparable magmatism of the same age: Cathaysia,
at the end of the Mesoproterozoic. This ‘hole’, in the eastern Congo, southern Amazonia and south-
present revised Rodinia model, is intriguing for western Baltica. If the revised Rodinia position for
several reasons. First, it predicts a Mediterranean- Australia þ Mawsonland (Fig. 4b) is correct, then
style slab rollback to account for arc magmatism the general proximity of 1.4 Ga A –type granite
and arc-continent collision in the Irumide belt at terrains in Congo and Baltica make them the most
c. 1050– 1020 Ma (De Waele et al. 2008) as well attractive candidates as the originally contiguous
as c. 1000–800 Ma tectonic events in Greenland extensions of the Antarctic magmatic province in
(Watt & Thrane 2001) and northern Norway pre-Rodinian times.
(Kirkland et al. 2006). Second, the large
c. 800 Ma evaporite basin hosting the Shaba–
Katanga copperbelt in southern Congo (Jackson Kalahari
et al. 2003) may have continued onto Laurentia as
the evaporitic upper part of the Amundsen basin As noted in a recent review of late Mesoproterozoic
and its correlative units in the Mackenzie Mountains palaeomagnetic data from southern Africa (Gose
(Rainbird et al. 1996). This composite evaporitic et al. 2006), the two anchors of the Kalahari APW
basin could represent a lithospheric sag precur- path are the c. 1100 Ma Umkondo grand-mean
sor to rifting and separation of ANACONDA – pole of highest reliability, followed by various mod-
accompanied by the Chuos, Grand Conglomerat erately well grouped poles from c. 1000 Ma terranes
and Rapitan glaciogenic deposits (Evans 2000) – within the Namaqua –Natal orogen. These two
between c. 750 and 700 Ma. Finally, it demonstrates anchor poles permit Kalahari to have been a
how the palaeomagnetic APW-matching method member of Rodinia throughout the intervening
can generate a more refined palaeogeographic time; they are separated by about 808 or 1008 of
framework for supercontinent reconstructions; great-circle arc, depending on relative geomagnetic
all previous models of Rodinia, using tectonostrati- polarity, the larger value being similar to that separ-
graphic comparisons or ‘closest-approach’ palaeo- ating Laurentian poles on the ‘master’ Rodinian
magnetic reconstructions, have placed the cratons APW path of the same pair of ages (Logan and
together as tightly as possible – essentially ruling Grenville loops, respectively). The precise dating
out even the possibility of Mediterranean-style of both Umkondo and early Keweenawan poles
remnant-ocean tectonism in the pre-Pangaean world. at 1108 Ma, with a pronounced geomagnetic
As a final note, recall that the palaeomagnetic polarity bias in both units, constrains the hemisphere
data from Australia were discordant in the present option of relative reconstructions between the
revised Rodinia model at c. 1140 Ma, requiring two cratons, such that the Grenville and Namaqua
collision of Australia þ Mawsonland to become a orogens cannot face each other (Hanson et al.
part of Rodinia at c. 1100 Ma. Proximity of Maw- 2004; Gose et al. 2006) as earlier proposed.
sonland to Tanzania in the Rodinia fit (Fig. 5b) In addition to this polarity constraint, if the
implies convergence and inferred collision there as c. 1000 Ma poles are aligned then Kalahari can
well. One difficulty with this inference is the lack occupy only one of two positions in relative palaeo-
of any direct evidence for c. 1100 Ma tectonism latitude and palaeolongitude to Laurentia plus
in the central Transantarctic Mountains (Goodge its surrounding cratons in Rodinia. The standard
et al. 2001), despite some tenuous Nd-isotopic depiction of Kalahari’s geon-10 APW path (e.g.
support for Mesoproterozoic activity there (Borg Powell et al. 2001; Evans et al. 2002; Meert &
& DePaolo 1994). However, if Kokonyangi et al. Torsvik 2003; Tohver et al. 2006; Gose et al.
(2006) are correct in proposing a c. 1080 Ma 2006; Li et al. 2008) is shown in Figure 6a in
suture between Tanzania þ Bangweulu and simplified form, using the two anchor poles plus
Congo cratons at the Kibaran orogen, then there that of the Kalkpunt redbeds of eastern Namaqua-
is the intriguing possibility presented by this land (Briden et al. 1979). Although Powell et al.
reconstruction, that Tanzania þ Bangweulu was (2001) suggested an age of 1065 Ma for the
originally a fragment of Australia þ Mawsonland, Kalkpunt red beds, recent U– Pb dating of a
Fig. 6. Superposition of Kalahari’s 1110 Ma Umkondo palaeomagnetic pole with polarity constraint tied to the early
Keweenawan Logan sills pole from Laurentia (Hanson et al. 2004; Gose et al. 2006), plus younger poles from the
Namaqua and Natal Provinces in various polarity options (poles listed in Table 2). Panel (a) shows a large distance
between Kalahari and the rest of Rodinia, whereas panel (b) illustrates the preferred reconstruction herein. Euler
parameters for the rotation of Kalahari and its poles from the present African reference frame to Laurentia, in present
North American coordinates, are (81.58N, 187.48E, 2163.48CCW) in panel (A) and (32.58N, 307.58E, þ76.58CCW) in
panel (B).
conformably underlying rhyolite indicates an age of vertical-axis rotation of the Koras Group, which is
only slightly younger than c. 1095 Ma (Pettersson only locally exposed within a region of large strike-
et al. 2007). The APW arc length transcribed by slip shear zones (Pettersson et al. 2007). About 30
this polarity option of the three poles is somewhat degrees of local rotation are required to bring the
shorter than that of the Keweenawan APW track Kalkpunt pole into alignment with poles from
of Laurentia, but dating of the Namaqua–Natal the younger Keweenawan lavas and intrusions at
poles is forgiving enough to allow for the precise c. 1095 Ma.
mismatch. However, the reconstruction derived by The preferred reconstruction of Kalahari shown
this APW superposition is one in which Kalahari in Figure 6b is thus the only possible way to
is widely separated from all other cratons in include this craton in the Rodinia assembly as
Rodinia – even if one were to choose more standard early as 1100 Ma according to existing palaeo-
models involving Australia and other cratons to the magnetic constraints. Other published solutions
SW of Laurentia. The Kalahari reconstruction involve late collision of Kalahari into Rodinia
shown in Figure 6a is, however, similar to that of at c. 1000 Ma (Pisarevsky et al. 2003a; Li et al.
Piper (2000, 2007), but as noted above this recon- 2008). The well-known Namaqua-Natal belt of
struction produces numerous palaeomagnetic southern Kalahari shows the main phases of defor-
mismatches when high-quality individual results mation at c. 1090 –1060 Ma (Jacobs et al. 2008),
are considered rather than broad means. which is typically correlated in an opposing colli-
The alternative polarity option for the Namaqua- sional sense to the Ottawan orogeny of the Grenville
Natal poles, while preserving the sanctity of geo- Province in Laurentia. These reconstructions,
magnetic polarity matching of Umkondo and early however, must either violate the geomagnetic
Keweenawan poles, produces the reconstruction polarity match between early Keweenawan and
shown in Figure 6b. The reconstruction juxtaposes Umkondo poles, or invoke an implausible 180-
Kalahari’s northern margin adjacent to the Vostok degree rotation of Kalahari relative to Laurentia as
(western) margin of Mawsonland (Fig. 6). This they approached each other in geon 10.
polarity option for Namaqua –Natal poles provides In the preferred reconstruction here (Fig. 6b),
a more acceptable APW arc length relative to the more proximal Mesoproterozoic margin to the
the Keweenawan APW track and Grenville APW Laurentian side of Rodinia is the present north-
loop from Laurentia, but it introduces a different western side of Kalahari. Along that margin, a
problem: the Kalkpunt pole now falls on the other single c. 1300 –1200 Ma orogen has been hypoth-
side of Umkondo, and reconstructs to a position esized (Singletary et al. 2003), and this orogen
slightly beyond the apex of the Logan APW was stabilized prior to widespread large igneous
loop in the Laurentian reference frame (Fig. 6b). province mafic magmatism at about 1110 Ma.
Rather than invoke an additional APW loop, this In the revised Rodinia reconstruction presented
discrepancy is perhaps best explained by local herein, the NW Kalahari orogen is proposed to
386 D. A. D. EVANS
record collision between Kalahari and the Vostok

margin of Australia þ Mawsonland (Fig. 6). A
complex collisional triple junction, suturing this
orogen, the Namaqua belt and the Albany belt in
Western Australia, would be partly reworked by
subsequent Pan-African (Damaride) and Pinjarran
tectonics, and partly buried by Antarctic ice;
testing this model by correlating the details of the
three collisions will be a challenging enterprise.
India, South China, Tarim

Relative to Australia, the reconstructed positions of
India, South China and Tarim are similar in this
Rodinia model to previously published versions
(Fig. 7). The reconstruction of India to Australia
essentially follows Torsvik et al. (2001), that of
South China relative to India follows Jiang et al.
(2003) and that of Tarim to Australia follows Li Fig. 7. Reconstruction of India (blue), South China
et al. (1996), Powell & Pisarevsky (2002) and Li (green), and Tarim (peach) near Australia as commonly
et al. (2008). Palaeomagnetic poles from these depicted in Rodinia models. As with Figures 3– 6, all
cratons are sparse, but they provide important cratons and poles have been rotated into the present
constraints on Rodinia configurations of these North American (Laurentian) reference frame (Table 3).
blocks and also support the existence of a large The Laurentian APW path is simplified into a curve
loop in the Rodinian APW path at c. 800 Ma. (light grey) for clarity. Also shown are three poles (red
colour) with poor absolute age control but correct
Discussing the results in detail, numerous Indian
stratigraphic order from Svalbard (Maloof et al. 2006;
poles (reviewed by Malone et al. 2008) are summar- Table 2), rotated to a modified position north of
ized by the three depicted in Figure 7. The most Greenland (Table 3) for better APW matching with both
reliable is from the Malani rhyolitic large igneous the Laurentian poles from geon 7 and the proposed APW
province of Rajasthan, with various U– Pb ages loop at c. 800 Ma. Palaeomagnetic poles (abbreviated as
centred around 770 Ma (Torsvik et al. 2001; in Table 2) are shown in the colours of their host cratons.
Malone et al. 2008). Aside from this result, which The location of Tarim poles is consistent with the craton
has been reproduced numerous times in the past reconstruction in darker colour; the lighter-coloured
few decades, the Indian palaeomagnetic poles alternative Tarim reconstruction aligns the Aksu dykes
pole (Ak) with c. 800 Ma poles from other cratons, but
from Rodinia times are questionable either in
results in a mismatch of the Beiyixi volcanics pole (Bei);
quality or in age. The oft-cited pole from Harohalli poles from this alternative reconstruction are not
alkaline dykes was previously assigned an age of illustrated, for sake of clarity. Queried ages of some of
c. 815 Ma, based on Ar/Ar ages (Radhakrishna & the Indian poles are discussed in text.
Mathew 1996). However, new U –Pb zircon data
from these dykes suggest a much older age of
c. 1190 Ma (cited in Malone et al. 2008). If the
latter age is correct, then the Harohalli dykes pole
is irrelevant for the present discussion, because interpretation could readily explain the large discor-
this portion of Rodinia is proposed herein to have dance between the Majhgawan virtual geomagnetic
assembled around 1100 Ma or younger. There is pole (VGP; not averaging geomagnetic secular vari-
also the pole from the well-dated Majhgawan ation due to brief emplacement of the kimberlite)
kimberlite (1074+14 Ma by Ar/Ar; Gregory and those from the nearly coeval Wajrakarur
et al. 2006), which is nearly identical to poles kimberlite field in south-central India (Miller &
from the nearby Rewa and Bhander sedimentary Hargraves 1994).
succession (Malone et al. 2008). The latter units If either of these kimberlite poles is primary,
have age uncertainties on the order of 500 Ma, then their significant distances from the Laurentian
as summarized by Malone et al. (2008). No field APW path would negate the proposed reconstruc-
stability tests have been performed on either the tion (Fig. 7) at that time. It is possible that collision
Majhgawan kimberlite or the Rewa/Bhander between India (plus attached Cathaysia block of
units, so there remains the possibility that South China) and NW Australia occurred after
these poles represent a two-polarity magnetic over- 1100–1075 Ma, accounting for the reconstructed
print across north-central India. Such an pole discrepancy.
Palaeomagnetic poles from China are more juxtapositions is correct for Tarim in Rodinia,
straightforward to interpret. In South China then Tarim would need to rift from that position
(Yangtze craton), the Xiaofeng dykes yield a high and re-collide with East Gondwanaland in its
palaeolatitude at 802+10 Ma (Li et al. 2004), and peri-Arabian position prior to mid-Cambrian
the Liantuo Formation red beds yield a moderate time. Neither Tarim nor northern India records
palaeolatitude at 748+ 12 Ma (Evans et al. 2000). Ediacaran-age orogenic activity that would docu-
Similarly, the Aksu dykes in Tarim were emplaced ment such convergence.
at high palaeolatitude at 807+ 12 Ma (Chen et al. Although the c. 800 Ma poles just described
2004), and the Beiyixi volcanics were erupted are far removed from the established Laurentian
at lower palaeolatitudes (Huang et al. 2005) at APW path in the proposed reconstruction, they con-
755+ 15 Ma (Xu et al. 2005). Matching these two stitute important independent support from several
pairs of poles from South China and Tarim, Rodinian cratons that they – if not the entire super-
however, results in a large distance between the continent – experienced an oscillatory pair of
cratons (not shown in Fig. 7), inconsistent with rotations at that time. The kinematic evidence for
their strongly compatible Sinian geological histories this proposed rotation does not specify a dynamic
(Lu et al. 2008a). Figure 7 shows two alternative cause, but inertial-interchange true polar wander
positions of Tarim relative to the cratons heretofore (IITPW) events are the most straightforward exp-
discussed. The preferred position is shown in a lanation (Li et al. 2004; Maloof et al. 2006).
darker colour, along with the properly rotated pair When the Svalbard magnetostratigraphic data of
of Tarim poles. In this position, where Tarim is Maloof et al. (2006) are considered (red colour in
directly adjacent to both South China and (present Fig. 7), they provide the hitherto unrecognized evi-
NW) India, the 755 Ma Beiyixi pole is aligned dence from Laurentia for the APW loop indicated
with middle-geon-7 poles from Laurentia and by India (if Harohalli dykes are c. 800 Ma), South
other cratons; however, the 807 Ma Aksu dykes China, Tarim and Congo (Fig. 5b). The precise
pole is discordant (Fig. 7). This could suggest reconstruction of Svalbard relative to Greenland is
post-800 Ma convergence between Tarim and uncertain, but direct connection between the two
Rodinia, or it could also be due to unrecognized areas of Laurentia are strongly supported by lithos-
local vertical-axis rotations of the Aksu area, as tratigraphy (Maloof et al. 2006). Also, because
suggested by Li et al. (2008) to be a general the Svalbard APW shift is recorded in several
problem for the minimally studied Tarim block. widely separated, continuously sampled magnetos-
Alternatively, the Aksu dykes pole could be tratigraphic sections, local vertical-axis rotations
aligned with the c. 800 Ma pole from the coeval cannot account for the directional shifts: the APW
Xiaofeng dykes in South China; in which case loop at c. 800 Ma is a genuine feature of the Lauren-
Tarim reconstructs next to northern Australia tian palaeomagnetic database that must be included
(lighter shade of peach colour in Fig. 7) in the in all Rodinia models.
same sense as Li et al. (1996, 2008). The 755 Ma The reconstruction of India, South China and
Beiyixi pole, however, is removed from the Rodi- Tarim, adjacent to Australia as shown in Figure 7,
nian APW path in this reconstruction. This would produces some intriguing tectonic juxtapositions,
suggest either early (pre-755 Ma) rifting of Tarim in which compatible histories can be considered as
from Rodinia, or local vertical-axis rotations of predictions of the model. First, the Sibao orogen
the Quruqtagh region where the Beiyixi volcanics in South China (Li et al. 1996, 2002) appears
are exposed. A third alternative reconstruction of to strike directly into northwestern India, where
Tarim – adjacent to eastern Australia, based on a earliest Neoproterozoic tectonomagmatic activity
proposed radiating dyke swarm at c. 820 Ma (Lu is postulated to be continuous with the Delhi fold-
et al. 2008a) – is broadly compatible with the belt in India (Deb et al. 2001), under Neoproterozoic
palaeomagnetic data from 755 Ma but, ironically, sedimentary cover of Rajasthan and north– central
not c. 800 Ma. Pakistan. If this represents a collisional orogen,
The present analysis leaves the position of then most of cratonic India should have more
Tarim somewhat uncertain, but the preferred pos- affinities with the Cathaysia block in South China,
ition is that described first, above, and illustrated colliding with the Yangtze þ Tarim craton during
with darker peach colour in Figure 7. The main final Rodinia assembly. The Tarimian orogeny
reason for this preference is that new palaeomag- of similar age (Lu et al. 2008a) could express a
netic results from Cambrian –Ordovician sedimen- poorly exposed continuation of this collisional belt.
tary rocks in the Quruqtagh area (Zhao et al. On the other side of India, the c. 1000–950 Ma
2008) are most compatible with the Gondwanaland Eastern Ghats orogen (Mezger & Cosca 1999) and
APW path if Tarim is reconstructed near Arabia, its continuation as the Rayner terrane in Antarctica
that is separated from Australia by India and South (Kelly et al. 2002), extends east of Prydz Bay
China. If either the northern or eastern Australian (Kinny et al. 1993; Wang et al. 2008), and according
388 D. A. D. EVANS
to this reconstruction splays into the Edmund (Zhang et al. 2006) from the Rodinia interval: the
foldbelt of Western Australia, which deformed Cuishang Formation (c. 950 Ma?) and the Nanfen
1070 Ma sills and their host Bangemall basin Formation (c. 790 Ma?). Ages from both of these
sedimentary rocks about tight NW–SE axes and formations are very tenuous.
led to moderate isotopic disturbance (Occhipinti & Pisarevsky and Natapov (2003) summarized the
Reddy 2009). The full extent of this orogen is prob- Meso-Neoproterozoic stratigraphic record across
ably hidden under the East Antarctic icecap (includ- the Siberian craton, as well as its palaeomagnetic
ing the Gamburtsev Subglacial Mountains; Veevers database. The most reliable palaeomagnetic poles
& Saeed 2008), and likely involves smaller define an APW trend that is supported by less
Archaean–Palaeoproterozoic cratonic fragments reliable results; only the three most reliable are
such as the Ruker terrane (Phillips et al. 2006). included in this synthesis, but the conclusion is not
The orogen is proposed here to involve collision affected by incorporating the others. The present
with Kalahari along the latter craton’s Namaqua analysis does not include the high-quality Linok
margin at c. 1090– 1060 Ma (Jacobs et al. 2008). Formation pole from the Turukhansk region
Tectonothermal events of similar age in the (Gallet et al. 2000), because it restores precisely
Central Indian Tectonic Zone (Chatterjee et al. atop that of the likely correlative Malgina Formation
2008; Maji et al. 2008) connect the Delhi and in the Uchur-Maya region marginal to the Aldan
Eastern Ghats/Rayner orogens in poorly shield, after restoration of the Devonian Vilyuy
understood ways. rift in central Siberia (Table 1).
The reconstruction also suggests that the pre- Matching of the Siberian APW path from the
cisely coeval igneous events recorded on several Uchur-Maya region with the Keweenawan APW
cratons at 755 Ma are genetically related: Mundine track to Grenville loop from Laurentia results in
Well dyke swarm in Australia (Wingate & Giddings two possibilities, because of geomagnetic polarity
2000), Malani large igneous province in India options. The first option (not shown) produces the
(Torsvik et al. 2001), Nanhua rift and related pro- typical reconstruction of Siberia with its southern
vinces in South China (Li et al. 2003) and Beiyixi margin in the vicinity of northern Laurentia
volcanics in Tarim (Xu et al. 2005). As will be (option ‘A’ of Pisarevsky & Natapov 2003;
shown below, western Siberia also reconstructs Pisarevsky et al. 2003a; Li et al. 2008; Pisarevsky
immediately adjacent to Tarim, and the Sharyzhal- et al. 2008), The hypothesized reconstruction of
gai massif contains mafic dykes of precisely the Siberia (Fig. 8) is essentially the same as ‘option
same age (Sklyarov et al. 2003). The Malani B’ of Pisarevsky & Natapov (2003) and the first
region in India is proposed here as the central option discussed by Meert & Torsvik (2003). Both
focus of a hotspot or mantle plume with radiating of those papers concluded that such a reconstruc-
arms extending across these cratons. tion would probably exclude Siberia because of
the great distance from Laurentia, but the present
revised Rodinia model covers this gap with
North China, Siberia Baltica, Australia, India, and North China.
Because there is scant to no evidence for a
Both North China and Siberia lack coherent late ‘Grenvillian’ orogen between India and North
Mesoproterozoic (‘Grenvillian’) orogens, although China in the proposed reconstruction (Fig. 8), a
possible vestiges can be found along the northern corollary of the model is that those two cratons
and southern margins of North China (Zhai et al. were joined in similar fashion since their Palaeo-
2003; but see also discussion in Lu et al. 2008b). proterozoic consolidations. Zhao et al. (2003)
Similar carbonate-dominated mid-Proterozoic plat- described a series of correlations between southern
form or passive-margin sedimentary successions India and eastern North China, and proposed
on both cratons have inspired, along with palaeo- four possible reconstructions in which those two
magnetic support, hypotheses of a close palaeo- regions could have been directly juxtaposed. In
geographic connection between the two blocks this paper I present a fifth alternative connection
in Rodinia and in earlier times (Zhai et al. 2003; (Fig. 8), which, unlike the previous four, is in
Wu et al. 2005; Zhang et al. 2006; Li et al. 2008). agreement with the Rodinia-era palaeomagnetic
A recent U –Pb age determination of c. 1380 Ma poles described above. There are no reliable and
on ash beds in the upper part of the North China precisely coeval palaeomagnetic results from the
cover succession (Su et al. 2008), however, shows two cratons yet available (Evans & Pisarevsky
that this succession is almost entirely older than 2008) to test their earlier hypothesised assembly.
the bulk of ‘Riphean’ sediments across Siberia Siberia is also proposed to have been connected
(Khudoley et al. 2007). The older age of the North to India and North China prior to Rodinia’s amal-
China succession also demonstrates that there are gamation in the late Mesoproterozoic. Pisarevsky
only two moderately reliable palaeomagnetic poles & Natapov (2003) summarized the Riphean
margin of Laurentia at the end of the Mesoprotero-

zoic Era (e.g. Bond et al. 1984; Dalziel 1991, 1997;
Hoffman 1991; Weil et al. 1998; Pisarevsky et al.
2003a; Tohver et al. 2006; Li et al. 2008). This
is perhaps surprising, given how many alternative
possibilities exist due to the global preponderance
of late Mesoproterozoic (‘Grenvillian’) orogens
among the world’s cratons (Fig. 1). The Laurentia þ
Amazonia connection is broadly supported by
Pb-isotopic signatures (Loewy et al. 2003), but
without a globally comprehensive dataset such
comparisons are merely indicative rather than diag-
nostic. If any craton-scale tectonic comparisons
are to be made, then the progressively younging,
accretionary character of Amazonia would fit
much better along strike of southwestern Laurentia
(Santos et al. 2008) rather than in the mirrored
configuration of the two cratons in typical Rodinia
models. The palaeomagnetic evidence in support
Fig. 8. Reconstruction of North China (tan) and Siberia
of the Laurentia þ Amazonia connection in
(pink) directly adjacent to India on the outer edge of Rodinia is not particularly strong, either. Based on
Rodinia. All cratons and their similarly coloured poles new palaeomagnetic data, Tohver et al. (2002,
have been rotated into the present North American 2004) and D’Agrella-Filho et al. (2003, 2008)
(Laurentian) reference frame (Table 3). The Laurentian have successively modified the kinematics of
APW path is simplified into a curve (light grey), with the the putative Laurentia þ Amazonia collision.
additional APW loop at c. 800 Ma (see text and Fig. 7) If confined to such a collisional model, the data
indicated by the dashed segment. Among the North now demand two unusual kinematic features:
China and Siberia poles shown here, all but the Kandyk (1) c. 5000 km of sinistral motion with Amazonia
Formation (Kan) are poorly constrained in age.
occupying positions adjacent to Texas and Labrador
at 1.2 and 1.0 Ga, respectively; and (2) 908 of
stratigraphic architecture of the present-day margins anticlockwise rotation of Amazonia relative to
of Siberia, demonstrating in many areas a clear Laurentia, so that Amazonia appears to roll like
thickening of strata away from the craton into a wheel along the Grenvillian margin during
deeper-water sedimentary facies. The long-lived its syncollisional sinistral odyssey. Such odd
Mesoproterozoic connections to North China kinematics could be avoided if the observations
and India (Fig. 8) would be inconsistent with the were not confined by the initial assumption of a
Siberian stratigraphic record if it could be demon- Laurentia þ Amazonia collision.
strated that the Turukhansk, Igarka, or northern Figure 9 shows the available palaeomagnetic
Siberian margins faced the open ocean through poles from Amazonia during the Rodinia interval.
the Meso-Neoproterozoic transition. However, in The Nova Floresta (NF) and Fortuna Formation
the best-documented areas of Turukhansk, there (FF) poles are fully published (Tohver et al. 2002;
is no preserved record of substantial westward D’Agrella-Filho et al. 2008), whereas the Aguapei
thickening of the Riphean stratigraphy as would be sills (Agua) result is presented in abstract only
expected for a continent– ocean crustal transition, (D’Agrella-Filho et al. 2003). This latter result
nor is there any preserved evidence of deep-water is important for constraining the possible position
facies in the middle Riphean succession (Bartley of Amazonia in Rodinia, however, because it, like
et al. 2001; Pisarevsky & Natapov 2003; Khudoley the Nova Floresta data, is from mafic igneous
et al. 2007). According to the available infor- rocks constrained in age by the Ar/Ar method.
mation, the present northwestern margin of Siberia The Fortuna Formation red beds are interpreted
is more likely a mid –late Neoproterozoic truncation as gaining their diagenetic hematite remanence at
of a more extensive Rodinian plate with widespread c. 1150 Ma, according to SHRIMP U –Pb dating
middle Riphean epicratonic cover. of xenotime (D’Agrella-Filho et al. 2008), but that
age assignment is questioned here because the
likely early-diagenetic xenotime U –Pb age may
Amazonia, West Africa, Plata have little bearing on the timing of hematite pig-
mentation in the studied sandstone. The reconstruc-
For the past 25 years, Amazonia has been the craton tion of Amazonia relative to Laurentia shown in
of choice for proposed colliders with the Grenville Figure 9 predicts a younger age of c. 1020 Ma for
390 D. A. D. EVANS
COBRA unites truncated Archean and Palaeo-

proterozoic basement provinces among these
cratons, suggesting that the amalgamation persisted
from the assembly of supercontinent Nuna at 1.8 Ga
(Hoffman 1996) until Rodinia fragmentation in
mid-Neoproterozoic times. In this reconstruction,
2.1–2.3 Ga terranes in subsurface Yukon-Alberta
(Ross 2002) continue into the Birimian (Gasquet
et al. 2004) and Maroni– Itacaiunas (Tassinari
et al. 2000) provinces of West Africa and Amazo-
nia, respectively. The Archean Wyoming/Medicine
Hat craton (Chamberlain et al. 2003) would have
been contiguous with the Nico Perez terrane in
Uruguay (Hartmann et al. 2001) and Luis Alvez
craton in southern Brazil (Sato et al. 2003), con-
stituting parts of an elongate collage of Archean
regions extending to the Leo massif in West
Africa (Thiéblemont et al. 2004) and the Carajas
block in Brazil (Tassinari et al. 2000). Palaeoproter-
Fig. 9. ‘COBRA’ reconstruction of Amazonia, West
ozoic accretion to the south of these provinces
Africa, and Plata near western Laurentia. Among these includes the Mojave province (Bennett & De
three cratons, only Amazonia has reliable palaeomag- Paolo 1987) as the orphaned edge of an extensive
netic constraints from the Rodinia time interval. Cratons region of juvenile 2.2–1.7 Ga terranes in South
are rotated into the present North American reference America (Tassinari et al. 2000; Santos et al. 2000,
frame according to Table 3. 2003), characterized by highly radiogenic
(207Pb-enriched) common-lead isotopic signatures
(Wooden & Miller 1990; Tosdal 1996). Detrital
the growth of remanence-bearing hematite pigments zircons of 1.5–1.9 Ga age in the Mesoproterozoic
in the Fortuna Formation. Belt-Purcell basin (Ross et al. 1992; Ross &
Because the three Amazonia poles fall roughly Villeneuve 2003) find numerous potential sources
along the same great circle, it is possible to consider in extensive granites of that age interval in South
the alternative polarity assignment relative to the America (Tassinari et al. 2000). The 1.3 –1.1 Ga
Laurentian APW path; in that case, however, Grenville orogen traces southwestward through
Amazonia reconstructs directly atop Australia and Sonora (Iriondo et al. 2004) and, according to
Kalahari (not shown in Fig. 9). the COBRA hypothesis, into Brazil and Bolivia,
As recently reviewed by Tohver et al. (2006), where it bifurcates into the Aguapei and Sunsas
palaeomagnetic results from the Meso-Neopro- belts (Sadowski & Bettencourt 1996). Direct juxta-
terozoic of West Africa are wholly unreliable. position of these provinces in Amazonia with SW
For the Plata Craton, Rapalini & Sánchez-Bettucci North America (Santos et al. 2008) is not allowable
(2008) similarly show that there are no reliable palaeomagnetically, by any of the three poles
Rodinian palaeomagnetic constraints. The separ- discussed above, regardless of their precise ages
ation between western Laurentia and Amazonia within the Meso-Neoproterozoic interval.
(Fig. 9), must be filled with cratonic fragments that COBRA is proposed to have begun rifting at
would form the conjugate rift margin of the Cor- 780 Ma, manifested by the Gunbarrel large
dilleran miogeocline in either mid-Neoproterzoic igneous province in North America (Harlan et al.
or terminal Neoproterozoic times (Bond 1997; 2003), and preceding highly oblique dextral separ-
Colpron et al. 2002; Harlan et al. 2003). Given that ation (Brookfield 1993) that prolonged rift mag-
all of the other large cratons of the world have matism to at least 685 Ma (Lund et al. 2003) and
been accounted for in the present Rodinia model, delayed passive-margin thermal subsidence to
the simplest placements of West Africa, Plata and latest Neoproterozoic time (Bond 1997). Precise
smaller cratonic fragments in South America (Fuck geochronology of the Gourma –Volta rift basins in
et al. 2008) are within the gap between Laurentia West Africa, presently lacking, could provide a
and Amazonia (Fig. 9). These juxtapositions are direct test of the proposed COBRA fit. Indications
collectively referred to as COBRA, named after the of c. 780 Ma mafic magmatism within a possible
general link between the proto-Cordilleran rifted West African craton fragment in the westernmost
margin of Laurentia with the proto-Brasiliano/ Hoggar shield (Caby 2003) and along the distal
Pharuside rifted margins of the West Gondwanaland western São Francisco margin in Brazil (Pimentel
cratons. et al. 2004) may extend the Gunbarrel province
into those regions. Proposing a sequence of rifts these relationships. Adding all other cratons’ reli-
in southern South America is difficult due to able palaeomagnetic data from 1100–750 Ma, in
Phanerozoic cover (compare Ramos 1988 and the ‘radically’ revised Rodinia proposed herein,
Cordani et al. 2003), but kinematic constraints has resulted in no additional APW loops.
on a COBRA –West Gondwanaland transition The complexity of Rodinia’s aggregate motion
require some events at c. 780 Ma and others is largely due to oscillatory swings in the APW
younger, represented by glaciogenic successions path between 1200 and 900 Ma, and the newly
on southernmost Amazonia (Trindade et al. 2003) recognized 800 Ma loop. The 1100–1000 Ma
and eastern Rio Plata (Gaucher et al. 2003) that segment, in particular, covers .10 000 km, thus
are correlated to the Marinoan ice age ending at averaging rates of latitudinal motion exceeding
635 Ma (Condon et al. 2005). 10 cm/yr. This would be fast enough for oceanic
plates of the modern world, but it is exceptionally
fast for a plate containing a supercontinent, presum-
Nuna to Rodinia to Gondwanaland ably with numerous lithospheric keels, to slide over
the asthenosphere. An alternative explanation for
Any palaeomagnetic reconstruction of a superconti-
nent requires concordance of data from constituent the majority of Rodinia’s motion, and one which
more easily accommodates its oscillatory nature, is
cratons into a coherent aggregate APW path. Such
that of TPW. Evans (2003) incorporated Rodinia’s
a path for the proposed long-lived Rodinia model
is shown in Figure 10. The numerous loops and latitude shifts as due to oscillatory TPW about a
prolate axis of the geoid inherited from the previous
turns could raise objection due to the implied
supercontinent, Nuna (a.k.a. Columbia).
complexity of the supercontinent’s motion through
its nearly 400 Ma of existence. Nonetheless, all Because the Siberian craton is surrounded on
many sides by c. 1700–1500 Ma rifted passive
of the loops are generated simply by joining the
Laurentian and Baltic APW paths in the proposed margins (Pisarevsky & Natapov 2003), it is likely
to have lain near the centre of Nuna. In contrast,
reconstruction (Fig. 3). The Laurentia þ Baltica
Figure 11 shows Siberia at the edge of Rodinia.
juxtapositions, before and after 1100 Ma, are the
least controversial aspect of the present Rodinia This would suggest that the kinematic evolution
between Nuna and Rodinia was partly ‘extroverted’
model, so the APW complexity of Rodinia will be
implied by any alternative model incorporating (Murphy & Nance 2003, 2005). However, proxi-
mity of the Amazonia, West Africa, Congo þ São
Francisco and Plata cratons in the proposed
Rodinia (Fig. 11) suggests long-lived connections
from the Palaeoproterozoic (similarities noted by
Rogers 1996, and inspiration for his conjectured
‘Atlantica’ assemblage of that age), rearranging
only moderately to form portions of Rodinia
and Gondwanaland. The relationships among this
group of cratons, as well as the longstanding pro-
ximity between Laurentia and Baltica (through
nearly the entire latter half of Earth history)
suggest a more ‘introverted’ kinematic style of
supercontinental evolution.
According to the Rodinia model proposed
herein, assembly took place rapidly at c. 1100 Ma,
although there are earlier collisions of cratons that
persisted into Rodinia time. Table 3 lists the postu-
lated ages of assembly for each craton for the
rotation parameters given relative to Laurentia.
Some of the proposed connections date from the
time of cratonization, typically the Palaeoprotero-
zoic amalgamations of Archean craton, set within
and among juvenile terranes. The relevant cratons
Fig. 10. Rodinia master apparent polar wander path in
in this category are those (West Africa, Plata, Ama-
the present North American (Laurentian) reference
frame. Palaeomagnetic poles are coloured according to zonia) proposed to reconstruct near Laurentia’s
the host cratons as depicted here and in Figure 1, and are proto-Cordilleran margin, where mid –late Neopro-
tabulated with abbreviations in Table 2. Late terozoic rifting cut across truncated basement
Mesoproterozoic (‘Grenvillian’) orogenic belts are fabrics. The hypothesized one-billion-year shared
shaded grey. history of these blocks constitutes a powerful
392 D. A. D. EVANS
Fig. 11. Rodinia radically revised, reconstructed to palaeolatitudes soon after assembly and shortly prior to breakup.
(a) 1070 Ma reconstruction showing extents of late Mesoproterozoic (‘Grenvillian’) orogens, both exposed and inferred
(dark grey), and the earlier collision between Laurentia and Congo (light grey). The figure is slightly anachronistic, as
several blocks are proposed to have collided at 1050 Ma, and South China and Tarim could have joined as late as
c. 850 Ma (Table 3); however, palaeogeographic reconstructions for those younger ages would involve substantial
components of Rodinia over the poles, rendering the Mollweide projection uninformative. (b) 780 Ma reconstruction
showing incipient breakup rift margins (red) and transform offsets (black). Ridge segments are dashed where not
precisely constrained in location.
prediction for future palaeomagnetic tests among reconstructions listed for India, North China and
these data-scarce blocks. Siberia are an example of this latter case; all these
Because Table 3 lists all rotations relative cratons are proposed to have sutured, as a unified
to Laurentia, it does not provide information on plate, to the Rodinia assemblage along the Eastern
relationships between non-Laurentian cratons or Ghats –Rayner orogen.
portions thereof, yet some of these may have simi- Although this model of Rodinia includes
larly ancient connections. For example, as dis- widespread assembly of the supercontinent at
cussed above, one hypothesis resulting from the 1100–1050 Ma, probably the most contentious of
proposed configuration of Congo and Mawsonland, its implications is that all the large cratons are
is that the Tanzania –Bangweulu block was orig- accounted for, and there are no sizable blocks left
inally part of Mawsonland and transferred to to play colliding roles in any of the Sveconorwe-
Congo via collision at c. 1100 Ma and rifting in gian, Grenville and Sunsas orogens. Instead,
the mid-Neoproterozoic. Palaeomagnetism of the these three orogens are placed along strike of each
three Palaeoproterozoic blocks (Congo, Tanzania, other, facing the Mirovian Ocean. All three
Mawsonland) can ultimately test this propo- orogens are characterized by an extensive prehistory
sition. Another example is the proposed Palaeo- of accretionary tectonism along the same margins,
Mesoproterozoic connection among India, North with successively younger age provinces pro-
China, and perhaps also Siberia. gressing outward from Archaean cratonic nuclei.
Where Table 3 presents a range of ages, these The great width and longevity of these three accre-
are set by the limits of palaeomagnetic concordance tionary systems is reminiscient of Panthalassan or
versus discordance when rotated by the given circum-Pangaean orogens of the Phanerozoic. The
Euler parameters. Parenthetical values indicate a model proposed here requires originally farther
best estimate based on geological histories of oceanward extents of the three orogens as younger
either collision or rifting, or by ‘piggy-back’ of an juvenile material would have accreted during the
intervening collision or rift with Laurentia. The early Neoproterzoic. Then, mid-Mirovian spreading
identical-aged c. 1050 Ma onset of proposed Euler ridges would have propagated into the orogens and
Table 3. Rotation parameters to Laurentia: Rodinia radically revised
Craton Maximum age Minimum age Euler rotn.
8N 8E 8CCW
†
Baltica 1120 –1070 615–555 (610) 81.5 250.0 250.0
Australia 1140 –1070 755 (720) 212.5 064.0 134.5
Congo (ex.Tanz) 1375 (1500) 755–550 (720) 224.0 249.0 128.5
Kalahari 1235 –1110 1000 (790?) 32.5 307.5 76.5
India 1100 –770 (1050) 770 (750) 04.0 066.0 93.0
South China 800 (850) 750 (750) 17.5 067.0 178.0
Tarim 755 (850?) 755 (750) 241.5 161.0 57.0
Svalbard cratonization Devonian? 86.0 120.0 248.0
North China 950 (1050) 790 (750?) 42.0 072.0 2113.0
Siberia (Aldan)‡ 1040 (1050) 990 (750?) 32.5 002.5 291.5
Amazonia 1200 (crat.) 980 (630?) 63.0 313.0 2139.0
West Africa cratonization (780) 254.0 246.5 115.0
Plata cratonization (780) 218.5 239.0 125.5
†
Alignments are similar to those discussed in Buchan et al. (2000) and Pisarevsky et al. (2003a).
‡
Reconstruction is similar to option ‘B’ of Pisarevsky & Natapov (2003) and that presented by Meert & Torsvik (2003).
removed the outboard, youngest terrains as ‘ribbon’ 2003; Valeriano et al. 2004; John et al. 2004;
continents. The protracted record of tectonism in Tohver et al. 2006). Although the interval between
the Scottish Highlands, inlcuding the Knoydartian rifting and collision in the Brasiliano foldbelts was
orogeny at c. 850 –800 Ma with further phases brief (‘young, short-lived’ orogenic cycle of Tromp-
possibly as young as c. 750 –700 Ma (reviewed ette 1997), the predominant strike-slip component
by Cawood et al. 2007) could represent the only of motion during assembly allowed those belts to
intact remnants of a once-extensive accretionary contain oceanic (Mirovian) terranes as old as
orogenic belt that lay outboard of the present c. 900–750 Ma (Pimentel & Fuck 1992; Babinski
Grenville orogen. et al. 1996). The prominent dextral shear zones
The present location of these postulated ribbon of the Borborema Province in northeastern Brazil,
terrains is unknown, but the kinematic histories continue into west –central Africa (Vauchez et al.
of more recent examples suggest that they would 1995; Cordani et al. 2003). These bound enigmatic
either be transported strike-slip along the cir- terranes recording unusual ‘Grenvillian’ tecto-
cum-Mirovian subduction girdle around Rodinia nothermal events that are otherwise largely absent
(such as present-day Baja California or the more in cratonic South America (Fuck et al. 2008),
extreme possibility of thousands of kilometres in bearing witness to the large amount of strike-slip
a ‘Baja –British Columbia’ evolution), or separated offset accommodating the assembly of West
far into Mirovia toward an unprescribed fate (such Gondwanaland.
as present Zealandia). Using these analogies, we On the other side of the proposed Rodinia, the
might expect to find them today as dismembered kinematic evolution toward East Gondwanaland
basement units within the Avalonian–Cadomian follows more conventional reconstructions, which
orogen (Evans 2005; Murphy et al. 2000; Keppie comes as little surprise because the relative pos-
et al. 2003) or Borborema –Pharuside strike-slip- itions of Australia, India, South China, and Tarim
dominated orogenic system (Caby 2003), or are similar to those earlier models. India migrated
perhaps partly to completely recycled into the sinistrally along the Pinjarra orogen to arrive at its
mantle by subduction-erosion (Scholl & von Gondwanaland position relative to Australia by
Huene 2007). c. 550 Ma (Powell & Pisarevsky 2002). South
Figure 11b shows the incipient breakup of China and Tarim would have lain along the same
Rodinia at 780 Ma, according to the revised tectonic plate during that time, arriving at accepta-
Rodinia model. A first stage of disaggregation at ble positions for their palaeomagnetic recon-
c. 780 –720 Ma around the western and northern struction into Gondwanaland (Zhang 2004; Zhao
margins of Laurentia, liberated the Congo, West et al. 2008). In the palaeogeographic co-ordinate
African, Amazonian and Plata cratons that would system of 780 Ma (Fig. 11b), Siberia would have
eventually recombine to form West Gondwanaland rifted to the east, separating from East Gondwana-
between c. 640 and c. 530 Ma (Trompette land fragments. It is debatable whether North
1997; Brito Neves et al. 1999; Piuzana et al. China was part of Palaeozoic Gondwanaland; if
394 D. A. D. EVANS
not, it too may have rifted far away with Siberia. a general model for Rodinia was conceived
Kalahari would have migrated to the north in the (Hoffman 1991), which, despite numerous chal-
reconstructed co-ordinate system of Figure 11b, lenges from both tectonostratigraphy and palaeo-
joining the West Gondwanaland cratons as they magnetism, has remained largely intact in the
drifted away from Laurentia þ Baltica. Final dis- latest consensus model (Li et al. 2008). However,
aggregation of Rodinia occurred c. 610– 550 Ma, in order to accommodate the palaeomagnetic
the age of extensive mafic magmatism in eastern data in particular, this latest model has shortened
Laurentia (reviewed by Cawood et al. 2001; the duration of Rodinia’s existence to merely
Puffer 2002) and Norway (Svenningsen 2001). 75 Ma (900– 825 Ma). Such brevity is acceptable
Global palaeogeography at the end of the Neo- and actualistic in terms of the short-lived Pangaea
proterozoic Era remains one of the most challenging landmass, but it appears at odds with Hoffman’s
problems in palaeomagnetic reconstruction, more (1991) original implication of globally widespread
difficult even than the quest for Rodinia. This is 1300–1000 Ma orogens and c. 750 Ma rifted
due to four factors: (1) lack of high-precision bio- margins, as representing Rodinia’s assembly and
stratigraphy in the Precambrian to correlate suc- breakup, respectively.
cessions and to date palaeomagnetic poles from Herein, I have proposed a Rodinia model that is
sedimentary rocks; (2) scarcity of datable volcanic both long-lived according to the original concept,
successions on the large cratons, relative to geon and compatible with the most reliable palaeomag-
7; (3) likelihood that most cratons were travell- netic data from the Meso-Neoproterozoic interval,
ing independently during the transition between with minimal number of APW loops. My model is
Rodinia and Gondwanaland, thus disallowing the a radical departure from all previous models (e.g.
APW superposition method used in this paper; and Li et al. 2008). Which existing Rodinia model,
(4) abnormally high dispersion of palaeomagnetic if any, will approximate the ‘true’ form of the
poles from each craton indicating either rapid Neoproterozoic supercontinent? As Wegener
plate tectonics, rapid TPW, or a non-uniformitarian [1929 (1966, p. 17)] wrote: ‘the earth at any one
geomagnetic field during that time. The most com- time can only have had one configuration.’ How
plete model incorporating the global tectonic will we test the current Rodinia models and
record and palaeomagnetic data is by Collins & achieve a long-lasting consensus that converges
Pisarevsky (2005), but this model still needed to toward the true palaeogeography?
resort to separate options of a low- versus high- Dalziel (1999) identified six criteria for validity
latitude subset of the Laurentian palaeomagnetic of a ‘credible’ supercontinent: (1) account for
data. If TPW is responsible for the large dispersions all rifted passive margins at the time of breakup;
in palaeomagnetic poles, which if read literally (2) accurately map continental promontories
would typically imply oscillatory motions conform- and embayments, that is in spherical geometry; (3)
ing to the IITPW model of Evans (2003), then there display sutures related to assembly; (4) match
is some hope to produce reconstructions using the older tectonic fabrics where appropriate; (5) show
long-lived prolate nonhydrostatic geoid as the refer- compatibility with palaeomagnetic data; and (6)
ence axis, rather than the geomagnetic-rotational be compatible with realistic kinematic evolution
reference frame (Raub et al. 2007). This alternative forward in time toward Pangaea. The revised
method, however, produces reconstructions that Rodinia model proposed herein satisfies all six of
are highly sensitive to small errors in magnetiza- these conditions, if one allows for a special
tion ages, depending on the rapidity of the putative consideration involving conjugate rifted and colli-
TPW oscillations. Regardless of which class of sional margins, as follows: the past few years of
interpretations will ultimately prove valid, questions palaeogeographic reconstruction of the Mesozoic –
such as the widths of Iapetan separation follow- Cenozoic world have led to increasing recogniztion
ing Rodinian juxtaposition of Amazonia with of ribbon-shaped continental fragments with lengths
eastern Laurentia (e.g. Cawood et al. 2001), must on the order of thousands of kilometres (e.g.
be considered premature until more precisely Lomonosov Ridge; Lawver et al. 2002; Lord Howe
dated palaeomagnetic poles are obtained. Rise/Zealandia; Gaina et al. 2003). Farther back in
time, the Cimmeride continental ribbon formed the
Permian rift conjugate to the .5000 km northern
Concluding Remarks passive margin of Gondwanaland (Stampfli &
Borel 2002). Tectonic shuffling and reworking of
Nearly two decades have elapsed since McMenamin Cimmeride blocks within the Alpine-Himalayan
& McMenamin (1990, p. 95) coined the name orogenic collage has largely obscured their original
‘Rodinia’ for the late Proterozoic supercontinent geometric continuity. The Lomonosov Ridge and
and ‘Mirovia’ for its encircling palaeo-ocean. Lord Howe Rise/Zealandia ribbons are yet to mig-
Within a year of these monikers’ establishment, rate to their final dispositions within accretionary
orogens, but they are unlikely to arrive in pristine B ERRY , R. F., H OLM , O. H. & S TEELE , D. A. 2005.
form. I propose that similar effects may hamper Chemical U– Th– Pb monazite dating and the Protero-
our ability to quantify the passive margin lengths zoic history of King Island, southeast Australia.
of any Precambrian continental ribbons. In the Australian Journal of Earth Sciences, 52, 461– 471.
B INGEN , B., D EMAIFFE , D. & VAN B REEMEN , O. 1998.
case of Zealandia, separation from Australia þ The 616 Ma old Egersund basaltic dike swarm, SW
Antarctica roughly followed the geometry of the Norway, and late Neoproterozoic opening of the
Terra Australis orogen (Cawood 2005), thus Iapetus Ocean. Journal of Geology, 106, 565– 574.
ending a Wilson cycle without a continent-continent B LEWETT , R. S. & B LACK , L. P. 1998. Structural and tem-
collision. In the Transantarctic Mountains, the rift poral framework of the Coen Region, north Queens-
propagated far enough inboard to bring some of land: implications for major tectonothermal events in
the oldest, most internal segments of the belt (Cam- east and north Australia. Australian Journal of Earth
brian –Ordovician Ross orogen) directly in contact Sciences, 45, 597–609.
with the oceanic passive margin. Would future B OGDANOVA , S. V., B INGEN , B., G ORBATSCHEV , R.,
K HERASKOVA , T. N., K OZLOV , V. I., P UCHKOV ,
palaeogeographers interpret this record as one of V. N. & V OLOZH , YU . A. 2008. The East European
Cambrian –Ordovician continent–continent col- Craton (Baltica) before and during the assembly of
lision, followed by tectonic stability inside a super- Rodinia. Precambrian Research, 160, 23–45.
continent, and subsequent Mesozoic breakup of that B OND , G. C. 1997. New constraints on Rodinia breakup
supercontinent? This example highlights the diffi- ages from revised tectonic subsidence curves. Geo-
culty in robustly characterizing tectonic histories logical Society of America Abstracts with Programs,
of Precambrian orogens without a palaeogeographic 29(6), 280.
framework. With focused effort on obtaining key B OND , G. C., N ICKESON , P. A. & K OMINZ , M. A. 1984.
geochronologic and palaeomagnetic data from the Breakup of a supercontinent between 625 Ma and
555 Ma: new evidence and implications for continental
Rodinian time interval, we may be able to provide histories. Earth and Planetary Science Letters, 70,
that framework. 325– 345.
B ORG , S. G. & D E P AOLO , D. J. 1994. Laurentia,
These ideas have evolved through the last five years, Australia, and Antarctica as a Late Proterozoic super-
benefiting from discussions with S. Bryan, P. Cawood, continent: constraints from isotopic mapping.
A. Collins, B. Collins, D. Corrigan, I. Dalziel, B. de Geology, 22, 307–310.
Waele, R. Ernst, I. Fitzsimons, R. Fuck, C. Gower, B RIDEN , J. C., D UFF , B. A. & K RÖNER , A. 1979. Palaeo-
P. Hoffman, K. Karlstrom, A. Kröner, Z.-X. Li, A. magnetism of the Koras Group, Northern Cape
Maloof, B. Murphy, V. Pease, M. Pimentel, S. Pisarevsky, Province, South Africa. Precambrian Research, 10,
the late C. Powell, T. Rivers, D. Thorkelson, E. Tohver 43–57.
and M. Wingate. P. Hoffman, I. Dalziel, C. MacNiocaill, B RITO N EVES , B. B., C AMPOS N ETO , M. C. & F UCK ,
and R. Van der Voo provided helpful criticisms on R. A. 1999. From Rodinia to Western Gondwana: an
earlier versions of the manuscript; P. Cawood and approach to the Brasiliano-Pan African Cycle and
S. Pisarevsky gave expedient and constructive reviews of orogenic collage. Episodes, 22, 155– 166.
this iteration. This work was supported by a Fellowship B ROOKFIELD , M. E. 1993. Neoproterozoic Laurentia-
in Science and Engineering from the David and Lucile Australia fit. Geology, 21, 683– 686.
Packard Foundation. B ROWN , L. L. & M C E NROE , S. A. 2004. Palaeomagnet-
ism of the Egersund-Ogna anorthosite, Rogaland,
Norway, and the position of Fennoscandia in the Late
Proterozoic. Geophysical Journal International, 158,
References 479– 488.
B UCHAN , K. L. & H ALLS , H. C. 1990. Paleomagnetism of
A LKMIM , F. F., M ARSHAK , S. & F ONSECA , M. A. 2001. Proterozoic mafic dyke swarms of the Canadian Shield.
Assembling West Gondwana in the Neoproterozoic: In: P ARKER , A. J., R ICKWOOD , P. C. & T UCKER ,
clues from the São Francisco craton region, Brazil. D. H. (eds) Mafic dykes and emplacement mechanisms.
Geology, 29, 319– 322. Balkema, Rotterdam, 209– 230.
B ABINSKI , M., C HEMALE , F., J R ., H ARTMANN , L. A., B UCHAN , K. L., M ERTANEN , S., P ARK , R. G., P ESONEN ,
V AN S CHMUS , W. R. & DA S ILVA , L. C. 1996. L. J., E LMING , S.-Å., A BRAHAMSEN , N. & B YLUND ,
Juvenile accretion at 750– 700 Ma in southern Brazil. G. 2000. Comparing the drift of Laurentia and Baltica
Geology, 24, 439– 442. in the Proterozoic: the importance of key palaeomag-
B ARTLEY , J. K., S EMIKHATOV , M. A., K AUFMAN , A. J., netic poles. Tectonophysics, 319, 167 –198.
K NOLL , A. H., P OPE , M. C. & J ACOBSEN , S. B. B UCHAN , K. L., E RNST , R. E., H AMILTON , M. A.,
2001. Global events across the Mesoproterozoic- M ERTANEN , S., P ESONEN , L. J. & E LMING , S.-Å.
Neoproterozoic boundary: C and Sr isotopic evidence 2001. Rodinia: the evidence from integrated palaeo-
from Siberia. Precambrian Research, 111, 165– 202. magnetism and U–Pb geochronology. Precambrian
B ENNETT , V. C. & D E P AOLO , D. J. 1987. Proterozoic Research, 110, 9– 32.
crustal history of the western United States as deter- B URRETT , C. & B ERRY , R. 2000. Proterozoic Australia-
mined by neodymium isotopic mapping. Geological Western United States (AUSWUS) fit between
Society of America Bulletin, 99, 674–685. Laurentia and Australia. Geology, 28, 103 –106.
396 D. A. D. EVANS
C ABY , R. 2003. Terrane assembly and geodynamic evol- Neoproterozoic Doushantuo Formation, China.
ution of central-western Hoggar: a synthesis. Journal Science, 308, 95– 98.
of African Earth Sciences, 37, 133–159. C ORDANI , U. G., D’A GRELLA -F ILHO , M. S., B RITO -
C ATELANI , E. L., E VANS , D. A. D., S MIRNOV , A. V., N EVES , B. B. & T RINDADE , R. I. F. 2003. Tearing
T RINDADE , R. I. F., D’A GRELLA -F ILHO , M. S. & up Rodinia: the Neoproterozoic palaeogeography of
H EAMAN , L. M. 2007. Revisiting the Meso- South American cratonic fragments. Terra Nova, 15,
Neoproterozoic dykes of Bahia, Brazil. EOS, Trans- 350–359.
actions of the American Geophysical Union, Joint D’A GRELLA -F ILHO , M. S., P ACCA , I. G., R ENNE , P. R.,
Assembly, Acapulco, Mexico. O NSTOTT , T. C. & T EIXEIRA , W. 1990. Paleomagnet-
C AWOOD , P. A. 2005. Terra Australis Orogen: rodinia ism of Middle Proterozoic (1.01 to 1.08 Ga) mafic
breakup and development of the Pacific and Iapetus dykes in southeastern Bahia State – São Francisco
margins of Gondwana during the Neoproterozoic and Craton, Brazil. Earth and Planetary Science Letters,
Paleozoic. Earth-Science Reviews, 69, 249– 279. 101, 332 –348.
C AWOOD , P. A. & K ORSCH , R. J. 2008. Assembling D’A GRELLA -F ILHO , M. S., T RINDADE , R. I. F.,
Australia: Proterozoic building of a continent. Precam- S IQUEIRA , R., P ONTE -N ETO , C. F. & P ACCA ,
brian Research, 166, 1 –38. I. I. G. 1998. Paleomagnetic constraints on the
C AWOOD , P. A. & P ISAREVSKY , S. A. 2006. Was Baltica Rodinia supercontinent: implications for its Neopro-
right-way-up or upside-down in the Neoproterozoic? terozoic break-up and the formation of Gondwana.
Journal of the Geological Society, London, 163, 1–7. International Geology Review, 40, 171–188.
C AWOOD , P. A., M C C AUSLAND , P. J. A. & D UNNING , G. D’A GRELLA -F ILHO , M. S., P ACCA , I. G., E LMING , S.-Å.,
R. 2001. Opening Iapetus: Constraints from the Laur- T RINDADE , R. I., G ERALDES , M. C. & T EIXEIRA , W.
entian margin of Newfoundland. Geological Society 2003. Paleomagnetic and rock magnetic studies of
of America Bulletin, 113, 443–453. Proterozoic mafic rocks from western Mato Grosso
C AWOOD , P. A., N EMCHIN , A. A., S TRACHAN , R., State, Amazonian craton. Geophysical Research
P RAVE , T. & K RABBENDAM , M. 2007. Sedimentary Abstracts, 5, no. 05619.
basin and detrital zircon record along East Laurentia D’A GRELLA -F ILHO , M. S., P ACCA , I. I. G., T RINDADE ,
and Baltica during assembly and breakup of Rodinia. R. I. F., T EIXEIRA , W., R APOSO , M. I. B. &
Journal of the Geological Society, London, 164, O NSTOTT , T. C. 2004. Paleomagnetism and
40
257– 275. Ar/39Ar ages of mafic dikes from Salvador
C HAMBERLAIN , K. R., F ROST , C. D. & F ROST , B. R. (Brazil): new constraints on the São Francisco craton
2003. Early Archean to Mesoproterozoic evolution of APW path between 1080 and 1010 Ma. Precambrian
the Wyoming Province: archaean origins to modern Research, 132, 55– 77.
lithospheric architecture. Canadian Journal of Earth D’A GRELLA -F ILHO , M. S., T OHVER , E., S ANTOS ,
Sciences, 40, 1357– 1374. J. O. S., E LMING , S.-Å., T RINDADE , R. I. F.,
C HATTERJEE , N., C ROWLEY , J. L. & G HOSE , N. C. 2008. P ACCA , I. I. G. & G ERALDES , M. C. 2008. Direct
Geochronology of the 1.55 Ga Bengal anorthosite and dating of paleomagnetic results from Precambrian
Grenvillian metamorphism in the Chhotanagpur gneis- sediments in the Amazon craton: evidence for Grenvil-
sic complex, eastern India. Precambrian Research, lian emplacement of exotic crust in SE Appalachians of
161, 303– 316. North America. Earth and Planetary Science Letters,
C HEN , Y., X U , B., Z HAN , S. & L I , Y. 2004. First mid- 267, 188 –199.
Neoproterozoic paleomagnetic results from the Tarim D ALZIEL , I. W. D. 1991. Pacific margins of Laurentia and
Basin (NW China) and their geodynamic implications. East Antarctica-Australia as a conjugate rift pair:
Precambrian Research, 133, 271–281. evidence and implications for an Eocambrian super-
C LARK , D. J., H ENSEN , B. J. & K INNY , P. D. 2000. Geo- continent. Geology, 19, 598– 601.
chronological constraints for a two-stage history of the D ALZIEL , I. W. D. 1997. Neoproterozoic-Paleozoic
Albany-Fraser Orogen, Western Australia. Precam- geography and tectonics: review, hypothesis, environ-
brian Research, 102, 155 –183. mental speculation. Geological Society of America
C OGNÉ , J. P. 2003. PaleoMac: a MacintoshTM application Bulletin, 109, 16– 42.
for treating paleomagnetic data and making plate D ALZIEL , I. W. D. 1999. Global paleotectonics; recon-
reconstructions. Geochemistry Geophysics Geosys- structing a credible supercontinent. Geological
tems, 4(1), doi: 10.1029/2001GC000227. Society of America Abstracts with Programs,
C OLLINS , A. S. & P ISAREVSKY , S. A. 2005. Amalgamat- 31(7), 316.
ing eastern Gondwana: the evolution of the Circum- D ALY , M. C., L AWRENCE , S. R., D IEMU -T SHIBAND , K.
Indian Orogens. Earth-Science Reviews, 71, 229–270. & M ATOUANA , B. 1992. Tectonic evolution of the
C OLPRON , M., L OGAN , J. M. & M ORTENSEN , J. K. 2002. Cuvette Centrale, Zaire. Journal of the Geological
U– Pb zircon age constraint for late Neoproterozoic Society, London, 149, 539– 546.
rifting and initiation of the lower Paleozoic passive D EB , M., T HORPE , R. I., K RSTIC , D., C ORFU , F. &
margin of western Laurentia. Canadian Journal of D AVIS , D. W. 2001. Zircon U–Pb and galena Pb
Earth Sciences, 39, 133– 143. isotope evidence for an approximate 1.0 Ga terrane
C ONDIE , K. C. 2002. The supercontinent cycle: are there constituting the western margin of the Aravalli-Delhi
two patterns of cyclicity? Journal of African Earth orogenic belt, northwestern India. Precambrian
Sciences, 35, 179–183. Research, 108, 195–213.
C ONDON , D., Z HU , M., B OWRING , S., W ANG , W., Y ANG , D EBLOND , A., P UNZALAN , L. E., B OVEN , A. & T ACK , L.
A. & J IN , Y. 2005. U–Pb ages from the 2001. The Malagarazi Supergroup of southeast
Burundi and its correlative Bukoba Supergroup of F IORETTI , A. M., B LACK , L. P., F ODEN , J. & V ISONA , D.
northwest Tanzania: neo- and Mesoproterozoic chron- 2005. Grenville-age magmatism at the South Tasman
ostratigraphic constraints from Ar– Ar ages on mafic Rise (Australia): a new piercing point for the recon-
intrusive rocks. Journal of African Earth Sciences, struction of Rodinia. Geology, 33, 769– 772.
32, 435–449. F ITZSIMONS , I. C. W. 2003. Proterozoic basement
D E W AELE , B., J OHNSON , S. P. & P ISAREVSKY , S. A. provinces of southern and southwestern Australia,
2008. Palaeoproterozoic to Neoproterozoic growth and their correlation with Antarctica. In: Y OSHIDA ,
and evolution of the eastern Congo Craton: its role in M., W INDLEY , B. F. & D ASGUPTA , S. (eds) Protero-
the Rodinia puzzle. Precambrian Research, 160, zoic East Gondwana: Supercontinent Assembly and
127–141. Breakup. Geological Society, London, Special Publi-
D IEHL , J. F. & H AIG , T. D. 1994. A paleomagnetic study cations, 206, 35– 55.
of the lava flows within the Copper Harbor Conglom- F UCK , R. A., B RITO N EVES , B. B. & S CHOBBENHAUS , C.
erate, Michigan: new results and implications. Cana- 2008. Rodinia descendants in South America. Precam-
dian Journal of Earth Sciences, 31, 369–380. brian Research, 160, 108– 126.
D RÜPPEL , K., L ITTMANN , S., R OMER , R. L. & G AINA , C., M ÜLLER , R. D., B ROWN , B. J. & I SHIHARA ,
O KRUSCH , M. 2007. Petrology and isotope geochem- T. 2003. Microcontinent formation around Australia.
istry of the Mesoproterozoic anorthosite and related In: H ILLS , R. R. & M ÜLLER , R. D. (eds) Evolution
rocks of the Kunene Intrusive Complex, NW and Dynamics of the Australian Plate. Geological
Namibia. Precambrian Research, 156, 1– 31. Society of America, Special Paper, 372, 405–416.
E RNST , R. E. & B UCHAN , K. L. 1993. A paleomagnetic G ALLET , Y., P AVLOV , V. E., S EMIKHATOV , M. A. &
study of the lava flows within the Copper Harbor P ETROV , P. YU . 2000. Late Mesoproterozoic
Conglomerate, Michigan: new results and impli- magnetostratigraphic results from Siberia: paleogeo-
cations. Canadian Journal of Earth Sciences, 30, graphic implications and magnetic field behavior.
1886–1897. Journal of Geophysical Research, 105(B7),
E RNST , R. E. & B UCHAN , K. L. 2002. Maximum size and 16481–16499.
distribution in time and space of mantle plumes: G ASQUET , D., B ARBEY , P., A DOU , M. & P AQUETTE ,
evidence from large igneous provinces. Journal of J. L. 2004. Structure, Sr-Nd isotope geochemistry
Geodynamics, 34, 309–342. and zircon U– Pb geochronology of the granitoids of
E VANS , D. A. D. 2000. Stratigraphic, geochronological, the Dabakala area (Côte d’Ivoire): evidence for a
and paleomagnetic constraints upon the Neoprotero- 2.3 Ga crustal growth event in the Palaeoproterozoic
zoic climatic paradox. American Journal of Science, of West Africa? Precambrian Research, 127,
300, 347– 433. 329– 354.
E VANS , D. A. D. 2003. True polar wander and superconti- G AUCHER , C., B OGGIANI , P. C., S PRECHMANN , P., S IAL ,
nents. Tectonophysics, 362, 303–320. A. N. & F AIRCHILD , T. 2003. Integrated correlation of
E VANS , D. A. D. 2005. Do Avalonian terranes contain the the Vendian to Cambrian Arroyo del Soldado and
lost Grenvillian hinterland? Geological Association of Corumbá Groups (Uruguay and Brazil): palaeo-
Canada Abstracts, 30, 54– 55. geographic, palaeoclimatic and palaeobiologic impli-
E VANS , D. A. D. 2006. Proterozoic low orbital obliquity cations. Precambrian Research, 120, 241– 278.
and axial-dipolar geomagnetic field from evaporite G IDDINGS , J. W. 1976. Precambrian palaeomagnetism in
palaeolatitudes. Nature, 444, 51–55. Australia I: basic dykes and volcanics from the
E VANS , D. A. D. & P ISAREVSKY , S. A. 2008. Plate tec- Yilgarn Block. Tectonophysics, 30, 91–108.
tonics on early Earth? Weighing the paleomagnetic G LASMACHER , U. A., B AUER , W. ET AL . 2001. The meta-
evidence. In: C ONDIE , K. C. & P EASE , V. (eds) morphic complex of Beloretzk, SW Urals, Russia—a
When Did Plate Tectonics Begin on Earth? Geological terrane with a polyphase Meso- to Neoproterozoic
Society of America, Special Paper, 440, 249– 263. thermo-dynamic evolution. Precambrian Research,
E VANS , D. A. D., L I , Z. X., K IRSCHVINK , J. L. & 110, 185–213.
W INGATE , M. T. D. 2000. A high-quality mid- G OODGE , J. W., F ANNING , C. M. & B ENNETT , V. C.
Neoproterozoic paleomagnetic pole from South 2001. U–Pb evidence of c. 1.7 Ga crustal tectonism
China, with implications for ice ages and the breakup during the Nimrod Orogeny in the Transantarctic
configuration of Rodinia. Precambrian Research, Mountains, Antarctica: implications for Proterozoic
100, 313– 334. plate reconstructions. Precambrian Research, 112,
E VANS , D. A. D., B EUKES , N. J. & K IRSCHVINK , J. L. 261– 288.
2002. Paleomagnetism of a lateritic paleo-weathering G OODGE , J. W., V ERVOORT , J. D. ET AL . 2008. A posi-
horizon and overlying Paleoproterozoic redbeds from tive test of East Antarctica – Laurentia juxtaposition
South Africa: implications for the Kaapvaal apparent within the Rodinia supercontinent. Science, 321,
polar wander path and a confirmation of atmospheric 235– 240.
oxygen enrichment. Journal of Geophysical Research, G OSE , W. A., H ANSON , R. E., D ALZIEL , I. W. D.,
107(B12), doi: 10.1029/2001JB000432. P ANCAKE , J. A. & S EIDEL , E. K. 2006. Paleomagnet-
F AHRIG , W. F., C HRISTIE , K. W. & J ONES , D. L. 1981. ism of the 1.1 Ga Umkondo large igneous province in
Paleomagnetism of the Bylot basins: evidence for southern Africa. Journal of Geophysical Research,
Mackenzie continental tensional tectonics. In: 111, B09101, doi: 10.1029/2005JB03897.
C AMPBELL , F. H. A. (ed.) Proterozoic Basins of G OWER , C. F., K AMO , S. & K ROGH , T. E. 2008. Indentor
Canada. Geological Survey of Canada, Paper, 81-10, tectonism in the eastern Grenville Province. Precam-
303–312. brian Research, 167, 201– 212.
398 D. A. D. EVANS
G OWER , C. F., R YAN , A. B. & R IVERS , T. 1990. Mid- H OFFMAN , P. F. 1996. Tectonic genealogy of North
Proterozoic Laurentia-Baltica: an overview of its geo- America. In: V AN DER P LUIJM , B. A. & M ARSHAK ,
logical evolution and a summary of the contributions S. (eds) Earth Structure: an Introduction to Structural
made by this volume. In: G OWER , C. F., R IVERS , T. Geology and Tectonics. McGraw-Hill, New York,
& R YAN , A. B. (eds) Mid-Proterozoic Laurentia- 459–464.
Baltica. Geological Association of Canada, Special H UANG , B., X U , B., Z HANG , C., L I , Y. & Z HU , R. 2005.
Paper, 38, 1 –20. Paleomagnetism of the Baiyisi volcanic rocks (c.
G REGORY , L. C., M EERT , J. G., P RADHAN , V., P ANDIT , 740 Ma) of Tarim, Northwest China: a continental
M. K., T AMRAT , E. & M ALONE , S. J. 2006. A paleo- fragment of Neoproterozoic Western Australia? Pre-
magnetic and geochronologic study of the Majhgawan cambrian Research, 142, 83– 92.
kimberlite, India: implications for the age of the Upper H YODO , H. & D UNLOP , D. J. 1993. Effect of anisotropy
Vindhyan Supergroup. Precambrian Research, 149, on the paleomagnetic contact test for a Grenville
65–75. dike. Journal of Geophysical Research, 98,
G RUNOW , A. M., K ENT , D. V. & D ALZIEL , I. W. D. 1991. 7997– 8017.
New paleomagnetic data from Thurston Island: impli- I DNURM , M. & G IDDINGS , J. W. 1995. Paleoproterozoic-
cations for the tectonics of West Antarctica and Neoproterozoic North America-Australia link:
Weddell Sea opening. Journal of Geophysical new evidence from paleomagnetism. Geology, 23,
Research, 96(B11), 17935–17954. 149–152.
H ALLS , H. C. 1974. A paleomagnetic reversal in the Osler I RIONDO , A., P REMO , W. R. ET AL . 2004. Isotopic, geo-
Volcanic Group, northern Lake Superior. Canadian chemical, and temporal characterization of Proterozoic
Journal of Earth Sciences, 11, 1200–1207. basement rocks in the Quitovac region, northwestern
H ALLS , H. C. & P ESONEN , L. J. 1982. Paleomagnetism of Sonora, Mexico: implications for the reconstruction
Keweenawan rocks. In: W OLD , R. J. & H INZE , W. J. of the southwestern margin of Laurentia. Geological
(eds) Geology and Tectonics of the Lake Superior Society of America Bulletin, 116, 154 –170.
Basin. Geological Society of America, Memoir, 156, I RVING , E., B AKER , J., H AMILTON , M. & W YNNE , P. J.
173– 201. 2004. Early Proterozoic geomagnetic field in western
H AMILTON , M. A. & B UCHAN , K. L. 2007. U –Pb geo- Laurentia: implications for paleolatitudes, local
chronology of the Western Channel Diabase, rotations and stratigraphy. Precambrian Research,
Wopmay Orogen: implications for the APWP for Laur- 129, 251 –270.
entia in the earliest Mesoproterozoic. Geological J ACKSON , G. D. & I ANNELLI , T. R. 1981. Rift-related
Association of Canada Abstracts, 32, 35–36. cyclic sedimentation in the Neohelikian Borden
H ANSON , R. E., C ROWLEY , J. L. ET AL . 2004. Coeval Basin, northern Baffin Island. In: C AMPBELL ,
large-scale magmatism in the Kalahari and Laurentian F. H. A. (ed.) Proterozoic basins of Canada. Geologi-
cratons during Rodinia assembly. Science, 304, cal Survey of Canada, Paper, 81-10, 269– 302.
1126–1129. J ACKSON , M. P. A., W ARIN , O. N., W OAD , G. M. &
H ARLAN , S. S., G EISSMAN , J. W. & S NEE , L. W. 1997. H UDEC , M. R. 2003. Neoproterozoic allochthonous
Paleomagnetic and 40Ar/39Ar geochronologic data salt tectonics during the Lufilian orogeny in the Katan-
from late Proterozoic mafic dykes and sills, Montana gan Copperbelt, central Africa. Geological Society of
and Wyoming. USGS Professional Paper, 1580, 1– 16. America Bulletin, 115, 314–330.
H ARLAN , S. S., H EAMAN , L., L E C HEMINANT , A. N. & J ACOBS , J., P ISAREVSKY , S., T HOMAS , R. J. & B ECKER ,
P REMO , W. R. 2003. Gunbarrel mafic magmatic T. 2008. The Kalahari Craton during the assembly and
event: a key 780 Ma time marker for Rodinia plate dispersal of Rodinia. Precambrian Research, 160,
reconstructions. Geology, 31, 1053–1056. 142–158.
H ARTMANN , L. A., C AMPAL , N., S ANTOS , J. O. S., J ACOBS , J. & T HOMAS , R. J. 2004. Himalayan-type
M C N AUGHTON , N. J., B OSSI , J., S CHIPILOV , A. & indenter-escape tectonics model for the southern part
L AFON , J.-M. 2001. Archean crust in the Rio de la of the late Neoproterozoic– early Paleozoic East
Plata Craton, Uruguay—SHRIMP U–Pb zircon recon- African-Antarctic orogen. Geology, 32, 721–724.
naissance geochronology. Journal of South American J IANG , G., S OHL , L. E. & C HRISTIE -B LICK , N. 2003.
Earth Sciences, 14, 557– 570. Neoproterozoic stratigraphic comparison of the
H ARTZ , E. H. & T ORSVIK , T. H. 2002. Baltica upside lesser Himalaya (India) and Yangtze block (south
down: a new plate tectonic model for Rodinia and China): paleogeographic implications. Geology, 31,
the Iapetus Ocean. Geology, 30, 255– 258. 917–920.
H EAMAN , L. M., L E C HEMINANT , A. N. & R AINBIRD , J OHANSSON , Å, G EE , D. G., L ARIONOV , A. N., O HTA , Y.
R. H. 1992. Nature and timing of Franklin igneous & T EBENKOV , A. M. 2005. Grenvillian and Caledo-
events, Canada: implications for a Late Proterozoic nian evolution of eastern Svalbard–a tale of two oroge-
mantle plume and the breakup of Laurentia. Earth nies. Terra Nova, 17, 317– 325.
and Planetary Science Letters, 109, 117 –131. J OHN , T., S CHENK , V., M EZGER , K. & T EMBO , F. 2004.
H ENRY , S. G., M AUK , F. J. & V AN DER V OO , R. 1977. Timing and PT evolution of whiteschist metamorphism
Paleomagnetism of the upper Keweenawan sediments: in the Lufilian Arc-Zambezi Belt orogen (Zambia):
the Nonesuch Shale and Freda Sandstone. Canadian implications for the assembly of Gondwana. Journal
Journal of Earth Sciences, 14, 1128–1138. of Geology, 112, 71–90.
H OFFMAN , P. F. 1991. Did the breakout of Laurentia J OHNSON , S. P., R IVERS , T. & D E W AELE , B. 2005. A
turn Gondwanaland inside-out? Science, 252, review of the Mesoproterozoic to early Palaeozoic
1409–1412. magmatic and tectonothermal history of south-central
Africa: implications for Rodinia and Gondwana. K RÖNER , A. & C ORDANI , U. G. 2003. African, southern
Journal of the Geological Society, London, 162, Indian and South American cratons were not part of
433–450. the Rodinia supercontinent: evidence from field
K AMO , S. L. & G OWER , C. F. 1994. Note: U– Pb badde- relationships and geochronology. Tectonophysics,
leyite dating clarifies age of characteristic paleomag- 375, 325–352.
netic remanence of Long Range dykes, southeastern K UMAR , A., H EAMAN , L. M. & M ANIKYAMBA , C. 2007.
Labrador. Atlantic Geology, 30, 259– 262. Mesoproterozoic kimberlites in south India: a possible
K AMO , S. L., G OWER , C. F. & K ROGH , T. E. 1989. Birth- link to c 1.1 Ga global magmatism. Precambrian
date for the Iapetus Ocean? A precise U–Pb zircon and Research, 154, 192– 204.
baddeleyite age for the Long Range dikes, southeast L AWVER , L. A., G RANTZ , A. & G AHAGAN , L. M. 2002.
Labrador. Geology, 17, 602– 605. Plate kinematic evolution of the present Arctic region
K ARLSTROM , K. E., H ARLAN , S. S., W ILLIAMS , M. L., since the Ordovician. In: M ILLER , E. L., G RANTZ ,
M C L ELLAND , J., G EISSMAN , J. W. & A HALL , K. I. A. & K LEMPERER , S. L. (eds) Tectonic Evolution of
1999. Refining Rodinia: geologic evidence for the the Bering Shelf–Chukchi Sea– Arctic Margin and
Australia-Western U.S. connection in the Proterozoic. Adjacent Landmasses. Geological Society of
GSA Today, 9, 1 –7. America, Special Paper, 360, 333– 358.
K ARLSTROM , K. E., Å HALL , K.-I., H ARLAN , S. S., L E C HEMINANT , A. N. & H EAMAN , L. M. 1989. Macken-
W ILLIAMS , M. L., M C L ELLAND , J. & G EISSMAN , zie igneous events, Canada: middle Proterozoic
J. W. 2001. Long-lived (1.8– 1.0 Ga) convergent hotspot magmatism associated with ocean opening.
orogen in southern Laurentia, its extensions to Earth and Planetary Science Letters, 96, 38–48.
Australia and Baltica, and implications for refining L I , Z.-X. 2000. Palaeomagnetic evidence for unification of
Rodinia. Precambrian Research, 111, 5 –30. the North and West Australian cratons by c. 1.7 Ga:
K ELLY , N. M., C LARKE , G. L. & F ANNING , C. M. 2002. new results from the Kimberley Basin of northwestern
A two-stage evolution of the Neoproterozoic Rayner Australia. Geophysical Journal International, 142,
Structural Episode: new U– Pb sensitive high resol- 173– 180.
ution ion microprobe constraints from the Oygarden L I , Z.-X. & P OWELL , C. MC A. 2001. An outline of
Group, Kemp Land, East Antarctica. Precambrian the palaeogeographic evolution of the Australasian
Research, 116, 307 –330. region since the beginning of the Neoproterozoic.
K EPPIE , J. D., N ANCE , R. D., M URPHY , J. B. & D OSTAL , Earth-Science Reviews, 53, 237–277.
J. 2003. Tethyan, Mediterranean, and Pacific analogues L I , Z.-X., Z HANG , L. & P OWELL , C. MC A. 1995. South
for the Neoproterozoic– Paleozoic birth and develop- China in Rodinia: part of the missing link between
ment of peri-Gondwanan terranes and their transfer Australia-East Antarctica and Laurentia? Geology,
to Laurentia and Laurussia. Tectonophysics, 365, 23, 407 –410.
195–219. L I , Z.-X., Z HANG , L. & P OWELL , C. MC A. 1996. Pos-
K HUDOLEY , A. K., K ROPACHEV , A. P., T KACHENKO , itions of the East Asian cratons in the Neoproterozoic
V. I., R UBLEV , A. G., S ERGEEV , S. A., M ATUKOV , supercontinent Rodinia. Australian Journal of Earth
D. I. & L YAHNITSKAYA , O. YU . 2007. Mesoprotero- Sciences, 43, 593–604.
zoic to Neoproterozoic evolution of the Siberian L I , Z.-X., L I , X.-H., Z HOU , H. & K INNY , P. D. 2002.
craton and adjacent microcontinents: an overview Grenvillian continental collision in south China: new
with constraints for a Laurentian connection. In: SHRIMP U –Pb zircon results and implications for
L INK , P. K. & L EWIS , R. S. (eds) Proterozoic the configuration of Rodinia. Geology, 30, 163–166.
Geology of Western North America and Siberia. L I , Z.-X., L I , X. H., K INNY , P. D., W ANG , J., Z HANG , S.
SEPM (Society for Sedimentary Geology) Special & Z HOU , H. 2003. Geochronology of Neoproterozoic
Publication, 86, 209–226. syn-rift magmatism in the Yangtze Craton, South
K INNY , P. D., B LACK , L. P. & S HERATON , J. W. 1993. China and correlations with other continents: evidence
Zircon ages and the distribution of Archaean and Pro- for a mantle superplume that broke up Rodinia.
terozoic rocks in the Rauer Islands. Antarctic Science, Precambrian Research, 122, 85– 109.
5, 193–206. L I , Z.-X., E VANS , D. A. D. & Z HANG , S. 2004. A 908 spin
K IRKLAND , C. L., D ALY , J. S. & W HITEHOUSE , M. J. on Rodinia: possible causal links between the Neopro-
2006. Granitic magmatism of Grenvillian and late terozoic supercontinent, superplume, true polar wander
Neoproterozoic age in Finnmark, Arctic Norway – and low-latitude glaciation. Earth and Planetary
Constraining pre-Scandian deformation in the Science Letters, 220, 409– 421.
Kalak Nappe Complex. Precambrian Research, 145, L I , Z.-X., B OGDANOVA , S. V. ET AL . 2008. Assembly,
24– 52. configuration, and break-up history of Rodinia: a syn-
K LERKX , J., L IÉGEOIS , J. P., L AVREAU , J. & C LAESSEN , thesis. Precambrian Research, 160, 179–210.
W. 1987. Crustal evolution of northern Kibaran belt L INK , P. K., C HRISTIE -B LICK , N. ET AL . 1993. Middle
in eastern and central Africa. In: K RÖNER , A. (ed.) and Late Proterozoic stratified rocks of the western
Proterozoic Lithospheric Evolution. American Geo- US Cordillera, Colorado Plateau, and Basin and
physical Union, Geodynamics Series, 17, 217–233. Range province. In: R EED , J. C., J R ., B ICKFORD , M.
K OKONYANGI , J. W., K AMPUNZU , A. B., A RMSTRONG , E., H OUSTON , R. S., L INK , P. K., R ANKIN , D. W.,
R., Y OSHIDA , M., O KUDAIRA , T., A RIMA , M. & S IMS , P. K. & V AN S CHMUS , W. R. (eds) Precam-
N GULUBE , D. A. 2006. The Mesoproterozoic Kibaride brian: Conterminous U.S. Boulder, Colorado, Geologi-
belt (Katanga, SE D.R. Congo). Journal of African cal Society of America, The Geology of North
Earth Sciences, 46, 1 –35. America, C-2, 463– 595.
400 D. A. D. EVANS
L OEWY , S. L., C ONNELLY , J. N., D ALZIEL , I. W. D. & Sandstone: implications for red bed diagenesis and
G OWER , C. F. 2003. Eastern Laurentia in Rodinia: Late Precambrian apparent polar wander of North
constraints from whole-rock Pb and U/Pb geochronol- America. Canadian Journal of Earth Sciences, 20,
ogy. Tectonophysics, 375, 169– 197. 105–112.
L U , S., L I , H., Z HANG , C. & N IU , G. 2008a. Geological M C E LHINNY , M. W., P OWELL , C. MC A. & P ISAREVSKY ,
and geochronological evidence for the Precambrian S. A. 2003. Paleozoic terranes of eastern Australia
evolution of the Tarim Craton and surrounding and the drift history of Gondwana. Tectonophysics,
continental fragments. Precambrian Research, 160, 362, 41– 65.
94–107. M C M ENAMIN , M. A. S. & M C M ENAMIN , D. L. S. 1990.
L U , S., Z HAO , G., W ANG , H. & H AO , G. 2008b. Precam- The Emergence of Animals: The Cambrian Break-
brian metamorphic basement and sedimentary cover of through. New York, Columbia University Press,
the North China Craton: a review. Precambrian 1– 217.
Research, 160, 77–93. M EERT , J. G. & S TUCKEY , W. 2002. Revisiting the paleo-
L UND , K., A LEINIKOFF , J. N., E VANS , K. V. & F ANNING , magnetism of the 1.476 Ga St. Francois Mountains
C. M. 2003. SHRIMP U –Pb geochronology of Neo- igneous province, Missouri, Tectonics, 21(2), doi:
proterozoic Windermere Supergroup, central Idaho: 10.1029/2000TC001265.
implications for rifting of western Laurentia and syn- M EERT , J. G. & T ORSVIK , T. H. 2003. The making and
chroneity of Sturtian glacial deposits. Geological unmaking of a supercontinent: rodinia revisited. Tecto-
Society of America Bulletin, 115, 349–372. nophysics, 375, 261 –288.
M AC L EAN , B. C. & C OOK , D. G. 2004. Revisions to the M EERT , J. G., H ARGRAVES , R. B., V AN DER V OO , R.,
Paleoproterozoic Sequence A, based on reflection H ALL , C. H. & H ALLIDAY , A. N. 1994. Paleo-
seismic data across the western plains of the Northwest magnetism and 40Ar/39Ar studies of Late Kibaran
Territories, Canada. Precambrian Research, 129, intrusives in Burundi, East Africa: implications for
271– 289. Late Proterozoic supercontinents. Journal of
M AIER , W. D., P ELTONEN , P. & L IVESEY , T. 2007. The Geology, 102, 621–637.
ages of the Kabanga North and Kapalagulu intrusions, M EERT , J. G., V AN DER V OO , R. & A YUB , S. 1995.
western Tanzania: a reconnaissance study. Economic Paleomagnetic investigation of the Late Proterozoic
Geology, 102, 147– 154. Gagwe lavas and Mbozi complex, Tanzania and the
M AJI , A. K., G OON , S., B HATTACHARYA , A., M ISHRA , assembly of Gondwana. Precambrian Research, 74,
B., M AHATO , S. & B ERNHARDT , H.-J. 2008. Protero- 225–244.
zoic polyphase metamorphism in the Chhotanagpur M ERTANEN , S., P ESONEN , L. J. & H UHMA , H. 1996.
Gneissic Complex (India), and implication for trans- Palaeomagnetism and Sm– Nd ages of the Neopro-
continental Gondwanaland correlation. Precambrian terozoic diabase dykes in Laanila and Kautokeino,
Research, 162, 385– 402. northern Fennoscandia. In: B REWER , T. S. (ed.) Pre-
M ALONE , S. J., M EERT , J. G. ET AL . 2008. Paleomagnet- cambrian Crustal Evolution in the North Atlantic
ism and detrital zircon geochronology of the upper Region. Geological Society, London, Special Publi-
Vindhyan sequence, Son Valley and Rajasthan, India: cations, 112, 331– 358.
a c. 1000 Ma closure age for the Purana basins? Pre- M EZGER , K. & C OSCA , M. A. 1999. The thermal history
cambrian Research, 164, 137–159. of the Eastern Ghats Belt (India) as revealed by
M ALOOF , A. C., H ALVERSON , G. P., K IRSCHVINK , J. L., U–Pb and 40Ar/39Ar dating of metamorphic and
S CHRAG , D. P., W EISS , B. P. & H OFFMAN , P. F. 2006. magmatic minerals: implications for the SWEAT
Combined paleomagnetic, isotopic, and stratigraphic correlation. Precambrian Research, 94, 251–271.
evidence for true polar wander from the Neoprotero- M ILLER , K. C. & H ARGRAVES , R. B. 1994. Paleomag-
zoic Akademikerbreen Group, Svalbard. Geological netism of some Indian kimberlites and lamproites.
Society of America Bulletin, 118, 1099– 1124. Precambrian Research, 69, 259–267.
M ASLOV , A. V., E RDTMANN , B.-D., I VANOV , K. S., M OORES , E. M. 1991. Southwest US–East Antarctic
I VANOV , S. N. & K RUPENIN , M. T. 1997. The main (SWEAT) connection: a hypothesis. Geology, 19,
tectonic events, depositional history and the palaeogeo- 425–428.
graphy of the southern Urals during the Riphean– early M URPHY , J. B. & N ANCE , R. D. 2003. Do supercontinents
Palaeozoic. Tectonophysics, 276, 313– 335. introvert or extrovert?: Sm–Nd isotope evidence.
M AYER , A., H OFMANN , A. W., S INIGOI , S. & M ORAIS , Geology, 31, 873 –876.
E. 2004. Mesoproterozoic Sm– Nd and U–Pb ages M URPHY , J. B. & N ANCE , R. D. 2005. Do supercontinents
for the Kunene Anorthosite Complex of SW Angola. turn inside-in or inside-out? International Geology
Precambrian Research, 133, 187–206. Review, 47, 591–619.
M BEDE , E. I., K AMPUNZU , A. B. & A RMSTRONG , R. A. M URPHY , J. B., S TRACHAN , R. A., N ANCE , R. D.,
2004. Neoproterozoic inheritance during Cainozoic P ARKER , K. D. & F OWLER , M. B. 2000. Proto-
rifting in the western and southwestern branches Avalonia: a 1.2– 1.0 Ga tectonothermal event and
of the East African rift system: evidence from carbona- constraints for the evolution of Rodinia. Geology, 28,
tite and alkaline intrusions. Conference abstract, The 1071– 1074.
East African Rift System: Development, Evolution M URTHY , G., G OWER , C., T UBRETT , M. & P ÄTZOLD , R.
and Resources, Addis Ababa, Ethiopia, June 20– 24, 1992. Paleomagnetism of Eocambrian Long Range
2004. dykes and Double Mer Formation from Labrador,
M C C ABE , C. & V AN DER V OO , R. 1983. Paleomagnetic Canada. Canadian Journal of Earth Sciences, 29,
results from the upper Keweenawan Chequamegon 1224– 1234.
M YERS , J. S., S HAW , R. D. & T YLER , I. M. 1996. Tec- P IMENTEL , M. M., F ERREIRA F ILHO , C. F. & A RMELE ,
tonic evolution of Proterozoic Australia. Tectonics, A. 2006. Neoproterozoic age of the Niquelândia
15, 1431– 1446. Complex, central Brazil: further ID-TIMS U– Pb and
O CCHIPINTI , S. A. & R EDDY , S. M. 2009. Neoproterozoic Sm–Nd isotopic evidence. Journal of South American
reworking of the Palaeoproterozoic Capricorn Orogen Earth Sciences, 21, 228–238.
of Western Australia and implications for the amalga- P IPER , J. D. A. 1974. Magnetic properties of the Cunene
mation of Rodinia. In: M URPHY , J. B. (ed.) Ancient anorthosite Complex, Angola. Physics of the Earth
Orogens and Modern Analogues. Geological Society, and Planetary Interiors, 9, 353–363.
London, Special Publications, 327, 445–456. P IPER , J. D. A. 1980. Analogous Upper Proterozoic appar-
P ALMER , H. C., M ERZ , B. A. & H AYATSU , A. 1977. The ent polar wander loops. Nature, 283, 845– 847.
Sudbury dikes of the Grenville Front region: paleo- P IPER , J. D. A. 2000. The Neoproterozoic supercontinent:
magnetism, petrochemistry, and K– Ar age studies. rodinia or Palaeopangaea? Earth and Planetary
Canadian Journal of Earth Sciences, 14, 1867–1887. Science Letters, 176, 131– 146.
P ARK , J. K. & J EFFERSON , C. W. 1991. Magnetic and P IPER , J. D. A. 2007. The Neoproterozoic supercontinent
tectonic history of the Late Proterozoic Upper Little Palaeopangaea. Gondwana Research, 12, 202– 227.
Dal and Coates Lake Groups of northwestern P ISAREVSKY , S. A. & B YLUND , G. 1998. Palaeomagnet-
Canada. Precambrian Research, 52, 1– 35. ism of a key section of the Protogine Zone, southern
P ARK , J. K., N ORRIS , D. K. & L AROCHELLE , A. 1989. Sweden. Geophysical Journal International, 133,
Paleomagnetism and the origin of the Mackenzie Arc 185– 200.
of northwestern Canada. Canadian Journal of Earth P ISAREVSKY , S. A. & B YLUND , G. 2006. Palaeomagnet-
Sciences, 26, 2194– 2203. ism of 935 Ma mafic dykes in southern Sweden and
P ATCHETT , P. J., B YLUND , G. & U PTON , B. G. J. 1978. implications for the Sveconorwegian Loop. Geophysi-
Palaeomagnetism and the Grenville orogeny: new cal Journal International, 166, 1095– 1104.
Rb–Sr ages from dolerites in Canada and Greenland. P ISAREVSKY , S. A. & N ATAPOV , L. M. 2003. Siberia and
Earth and Planetary Science Letters, 40, 349–364. Rodinia. Tectonophysics, 375, 221–245.
P AVLOV , V. E., G ALLET , Y., P ETROV , P. YU ., P ISAREVSKY , S. A., W INGATE , M. T. D. & H ARRIS , L. B.
Z HURAVLEV , D. Z. & S HATSILLO , A. V. 2002. Uy 2003a. Late Mesoproterozoic (c. 1.2 Ga) palaeomag-
series and late Riphean sills of the Uchur-Maya area: netism of the Albany– Fraser orogen: no pre-Rodinia
isotopic and palaeomagnetic data and the problem of Australia– Laurentia connection. Geophysical Journal
the Rodinia supercontinent. Geotectonics, 36, 278–292. International, 155, F6–F11.
P AYNE , J. L., H AND , M., B AROVICH , K. M., R EID , A. & P ISAREVSKY , S. A., W INGATE , M. T. D., P OWELL , C.
E VANS , D. A. D. 2009. Correlations and recon- MC A., J OHNSON , S. & E VANS , D. A. D. 2003b.
struction models for the 2500–1500 Ma evolution of Models of Rodinia assembly and fragmentation. In:
the Mawson Continent. In: R EDDY , S. M., Y OSHIDA , M., W INDLEY , B. F. & D ASGUPTA , S.
M AZUMDER , R., E VANS , D. A. D. & C OLLINS , (eds) Proterozoic East Gondwana: Supercontinent
A. S. (eds) Palaeoproterozoic Supercontinents and Assembly and Breakup. Geological Society, London,
Global Evolution. Geological Society, London, Special Publications, 206, 35– 55.
Special Publications, 323, 319– 355. P ISAREVSKY , S. A., W INGATE , M. T. D., S TEVENS ,
P ESONEN , L. J. & N EVANLINNA , H. 1981. Late Precam- M. K. & H AINES , P. W. 2007. Palaeomagnetic
brian Keweenawan asymmetric reversals. Nature, results from the Lancer 1 stratigraphic drillhole,
294, 436– 439. Officer Basin, Western Australia, and implications
P ESONEN , L. J., E LMING , S.-Å. ET AL . 2003. Palaeomag- for Rodinia reconstructions. Australian Journal of
netic configuration of continents during the Protero- Earth Sciences, 54, 561–572.
zoic. Tectonophysics, 375, 289– 324. P ISAREVSKY , S. A., N ATAPOV , L. M., D ONSKAYA , T. V.,
P ETTERSSON , Å., C ORNELL , D. H., M OEN , H. F. G., G LADKOCHUB , D. P. & V ERNIKOVSKY , V. A. 2008.
R EDDY , S. & E VANS , D. 2007. Ion-probe dating of Proterozoic Siberia: a promontory of Rodinia.
1.2 Ga collision and crustal architecture in the Precambrian Research, 160, 66– 76.
Namaqua-Natal Province of southern Africa. Precam- P IUZANA , D., P IMENTEL , M. M., F UCK , R. A. &
brian Research, 158, 79–92. A RMSTRONG , R. 2003. SHRIMP U –Pb and Sm– Nd
P HILLIPS , G., W ILSON , C. J. L., C AMPBELL , I. H. & data for the Araxa Group and associated mag-
A LLEN , C. M. 2006. U–Th– Pb detrital zircon geo- matic rocks: constraints for the age of sedimentation
chronology from the southern Prince Charles Moun- and geodynamic context of the southern Brasilia
tains, East Antarctica—Defining the Archaean to Belt, central Brazil. Precambrian Research, 125,
Neoproterozoic Ruker Province. Precambrian 139– 160.
Research, 148, 292 –306. P OWELL , C. MC A. & P ISAREVSKY , S. A. 2002. Late
P IMENTEL , M. M. & F UCK , R. A. 1992. Neoproterozoic Neoproterozoic assembly of East Gondwana.
crustal accretion in central Brazil. Geology, 20, Geology, 30, 3– 6.
375–379. P OWELL , C. MC A., L I , Z. X., M C E LHINNY , M. W.,
P IMENTEL , M. M., F ERREIRA F ILHO , C. F. & M EERT , J. G. & P ARK , J. K. 1993. Paleomagnetic
A RMSTRONG , R. A. 2004. SHRIMP U–Pb and Sm– constraints on timing of the Neoproterozoic breakup
Nd ages of the Niquelândia layered complex: meso- of Rodinia and the Cambrian formation of Gondwana,
(1.25 Ga) and Neoproterozoic (0.79 Ga) extensional Geology, 21, 889–892.
events in central Brazil. Precambrian Research, 132, P OWELL , C. MC A., J ONES , D. L., P ISAREVSKY , S. &
133–153. W INGATE , M. T. D. 2001. Palaeomagnetic constraints
402 D. A. D. EVANS
on the position of the Kalahari craton in Rodinia. R OY , J. L. & R OBERTSON , W. A. 1978. Paleomagnetism
Precambrian Research, 110, 33– 46. of the Jacobsville Formation and the apparent polar
P UFFER , J. H. 2002. A late Neoproterozoic eastern Lauren- wander path for the interval 1100 to 670 m.y. for
tian superplume: location, size, chemical composition, North America. Journal of Geophysical Research,
and environmental impact. American Journal of 83, 1289–1304.
Science, 302, 1– 27. R OYER , J.-Y., M ÜLLER , R. D., G AHAGAN , L. M.,
R ADHAKRISHNA , T. & M ATHEW , J. 1996. Late Pre- L AWVER , L. A., M AYES , C. L., N ÜRNBERG , D. &
cambrian (850–800 Ma) palaeomagnetic pole for the S CLATER , J. G. 1992. A global isochron chart.
south Indian shield from the Harohalli alkaline University of Texas Institute for Geophysics Technical
dykes: geotectonic implications for Gondwana recon- Report, no. 117.
structions. Precambrian Research, 80, 77–87. S ADOWSKI , G. R. & B ETTENCOURT , J. S. 1996. Mesopro-
R AINBIRD , R. H., J EFFERSON , C. W. & Y OUNG , G. M. terozoic tectonic correlations between eastern Lauren-
1996. The early Neoproterozoic sedimentary Succes- tia and the western border of the Amazon Craton.
sion B of northwestern Laurentia: correlations and Precambrian Research, 76, 213–227.
paleogeographic significance. Geological Society of S ALMINEN , J., P ESONEN , L. J., M ERTANEN , S., V UOLLO ,
America Bulletin, 108, 454– 470. J. & A IRO , M.-L. 2009. Palaeomagnetism of the Salla
R AINBIRD , R. H., S TERN , R. A., K HUDOLEY , A. K., Diabase Dyke, northeastern Finland, and its
K ROPACHEV , A. P., H EAMAN , L. M. & implication for the Baltica– Laurentia entity during
S UKHORUKOV , V. I. 1998. U– Pb geochronology the Mesoproterozoic. In: R EDDY , S. M., M AZUMDER ,
of Riphean sandstone and gabbro from southeast R., E VANS , D. A. D. & C OLLINS , A. S. (eds)
Siberia and its bearing on the Laurentia-Siberia con- Palaeoproterozoic Supercontinents and Global
nection. Earth and Planetary Science Letters, 164, Evolution. Geological Society, London, Special
409– 420. Publications, 323, 199–218.
R AMOS , V. A. 1988. Late Proterozoic-Early Paleozoic S ANTOS , J. O. S., H ARTMANN , L. A., G AUDETTE , H. E.,
of South America—a collisional history. Episodes, G ROVES , D. I., M C N AUGHTON , N. J. & F LETCHER ,
11, 168–174. I. R. 2000. A new understanding of the provinces
R APALINI , A. E. & S ÁNCHEZ -B ETTUCCI , L. 2008. Wide- of the Amazon craton based on integration of field
spread remagnetization of late Proterozoic sedimen- mapping and U– Pb and Sm–Nd geochronology.
tary units of Uruguay and the apparent polar wander Gondwana Research, 3, 453 –488.
path for the Rio de La Plata craton. Geophysical S ANTOS , J. O. S., H ARTMANN , L. A., B OSSI , J., C AMPAL ,
Journal International, 174, 55– 74. N., S CHIPILOV , A., P IÑEYRO , D. & M C N AUGHTON ,
R AUB , T. D., K IRSCHVINK , J. L. & E VANS , D. A. D. N. J. 2003. Duration of the Trans-Amazonian cycle
2007. True polar wander: linking deep and shallow and its correlation within South America based
geodynamics to hydro- and bio-spheric hypotheses. on U–Pb SHRIMP geochronology of the La Plata
In: K ONO , M. (ed.) Treatise on Geophysics, Volume craton, Uruguay. International Geology Review, 45,
5: Geomagnetism. Amsterdam, Elsevier, 565– 589. 27–48.
R ENNE , P. R., O NSTOTT , T. C., D’A GRELLA -F ILHO , S ANTOS , J. O. S., R IZZOTTO , G. J. ET AL . 2008. Age and
M. S., P ACCA , I. G. & T EIXEIRA , W. 1990. autochthonous evolution of the Sunsás Orogen in West
40
Ar/39Ar dating of 1.0– 1.1 Ga magnetizations from Amazon Craton based on mapping and U–Pb geochro-
the São Francisco and Kalahari cratons: tectonic nology. Precambrian Research, 165, 120– 152.
implications for Pan-African and Brasiliano mobile S ATO , K., S IGA , O., J R ., N UTMAN , A. P., B ASEI , M. A. S.,
belts. Earth and Planetary Science Letters, 101, M C R EATH , I. & K AULFUSS , G. 2003. The Atuba
349– 366. complex, southern South American platform: archean
R OEST , W. R. & S RIVASTAVA , S. P. 1989. Seafloor components and Paleoproterozoic to Neoproterozoic
spreading in the Labrador Sea: a new reconstruction. tectonothermal events. Gondwana Research, 6,
Geology, 17, 1000– 1004. 251–263.
R OGERS , J. J. W. 1996. A history of continents in the S CHMIDT , P. W., W ILLIAMS , G. E., C AMACHO , A. &
past three billion years. Journal of Geology, 104, L EE , J. K. W. 2006. Assembly of Proterozoic
91–107. Australia: implications of a revised pole for the
R OSS , G. M. 2002. Evolution of Precambrian continental c. 1070 Ma Alcurra Dyke Swarm, central Australia.
lithosphere in Western Canada: results from Litho- Geophysical Journal International, 167, 626 –634.
probe studies in Alberta and beyond. Canadian S CHOLL , D. W. & VON H UENE , R. 2007. Crustal recycling
Journal of Earth Sciences, 39, 413– 437. at modern subduction zones applied to the past issues
R OSS , G. M. & V ILLENEUVE , M. 2003. Provenance of of growth and preservation of continental basement
the Mesoproterozoic (1.45 Ga) Belt basin (western crust, mantle geochemistry, and supercontinent recon-
North America): another piece in the pre-Rodinia struction. In: H ATCHER , R. D., J R ., C ARLSON , M. P.,
paleogeographic puzzle. Geological Society of M C B RIDE , J. H. & M ARTINEZ -C ATALAN , J. R. (eds)
America Bulletin, 115, 1191–1217. 4-D Framework of Continental Crust. Geological
R OSS , G. M., P ARRISH , R. R. & W INSTON , D. 1992. Pro- Society of America Memoir, 200, 9– 32.
venance and U –Pb geochronology of the Mesoproter- S EARS , J. W. & P RICE , R. A. 2003. Tightening the Siber-
ozoic Belt Supergroup (northwestern United States): ian connection to western Laurentia. Geological
implications forage of deposition and pre-Panthalassa Society of America Bulletin, 115, 943 –953.
plate reconstructions. Earth and Planetary Science S ETH , B., A RMSTRONG , R. A., B RANDT , S., V ILLA , I. M.
Letters, 113, 57–76. & K RAMERS , J. D. 2003. Mesoproterozoic U–Pb and
Pb–Pb ages of granulites in NW Namibia: reconstruct- T HORKELSON , D. J., A BBOTT , J. A., M ORTENSEN , J. K.,
ing a complete orogenic cycle. Precambrian Research, C REASER , R. A., V ILLENEUVE , M. E., M C N ICOLL ,
126, 147– 168. V. J. & L AYER , P. W. 2005. Early and Middle Proter-
S INGLETARY , S. J., H ANSON , R. E. ET AL . 2003. ozoic evolution of Yukon, Canada. Canadian Journal
Geochronology of basement rocks in the Kalahari of Earth Sciences, 42, 1045–1071.
Desert, Botswana, and implications for regional T OHVER , E., VAN DER P LUIJM , B. A., V AN DER V OO , R.,
Proterozoic tectonics. Precambrian Research, 121, R IZZOTTO , G. & S CANDOLARA , J. E. 2002. Paleo-
47– 71. geography of the Amazon craton at 1.2 Ga: early
S KLYAROV , E. V., G LADKOCHUB , D. P., M AZUKABZOV , Grenvillian collision with the Llano segment of
A. M., M ENSHAGIN , YU . V., W ATANABE , T. & Laurentia. Earth and Planetary Science Letters, 199,
P ISAREVSKY , S. A. 2003. Neoproterozoic mafic dike 185– 200.
swarms of the Sharyzhalgai metamorphic massif, T OHVER , E., B ETTENCOURT , J. S., T OSDAL , R.,
southern Siberian craton. Precambrian Research, M EZGER , K., L EITE , W. B. & P AYOLLA , B. L.
122, 359– 376. 2004. Terrane transfer during the Grenville orogeny:
S METHURST , M. A., K HRAMOV , A. N. & T ORSVIK , T. H. tracing the Amazonian ancestry of southern Appala-
1998. The Neoproterozoic and Palaeozoic palaeomag- chian basement through Pb and Nd isotopes. Earth
netic data for the Siberian platform: from Rodinia to and Planetary Science Letters, 228, 161– 176.
Pangea. Earth-Science Reviews, 43, 1 –24. T OHVER , E., D ’A GRELLA -F ILHO , M. S. & T RINDADE ,
S RIVASTAVA , S. P. & R OEST , W. R. 1989. Seafloor R. I. F. 2006. Paleomagnetic record of Africa and
spreading history II–IV. In: Map sheets L17-2– South America for the 1200–500 Ma interval, and
L17-6 (Atlantic Geoscience Centre, Geological evaluation of Rodinia and Gondwana assemblies.
Survey of Canada, Ottawa). Precambrian Research, 147, 193– 222.
S TAMPFLI , G. M. & B OREL , G. D. 2002. A plate tectonic T ORSVIK , T. H., C ARTER , L. M., A SHWAL , L. D.,
model for the Paleozoic and Mesozoic constrained B HUSHAN , S. K., P ANDID , M. K. & J AMTVEIT , B.
by dynamic plate boundaries and restored synthetic 2001. Rodinia refined or obscured: palaeomagnetism
oceanic isochrons. Earth and Planetary Science of the Malani igneous suite (NW India). Precambrian
Letters, 196, 17–33. Research, 108, 319– 333.
S TEARN , J. E. F. & P IPER , J. D. A. 1984. Palaeomagnetism T ORSVIK , T. H., M ÜLLER , R. D., V AN DER V OO , R.,
of the Sveconorwegian mobile belt of the Fennoscan- S TEINBERGER , B. & G AINA , C. 2008. Global
dian Shield. Precambrian Research, 23, 201 –246. plate motion frames: toward a unified model.
S U , W., Z HANG , S. ET AL . 2008. SHRIMP U–Pb ages of Reviews of Geophysics, 46, RG3004, doi: 10.1029/
K-bentonite beds in the Xiamaling Formation: impli- 2007RG000227.
cations for revised subdivision of the Meso- to Neopro- T OSDAL , R. M. 1996. The Amazon-Laurentian connection
terozoic history of the North China Craton. Gondwana as viewed from the Middle Proterozoic rocks in the
Research, 14, 543– 553. central Andes, western Bolivia and northern Chile.
S VENNINGSEN , O. M. 2001. Onset of seafloor spreading in Tectonics, 15, 827–842.
the Iapetus Ocean at 608 Ma: precise age of the Sarek T RINDADE , R. I. F., F ONT , E., D’A GRELLA -F ILHO ,
Dyke Swarm, northern Swedish Caledonides. Precam- M. S., N OGUEIRA , A. C. R. & R ICCOMINI , C. 2003.
brian Research, 110, 241–254. Low-latitude and multiple geomagnetic reversals
T ACK , L., L IÉGEOIS , J. P., D EBLOND , A. & D UCHESNE , in the Neoproterozoic Puga cap carbonate, Amazon
J. C. 1994. Kibaran A-type granitoids and mafic rocks craton. Terra Nova, 15, 441– 446.
generated by two mantle sources in a late orogenic T ROMPETTE , R. 1997. Neoproterozoic (c. 600 Ma) aggre-
setting (Burundi). Precambrian Research, 68, gation of Western Gondwana: a tentative scenario.
323–356. Precambrian Research, 82, 101– 112.
T ANAKA , H. & I DNURM , M. 1994. Palaeomagnetism of U PTON , B. G. J., R ÄMÖ , O. T., H EAMAN , L. M.,
Proterozoic mafic intrusions and host rocks of the B LICHERT -T OFT , J., K ALSBEEK , F., B ARRY , T. L.
Mount Isa Inlier, Australia: revisited. Precambrian & J EPSEN , H. F. 2005. The Mesoproterozoic Zig-Zag
Research, 69, 241– 258. Dal basalts and associated intrusions of eastern
T ASSINARI , C. C. G., B ETTENCOURT , J. S., G ERALDES , North Greenland: mantle plume-lithosphere inter-
M. C., M ACAMBIRA , M. J. B. & L AFON , J. M. 2000. action. Contributions to Mineralogy and Petrology,
The Amazonian Craton. In: C ORDANI , U. G., M ILANI , 149, 40–56.
E. J., T HOMAZ F ILHO , A. & C AMPOS , D. A. (eds) V ALERIANO , C. M., M ACHADO , N., S IMONETTI , A.,
Tectonic evolution of South America. 31st International V ALLADARES , C. S., S EER , H. J. & S IMÕES ,
Geological Congress, Rio de Janeiro, 41– 95. L. S. A. 2004. U– Pb geochronology of the southern
T HIÉBLEMONT , D., G OUJOU , J. C., E GAL , E., C OCHERIE , Brasilia belt (SE-Brazil): sedimentary provenance,
A., D ELOR , C., L AFON , J. M. & F ANNING , C. M. Neoproterozoic orogeny and assembly of West Gond-
2004. Archean evolution of the Leo Rise and its wana. Precambrian Research, 130, 27– 55.
Eburnean reworking. Journal of African Earth V AN DER V OO , R. 1990. The reliability of paleomagnetic
Sciences, 39, 97–104. data. Tectonophysics, 184, 1– 9.
T HORKELSON , D. J., M ORTENSEN , J. K., C REASER , V AUCHEZ , A., N EVES , S., C ABY , R., C ORSINI , M.,
R. A., D AVIDSON , G. J. & A BBOTT , J. G. 2001. E GYDIO -S ILVA , M., A RTHAUD , M. & A MARO , V.
Early Proterozoic magmatism in Yukon, Canada: con- 1995. The Borborema shear zone system, NE Brazil.
straints on the evolution of northwestern Laurentia. Journal of South American Earth Sciences, 8,
Canadian Journal of Earth Sciences, 38, 1479–1494. 247– 266.
404 D. A. D. EVANS
V EEVERS , J. J. & S AEED , A. 2008. Gamburtsev Subgla- Uinta Mountain supergroup, Central Rocky Mountain
cial Mountains provenance of Permian-Triassic sand- States, USA. Precambrian Research, 147, 234– 259.
stones in the Prince Charles Mountains and offshore W HITE , R. W., C LARKE , G. L. & N ELSON , D. R. 1999.
Prydz Bay: Integrated U– Pb and TDM ages and SHRIMP U– Pb zircon dating of Grenville-age events
host-rock affinity from detrital zircons. Gondwana in the western part of the Musgrave Block, central
Research, 14, 316–342. Australia. Journal of Metamorphic Geology, 17,
V OLOZH , YU . A., A NTIPOV , M. P., B RUNET , M.-F., 465–481.
G ARAGASH , I. A., L OBKOVSKII , L. I. & C ADET , W INGATE , M. T. D. & E VANS , D. A. D. 2003. Palaeomag-
J.-P. 2003. Pre-Mesozoic geodynamics of the Precas- netic constraints on the Proterozoictectonic evolution
pian Basin (Kazakhstan). Sedimentary Geology, 156, of Australia. In: Y OSHIDA , M., W INDLEY , B. &
35–58. D ASGUPTA , S. (eds) Proterozoic East Gondwana:
V RANA , S., K ACHLIK , V., K RÖNER , A., M ARHEINE , D., Super Continent Assembly and Break-up. Geological
S EIFERT , A. V., Z ACEK , V. & B ABUREK , J. 2004. Society, London, Special Publications, 206, 77–91.
Ubendian basement and its late Mesoproterozoic W INGATE , M. T. D. & G IDDINGS , J. W. 2000. Age and
and early Neoproterozoic structural and metamorphic palaeomagnetism of the Mundine Well dyke swarm,
overprint in northeastern Zambia. Journal of African Western Australia: implications for an Australia-
Earth Sciences, 38, 1 –21. Laurentia connection at 755 Ma. Precambrian
W ALDERHAUG , H. J., T ORSVIK , T. H., E IDE , E. A., Research, 100, 335–357.
S UNDVOLL , B. & B INGEN , B. 1999. Geochronology W INGATE , M. T. D., P ISAREVSKY , S. A. & E VANS ,
and palaeomagnetism of the Hunnedalen dykes, SW D. A. D. 2002. Rodinia connections between Australia
Norway: implications for the Sveconorwegian appar- and Laurentia: no SWEAT, no AUSWUS? Terra
ent polar wander loop. Earth and Planetary Science Nova, 14, 121– 128.
Letters, 169, 71–83. W OODEN , J. L. & M ILLER , D. M. 1990. Chronologic and
W ALDERHAUG , H. J., T ORSVIK , T. H. & H ALVERSON , E. isotopic framework for Early Proterozoic crustal evol-
2007. The Egersund dykes (SW Norway): a robust ution in the eastern Mojave Desert region. Journal of
Early Ediacaran (Vendian) palaeomagnetic pole Geophysical Research, 95, 20133– 20146.
from Baltica. Geophysical Journal International, W U , H., Z HANG , S., L I , Z.-X., L I , H. & D ONG , J. 2005.
168, 935– 948. New paleomagnetic results from the Yangzhuang
W ANG , Y., L IU , D., C HUNG , S.-L., T ONG , L. & R EN , L. Formation of the Jixian System, North China, and
2008. SHIMP zircon ageconstraints from the Larse- tectonic implications. Chinese Science Bulletin, 50,
mann Hills region, Prydz Bay, for a late Mesoprotero- 1483– 1489.
zoic to early Neoproterozoic tectono-thermal event in X U , B., J IAN , P., Z HENG , H., Z OU , H., Z HANG , L. & L IU ,
East Antarctica. American Journal of Science, 308, D. 2005. U–Pb zircon geochronology and geochemis-
573– 617. try of Neoproterozoic volcanic rocks in the Tarim
W ARNOCK , A. C., K ODAMA , K. P. & Z EITLER , P. K. Block of northwest China: implications for the
2000. Using thermochronometry and low-temperature breakup of Rodinia supercontinent and Neoproterozoic
demagnetization to accurately date Precambrian paleo- glaciations. Precambrian Research, 136, 107– 123.
magnetic poles. Journal of Geophysical Research, 105, Y ANG , Z. & B ESSE , J. 2001. New Mesozoic apparent
19435– 19453. polar wander path for south China: tectonic con-
W ATT , G. R. & T HRANE , K. 2001. Early Neoproterozoic sequences. Journal of Geophysical Research, 106,
events in East Greenland. Precambrian Research, 110, 8493– 8520.
165– 184. Z HAI , M., S HAO , J., H AO , J. & P ENG , P. 2003. Geological
W EGENER , A. 1929. The Origin of Continents and signature and possible position of the North China
Oceans. [English translation by J. BIRAM , 1966. Block in the supercontinent Rodinia. Gondwana
New York, Dover]. Research, 6, 171– 183.
W EIL , A. B., V AN DER V OO , R., M AC N IOCAILL , C. & Z HANG , S. 2004. South China’s Gondwana connection in
M EERT , J. G. 1998. The Proterozoic supercontinent the Paleozoic: paleomagnetic evidence. Progress in
Rodinia: paleomagnetically derived reconstructions Natural Science, 14, 85– 90.
for 1100 to 800 Ma. Earth and Planetary Science Z HANG , S., L I , Z.-X. & W U , H. 2006. New Precambrian
Letters, 154, 13–24. palaeomagnetic constraints on the position of the
W EIL , A. B., G EISSMAN , J. W., H EIZLER , M. & V AN DER North China Block in Rodinia. Precambrian Research,
V OO , R. 2003. Paleomagnetism of Middle Proterozoic 144, 213 –238.
mafic intrusions and Upper Proterozoic (Nankoweap) Z HAO , G., S UN , M. & W ILDE , S. A. 2003. Correlations
red beds from the Lower Grand Canyon Supergroup, between the Eastern Block of the North China Craton
Arizona. Tectonophysics, 375, 199 –220. and the South Indian Block of the Indian Shield: an
W EIL , A. B., G EISSMAN , J. W. & V AN DER V OO , R. 2004. Achaean to Palaeoproterozoic link. Precambrian
Paleomagnetism of the Neoproterozoic Chuar Group, Research, 122, 201–233.
Grand Canyon Supergroup, Arizona: implications for Z HAO , X. X., P ARK , Y. H. & W ANG , C. S. 2008. Tarim
Laurentia’s Neoproterozoic APWP and Rodinia affinities of the Western Australia and Gondwana:
break-up. Precambrian Research, 129, 71– 92. new constraints from Late Cambrian and Middle Ordo-
W EIL , A. B., G EISSMAN , J. W. & A SHBY , J. M. 2006. vician paleomagnetic poles. The Gondwana 13
A new paleomagnetic pole for the Neoproterozoic Program and Abstracts, 267.
The Grenville Province as a large hot long-duration collisional
orogen – insights from the spatial and thermal evolution of its
orogenic fronts
TOBY RIVERS
Department of Earth Sciences, Memorial University of Newfoundland, PO Box 4200,
St. John’s NL, A1B 3X5, Canada (e-mail: trivers@mun.ca)
Abstract: The proposition that the Grenville Province is a remnant of a large hot long-duration
collisional orogen is examined through a comparative study of its present orogenic front, the Gren-
ville Front, and a former front, the Allochthon Boundary Thrust. Structural, metamorphic and geo-
chronologic data for both boundaries and their hanging walls from the length of the Grenville
Province are compared. Cumulative displacement across the Grenville Front was minor (10 s of
km) whereas that across the Allochthon Boundary Thrust was major (100 s of km), consistent
with the observation that the latter boundary separates rocks with a different age, and P –T
character, of metamorphism.
On an orogen scale, Grenvillian metamorphism can be subdivided into two spatially and tem-
porally distinct orogenic phases, a relatively high T Ottawan (c. 1090– 1020 Ma) phase in the
hanging wall of the Allochthon Boundary Thrust, and a relatively lower T Rigolet (c. 1000–
980 Ma) phase in the hanging wall of the Grenville Front. It is argued that the structural setting
and 50 My duration of Ottawan metamorphism are compatible with some form of channel
flow beneath an orogenic plateau, with the Allochthon Boundary Thrust forming the base of the
channel. Channel flow ceased at c. 1020 Ma when the Allochthon Boundary Thrust was reworked
as part of a system of normal-sense shear zones, and following a hiatus of c. 20 My the short-lived
Rigolet metamorphism took place in the former foreland and involved the development of a new
orogenic front, the Grenville Front. Taken together, this suggests the Grenville Orogen developed
as a large hot long-duration orogen during the Ottawan orogenic phase, but following gravitational
collapse of the plateau the locus of thickening migrated into the foreland and active tectonism was
restricted to a subjacent small cold short-duration orogen. The foreland-ward migration of the oro-
genic front from the Allochthon Boundary Thrust to the Grenville Front, the contrasting P –T –t
character of the metamorphic rocks in their hanging walls, and the evidence for orogenic collapse
followed by renewed growth, provide insights into the complex evolution of a long-duration
collisional orogen.
Introduction parameters, for example, crustal rheology, radio-

active heat-producing potential, convergence vel-
From a mechanical perspective, the evolution of a ocity, erosion rate, mantle heat flow, etc., orogen
collisional orogen can be conceptually divided width, duration of orogenesis and the maximum
into constructive (or growth) and destructive (or temperature attained are all positively correlated.
decay) phases separated by a climax representing On this basis, Beaumont et al. (2006) suggested
the time of maximum volume. The constructive that collisional orogens form a continuum between
phase is characterized by increasing crustal thick- two end-members: (i) small cold short-duration
ness and potential energy (Hodges 1998), whereas orogens (SCOs), for which the essential structural
during the destructive phase the stored potential characteristics can be simulated with mechanical
energy gradually decays as the orogenic crust thins models; and (ii) large hot long-duration orogens
by gravitational collapse and erosion (Dewey (LHOs), for which coupled thermal-mechanical
1988; Rey et al. 2001). models are necessary to produce realistic simu-
Modern collisional orogens vary greatly in width lations. Moreover, as is implicit in the experiments
and duration, for example from ,150 km and and is developed here with respect to the Grenville
,30 My for the Pyrenees, to .1000 km and Province, LHOs also become increasingly
.50 My for the Himalaya –Tibet Orogen. This vari- complex over time. This is a result of the protracted
ation has also been reproduced in thermal- constructive phase, which leads to widening of the
mechanical experiments (e.g. Beaumont et al. orogen, and because the accompanying regional
2001a, b, 2004, 2006; Jamieson et al. 2002, 2004, metamorphism is of high temperature due to the
2007), which show that for similar model long time available for conductive and radioactive
406 T. RIVERS
heating, thereby promoting rheological weakening Orogenic fronts of the Grenville Province
and the potential for mid-crustal flow.
Orogenic fronts define the contemporary limits The Grenville Orogen, part of which is exposed in
of orogens, their contractional character providing the Grenville Province in the SE Canadian Shield,
linkage to the hinterland where the greatest thicken- is a late Mesoproterozoic–early Neoproterozoic
ing and heating occur. This paper focuses on the feature widely interpreted to have developed
structural/metamorphic character and timing of during assembly of Rodinia as a result of collision
two orogenic fronts in the Grenville Province and between Laurentia and Amazonia (Li et al. 2008).
the insight they provide on the evolution of large It is a prime candidate for a large hot long-duration
hot long-duration collisional orogens. orogen, because it is .600 km wide (minimum
estimate because the SE margin was removed by
Neoproterozoic rifting), largely underlain by Gren-
Factors affecting orogenic fronts in villian granulite- and upper-amphibolite-facies
rocks at the present erosion surface implying it was
large hot long-duration orogens hot, and because the Grenvillian Orogeny lasted
As inferred for the Himalaya –Tibet Orogen (e.g. for 100 My (from c. 1090–980 Ma; Rivers 1997;
Clark & Royden 2000; Grujic et al. 2002; Godin Carr et al. 2000; Rivers 2008). The orogenic fronts
et al. 2006a; Searle et al. 2006) and modelled in of the Grenville Province discussed in this paper
numerical experiments, LHOs develop in orthog- are an older, more internal boundary known as the
onal collisional settings and are characterized by a Allochthon Boundary Thrust and a younger external
plateau in the orogenic hinterland that forms when boundary known as the Grenville Front. Their P–T
the crust is thickened beyond its load-bearing evolution and crustal-scale architecture are used to
capacity. The resultant forced flow of melt- develop a tectonic model for the Grenville Orogen
weakened crust reduces local topography on the that is compared to a thermal-mechanical model
plateau and under appropriate erosional conditions for the western Grenville Province developed by
leads to exhumation of a mid-crustal ‘channel’ at Jamieson et al. (2007).
the orogenic front. On the basis of their numerical
models, Beaumont et al. (2006) distinguished
three crustal flow modes: (i) homogeneous channel General features of the Grenville Province
flow of low-viscosity mid crust; (ii) heterogeneous The Grenville Province is composed of rocks of
channel flow of low-viscosity mid-crust and weak Archaean to Mesoproterozoic age that generally
segments of lower crust, both modes driven by grav- young towards the SE, reflecting the accretionary
itational forcing; and (iii) extrusion of higher vis- growth of southeastern Laurentia during the
cosity hot fold nappes by tectonic forcing. Palaeo- and Mesoproterozoic (Hoffman 1989;
It follows from the modelling of Beaumont et al. Rivers 1997; Karlstrom et al. 2001). Thus Archaean
(2006) that in a stable LHO, defined as one in which rocks derived from the Superior Province extend
the volume of crust remains constant, the amount of into the northern Grenville Province and underlie
new crust added by thrusting is balanced by that much of it at depth and some Palaeoproterozoic
removed by erosion. Since plateaux are regions of units can be traced across the Grenville Front,
low erosion rate, most of the mass is removed at oro- whereas the youngest units, c. 1.3–1.2 Ga accreted
genic fronts where rates are higher (Grujic et al. terranes, occur in the upper and mid-crust in the
2006). Thus, emergence of the mid-crustal channel SE (Figs 1 & 2; e.g. Gower et al. 1980; Rivers &
at the orogenic front is the principal method by Chown 1986; Rivers et al. 1989). The collisional
which stable LHOs reduce their mass. Destabiliza- suture with Amazonia does not outcrop in the
tion of the mass balance would occur if, for instance, exposed Grenville Province, but on the basis of Pb
the volume of crust added to the orogen exceeded isotopic data it may lie close to the Grenville
the volume flowing through the channel, or if the inliers in the SE Appalachians (Loewy et al. 2003;
viscosity of the flowing mid-crust in the channel Fig. 1 inset). The Interior Magmatic Belt (Fig. 1;
was changed beyond critical values by heating or Rivers 1997), characterized by syn- to post-
cooling. Such scenarios could lead to overthicken- Grenvillian plutons, defines the thermal core of
ing and orogenic collapse, and/or widening and the orogen.
the formation of a new orogenic front. In this contri-
bution, it is argued that a record of these processes is
preserved in the Grenville Province where there is Crustal and metamorphic architecture and
evidence for collapse following thickening in the timing of metamorphism
orogenic interior, and for outward growth of the
orogen into its former foreland and the formation The Grenville Orogen is a crustal-scale thrust stack
of a new orogenic front. (modified by later extension – see below) composed
THE GRENVILLE PROVINCE AS A LARGE HOT LONG-DURATION COLLISIONAL 407
Fig. 1. Sketch map of the Grenville Province showing age of pre-Grenvillian rocks (modified from Rivers et al. in
review). Note continuity of Archaean and Palaeoproterozoic units from the foreland into the Grenville Orogen and
presence of Mesoproterozoic accreted terranes in the southwestern Grenville Province. Inset: sketch map of North
America showing that the Grenville Front truncates older orogenic provinces (after Hoffman 1989) and possible
location of the orogenic suture. ABT, Allochthon Boundary Thrust; GF, Grenville Front; IMB, Interior Magmatic Belt.
of structurally-bound terranes and domains that are roof (Fig. 2). As elaborated below, the Grenville
grouped into belts on a regional scale (the term Front is a contractional structure that carries the
terrane is used in the sense of a metamorphic Parautochthonous Belt over the adjacent foreland,
terrane with a distinctive Grenvillian metamorphic but continuity of several units across it indicates it
history; it does not necessarily connote an exotic is not the site of major tectonic transport (Rivers
origin with respect to Laurentia). The lowest belt et al. 1989). The Allochthon Boundary Thrust is a
in the stack, the Parautochthonous Belt (Rivers major ductile shear zone that, as its name implies,
et al. 1989), is bounded by the Grenville Front at was initiated as a contractional structure, but was
its base and the Allochthon Boundary Thrust at its reworked in extension locally (Culshaw et al.
408 T. RIVERS
Fig. 2. Sketch map of the Grenville Province showing the distribution of late Mesoproterozoic metamorphism (modified
from Rivers et al. submitted). Grenvillian metamorphism is subdivided into Ottawan (c. 1190– 1120 Ma) and Rigolet
(c. 1005– 980 Ma) orogenic phases. Metamorphism associated with accretion of the Composite Arc Belt (CAB; c. 1240–
1225 Ma) and Frontenac– Adirondack Belt (F –AB; c. 1190–1140 Ma) was pre-Grenvillian. Blue lines labelled 4A–B
are locations of Lithoprobe deep-seismic reflection experiments shown in Figure 4; boxes labelled 5A–H and 7A–D
show locations of detailed studies of the Grenville Front and Allochthon Boundary Thrust in Figures 5 and 7.
1997; Ketchum et al. 1998). As shown below, the 1997, 2008; Rivers et al. 2002). Apart from local
Allochthon Boundary Thrust marks a major break accretionary metamorphic events that pre-dated col-
in the distribution of lithologic units, structural lisional orogenesis and are not considered in detail
fabrics, and in the age and character of metamorph- here, the main regional metamorphism and associ-
ism. In order to emphasise these contrasts, Rivers ated crustal thickening in the allochthonous terranes
et al. (1989) referred to the terranes in its hanging in the interior Grenville Province took place from c.
wall as allochthonous, indicating that they are far- 1090–1020 Ma, referred to as the Ottawan event,
travelled (but not exotic to Laurentia). whereas in the Para-autochthonous Belt at the north-
The Grenville Province is largely underlain by western margin the main metamorphism took place
high-grade rocks formed during the Grenvillian from c. 1005 –980 Ma (Rigolet event; Fig. 2). Both
Orogeny, implying the present erosion surface pro- these metamorphic events are of regional extent but
vides a view of the mid- and lower orogenic crust. In for the most part their effects are geographically
order to analyse its crustal-scale structure, Grenvil- separated, and moreover there was a hiatus
lian metamorphism is subdivided on the basis of age between them following an important period of
and baric character (Fig. 2). extension at c. 1020 Ma at the end of the Ottawan
phase. The tectonic significance of these two
Age of Grenvillian metamorphism. Grenvillian regional metamorphisms and associated contrac-
metamorphism took place from c. 1090– 980 Ma, tional events has spawned a confusing terminology.
but there are important temporal variations that cor- It is possible they represent independent collisions,
respond to the spatial division into the para- in which case each would merit the status of an
autochthonous and allochthonous belts (Rivers orogeny, as proposed by McLelland et al. (2001).
Alternatively they may be manifestations of a single corresponding to anorthosite complexes, much of

long-duration collisional orogeny, an interpretation the interior Grenville Province is in approximate
supported by comparison with models of LHOs isostatic equilibrium, confirming seismic studies
and preferred by Rivers (1997, 2008), Gower & that it is underlain by crust of average thickness
Krogh (2002), Rivers et al. (2002), and further devel- (c. 35 km). In contrast, the Bouguer signature of
oped herein. Thus in this paper the Ottawan and the Parautochthonous Belt varies from positive in
Rigolet are referred to as orogenic phases of the the west to negative under the c. 1000 km-long
long-duration Grenvillian Orogeny. In any case, Grenville Front Gravity Low (GFGL). Assuming
the distribution of metamorphic ages in Figure 2 the underlying crust is of average density, this
implies that in Ottawan times the contemporary oro- implies there are important variations in crustal
genic front lay close to the Allochthon Boundary thickness under the northern Grenville Province,
Thrust and that in Rigolet times the orogen advanced from c. 30 km in the west to 50 km under the
into its foreland and the Grenville Front was GFGL (Hynes 1994), estimates independently con-
established. firmed by seismic reflection studies (Green et al.
1988; Martignole & Calvert 1996; Hynes et al.
Baric subdivisions. In light of their different ages, 2000). The significance of this is discussed later.
metamorphism in the Parautochthonous Belt and
allochthonous belts is considered separately, symbo-
lized by the prefixes p and a respectively. With The northwestern Grenville Province –
respect to the Parautochthonous Belt, the Rigolet Rigolet deformation and metamorphism
metamorphism is principally medium pressure
(pMP) or Barrovian in character, but there are two In this section, the deep seismic, structural, meta-
known areas of high pressure metamorphism (pHP, morphic and age characteristics of the Grenville
.1200 MPa) characterized by eclogite- and HP Front and Parautochthonous Belt are assembled
granulite-facies assemblages. With respect to the from published studies. The results of two Litho-
allochthonous belts, Ottawan metamorphism is sub- probe deep seismic experiments in Ontario and
divided into a High-Pressure (aHP) Belt, also charac- western Québec are reviewed first, followed by a
terized by eclogite, relict eclogite and HP granulite discussion of the surface geology from case
assemblages, a Medium- to Low-Pressure (aM-LP) studies along the length of the Grenville Province.
Belt underlain by granulite- to amphibolite-facies
and locally greenschist-facies assemblages, and an Seismic images of the Grenville Front
Orogenic Lid that largely escaped metamorphic
reworking (White et al. 2000; Rivers et al. 2002; The Lithoprobe seismic study over Lake Huron,
Rivers & Indares 2006; Rivers 2008). More detailed Ontario (Green et al. 1988; White et al. 2000) sup-
mapping is necessary to subdivide the M-LP belt into ported field evidence indicating that the Grenville
discrete MP and LP segments. Front in this region marks the northern limit of a
30 km-wide, SE-dipping, crustal-scale shear zone,
Geophysical character the Grenville Front Tectonic Zone (GFTZ;
Wynne-Edwards 1972). The seismic section
The levelled magnetic and Bouguer gravity (Fig. 4a) shows the front truncates a prominent sub-
anomaly maps of the Grenville Province are horizontal mid-crustal reflector in the foreland and
shown in Figure 3. In terms of magnetic expression reaches c. 30 km depth at the SE end of the line.
(Fig. 3a), the Parautochthonous Belt has a distinc- The straightened granitoid and mafic gneisses
tive subdued signature in the central Grenville exposed in the GFTZ exhibit upper amphibolite-
Province, contrasting with the E-trending pattern in facies assemblages, reverse-sense kinematic indi-
the adjacent Superior Province despite the continuity cators and down-dip elongation lineations, implying
of Archaean crust across the Grenville Front. exhumation from the SE.
Although not studied in detail, the contrast may be In the seismic experiment near Témiskamingue,
related to the breakdown of magnetite during the western Québec (Fig. 4b; Kellett et al. 1994), not
Rigolet metamorphic overprint (Toft et al. 1993). only was reflectivity much weaker but a very differ-
In contrast to the Parautochthonous Belt, the over- ent image of the crustal structure was obtained. The
lying allochthonous belts exhibit a short-wavelength, Grenville Front in this area is located at a steep fault,
variable-intensity magnetic signature that defines one of several reworked Archaean structures with
the complex regional structure, the high intensity sinistral splays, and exhibits a ramp-flat profile
components of the signal suggesting magnetite was with prominent flats at c. 10 and 20 km depth.
stable during the Ottawan metamorphism. Reflectors were poorly imaged near the surface,
With respect to the Bouguer gravity map (Fig. probably due to the steep dips, and moreover
3b), apart from several large negative anomalies Kellett et al. (1994) noted that Grenvillian fabrics
410 T. RIVERS
Fig. 3. Levelled potential field maps of the Grenville Province (after Rivers et al. 1989); grey scales indicate relative
intensities. (a) aeromagnetic intensity; red lines in Archaean Superior Province highlight E– W structural trends;
(b) Bouguer gravity anomaly. ABT, Allochthon Boundary Thrust; GF, Grenville Front; GFGL, Grenville Front
Gravity Low.
at the surface were limited, except in the vicinity illustrate the structural architecture, metamorphic
of faults. character and thermal history of the Grenville Front
The two seismic sections are located c. 450 km and northern Parautochthonous Belt. Emphasis is
apart (Fig. 2), implying the tectonic character of placed on units that can be traced across the Grenville
the Grenville Front and northern Parautochthonous Front and only experienced Grenvillian orogenesis,
Belt is heterogeneous at this scale. For instance, and for one area two maps are presented to illustrate
there is no structure comparable to the GFTZ at the range of features in different units. Study areas
the surface in the western Québec section, the archi- are discussed from west to east, locations are
tecture of the Grenville Front changes from shown in Figure 2, and details of the P–T estimates
moderately-dipping to a ramp-flat geometry, and and geochronologic data are given in Table 1.
reaches c. 30 km depth at the SE end of the Ontario
section, but only c. 20 km in western Québec.
Grenville Front Tectonic Zone near Killarney,
Ontario. Part of the Mesoproterozoic (c. 1238 Ma)
Grenville Front and Parautochthonous Sudbury dyke swarm crosses the Grenville Front
Belt – Case studies from the foreland into the GFTZ near Killarney,
Ontario, making it a suitable marker for Grenvillian
Maps and cross-sections at various scales are used deformation and metamorphism (Fig. 5a; Bethune
with relevant P –T and geochronological data to 1997). The foreland in this part of the Southern
Fig. 4. Lithoprobe deep seismic reflection transects across the Grenville Front and Allochthon Boundary Thrust.
(a) part of GLIMPCE-J line in Ontario. (b) part of line 15 in western Québec. The Grenville Front (GF) and Grenville
Front Tectonic Zone (GFTZ) are crustal-scale contractional structures reaching depths of 25–30 km. The Allochthon
Boundary Thrust (ABT) is a sub-horizontal contractional structure that was reworked in extension locally.
(Penokean) Province is underlain by the Palaeopro- Petrographic analysis of Sudbury dykes in the
terozoic Huron Supergroup and Nipissing gabbro GFTZ has shown that metamorphic mineralogy
sills with E-trending structures and greenschist- in coronas separating primary olivine and plagio-
facies assemblages cut by a younger NE-trending clase changes from amphibolite- to granulite-facies
belt of Palaeoproterozoic granite. Reworked equiva- with distance from the Grenville Front. Peak P–T
lents of all these units occur in the GFTZ, the estimates (i.e. P at maximum T ) for garnet-bearing
location of the Grenville Front in the granitoid belt coronas range from c. 630 MPa/710 8C for
suggesting it followed a pre-existing discontinuity. amphibolite-facies assemblages c. 7 km SE of the
The SE-trending Sudbury dykes exhibit straight seg- Grenville Front to c. 830 MPa/730 8C for granulite-
ments and are undeformed in the foreland, some facies assemblages c. 16 km SE of the Front, imply-
are truncated at the Grenville Front and those that ing the depth of exhumation increases towards
continue into the GFTZ are disrupted, folded and the SE (Bethune 1997; Bethune & Davidson
partially transposed into NE trends (Fig. 5a). 1997). Circa 1 Ga zircon overgrowths on igneous
412
Table 1. Compilation of Rigolet mineral assemblages and estimated metamorphic conditions, U – Pb ages of metamorphism, and 40Ar/39Ar Hbl cooling ages in the
Parautochthonous Belt (PB)
Location W ! E, Assemblages Estimated P –T– t of References Comments

pMP/pHP metamorphism
20 km transect from Granitic ogn, pgn, amphibolite – Ttn 11 U– Pb Ttn analyses lie on chord Haggart et al. 1993; Krogh c. 0.98 Ga Ttn rims on
GF near Killarney and Hbl-bearing assemblages with UI at 1450 Ma at GF and 1994 1.45 Ga Ttn (or Pb loss?)
(Ontario) SE into LI at 978 Ma 20 km SE of GF. Hbl plateau ages 15 km
GFTZ [pMP] Ar Hbl plateau ages from SE of GF (partially reset
993 – 979 Ma older ages near GF)
c. 30 km SE GFTZ, Garnet amphibolite: Grt-Pl-Hbl- 1100 MPa/750 8C c. 990 – Culshaw et al. 1991; Jamieson Quasi-isothermal
Georgian Bay, Bt-Qtz-Ilm-Spl 980 Ma. Ar Hbl ages from et al. 1995; Reynolds et al. decompression path in
(Ontario) [pMP] 980 – 960 Ma 1995 southern GFTZ
NW GFTZ (between Metabasite: Grt-Cpx-Pl-Opx- 630 MPa/710 8C c. 7 km SE of Davidson & van Breemen Metamorphic field gradient
Killarney and Qtz-Hbl-Spl GF, 830 MPa/730 8C c. 16 km 1988; Dudás et al. 1994; SE of GF
Sudbury, Ontario) SE of GF. 1.0 Ga overgrowths Bethune 1997; Bethune &
[pMP] of metamorphic Zrn on Davidson 1997
igneous Bdy
T. RIVERS
GFTZ SE of Val d’Or Metabasite: Grt-Cpx-Pl- 1200– 1500 MPa/800 8C to Indares & Martignole 1989; Quasi-isothermal
(W. Québec) [pMP Opx-Qtz-Hbl 900 MPa/700 8C c. 1 Ga (Mnz) Gariépy et al. 1990; Childe decompression path.
and pHP] Ar Hbl ages c. 995 Ma et al. 1993; Martignole & Grenvillian Mnz in
Reynolds 1997; Martignole southern PB
& Martelat 2005
Gagnon terrane Pelite: Qtz-Ms-Bt-Grt-Ky-Pl-L c. 1400 MPa/.800 8C 995 – Indares 1995; Jordan et al. Bt dehydration melting
(E. Québec) [pMP 985 Ma 2006 followed by melt
and pHP] crystallization
Gagnon terrane (near Pelite: Mfg from Qtz-Chl-Bt to c. 600 MPa/450 8C 10 km SE of Rivers 1983a, b; Dallmeyer & Inverted metamorphic field
Wabush, W. Labrador) Qtz-Ms-Bt-Grt-Ky-Pl-L GF to c. 1100 MPa/750 8C Rivers 1983; van Gool et al. gradient in thrust stack
[pMP] 30 km SE of GF. c. 1 Ga 2008; Rivers et al. 1995, SE of GF
(Zrn, Mnz) Ar Hbl ages 2002; Cox & Rivers in press
c. 960– 940 Ma
Molson Lake terrane Prg-Cpx-Opx-Grtcoronas separating 1200– 1400 MPa, 700– 750 8C in Dallmeyer & Rivers 1983; Coronitic textures.
[pHP] Ol and Pl (Ky inclusions in NW, 1600 – 1800 MPa, Rivers & Mengel 1988; Metamorphic field
Pl; 51% Jd in Cpx) 800 – 850 8C in SE. Zrn LI Connelly & Heaman 1993; gradient from NW to SE.
c. 1005 Ma; Ar Hbl cooling Indares & Rivers 1995;
ages 940 – 905 Ma Connelly et al. 1995
Smokey Archipelago Mafic to intermediate 560– 630 8C, P undetermined Owen et al. 1986, 1988; Negligible Grenvillian
(E. Labrador) [pMP] orthogneiss: Grt-Hbl-Bt-Pl-Qtz c. 1040 – 1030 Ma (Zrn: LI) Krogh et al. 2002 Zrn and Mnz near GF
Abbreviations: pMP, medium pressure; pHP, high pressure; GF, Grenville Front; GFTZ, Grenville Front Tectonic Zone; mfg, metamorphic field gradient; ogn, orthogneiss; pgn, paragneiss; Bdy, Baddeleyite;
Bt, biotite; Cpx, clinopyroxene; Grt, garnet; Hbl, hornblende; Ilm, ilmenite; Jd, jadeite; Ky, kyanite; L, leucosome (granitic liquid); Mnz, monazite; Ms, muscovite; Ol, olivine; Opx, orthopyroxene;
Pl, plagioclase; Qtz, quartz; Spl, spinel; Ttn, titanite; Zrn, zircon. LI, UI, lower and upper intercepts on concordia.
Fig. 5. Geological maps and cross-sections illustrating features of the Grenville Front and adjacent
Parautochthonous Belt, northwestern Grenville Province; for locations see Figure 2. (a) Sudbury dykes (SD) in
the Grenville Front Tectonic Zone (GFTZ) of Ontario. (b) Truncation of the Huron Supergroup at the Grenville
Front near Coniston Ontario. (c) Mafic dykes in the Grenville foreland and the GFTZ in western Québec.
414 T. RIVERS
Fig. 5. (Continued) (d) Regional map of the Grenville Front near Forsythe, Québec (redrawn from 1:400 000 scale
on-demand map of the Ministère des richesses naturelles, QC). (e) Grenville Front near Chibougamau, Québec.
(f) Grenville Front near Lac Mistassini, Québec (redrawn from 1:400 000 scale map of the Ministère des richesses
naturelles, QC).
baddeleyite indicate corona reactions took place with Nipissing gabbro sills is also present in the
during the Rigolet orogenic phase (Davidson & Grenville foreland near Coniston, some 80 km
van Breemen 1988). farther NE (Fig. 5b). However, at this location
both units and the Sudbury dykes are essentially
Grenville Front Tectonic Zone near Coniston, truncated close to the Grenville Front, which is the
Ontario. The greenschist-facies Huron Supergroup site of an abrupt rise in metamorphic grade indicated
Fig. 5. (Continued) (g) Grenville Front and Para-autochthonous Belt in Gagnon terrane, western Labrador; cross-
sectional model shows inferred tectonic architecture before erosion. (h) Grenville Front near Smokey, eastern Labrador.
by staurolite and sillimanite isograds mapped in thin Palaeoproterozoic Murray Fault system, indicating
slivers of pelitic schist derived from the Huron the Grenville Front followed a pre-existing structure
Supergroup (Davidson 1998). Farther SE, the at this locality too. The ‘disappearance’ of a thick
Huron Supergroup and Nipissing gabbro have not supracrustal sequence together with the rapid rise
been identified, their absence first documented by in metamorphic grade implies the Grenville Front
Quirke & Collins (1930) in their paper entitled is the site of profound exhumation over a narrow
‘The disappearance of the Huronian’ and now inter- zone a few kilometres wide at this locality.
preted to be due to exhumation above the present
erosion surface. Davidson (1998) located the Gren- Grenville Front Tectonic Zone near Grand Lac
ville Front at the steep Wanapitei Fault (Fig. 5b), in Victoria, western Québec. The Grenville foreland
which slivers of quartzite of the Huron Supergroup in western Québec is composed of belts of Neoarch-
exhibit mylonitic and ultramylonitic microstruc- aean supracrustal and granitoid rocks cut by Proter-
tures. The Wanapitei Fault reworks part of the ozoic mafic dykes, including the Senneterre and
416 T. RIVERS
Preissac (c. 2214 Ma) and Abitibi (c. 1140 Ma) and penetrative ductile Grenvillian reworking
swarms, the deformed and disaggregated remnants within the Parautochthonous Belt.
of which continue into the Grenvillian Parautoch-
thonous Belt (Fig. 5c; Martignole & Martelat Grenville Front Tectonic Zone near Chibougamau,
2005). In this area, the GFTZ consists of stacked Québec. Another strand of the Abitibi greenstone
slices of reworked Archaean rocks, including belt intersects the Grenville Front near Chibouga-
amphibolite-facies metasediments (in Lac Témiska- mau (Fig. 5e). At this location, in contrast to near
mingue terrane) structurally overlain by granulite- Forsythe, the greenstone belt is effectively termi-
facies migmatite and orthogneiss (in X-terrane). nated at the Front, replaced to the SE by km- to
X-terrane is itself structurally overlain by dm-scale pods and lenses of chemically similar
amphibolite-facies gneisses of the Réservoir amphibolite, meta-anorthosite and metaperidotite
Dozois terrane along the Dorval Detachment. In in tonalitic to dioritic gneiss (Bandyayera et al.
terms of their Grenvillian structure, the Archaean 2006). Daigneault & Allard (1994) described the
rocks of the Parautochthonous Belt exhibit NE Grenville Front at this location as a 20 km-wide
trends parallel to the Grenville Front and gently to amphibolite-facies shear zone with SE-plunging
moderately SE-plunging elongation lineations mineral lineations and reverse (SE-side-up) ductile
developed during NW-directed thrusting. Lineations strain indicated by rotated porphyroblasts, s–c
rotate into ENE-trends in the vicinity of the Dorval fabrics, shear bands and asymmetric pressure
Detachment, a crustal-scale transtensional structure shadows. This shear zone is overprinted by a
that offsets the Moho at the SE end of the cross- family of narrow lower-grade shear zones (shown
section (Fig. 5c; Martignole & Martelat 2005). as faults on Fig. 5e), also with SE-plunging linea-
Grenvillian coronas in mafic dyke remnants tions indicating reverse and minor sinistral strike-
in Lac Témiskamingue terrane at the base of slip kinematics. Taken together, the two sets of
the thrust stack exhibit amphibolite-facies assem- structures document progressive exhumation of the
blages, whereas those in the overlying X-terrane greenstone belt towards the NW, with the result
are granulite-facies. Peak P–T estimates of c. 900 that the Parautochthonous Belt is underlain by its
MPa/700 8C and 1500 MPa/750 –800 8C, respect- reworked orthogneissic basement. Daigneault &
ively have been determined (Martignole & Allard (1994) defined a ‘Foreland –Parautochthon
Martelat 2005), implying the latter are high-pressure Transition Zone’ across which cumulative exhuma-
granulites, with variably retrograded samples defin- tion of at least 5–10 km was inferred. As at Killar-
ing a quasi-isothermal decompression path. Several ney and Grand Lac Victoria, Proterozoic mafic
geochronological studies of reworked gneisses dykes that cross-cut Archaean structures in the fore-
from this area using the ID-TIMS method have land are truncated, segmented and reoriented sub-
yielded Archaean ages (e.g. Gariépy et al. 1990; parallel to the Grenville Front.
Childe et al. 1993), but in-situ chemical dating of
monazite has provided evidence for 1 Ga over- Grenville Front near Lac Mistassini, Québec. Fifty
growths on Archaean grains and for Grenvillian kilometres NE of Chibougamau, the Grenville
monazite as young as c. 960 + 30 Ma (Martignole foreland is underlain by the Palaeoproterozoic Mis-
& Martelat 2005). tassini Group, an unmetamorphosed and essentially
flat-lying platformal carbonate sequence. As shown
Grenville Front Tectonic Zone near Forsythe, in the northern-most part of Figure 5e and in
Québec. The area shown in Figure 5d overlaps Figure 5f, the Mistassini Group is truncated at the
with that shown in Figure 5c, but focuses on the Grenville Front, which is placed locally at the
architecture of Neoarchaean mafic rocks where the Mistassini Fault, one of the late, narrow, reverse-
Val d’Or strand of the Abitibi greenstone belt is sense shear zones noted above, that at this location
intersected by the Grenville Front. The sketch-map exhumes the reworked gneissic basement of the
illustrates the dramatic change in outcrop pattern greenstone belt over the Mistassini Group
from wide WNW-trending greenschist-facies belts (Daigneault & Allard 1994). As shown in a cross-
on the Superior side to thin wispy, NE-trending section from the southern end of its exposure in a
layers, pods and lenses of polydeformed amphibo- former mine (Fig. 5f), folding in the Mistassini
lite on the Grenville side. Granitoid rocks on the Group is restricted to the immediate footwall of
Superior side commonly exhibit igneous textures the Mistassini Fault and the lack of metamorphism
and original contact relationships, whereas quartzo- implies a structurally high level of the foreland is
feldspathic gneisses on the Grenville side preserve preserved here.
little evidence of either. These features suggest the
exhumed, reworked roots of the greenstone belt Grenville Front near Wabush – Labrador City,
occur at the erosion surface SE of the Grenville western Labrador. In western Labrador, the Gren-
Front, implying significant reverse displacement ville foreland is underlain by a Palaeoproterozoic
continental-margin sequence (the Knob Lake Gabbro. Grenvillian metamorphic grade reached
Group, Fig. 2) and its Archaean basement. Both epidote-amphibolite facies in domain T and mid-
cover and basement lithologies can be traced amphibolite facies in domain G, where Michael
across the Grenville Front into the parautochtho- Gabbro dykes exhibit garnet amphibolite margins.
nous Gagnon terrane (Fig. 3), where they are imbri- Owen et al. (1988) estimated Grenvillian tempera-
cated in a metamorphic foreland-fold-thrust belt ture reached 560–630 8C in domain G. Krogh
(Fig. 5g; Rivers 1983a, b; Rivers et al. 1993; van et al. (2002) showed that zircon, monazite and tita-
Gool et al. 2008). The NE grain of the thrust belt, nite in the vicinity of the Grenville Front yield
the SE plunge of the mineral elongation lineations, pre-Grenvillian U – Pb ages, and Grenvillian lower
and the asymmetry of ductile microstructures all intercept ages of c. 1040–1030 Ma were obtained
indicate NW-vergence. The Gagnon terrane is over- in the vicinity of the Cut Throat Island Fault. This
lain by the Molson Lake terrane, which is character- suggests this shear zone was active in Ottawan
ized by a high-pressure Grenvillian metamorphic time, but whether it was also active in Rigolet
signature (Indares & Rivers 1995). Within the time, as elsewhere along the front, has not been
Gagnon terrane, thrusting was thin-skinned in the determined.
Palaeoproterozoic cover sequence and thick-
skinned in the underlying Archaean basement, Summary of characteristics
leading to a dual-level thrust belt (Fig. 5g). In this
part of the Gagnon terrane, grade of metamorphism The seismic sections indicate the Grenville Front is
ranges from greenschist facies near the Grenville the limit of SE-dipping Grenvillian fabrics and the
Front, to upper amphibolite facies farther SE structural/metamorphic studies imply reverse-sense
where partial melting was widespread, the higher kinematics consistent with the increasing grade
grade rocks structurally overlying the lower grade of Grenvillian metamorphism in that direction.
rocks in a classic inverted metamorphic sequence. However, as noted, the total amount of differential
P–T estimates range from c. 600 MPa/450 8C near exhumation across the front is limited by the conti-
the Grenville Front to c. 1100 MPa/750 8C some nuity of several supracrustal units across it, imp-
30 km from the front near the southeastern limit of lying the vertical component of transport could not
the Gagnon terrane (van Gool et al. 2008). Geochro- have exceeded a few tens of kilometres. This is com-
nological data, including metamorphic monazite patible with the observation that it follows and
ages, crystallization ages of cross-cutting pegmatite, reworks pre-existing structures locally and that the
and several zircon lower intercepts, indicate meta- intensity of Grenvillian strain was quite variable
morphism in both Gagnon and Molson Lake ter- (Fig. 6). In terms of its architecture, although the
ranes took place at c. 1 Ga (Rivers et al. 2002; van front is everywhere a moderately-dipping, reverse-
Gool et al. 2008; Cox & Rivers in press). About sense shear zone, its width and soling depth vary
200 km farther SW in the Gagnon terrane significantly and it feeds into a mid-crustal foreland
in eastern Québec, peak P –T rises to c. 1400 fold-thrust belt locally.
MPa/8008C and the metamorphism has been dated Also variable is the extent of exhumed high-
at 995 –985 Ma (Indares 1995; Jordan et al. 2006). grade Grenvillian metamorphic rocks from the
mid and lower crust. Such rocks are absent near
Grenville Front in the Smokey Archipelago, eastern Témiskamingue and in the Smokey Archipelago
Labrador. The Grenville Front in the Smokey (Fig. 6b, f ), suggesting these areas represent rela-
Archipelago at the east of the exposed Grenville tively high levels of the Parautochthonous Belt,
Province was studied by Owen et al. (1986, 1988). whereas a deeper crustal level is exposed in the
Figure 5h shows the NNE-trends in the Palaeopro- GFTZ. In detail, two styles of crustal architecture
terozoic Makkovik Province in the foreland are are apparent in the GFTZ: a crustal-scale shear
truncated at the Grenville Front by E-trending Gren- zone in which the depth of exhumation increases
villian structures. In detail, there are two narrow, systematically towards the SE (e.g. near Killarney
sub-parallel, thrust-sense shear zones, the Benedict and Chibougamau; Fig. 6a, d), and two lithologi-
and Cut Throat Island faults, the latter being con- cally- and metamorphically-distinct stacked terranes
sidered the local Grenville Front. North of the Ben- derived from discrete crustal levels (near Grand Lac
edict Fault (domain M), Makkovik structures are Victoria and Wabush; Fig. 6c, e).
essentially unaffected by Grenvillian overprinting; The SE-plunging elongation lineations and
in domain T (transition) between the Benedict and reverse-sense kinematics imply tectonic transport
Cut Throat Island faults both Makkovik and Gren- was approximately orthogonal to the Grenville
villian fabrics are present, and south of Cut Throat Front, except in local NNE-trending transfer zones
Island Fault Grenvillian structures become more where strike-slip structures record a component of
prominent (domain G), as exhibited by polyphase sinistral transpression. However, the depth from
folding and boudinage of the c. 1.43 Ga Michael which the overthrust slice was derived, the extent
418 T. RIVERS
Fig. 6. Summary figure showing schematic crustal-scale structural signatures of the Grenville Front.
Fig. 7. Geological maps and crustal-scale cross-sections illustrating features of the Allochthon Boundary Thrust
and adjacent terranes and domains. (a) In Ontario showing the distribution of mafic rocks on either side of the
Allochthon Boundary Thrust (1.24 Ga Sudbury diabase in footwall and 1.16 Ga Algonquin gabbro and retrogressed
eclogite in hanging wall). Note lobate shape of Muskoka domain (MD). (b) In western Québec. (c) In eastern Québec –
western Labrador showing the distribution of c. 1.43 Ga Shabogamo Gabbro in the footwall of the Allochthon Boundary
Thrust: the hanging wall Lac Joseph and Hart Jaune terranes (LJT, HJT) contain c. 1.65 and 1.45 Ma granite and
gabbronorite bodies not present in the footwall. (d) In eastern Labrador showing the distribution of c. 1.43 Ga Michael
Gabbro in the footwall of the Allochthon Boundary Thrust: GBT, Groswater Bay terrane; LMT, Lake Melville terrane;
MMT, Mealy Mountains terrane.
420 T. RIVERS
to which deformation propagated into the foreland, orogenic collapse (Fig. 6c, e). However, in contrast
the presence or absence of ramp-flat geometry and to the Allochthon Boundary Thrust (see below),
the intensity of Grenvillian strain are all variable, the Grenville Front itself was not reworked in
presumably controlled by local features such as extension.
footwall rheology, pre-existing structural weak- In summary, peak pressure estimates for meta-
nesses and the amount of shortening. For instance, morphic rocks in the hanging wall of the Grenville
where present, high-pressure metamorphic terranes Front imply exhumation from depths of c. 20 –
occur above the structural base of the thrust stack 45 km, zircon and monazite ages of c. 1000–
suggesting their emplacement drove thrust propa- 980 Ma indicate metamorphism was during the
gation into the foreland, and they are structurally Rigolet orogenic phase except possibly in the east-
overlain by normal-sense detachment zones that ernmost Grenville Province, and the 40Ar/39Ar
carry lower grade rocks in their hanging walls, hornblende ages of c. 990–940 Ma (Table 1)
implying the overthickened crust underwent imply rapid post-peak cooling.
The interior Grenville Province – Ottawan Davidson (2000), and is principally known from
deformation and metamorphism regional studies (e.g. Rivers et al. 1989, 2002;
Wardle et al. 1990; Culshaw et al. 1997; Carr
In this study, the interior Grenville Province is et al. 2000; Corrigan et al. 2000; Hynes et al.
defined as that part of the orogen in the hanging 2000; Martignole et al. 2000). Following an
wall of the Allochthon Boundary Thrust. However, examination of seismic images of the Allochthon
the thrust itself, despite its recognized impor- Boundary Thrust, the geological signature of
tance, has received little detailed examination rocks in its hanging wall is reviewed in four
except by Ketchum et al. (1998) and Ketchum & case studies.
422 T. RIVERS
Seismic images of the Allochthon Allochthon Boundary Thrust and the

Boundary Thrust allochthonous terranes – Case studies
Seismic images of the northwestern extremity of the Allochthon Boundary Thrust near Georgian Bay,
Allochthon Boundary Thrust in Figure 2 indicate it Ontario. In the western Grenville Province, the Allo-
is a gently SE-dipping structure located above the chthon Boundary Thrust exhibits a lobate map pattern
Grenville Front. In the seismic section in Ontario and is manifest in the field as a km-wide shear zone
(Fig. 4a) it is not the site of a change in reflectivity with amphibolite-facies high-strain fabrics and
and so it is not readily picked without surface infor- reverse-sense kinematics locally overprinted in
mation, whereas in that of western Québec (Fig. 4b) extension. Lithological contacts are ‘buried’ in the
it forms a strong package of reflectors at the base of wide high-strain zone (Culshaw et al. 1997), but are
a klippe. On the basis of seismic sections SE of those apparent at the regional scale. For instance,
shown, the Allochthon Boundary Thrust is inferred Figure 7a shows that it separates 1.24 Ga Sudbury
to be a crustal-scale structure that penetrates the dykes metamorphosed at c. 1000 Ma in its footwall
mid and lower crust in the SE Grenville Province. from 1.16 Ga Algonquin gabbro and relict eclogite
metamorphosed at c. 1090 Ma in its hanging wall Indares et al. 2000; Rivers et al. 2002). The original
(Ketchum & Davidson 2000). Moreover, it also thrust-sense boundary of the Allochthon Boundary
coincides with a change in Nd model ages (Dickin Thrust is preserved at the base of the aHP Lelukuau
& Guo 2001), all features consistent with it being terrane, but elsewhere it was extensively reworked
the site of major tectonic transport. On the basis of in extension associated with the emplacement of
structural constraints, Culshaw et al. (1997) esti- the Orogenic Lid (Hart Jaune and Lac Joseph ter-
mated it had accommodated a minimum of 100 km ranes). Well-preserved eclogite-facies rocks in the
displacement. The cross-section in Figure 7a shows Lelukuau terrane have yielded metamorphic ages
that it reaches the mid-crust, implying that the relict of c. 1060– 1040 Ma and subsequent extension has
eclogite-facies (lower-crustal) rocks in its hanging been dated at c. 1015 Ma (Indares et al. 1998,
wall must have been derived from farther SE. 2000). The seismic section shows that the Alloch-
The structural succession in the hanging wall of thon Boundary Thrust reaches the Moho in this area.
the Allochthon Boundary Thrust at this location
comprises: (i) the aHP Algonquin terrane character- Allochthon Boundary Thrust near Rigolet, eastern
ized by 1090 Ma relict eclogite, overlain by (ii) Labrador. The structural interpretation of the
the aM-LP Muskoka domain characterized by Allochthon Boundary Thrust in eastern Labrador
c. 1080 Ma upper amphibolite-facies and granulite- (Fig. 7d) is unconstrained by seismic information
facies assemblages, which in turn is overlain by (iii) and complicated by several ages of faulting. In
the Composite Arc Belt, part of which was reworked addition to brittle normal faults associated with the
under MP Ottawan metamorphic conditions and Neoproterozoic Lake Melville graben, the Rigolet
part of which comprises the Orogenic Lid. thrust, a ductile shear zone assumed to be the
As indicated in Figure 2, the Allochthon Boun- location of the Allochthon Boundary Thrust during
dary Thrust approximately coincides with the north- field studies, appears to be of Palaeoproterozoic
western limit of Ottawan metamorphism in much age (Corrigan et al. 2000; F. Korhonen, personal
of the Grenville Province, but Figure 7a shows that communication 2008), although many of the rocks
Ottawan metamorphism locally occurs NW of (struc- in its hanging wall have undergone penetrative
turally below) the Allochthon Boundary Thrust. Grenvillian metamorphism (see below). As a
Finally, a key feature of the Allochthon Boundary result, the exact location of the Allochthon Bound-
Thrust in this area was its profound ductile rework- ary Thrust is not well constrained, but it may lie
ing in extension at c. 1020 Ma as part of the Alloch- along an unnamed transpressional shear zone a
thon Boundary Detachment system of shear zones few kilometres south of the Rigolet Thrust.
(Ketchum et al. 1998). It appears to have been inac- Despite this uncertainty, its location can be approxi-
tive since that time. mately mapped from the distributions of the
c. 1.43 Ga Michael Gabbro with c. 1000 Ma meta-
Allochthon Boundary Thrust in western Québec. morphism in its footwall (in Groswater Bay
Although less studied than in Ontario, the Allochthon terrane) and Palaeoproterozoic gabbro bodies with
Boundary Thrust in western Québec (Fig. 7b) also c. 1080 Ma metamorphism in its hanging wall (in
exhibits a lobate outcrop pattern and separates Lake Melville terrane). In the south of the map
rocks with remnants of Proterozoic dyke swarms area, Lake Melville terrane is separated from the
present in the foreland in its footwall (Fig. 5c) from Mealy Mountains terrane, part of the Orogenic
mafic rocks with relict HP Ottawan assemblages in Lid, by the transtensional English River shear zone.
its immediate hanging wall (Lac Dumoine terrane Two features of this area set it apart from those
and a related klippe). Although the aM –LP terranes described previously: (i) high-pressure rocks do
do not physically overlie the aHP terranes in this not occur, the Lake Melville terrane being part of
map, they may do in three dimensions. Other features the aM-LP Belt; and (ii), regional E-trending
contrasting with Figure 7a include the presence of folds of ductile gneisses occur in the hanging wall
several transpressional terrane boundaries and the of the Allochthon Boundary Thrust (cross-section,
seismic interpretation that the Allochthon Boundary Fig. 7d). Corrigan et al. (2000) interpreted the latter
Thrust reaches the Moho in this section. structures to have formed when the early Ottawan
(c. 1080 Ma) gneissic fabric underwent shortening
Allochthon Boundary Thrust, eastern Québec – at c. 1050 Ma between two relatively rigid blocks,
western Labrador. In the vicinity of the Manico- the contemporaneous foreland to the north and the
uagan Reservoir (Fig. 7c), the Allochthon Boundary down-dropped Mealy Mountains terrane, part of
Thrust separates terranes containing the c. 1.43 Ga the Orogenic Lid to the south.
Shabogamo Gabbro metamorphosed at c. 1000 Ma
Summary of structural characteristics
in its footwall from terranes containing c. 1.65 Ga
gabbronorite and anorthosite metamorphosed at Its surface manifestation as a ductile shear zone
c. 1060 Ma in its hanging wall (Hynes et al. 2000; whose location is independent of the local footwall
424 T. RIVERS
Fig. 8. (a) Metamorphic map of the Grenville Province showing locations of detailed studies with P –T estimates linked
to geochronology of the metamorphic rocks. (b) P– T diagram showing estimated Ottawan P –T conditions in the
allochthonous belts and Rigolet metamorphic conditions in the Parautochthonous Belt. (c) Summary of Ottawan and
Rigolet metamorphic field gradients. Abbreviations in A and B: B, Berthé terrane; BR, Blair River outlier;
geology, together with the contrasting distributions Allochthonous High Pressure Belt. Three separate
of Mesoproterozoic dyke swarms on either side, areas of outcrop of the aHP Belt have been ident-
suggest the Allochthon Boundary Thrust is the site ified in the western, central and eastern Grenville
of a large, but unquantified amount of subhorizontal Province, in each of which the HP rocks lie close
displacement. This inference is supported by the to or in the immediate hanging wall of the Alloch-
presence of high-grade Ottawan metamorphic thon Boundary Thrust (Fig. 2). In addition as
rocks in its hanging wall, including some with noted, Ottawan HP rocks also occur locally in its
eclogite-facies assemblages implying substantial footwall in the western Grenville Province. Esti-
tectonic transport on such gently-dipping surfaces. mates of peak metamorphic conditions, time of
Although there is local evidence for Ottawan meta- metamorphism and references to the original litera-
morphism in the footwall of the Allochthon Bound- ture are given in Table 2. In the Cape Caribou River
ary Thrust, available evidence suggests that it marks Allochthon in the east, peak P–T estimates in
the approximate limit of high-grade metamorphism HP granulite-facies assemblages are c. 1400 MPa/
in Ottawan times. 8758C and the metamorphism is dated at
Following its initiation as a thrust-sense shear c. 1050 Ma, whereas eclogite and HP granulite
zone, the Allochthon Boundary Thrust was locally from Lelukuau terrane in the central Grenville Pro-
reworked as a normal-sense structure in late vince have yielded peak P–T conditions of c. 1700–
Ottawan times at c. 1020 Ma (Ketchum et al. 1900 MPa/850– 9208C at c. 1060 –1040 Ma. In
1998; Indares et al. 2000; Rivers et al. 2002), and Algonquin–Lac Dumoine terrane in the western
elsewhere normal-sense displacement occurred on Grenville Province, the eclogite assemblages were
shear zones a few kilometres to the SE. Considering pervasively overprinted during exhumation, their
the Grenville Province as a whole, the normal-sense former presence being inferred from textural and
displacements began locally as early as c. 1060 Ma mineralogical evidence (Davidson 1990). As a
but became widespread at c. 1020 Ma, signalling result, no P–T data are available, with the best esti-
gravitational collapse of the Ottawan crust, the mate for the age of eclogite metamorphism being
regional folding in eastern Labrador suggesting 1090 Ma. Finally the Ottawan metamorphic
this occurred in an overall compressional setting. rocks that occur locally in the footwall of the
Allochthon Boundary Thrust have yielded peak
P –T conditions of 1200–1400 MPa/850–870 8C
Metamorphic characteristics of hanging and a minimum monazite age of c. 1035 Ma.
wall rocks Taken together, these data imply differential exhu-
mation of the HP terranes and that timing of
In this section, the metamorphic signature of the Ottawan HP metamorphism was diachronous.
Allochthon Boundary Thrust and the overlying
allochthonous belts in the four areas are integrated Allochthonous Medium to Low Pressure Belt.
with data from elsewhere in the Grenville Province. Regional Grenvillian metamorphic conditions and
Locations mentioned in the text are shown in geochronological information for the aM–LP Belt
Figure 8a. are given in Table 3. With respect to the MP
segment of this belt, peak Ottawan P– T conditions
Allochthon Boundary Thrust. The wide shear zone for these upper amphibolite- and granulite-facies
defining the Allochthon Boundary Thrust exhibits assemblages cluster in the range c. 800–1100 MPa
upper amphibolite-facies assemblages formed by and 800–8508C suggesting the present erosion
retrogression and annealing of the eclogite- and surface exposes metamorphic rocks formed at 25 –
granulite-facies rocks in the adjacent hanging wall 33 km depth. The P –T data plot in the sillimanite
terranes. As a result, there is no obvious meta- field, compatible with its presence in assemblages
morphic contrast across it in the field. from this belt, and imply an elevated geothermal
Fig. 8. (Continued) BT, Britt domain; CAB, Composite Arc Belt; CCR, Cape Caribou River allochthon; FAB,
Frontenac– Adirondack Belt; GFTZ, Grenville Front Tectonic Zone (On, Ontario; Qc, western Québec); Gn, Gagnon
terrane (Qc, eastern Québec; WL, western Labrador); L, Lelukuau terrane; LD, Lac Dumoine terrane; LM, Lake
Melville terrane; LR, Long Range outlier; LaR, La Romaine domain; Ma, Mazinaw domain; MIZ, Manicouagan
Imbricate Zone; Mk, Muskoka domain; ML, Molson Lake terrane; MT, Mékinac–Taureau terrane; N, Natashquan
domain; P, Pinware terrane; PSM, Portneuf– St-Maurice domain; S, Shawanaga domain; SJ, St-Jean domain; T,
Tshenukutish terrane. Abbreviations and symbols in B and C: aHP and aM– LP, allochthonous High-Pressure and
Medium- to Low-Pressure belts with Ottawan metamorphism; OL, Orogenic Lid; pMP and pHP, parautochthonous
Medium-Pressure and High-pressure belts with Rigolet metamorphism. Circles in B represent approximate
uncertainties in P –T estimates, ellipses represent ranges of P– T conditions in metamorphic field gradients, arrows
indicate pro- or retrograde character.
426
Table 2. Compilation of Ottawan metamorphic conditions, U – Pb metamorphic ages, and 40Ar/39Ar hornblende and muscovite cooling ages in the aHP Belt
Location W ! E Assemblages Estimated P –T and time of References Comments

metamorphism
Shawanaga domain Cpx-Opx-Grt-Pl (retrograde Prg, U/Pb Zrn: Eclogite facies Davidson 1990, 1998; Granulite-facies overprint
[western HP segment] Spl, Spr, Crn replacing 1090 Ma? Granulite Bussy et al. 1995; associated with exhumation of
Grt-Omp) facies at c. 1080 Ma Ketchum & Krogh 1997, HP Belt; amphibolite-facies
Amphibolite facies at 1998; Ketchum et al. overprint with extension on
c. 1020 Ma. Ar Hbl 1998; Reynolds et al. Shawanaga shear zone
cooling age 970 Ma 1995
Southern GFTZ and Grt-Hbl-Pl + Qtz + Cpx + HP granulite; 1380 MPa/ Corrigan et al. 1994; Structurally beneath the ABT. Mnz
northern Britt domain Opx + Bt + Ilm + Mt 870 8C. U – Pb Mnz: Jamieson et al. 1995 dates are minimum ages for the
[western HP segment] 1037 + 1, 1035 + 1 Ma. granulite-facies metamorphism
T. RIVERS
Ar Hbl c. 975 – 965 Ma;
Ms c. 925 – 905 Ma
Lac Dumoine terrane Cpx-Opx-Grt-Pl (retrograde Prg, 1350 MPa, 7208C 1070 Ma Indares & Martignole Retrograde P– T conditions
[western HP segment] Spr, Crn replacing Grt-Omp) 1990a; Indares &
Dunning 1997
Manicouagan Imbricate Cpx-Opx-Grt-Pl (Ky inclusions in 1700– 1900 MPa, 750– Indares 1997, 2003; Cox Eclogite and HP granulite with
Zone (Lelukuau and Pl); Grt-Omp replacing igneous 920 8C U – Pb Zrn: 1060 – et al. 1998; Cox & coronitic and granoblastic
Tshenukutish terranes) Ol, Px and Pl 1040 Ma Indares 1999a, b; Indares textures. HP granulite in
[central HP segment] et al. 1998, 2000; lithologies with low
Indares & Dunning bulk Na2O. Steep retrograde
2004; Yang & Indares dP/dT paths
2005
Cape Caribou River Prg-Cpx-Opx-Grt-Pl 1400 MPa, 875 8C c. Krauss & Rivers 2004; Cox HP granulite assemblages
Allochthon [eastern 1050 Ma (U – Pb Zrn) et al. unpublished
HP segment]
ABT, Allochthon Boundary Thrust; Mineral abbreviations: Cpx, clinopyroxene; Crn, corundum; Grt, garnet; Hbl, hornblende; Ilm, ilmenite; Jd, jadeite; Ky, kyanite; Mnz, monazite; Ms, muscovite;
Mt, magnetite; Ol, olivine; Omp, omphacite; Opx, orthopyroxene; Pl, plagioclase; Prg, pargasite; Px, pyroxene; Spl, spinel; Spr, sapphirine; Zrn, zircon.
Table 3. Compilation of Ottawan metamorphic conditions, U – Pb metamorphic ages, and 40Ar/39Ar hornblende and muscovite cooling ages in the aM – LP Belt
Location W ! E; Lithology and assemblages Metamorphic facies, Estimated References Comments

MP/LP P – T and time of metamorphism
THE GRENVILLE PROVINCE AS A LARGE HOT LONG-DURATION COLLISIONAL

Parry Sound terrane Pelite, orthogneiss, marble Granulite- to Ur amphibolite; Reynolds et al. 1995; Wodicka Ottawan overprint on
MP c. 1080 Ma; Ar Hbl c. 890 – et al. 1996 Shawinigan metamorphism
880 Ma
Shawanaga domain Principally felsic orthogneiss, Ur amphibolite; c. 1050– Culshaw et al. 1994, 1997; Ottawan metamorphism at
MP minor paragneiss 1020 Ma. Ar Hbl c. 970 Ma Reynolds et al. 1995; c. 1050 Ma, extensional
Ketchum et al. 1998 shearing at c. 1020 Ma
Muskoka domain Metabasite: 1000– 1150 MPa, 750 – 850 8C; Cosca et al. 1991; Cross-cutting granulite vein
MP Pl-Grt-Cpx-Opx-Hbl-Qtz c. 1080– 1050 Ma. Ar Hbl Timmermann et al. 1997, yielded 1065 Ma (Zrn) and
Felsic granulite: c. 1000 and 970 Ma 2002; Slagstad et al. 2004 1045 Ma (Ttn) ages; implies
Kfs-Pl-Qtz-Opx-Hbl-Bt hot crust for 20 – 30 My.
Mazinaw terrane LP Pelite: Qtz-Bt-Ms-Ky-Sil-Pl-Trm Greenschist to Ur amphibolite; Moore & Thompson 1980; Metamorphic field gradient from
c. 300– 600 MPa, 550– 650 8C; Cosca et al. 1991; Corfu & greenschist to Ur amphibolite
1050 – 1000 Ma. Ar Hbl Easton 1995 facies
c. 940 Ma
Adirondack Pelite: Crd-Grt-Sil-Spl-Opx-Pl- Granulite; 700 – 800 MPa, Bohlen 1987; Onstott & Ottawan overprint on
Highlands terrane Bt-Kfs-Ilm-Trm-Psm-L 675 – 915 8C; Ottawan rims Peacock 1987; Mezger et al. Shawinigan (c. 1180 –
MP date anatexis at c. 1050 – 1991; McLelland et al. 1140 Ma) granulite-facies
1020 Ma (U – Pb Zrn). 1996, 2001; Spear & metamorphism. Ottawan
Local outermost rims yield Markussen 1997; Alcock anatexis at 1050 Ma, orogenic
1012 – 990 Ma ages. Ar Hbl et al. 2004; Darling et al. collapse on Carthage-Colton
c. 950– 900 Ma 2004; Johnson et al. 2004; shear zone pre-1020 Ma.
Bickford et al. 2008 Local Zrn growth syn-Rigolet
Morin terrane MP Pelite: Qtz-Pl-Kfs-Bt-Grt-Sil Granulite; 600 – 700 MPa, Indares & Martignole 1990b; Peck et al. (2005) argued
Metabasite: 650 – 775 8C Peak Martignole & Friedman assemblages are Ottawan
Pl-Hbl-Grt-Opx-Cpx-Qtz-Bt-L metamorphism not directly 1998; Peck et al. 2005; based on reactions in skarn
dated. Ar Hbl c. 1050– Martignole et al. 2006 adjacent to Morin anorthosite
990 Ma
Mékinac–Taureau Granulite-facies orthogneiss Granulite; 700 – 1150 MPa, Corrigan & van Breemen 1997 FW of extensional Tawachiche
terrane MP 725 – 900 8C; c. 1087 Ma shear zone
(Zrn; thrusting) 21050 Ma
(Zrn; extension)
Portneuf – Pelite: Opx-Ath-Crd-Bt-Pl, Ur amphibolite to granulite; Herd et al. 1986; Corrigan & Montauban Group, HW of
St.-Maurice LP Crd-Spl-Grt-Sil-Qtz 500 – 600 MPa, 750 8C; van Breemen 1997 extensional Tawachiche
c. 1056 Ma (Zrn) shear zone
(Continued)
427
428
Table 3. Continued
Location W ! E; Lithology and assemblages Metamorphic facies, Estimated References Comments

MP/LP P – T and time of metamorphism
Berthé terrane Pelite: Qtz-Pl-Grt-Sil-Kfs-Bt-L Granulite; 1000 – 1100 MPa Indares & Dunning 2004 Structurally juxtaposed against
(Gabriel Metabasite: Opx-Cpx-Pl-Grt c.1050– 1040 Ma (Zrn) Grenvillian HP Belt and
complex) MP Orogenic Lid
St-Jean domain Pelite: Qtz-Kfs-Grt-Sil-Bt Ur amphibolite to granulite Gobeil et al. 2003; Wodicka Structurally overlies Natashquan
MP? Metabasite: Opx-Cpx-Pl-Grt (P – T unquantified); 1080– et al. 2003 domain
1070 and 1060– 1045 Ma (Zrn
& Mnz), 950 Ma (Rt)
Natashquan domain Wakeham Group: pelite – Greenschist to amphibolite; Indares & Martignole 1993; P– T estimate from
LP Qtz-Ms-Grt-Bt-And + St + Sil 350 MPa/5508C; 1052 Ma Madore et al. 1999; Gobeil amphibolite-facies sample in
(Rt), c. 1030, 1010 – 990 Ma et al. 2003; Wodicka et al. metamorphic field gradient
(Mnz), 972, 938 Ma (Ttn) 2003
T. RIVERS
Natashquan domain Wakeham Group: pelite (altered Ur amphibolite, (P – T Bonnet et al. 2005; Corriveau Ur amphibolite- to
– La Romaine volcanic rocks) – unquantified) c. 1019 Ma & Bonnet 2005; van granulite-facies in
segment LP? Grt-Bt-Sil-Pl-Kfs-L, (Zrn), 1010 –1000 (Mnz), Breemen & Corriveau 2005 continuation of Wakeham
Qtz-Crd-Kfs 990 Ma (Ttn) Group
Pinware terrane Orthogneiss Amphibolite; (P – T unquantified) Wasteneys et al. 1997; Gower Pinware terrane MP?
MP? 1036 – 1020 Ma (Zrn) & Krogh 2002
Lake Melville Pelitic paragneiss Granulite to amphibolite, Corrigan et al. 2000 Single-grain Mnz ages
terrane MP Grt-Bt-Ky-Sil-Kfs-L 800 MPa, 820 8C; 1088, 1057
and 1046 Ma (Mnz)
Blair River inlier, Mafic and felsic orthogneiss Granulite (P – T unquantified), Miller et al. 1996
Cape Breton, 1040 Ma (Zrn)
Nova Scotia MP?
Long Range inlier, Mafic and felsic orthogneiss Granulite to amphibolitefacies Heaman et al. 2002
west (P – T unquantified) 1032
Newfoundland and 1022 Ma (Zrn)
MP?
MP, medium pressure; LP, low pressure; ? indicates uncertain affinity; FW/HW, footwall/hanging wall; Lr, lower; Me, middle; Ur, upper. Mineral abbreviations: And, andalusite; Ath, anthophyllite; Bt,
biotite; Cpx, clinopyroxene; Crd, cordierite; Grt, garnet; Hbl, hornblende; Ilm, ilmenite; Ky, kyanite; Kfs, K feldspar; L, leucosome (former granitic liquid); Mnz, monazite; Ms, muscovite; Opx, orthopyroxene;
Pl, plagioclase; Qtz, quartz; Psm, prismatine; Rt, rutile; Sil, sillimanite; Spl, spinel; St, staurolite; Trm, tourmaline; Ttn, titanite; Zrn, zircon.
gradient. High-grade MP Ottawan metamorphism between U – Pb zircon and 40Ar/39Ar hornblende

took place from c. 1080– 1050 Ma at slightly differ- ages (Table 1). The different slopes of the P–T gra-
ent times in individual terranes (Table 3). dients in the Parautochthonous Belt may be a func-
With respect to the LP segments, the Portneuf– tion of the architecture of the Grenville Front, with
St.-Maurice domain exhibits LP granulite- the steeper gradients in the GFTZ related to a
facies assemblages, whereas the LP signature of crustal-scale ramp and the gentler gradients to the
Natashquan domain is indicated by the andalu- presence of flats in a fold-thrust belt where horizon-
site ! sillimanite transition and widespread tal tectonics dominated.
cordierite (Madore et al. 1999; Bonnet et al.
2005; Corriveau & Bonnet 2005). Although the Ottawan metamorphism
kyanite ! sillimanite transition occurs in Mazinaw
terrane, the metamorphic field gradient passes Quantitative estimates of peak Ottawan meta-
close to the Al-silicate triple point indicating a LP morphic conditions form three clusters in
evolution (Moore & Thompson 1980). Estimated Figure 8b, consistent with the aHP, aMP and aLP
peak P– T conditions range from c. 600– 700 MPa/ groups and implying derivation from relatively dis-
750 8C in the Portneuf– St.-Maurice domain, to crete crustal levels. The Orogenic Lid, which was
c. 350 –400 MPa/550 8C in Natashquan domain heated to , 5008C, is interpreted as Ottawan
and Mazinaw terrane, implying metamorphism at upper crust. With regard to the aHP Belt, the
depths of c. 18–20 and 10–12 km, respectively. highest peak P–T estimates are c. 1700–
1800 MPa/800– 8758C, the high temperatures at
Orogenic Lid. The Orogenic Lid is composed of ter- these pressures being characteristic of orogenic eclo-
ranes that lack evidence for penetrative Grenvillian gite formed near the base of doubly thickened crust
metamorphism and deformation, except in their (c. 50–55 km; O’Brien et al. 1990; Indares et al.
basal shear zones. Typical evidence includes pre- 2000; Rivers et al. 2002). With respect to the aMP
Grenvillian 40Ar/39Ar hornblende ages and miaro- segment, all P–T estimates fall in the sillimanite
litic cavities and chilled margins in c. 1080 Ma field, compatible with mineral assemblages
intrusions (Rivers 2008), implying the rocks were (Table 3), and the P estimates of c. 800–1100 MPa
in the upper crust during Ottawan times. The Oro- imply exhumation from 25 –33 km depth, that is
genic Lid occurs in two segments, in the NE Gren- near the middle of doubly thickened crust. The
ville Province where it is largely composed of peak T estimates cluster around 800–8758C, imply-
Palaeoproterozoic (Labradorian) gneisses, and in ing an elevated thermal regime in the Ottawan mid-
the SW within part of the accreted terranes. The crust. Moreover, they overlap with those for the aHP
presence of pre-Grenvillian 40Ar/39Ar hornblende segment, indicating that the temperature did not
ages in both segments implies it was not heated increase significantly from c. 30 km depth (in the
above c. 5008C during the Ottawan orogenic phase. aMP segment) to c. 55 km depth (in the aHP
segment). This may impute a role for advection of
Comparison of Rigolet and Ottawan magmatic heat, compatible with the location of the
aMP segment within the Interior Magmatic Belt
metamorphisms (Fig. 1) and local observations of syntectonic mafic
Estimated P –T conditions during the Rigolet and intrusions (Indares & Dunning 2004). Pressure esti-
Ottawan phases of the Grenvillian Orogeny are mates for the Ottawan LP terranes range from
shown in Figure 8b. c. 350–600 MPa, which when considered with
their associated peak temperatures imply formation
Rigolet metamorphism in a similar LP–HT metamorphic regime to the
MP segment, but at shallower depths.
Most peak P– T estimates for Rigolet metamorph-
ism fall in the kyanite field, compatible with Summary and implications for crustal
mineral assemblages and suggesting a relatively rheology
high-pressure Barrovian character. The data are sub-
divided into pMP and pHP groups with the most The P –T estimates for Ottawan and Rigolet assem-
deeply exhumed parts of the exposed Parautochtho- blages are compatible with the dominant Al-silicate
nous Belt formed at pressures of c. 1400–1700 MPa species present indicating they are probably robust
implying exhumation from 40–50 km depth. In and the differences between them are real. This
western Québec, retrogression from peak condi- has been quantified in Figure 8b by selecting
tions followed a steep dP/dT path suggesting 1000 MPa (c. 30 km depth) as a reference pressure,
rapid exhumation, an interpretation supported by from which it can be seen that the T at this depth
geochronological data from throughout the Parau- was c. 800–900 8C during the Ottawan metamor-
tochthonous Belt, indicating a small difference phism, but c. 650–750 8C during the Rigolet
430 T. RIVERS
metamorphism. This implies the two metamorph- Cross-sections of the Grenville Orogen
isms developed under contrasting geothermal gradi-
ents (Fig. 8c; i.e. c. 27 –308C km21 during the Several deep-seismic images of the Grenville
Ottawan orogenic phase versus c. 23 –278C km21 Orogen produced in the course of the Lithoprobe
during the Rigolet orogenic phase, assuming an program are shown in Figure 9, augmented by
average crustal density of 3000 kg m23). crustal-scale cross-sections derived from structural
Another important contrast between the two studies (modified from Rivers et al. 2002). The
metamorphisms is the rate of cooling. The large overall SE-dipping structural grain of the northern
difference ( 100 My) between U –Pb zircon/mon- Grenville Province, in part defined by the Grenville
azite ages of peak metamorphism and 40Ar/39Ar Front and Allochthon Boundary Thrust, is evident in
hornblende cooling ages for Ottawan metamorphic the cross-sections. Both structures exhibit ramp-flat
rocks (Tables 2 & 3) implies cooling in the interval profiles at the crustal scale and there is a correlation
c. 850 and 500 8C was slow. In contrast, the small between the presence of ramps and the occurrence
difference (15 My) for Rigolet metamorphic of Archaean rocks in the footwall, suggesting the
rocks (Table 1) implies much more rapid cooling. latter were competent and formed a buttress
In summary, the Ottawan metamorphism: (i) during crustal shortening.
took place in the orogen interior from c. 1090– The architecture of the Grenville Orogen is a
1020 Ma; (ii) was associated with stable magnetite; result of thrust- and normal-sense tectonics. The dis-
(iii) involved a relatively high geothermal gradient tribution of metamorphic rocks in Figure 9 suggests
that may have been enhanced by advected magmatic it principally developed during thrusting, but several
heat; and (iv) was followed by slow cooling. In con- contractional boundaries, including the Allochthon
trast, Rigolet metamorphism (v) took place at the Boundary Thrust, locally carry rocks with a lower
orogen margin from c. 1000–980 Ma; (vi) resulted grade of Grenvillian metamorphism in their hang-
in assemblages in which magnetite was not stable ing wall than footwall, consistent with normal-sense
(at least in the central Grenville Province); (vii) reworking. On the basis of available data, ductile
was characterized by a lower geothermal gradient normal-sense shearing in the interior Grenville
and the absence of advected magmatic heat (it is Province began as early as c. 1060 Ma, became
situated outside the Interior Magmatic Belt); and widespread in the late Ottawan (c. 1020 Ma) and
(viii) was followed by fast cooling. Ottawan and continued in narrow shear zones until late- to post-
Rigolet metamorphisms thus took place at different Rigolet times (Cosca et al. 1991; Corrigan & van
times, in geographically separate parts of the Gren- Breemen 1997; Ketchum et al. 1998). With regard
ville Orogen, were of different P–T character, to the northwestern margin of the province, there
developed under different oxygen fugacities and is no evidence the Grenville Front was reactivated
geothermal gradients, and followed different as a normal-sense structure, but the shear zone that
cooling paths. offsets the Moho beneath the Parautochthonous
The estimated peak temperatures in the Ottawan Belt in western Québec (Fig. 5c) implies crustal-
mid-crust are pertinent to the discussion of channel scale extension of post-Rigolet age. The signifi-
flow that follows. The low viscosities necessary cance of this is discussed later.
for channel flow require extensive dehydration
melting of micas in felsic lithologies (Rosenberg
& Handy 2005), a process initiated at temperatures Comparison with thermal –mechanical
of 750 –8008C. Figure 8b shows that temperatures models for large hot orogens
in this range or greater occurred throughout the
aMP segment, compatible with field evidence Jamieson et al. (2007) compared the crustal-scale
for abundant leucosome in felsic lithologies. For structure of the western Grenville Province with
instance, Slagstad et al. (2004) documented that simulated in numerical experiments in which
Ottawan T estimates of 750 –8508C and widespread the kinematic and intrinsic properties of the scale
field evidence for partial melting in the Muskoka models were specified and the thermal–mechanical
domain, the map pattern of which resembles a response of the crust was calculated. Their exper-
ductile nappe (Fig. 7a), leading them to suggest it iments were continued for over 100 Myemt
had undergone viscous flow. Since P –T conditions (elapsed model time) to simulate a LHO, and
in Muskoka domain were not atypical of the MP weak lower-crustal blocks were incorporated close
segment (Fig. 8b) and large-scale nappes also to the suture zone as a proxy for accreted terranes
formed elsewhere at this time (e.g. Adirondack at the continental margin. The experiments exhib-
Highlands terrane), it is likely that large segments ited plateau development in the hinterland after
of the Ottawan mid-crust were rheologically weak c. 40 Myemt due to decreased viscosity of the hot
and capable of flowing when subjected to a small mid-crust and initiation of heterogeneous channel
differential stress. flow shortly thereafter. The first-order crustal
Fig. 9. Crustal-scale cross-sections of the Grenville Orogen based on Lithoprobe deep seismic imaging (A, B, C, D, G),
COCORP deep seismic imaging (H), and inferred from surface geological data (E, F). Cross-sections are coloured
according to the inferred age of metamorphism (modified after Rivers et al. 2002).
structure and P –T distribution of the Georgian Bay may be analogous to ductile shear zones in nature.
and western Québec sections were closely simulated Thus, with regard to the two structural boundaries
after 97.5 Myemt and 82.5 Myemt respectively, that are the focus of this contribution, Jamieson
leading Jamieson et al. (2007) to conclude that the et al. (2007) equated the base of the mid-crustal
interior of the western Grenville Province was the channel in the experiments to the Allochthon Bound-
site of channel flow under an orogenic plateau, ary Thrust and a high-strain zone near the orogen
that is the Grenville Province is part of a LHO. margin to the Grenville Front. With respect to the
Although discontinuities cannot be simulated former boundary, examination of their results
with the finite-element method used in the model- shows that lower crustal rocks from the pro-side of
ling, zones of very high strain in the experiments the model orogen underwent tectonic transport of
432 T. RIVERS
several hundred km towards the retro-side before high-pressure thrust sheets of Rigolet age
partial exhumation into a mid-crustal channel, beneath the Allochthon Boundary Thrust; and
return flow and final exhumation at the pro-margin, (4) The contrasting rheological behaviour of the
which is compatible with the inference of major tec- western and central segments of the aHP
tonic transport on the Allochthon Boundary Thrust Belt is not explained.
(Fig. 7). With respect to the high-strain zone that These points are discussed in the next section.
may simulate the Grenville Front, it exhumes mid-
crustal rocks from 20–30 km depth compatible Ottawan orogenic collapse and formation
with the GFTZ in the western Grenville Province. of the Orogenic Lid
The effect of the variable lower crustal rheology
is well illustrated by these experiments, prompting a As noted above, normal-sense displacements on
search for specific analogues in the Grenville Pro- Ottawan shear zones began locally at c. 1060 Ma,
vince. For instance, weak lower crustal blocks that were widespread by 1020 Ma, and sole in the mid-
enter the model orogen early detach from the subja- crust. Assuming the high T (8008C) and abundant
cent mantle lithosphere, become exhumed to the leucosome of the aMP segment signal the Ottawan
mid-crust and are transported in a channel towards mid-crust was ductile, and considering numerical
the erosion front; such a scenario may be analogous simulations of LHOs ‘inevitably’ result in plateau
to the evolution of the aHP Algonquin–Lac formation, the normal-sense displacements are inter-
Dumoine terrane in the western Grenville Province. preted to imply gravitational collapse of the plateau
In contrast, strong lower crustal blocks that enter the on a mid-crustal detachment system. A direct con-
model orogen late remain attached to the subjacent sequence of collapse was juxtaposition of the upper
mantle lithosphere, driving reverse flow of the mid- and mid-crust (Orogenic Lid and aM-LP Belt). In
crust above it, providing a possible analogue for eastern Labrador, regional folding of the ductile
Archaean crust in the footwall of crustal-scale MP footwall of the detachment at c. 1050 Ma
ramps in the northern Grenville Province. Thus, as between the more rigid Orogenic Lid and con-
argued by Jamieson et al. (2007), these experiments temporaneous foreland implies compression con-
offer important insight into the causes and modes of tinued after normal faulting, suggesting collapse
flow and their role on the first-order architecture and took place in an overall contractional setting due
thermal structure of LHOs in general, and the Gren- to failure of melt-weakened mid-crust (fixed bound-
ville Orogen in particular. ary collapse; Rey et al. 2001). Moreover from an
The data assembled in this paper provide support orogen-scale perspective, the widespread reworking
for the interpretation of Jamieson et al. (2007) that of the Allochthon Boundary Thrust as part of a
the first-order architecture of the interior Grenville system of detachment zones at c. 1020 Ma appears
Orogen is a result of some form of channel flow to have halted channel flow, paving the way for a
beneath an orogenic plateau following prolonged new flow path during the Rigolet phase. Since there
crustal thickening. However, several features of is no evidence for reactivation of the Allochthon
the Grenville Front and the Allochthon Boundary Boundary Thrust after 1020 Ma, it is inferred that
Thrust discussed above are not readily incorporated orogenic collapse terminated the channel flow
into their model and point to the need for some regime that had existed since early Ottawan times,
second-order modifications. These include the thereby promoting a fundamental change in the
following: architectural and thermal structure of the orogen.
Preservation of the Orogenic Lid above the Alloch-
(1) The present architecture of the Grenville thon Boundary Thrust is a result of this change, its
Province is a result of both thrust- and normal- lack of penetrative Ottawan deformation suggesting
sense structures, but only thrust-sense it formed part of the orogenic suprastructure.
structures are simulated in their numerical
experiments. This omission is particularly Significance of hiatus between Ottawan and
important with respect to Allochthon Boun- Rigolet orogenic phases
dary Thrust and the formation of the
Orogenic Lid; The paucity of metamorphic ages between c. 1020
(2) Although there is a progressive younging of the and 1000 Ma points to a hiatus in the tectonic devel-
age of deformation and metamorphism tow- opment of the Grenville Orogen. Moreover, when
ards the foreland in the model orogen, there is tectonism resumed it was focused in the Parau-
no apparent division into two orogenic phases tochthonous Belt, the former orogenic foreland.
separated by a hiatus in their experiments; As noted, the significance of this hiatus has engen-
(3) The contrasting metamorphic field gradients dered discussion and the interpretation proposed
of the Ottawan and Rigolet orogenic here is in the context of the LHO paradigm. Given
phases are not explained, nor the presence of the role of the Allochthon Boundary Thrust as the
base of a mid-crustal channel for 70 My (from High-pressure thrust sheets in the

c. 1090–1020 Ma), its termination implies a funda- Parautochthonous Belt
mental change in the orogen. Specifically, it is
inferred that juxtaposition of upper- and mid-crustal In the model of Jamieson et al. (2007), the Rigolet
rocks during orogenic collapse cooled the mid- orogenic phase is associated with the detachment
crust and caused widespread freezing of leucosomes and exhumation of HP nappes above a strong lower-
rendering re-activation of the channel mechanically crustal indentor after 90– 100 Myemt. Exhumation
unfavourable. As a result, ongoing shortening led to occurs well inboard from the contemporary orogen
initiation of a new shear zone system and eventually margin (i.e. they would be aHP nappes in the termi-
the formation of the Grenville Front and related nology used in this paper) and they form as the
structures in the Parautochthonous Belt. In this channel propagates outwards into the contemporary
context, the c. 20 My hiatus may represent the foreland. Moreover the Grenville Front is the site of
time taken for migration of the locus of high-grade exhumed mid-crustal rocks; no HP rocks occur (i.e.
metamorphism from the hinterland to the foreland, pHP in the terminology of this paper). Thus, when
development of a crustal-scale critical wedge, and compared with the Grenville Province, their model
conductive heating of the thickened crust. raises two issues: (i) the timing does not fit that in
the Grenville Province, where aHP terranes in the
Significance of contrasting Ottawan and hanging wall of the Allochthon Boundary Thrust
Rigolet metamorphic field gradients were exhumed in mid- to late-Ottawan times, well
before the Rigolet stage; (ii) there is no mechanism
Evidence for the contrasting Ottawan and Rigolet in the experiments to cause the exhumation of pHP
metamorphic field gradients is summarized in terranes with Rigolet metamorphism. With respect
Figure 8. Concerning the relatively high Ottawan to the first issue, earlier exhumation of aHP terranes
gradient, the mid-crust in a LHO is hot principally could probably be achieved in the experiments by
because the low erosion rate on the plateau results inserting strong lower crustal blocks closer to the
in a protracted period of conductive heating, but orogenic suture, so this may not require fundamental
factors such as radioactive self-heating, mantle modification to their model. However, the second
heat flow and advection of magmatic heat also issue constitutes a more significant problem and it
contribute. In this case, conductive heating during remains unclear to this author whether a single-stage
the 70 My duration of the Ottawan orogenic phase model, in which the mid-crustal channel remains
was clearly critical, but radioactive heating may operational throughout contractional orogenesis,
have also been significant considering the hanging can explain the exhumation of pHP terranes near
wall of the Allochthon Boundary Thrust is largely the orogen margin late in the collisional process.
composed of mid- to late-Proterozoic (i.e. relatively
young) granitoid rocks. Moreover, the presence of Contrasting rheological behaviour of the
mantle- and crustal-derived intrusions of Ottawan western and central aHP belts
age in the Interior Magmatic Belt (Fig. 1) implies
the contemporary mantle heat flow and advected The contrasting rheological expression of the
magmatic heat would have been high at least aHP terranes in the western and central Grenville
locally and augmented the regional gradient. Province has been noted. In the central Grenville Pro-
With respect to the cooler Rigolet gradient, the vince, mafic lithologies in the Manicouagan Imbri-
incorporation of cold crust from the former foreland cate Zone form coherent units and their HP mineral
and the limited time for conductive heating were assemblages are well-preserved despite ductile
probably crucial factors. Peak temperatures at mid- deformation and locally intense shearing, whereas
crustal depths, although appropriate for partial in the western Grenville Province, mafic units in
melting, were too short-lived to initiate channel Algonquin-Lac Dumoine occur as disaggregated
flow. Ancillary factors include the inferred low self- pods and lenses, the terrane as a whole has undergone
heating potential of the Archaean crust, which pervasive ductile flow, and the HP minerals are over-
underlies a large part of the Parautochthonous printed by granulite-facies assemblages.
Belt, and the location of the latter outside the Since exhumation of HP rocks implies passage
Interior Magmatic Belt implying mantle heat flow through the mid-crust, its temperature and the dur-
and advected magmatic heat would have been low. ation of mid-crustal residence may be critical
The negligible role of advected magmatic heat on factors controlling preservation. In this case, T esti-
the Rigolet gradient is highlighted by the presence mates for the MP segments in the two areas are
of many small intrusions of late- to post-Rigolet similar (Fig. 8), whereas mid-crustal residence
age (c. 980 –950 Ma) in the Interior Magmatic time was inferred to have been short for the Mani-
Belt (e.g. Gower et al. 1991), outside the area of couagan Imbricate Zone, but prolonged for the
penetrative Rigolet metamorphism. Algonquin–Lac Dumoine terrane (Davidson 1990;
434 T. RIVERS
Indares et al. 2000). The heterogeneous flow model rapid exhumation of parautochthonous lower- and
of Jamieson et al. (2007) provides additional context mid-crustal rocks in the hanging wall of the Gren-
for this contrast. In the central Grenville Province, ville Front. Metamorphism and deformation at this
Archaean crust penetrates over 100 km into the time were focused between the Grenville Front
orogen at the surface and over twice that distance and Allochthon Boundary Detachment, with the
at depth (Fig. 1). Assuming it behaved in a similar hinterland, although still hot, being essentially inac-
manner to a strong indentor in the experiments of tive except for local small-volume plutonism, minor
Jamieson et al. (2007), it would have promoted normal-sense adjustments in the upper crust, and
rapid exhumation of weaker, inboard lower crustal local growth of metamorphic rims on zircon and
blocks by reverse flow, as hot nappes above it. On monazite. There was thus an evolution over an inter-
the other hand, the absence of Archaean rocks, and val of c. 20 My from a large hot long-duration
hence a strong indentor, in the western Grenville orogen in the Ottawan phase to a subjacent, small
Province would have suppressed exhumation, but cold short-duration orogen in the Rigolet phase.
promoted heterogeneous flow in a mid-crustal Schematic particle and P–T –t paths are shown in
channel, compatible with the granulite-facies over- Figure 11. The particle paths (Fig. 11a) illustrate
print and ductile disaggregation of the mafic units. transport of crust from the hinterland towards the
Hence it is suggested that the contrasting evolutions foreland in Ottawan times by viscous flow in a mid-
of aHP terranes do not relate to intrinsic differences crustal channel and as hot lower-crustal nappes.
between them, but rather to the rheological charac- Lithologic linkages between the Orogenic Lid and
ter of the crust that entered the orogen behind them. adjacent mid-crustal terranes suggest independent
transport of the upper crust was limited. After oro-
genic collapse terminated channel flow, a shorter
Refinements to the heterogeneous channel particle path developed during the Rigolet phase
flow model for the Grenville Orogen in the former foreland, bypassing the hinterland
beneath the Allochthon Boundary Detachment.
In light of the above, several second-order refine- This model, in addition to accommodating the differ-
ments to the heterogeneous channel flow model of ent times, P–T gradients and particle paths of the
Jamieson et al. (2007) are proposed. As discussed, Ottawan and Rigolet metamorphisms, is compatible
the preferred model involves two stages of thrusting with evidence for greater horizontal displacement on
and crustal thickening in different parts of the the Allochthon Boundary Thrust than the Grenville
orogen separated by an important episode of exten- Front. Moreover, the change in particle paths not
sion and crustal thinning, all of which took place in only signalled destabilization of Ottawan channel
an overall compressional orogen. flow, but is also compatible with migration of the
The model is illustrated schematically in metamorphic core of the orogen into its former fore-
Figure 10. The Ottawan stage (Fig. 10a) involved land. In numerical experiments, such ‘tunnelling’ of
growth of an orogenic plateau with channel flow the mid-crust towards the foreland is enhanced when
of low-viscosity mid-crust and extrusion of hot there is no erosion at the orogen margin (thereby sup-
lower-crustal nappes, both of which were exhu- pressing exhumation; Beaumont et al. 2006)
med at the orogenic front. This stage closely res- suggesting this may have been a time of relatively
embles the model of Jamieson et al. (2007) with subdued topography in the Grenville Orogen.
the caveats that criteria to distinguish between There are few published P –T–t paths for the
heterogeneous channel flow and homogeneous Grenville Province, so the paths in Figure 11b
channel flow with coeval extrusion of hot nappes were constrained by: (i) P–T estimates of peak
are not presently available, and that there is evi- metamorphic conditions; (ii) U – Pb zircon and mon-
dence for local normal-sense shear zones in the oro- azite determinations of the time of peak metamorph-
genic hinterland at this time. The first significant ism; (iii) compatibility with the Al-silicate phase
deviation from the model of Jamieson et al. (2007) diagram; and (iv) the 40Ar/39Ar hornblende
occurred at c. 1020 Ma (Fig. 10b), when the Alloch- cooling ages (data in Tables 1–3). The Ottawan
thon Boundary Thrust was reworked as part of the and Rigolet paths in Figure 11b are distinctive, sup-
normal-sense Allochthon Boundary Detachment porting the inference they developed under different
system of shear zones. It is inferred that this resulted geothermal gradients. The 40Ar/39Ar hornblende
in cooling the channel, increasing its viscosity cooling ages, representing the time of cooling
above the critical value for flow thereby terminating through c. 5008C, are mostly in the range 980–
channel flow despite ongoing subduction of Lauren- 940 Ma throughout the northern Grenville Province,
tian sub-continental lithospheric mantle. In the implying slow cooling in the allochthonous belts
Rigolet stage (Fig. 10c), the orogen adjusted to after the Ottawan orogenic phase, but much faster
loss of the material flow-path through the hinterland cooling in the Parautochthonous Belt after the
by advancing into its foreland, leading to burial and Rigolet orogenic phase. Slow cooling following
Fig. 10. Schematic diagram illustrating inferred distribution of metamorphic facies and tectonic evolution of the
Grenville Province. (a) early Ottawan, formation of orogenic plateau with some form of channel flow and/or extrusion
of hot nappes on Allochthon Boundary Thrust (ABT), local orogenic collapse. (b) late Ottawan, widespread orogenic
collapse, reworking of ABT as Allochthon Boundary detachment (ABD), cessation of channel flow. (c) Rigolet, orogen
migrates into foreland, formation of Grenville Front (GF), hinterland becomes tectonically inactive. See text for an
additional explanation.
436 T. RIVERS
Fig. 11. Schematic particle paths and P –T –t paths for the Grenville Orogen (adapted from a figure in Jamieson et al.
2004). (a) Solid red arrows labelled A –C represent particle paths for the aHP, aMP, aLP segments, dashed red arrow
labelled D represents the path for the pMP and pHP segments. Path A probably involved extrusion of a lower-crustal
nappe whereas paths B– C may have involved some form of mid-crustal channel flow. Path D took place beneath the
former channel and involved rapid burial and thrust exhumation. ABT, ABD and GF are locations of Allochthon
Boundary Thrust, Allochthon Boundary Detachment and Grenville Front. (b) P– T –t diagrams are based on peak P– T
estimates, U –Pb zircon or monazite ages of peak metamorphism, Al-silicate phase present, and 40Ar/39Ar hornblende
cooling ages (Tc 5008C; Tables 1– 3).
the Ottawan phase took place beneath the residual model proposed here, is based on field evidence
orogenic plateau, whereas fast cooling in the Parau- and empirical deduction, but was not predicted by
tochthonous Belt is compatible with a short particle published numerical simulations of LHOs. In
path that bypassed the warm hinterland (Fig. 11a). effect, it is inferred that widespread fixed boundary
collapse in a convergent setting in late Ottawan
Discussion times paved the way for a new tectonic style domi-
nated by thrusting rather than flow in the Rigolet
In the introduction it was proposed that evolution of orogenic phase. Although normal-sense structures
a collisional orogen could be conceptually divided have been identified in the field, the gravitational
into constructive and destructive stages separated collapse aspect of orogenic evolution remains
by a climax representing the time of maximum understudied given its significance in the proposed
size. This study suggests the Grenville Orogen model. Another problematic issue concerns the
underwent a more nuanced evolution with interrup- mechanism by which the Allochthon Boundary
tion of the long-duration constructive stage by oro- Thrust was exhumed at the orogenic front during
genic collapse due to mid-crustal weakening, the Ottawan orogenic phase. Numerical models of
followed by internal reorganization and renewed channel flow suggest that elevated rates of erosion
growth near the orogen margin before the orogen are required to exhume the channel, such as result
as a whole entered its destructive stage. Imaging from the monsoons in the Himalaya –Tibet Orogen
this evolution through the Grenville Front and (e.g. Beaumont et al. 2004; Jamieson et al. 2004),
Allochthon Boundary Thrust and their immediate but it may be problematic to extend a rainfall-driven
hanging walls has highlighted several critical fea- erosion model to a mountain chain located in the
tures. In particular, the evidence for long-distance middle of a supercontinent, as implied by the
transport of mid and lower crust on the Allochthon Rodinia reconstruction of Li et al. (2008).
Boundary Thrust for 50 My during the Ottawan Turning to the decay phase of orogenesis, the
orogenic phase by some form of channel flow presence of crust of average thickness (c. 35 km)
and/or as hot nappes suggests a mass balance throughout the present-day hinterland of the Gren-
between the material added to the orogen as a ville Province implies isostatic equilibrium was
result of collision and that removed by erosion at eventually attained as the mantle lithosphere under-
the orogen margin was established during this went visco-elastic rebound to a gravitationally
period. Destabilization of this mass balance at the stable configuration (Fig. 3). However, the Parau-
end of the Ottawan phase, probably due to rheologi- tochthonous Belt near the orogen margin is under-
cal weakening of the mid-crust, led to orogenic col- lain by anomalously thick (c. 50 km) crust in the
lapse and initiated a fundamental change in east and anomalously thin (c. 30 km) crust in the
structural style as crustal thickening migrated into west, indicating equilibrium was not established at
the former foreland. New and shorter particle these locations. In the case of thick crust under the
paths were established in a subjacent small cold Grenville Front Gravity Low, the lack of isostatic
orogen, resulting in formation of the Grenville compensation has been attributed to suppression of
Front and related structures in the Parautochthonous flexural rebound due to cool mantle lithosphere
Belt, and the hinterland essentially became tectoni- (Hynes 1994), although further study is needed to
cally inactive. Thus, despite their superficially determine whether a correlation exists between
similar appearances in the field and on seismic sec- crustal thickness and exposed structural level at
tions, the tectonic signatures of the Grenville Front the Grenville Front. On the other hand, the presence
and the Allochthon Boundary Thrust are fundamen- of thin crust under the Parautochthonous Belt in the
tally different in several important respects and western Grenville Province suggests the mantle
together they encapsulate much of the first-order lithosphere was sufficiently hot to undergo flexural
tectonic evolution of the Grenville Orogen. rebound there. In this case, a causal relationship
Many aspects of this model require additional with the crustal-scale Dorval Detachment in cross-
documentation and testing. The difficulty in assem- section C discussed previously and the Boundary
bling unambiguous evidence for the former exist- Shear in cross-section A (the two comprising the
ence of an orogenic plateau and channel flow in Mid-Parautochthon Detachment; Fig. 9; Rivers
ancient orogens was addressed by Godin et al. et al. 2002) is likely, specifically that rebound led
(2006a). In this case it is argued that compatible fea- to extension and crustal thinning. In summary, the
tures include the long duration of high temperature regionally compensated gravity signature beneath
metamorphism in the orogenic interior and preser- the hinterland where the mantle lithosphere was
vation of the rheologically strong orogenic supra- hot and cooling was slow, and the non-uniform
structure in crustal-scale graben. The inference gravitational compensation beneath the Parau-
that orogenic collapse triggered a fundamental tochthonous Belt where the temperature of the
change in crustal flow paths, a central tenet of the mantle lithosphere was variable and cooling was
438 T. RIVERS
rapid provides another contrast between the Although the duration of both metamorphisms was
Ottawan and Rigolet orogenic phases. shorter than in the Grenville Orogen and there was
Finally, it is appropriate to briefly compare the no inference they were separated by a period of oro-
proposed tectonic model for the Grenville Orogen genic collapse, the parallels are nonetheless striking
with that for the Himalaya –Tibet Orogen, the and suggest that successive metamorphic regimes
most intensely investigated LHO. In addition to evi- with contrasting P– T–t paths in the orogenic hinter-
dence for the long duration of orthogonal collision land and foreland may be a fundamental feature of
and the striking similarities between crustal-scale LHOs.
structures, such as the Allochthon Boundary
Thrust and High Himal Thrust, and the Grenville
Front and Main Boundary Thrust, both orogens
grew by advancing into their respective forelands. Conclusions
In the case of the Himalaya– Tibet Orogen the tec- The aim of this paper has been to show that an exam-
tonic evolution has been interpreted in terms of criti- ination of orogenic fronts can provide unique insight
cal wedge theory, with plateau formation and into the evolution of the orogen as a whole. In the
channel flow taking place in a stable wedge with a case of the Grenville Province, an archetypical can-
critical angle of taper (wc), extension on the South didate for a large hot long-duration orogen, the fronts
Tibetan Detachment leading to the wedge becoming associated with the Ottawan and Rigolet orogenic
sub-critical (w , wc), and renewed thickening of the phases document two contrasting tectonic styles
hinterland by thrusting and crustal-scale folding that took place at different times and in different
leading to re-establishment of critical taper places, pointing to an evolution from a large hot
(w ¼ wc), which in turn permitted advance of the long-duration orogen with some form of channel
wedge into its foreland (Hodges 1996, 2006; flow under a plateau during the Ottawan phase to a
Godin et al. 2006b). However, in the case of the subjacent small cold short-duration orogen during
Grenville Province, the apparent lack of evidence the Rigolet phase. The two orogenic phases are
for penetrative deformation in the hinterland after components of a prolonged collision that lasted
orogenic collapse (Fig. 10c) suggests a comparable c. 110 My and are separated by a period of mid-
model may not be viable. Much depends on the tec- crustal collapse that terminated channel flow and
tonic significance of the sparse zircon and monazite caused the orogen to migrate into its former foreland.
rim ages in the aM–LP Belt. If the hinterland did This evolution is qualitatively compatible with
not undergo renewed thickening after 1020 Ma, observed differences in P–T –t paths and cooling
but essentially became tectonically inactive as histories between the allochthonous and parau-
suggested by the available evidence and indicated tochthonous belts, and with the inference that the
in Figure 10c, the advance into the foreland at c. Allochthon Boundary Thrust formed the base of a
1000 Ma must have had a different driving force. mid-crustal channel in the orogenic hinterland
One possible explanation is that the basal friction whereas the Grenville Front represents the limit of
and internal strength of the lower and mid crust in orogenic crust derived from the immediate foreland
the orogenic wedge increased significantly due to and exhumed by thrusting. Although exhibiting first-
freezing of the abundant leucosome and subsequent order compatibility with a published heterogeneous
escape of magmatic fluids, rendering the wedge channel flow model for the western Grenville Pro-
stronger than the adjacent hot foreland. Such an vince (Jamieson et al. 2007), the model proposed
interpretation would be compatible with establish- in this paper incorporates several important modifi-
ment of the inferred short particle paths during the cations based on diverse second-order geological
Rigolet orogenic phase (Fig. 11). constraints from the length of the Grenville
The metamorphic evolution provides another Province.
relevant point of comparison. Goscombe et al. (2006)
showed that metamorphism in the Himalaya–Tibet The field data on which this paper is based were produced
Orogen can be subdivided into two temporally and by a generation of students of the Grenville Orogen. In
geographically distinct regimes, an older MP –HT addition, there is an obvious intellectual debt to the
regime in the orogenic hinterland and a younger finite-element modelling of C. Beaumont, R. Jamieson
HP–MT regime nearer the foreland. As in the Gren- and colleagues and the unique insight it provides into the
ville Province, the older, higher T regime in the hin- physical basis of orogenesis. This is acknowledged with
terland structurally overlies the younger, lower T gratitude, but in no way implies they agree with everything
in this paper. I thank L. Godin and A. Hynes for construc-
regime nearer the foreland along a crustal-scale tive reviews, R. Law for pointing out the similar two-stage
thrust (the High Himal Thrust), and Goscombe metamorphic evolution of the Himalaya–Tibet Orogen,
et al. (2006) argued that the contrasting P –T charac- and B. Murphy, D. Keppie and A. Hynes for their editorial
ter of the two metamorphic regimes resulted from handiwork. My research has benefitted from the Litho-
development under different geothermal gradients. probe program and many years of funding by NSERC.
References B USSY , F., K ROGH , T. E., K LEMENS , W. P. &

S CHWERDTNER , W. M. 1995. Tectonic and meta-
A LCOCK , J., I SACHSEN , C., L IVI , K. & M ULLER , P. morphic events in the westernmost Grenville Province,
2004. Unraveling growth history of zircon in anatec- central Ontario: new results from high-precision geo-
tites from the northeast Adirondack Highlands, chronology. Canadian Journal of Earth Sciences, 32,
New York: constraints on pressure-temperature-time 660– 671.
paths. In: T OLLO , R. P., C ORRIVEAU , L., C ARR , S. D., E ASTON , R. M., J AMIESON , R. A. &
M C L ELLAND , J. & B ARTHOLOMEW , M. J. (eds) C ULSHAW , N. G. 2000. Geologic transect across the
Proterozoic Tectonic Evolution of the Grenville Grenville orogen of Ontario and New York. Canadian
Orogen in North America. Geological Society of Journal of Earth Sciences, 37, 193– 216.
America Memoir, 197, 267– 284. C HILDE , F., D OIG , R. & G ARIÉPY , C. 1993. Monazite as a
B ANDYAYERA , D., C ADÉRON , S., R OY , P. & R IVERS , T. metamorphic chronometer, south of the Grenville
2006. Archaean volcano-sedimentary formations in the Front, western Quebec. Canadian Journal of Earth
Chibougamau and Caopatina segments and their exten- Sciences, 30, 1056–1065.
sions in the Grenvillian Parautochtonous Belt, Québec, C LARK , M. K. & R OYDEN , L. H. 2000. Topographic ooze:
Canada. Geological Association of Canada – Minera- building the eastern margin of Tibet by lower crustal
logical Association of Canada, Annual Meeting Mon- flow. Geology, 28, 703–706.
tréal, Abstract volume, 9 –10. C ONNELLY , J. N. & H EAMAN , L. M. 1993. U–Pb geo-
B EAUMONT , C., J AMIESON , R. A., N GUYEN , M. H. & chronological constraints on the tectonic evolution of
L EE , B. 2001a. Himalayan tectonics explained by the Grenville Province, western Labrador. Precam-
extrusion of a low-viscosity crustal channel coupled brian Research, 63, 123–142.
to focused surface denudation. Nature, 414, 738– 742. C ONNELLY , J. N., R IVERS , T. & J AMES , D. T. 1995. Ther-
B EAUMONT , C., J AMIESON , R. A., N GUYEN , M. H. & motectonic evolution of the Grenville Province of
L EE , B. 2001b. Mid-crustal channel flow in large hot western Labrador. Tectonics, 14, 202– 217.
orogens: results from coupled thermal-mechanical C ORFU , F. & E ASTON , R. M. 1995. U– Pb geochronology
models. In: Slave – Northern Cordillera Lithospheric of the Mazinaw terrane, an imbricate segment of the
Evolution (SNORCLE) and Cordilleran Tectonics Central Metasedimentary Belt, Grenville Province,
Workshop, Lithoprobe Secretariat, UBC, Vancouver, Ontario. Canadian Journal of Earth Sciences, 32,
BC, Lithoprobe Report, 79, 112–170. 959– 976.
B EAUMONT , C., J AMIESON , R. A., N GUYEN , M. H. & C ORRIGAN , D. & VAN B REEMEN , O. 1997. U –Pb age
M EDVEDEV , S. 2004. Crustal channel flows: constraints for the lithotectonic evolution of the Gren-
1. Numerical models with applications to the tectonics ville Province along the Mauricie transect, Québec.
of the Himalayan-Tibetan orogen. Journal of Geophy- Canadian Journal of Earth Sciences, 34, 299–316.
sical Research, 109, B06406; doi: 10.1029/ C ORRIGAN , D., C ULSHAW , N. G. & M ORTENSEN , J. K.
2003JB002809. 1994. Pre-Grenvillian evolution and Grenvillian
B EAUMONT , C., N GUYEN , M. H., J AMIESON , R. A. & overprinting of the Parautochthonous Belt in the Key
E LLIS , S. 2006. Crustal flow modes in large hot Harbour area, Ontario: U–Pb and field constraints.
orogens. In: L AW , R. D., S EARLE , M. P. & G ODIN , Canadian Journal of Earth Sciences, 31, 583–596.
L. (eds) Channel Flow, Ductile Extrusion and Exhuma- C ORRIGAN , D., R IVERS , T. & D UNNING , G. 2000. U–Pb
tion in Continental Collision Zones. Geological constraints for the plutonic and tectonometamorphic
Society, London, Special Publications, 268, 91–145. evolution of Lake Melville terrane, Labrador and
B ETHUNE , K. M. 1997. The Sudbury dyke swarm and its implications for basement reworking in the northeast-
bearing on the tectonic development of the Grenville ern Grenville Province. Precambrian Research, 99,
Front, Ontario, Canada. Precambrian Research, 85, 65–90.
117–146. C ORRIVEAU , L. & B ONNET , A.-L. 2005. Pinwarian
B ETHUNE , K. M. & D AVIDSON , A. 1997. Grenvillian (1.50 Ga) volcanism and hydrothermal activity at the
metamorphism of the Sudbury diabase dyke-swarm: eastern margin of the Wakeham Group, Grenville
from protolith to two-pyroxene– garnet coronite. Province, Quebec. Canadian Journal of Earth
Canadian Mineralogist, 35, 1191– 1220. Sciences, 42, 1749–1782.
B ICKFORD , M. E., M C L ELLAND , J. M., S ELLECK , B. W., C OSCA , M. A., S UTTER , J. F. & E SSENE , E. J. 1991.
H ILL , B. M. & H EUMANN , M. J. 2008. Timing of ana- Cooling and inferred uplift/erosion history of the
texis in the eastern Adirondack Highlands: impli- Grenville orogen, Ontario: constraints from
40
cations for tectonic evolution during c. 1050 Ma Ar/39Ar thermochronology. Tectonics, 10, 959–977.
Ottawan orogenesis. GSA Bulletin, 120, 950– 961, C OX , R. & I NDARES , A. 1999a. High pressure and temp-
doi: 10.1130/B26309.1. erature metamorphism of the mafic and ultramafic Lac
B OHLEN , S. R. 1987. Pressure, temperature, time paths, Espadon suite, Manicouagan Imbricate Zone, eastern
and a tectonic model for the evolution of granulites. Grenville Province. Canadian Mineralogist, 37,
Journal of Geology, 95, 617– 632. 335– 357.
B ONNET , A.-L., C ORRIVEAU , L. & L A F LÈCHE , M. R. C OX , R. & I NDARES , A. 1999b. Transformation of Fe-Ti
2005. Chemical imprint of highly metamorphosed gabbro to coronite, eclogite and amphibolite in the
volcanic-hosted hydrothermal alterations in the La Baie du Nord segment, eastern Grenville Province.
Romaine supracrustal belt, eastern Grenville Province, Journal of Metamorphic Geology, 17, 537– 555.
Quebec. Canadian Journal of Earth Sciences, 42, C OX , R. & R IVERS , T. in press. Reconnaissance laser abla-
1783–1814. tion (LA)-ICP-MS U–Pb zircon and monazite
440 T. RIVERS
geochronology in reworked Archaean basement and D UDÁS , F. Ö., D AVIDSON , A. & B ETHUNE , K. M.
Palaeoproterozoic supracrustal rocks, Gagnon 1994. Age of the Sudbury diabase dykes and their
terrane, Grenville Province of Québec and Labrador. metamorphism in the Grenville Province. In: Radio-
Precambrian Research. genic age and isotope studies: Report 8. Geological
C OX , R., I NDARES , A. & D UNNING , G. 1998. Petrology Survey of Canada, Current Research, 1994–F,
and U–Pb geochronology of mafic, high-P meta- 97–106.
morphic coronites from the Tshenukutish domain, G ARIÉPY , C., V ERNER , D. & D OIG , R. 1990. Dating
eastern Grenville Province. Precambrian Research, Archaean metamorphic minerals southeast of the Gren-
90, 59– 83. ville Front, western Quebec, using Pb isotopes.
C ULSHAW , N. G., R EYNOLDS , P. H. & C HECK , G. 1991. Geology, 18, 1078–1081.
A 40Ar/39Ar study of post-tectonic cooling in the Britt G OBEIL , A., B RISEBOIS , D., C LARK , T., V ERPAELST , P.,
domain of the Grenville Province, Ontario. Earth & M ADORE , L., W ODICKA , N. & C HEVÉ , S. 2003.
Planetary Science Letters, 105, 405–415. Géologie de la moyenne Côte-Nord. In: B RISEBOIS ,
C ULSHAW , N. G., K ETCHUM , J. W. F., W ODICKA , N. & D. & C LARK , T. (eds) Géologie et resources minerales
W ALLACE , P. 1994. Ductile extension following thrust- des la partie est de la Province de Grenville. Ministère
ing in the southwestern Grenville Province, Ontario. de Richesses naturelles, Québec, DV2002–2003,
Canadian Journal of Earth Sciences, 31, 160– 175. 59–117.
C ULSHAW , N. G., J AMIESON , R. A., K ETCHUM , J. W. F., G ODIN , L., G RUJIC , D., L AW , R. D. & S EARLE , M. P.
W ODICKA , N., C ORRIGAN , D. & R EYNOLDS , P. H. 2006a. Channel flow, ductile extrusion and exhuma-
1997. Transect across the northwestern Grenville tion in continental collision zones: an introduction.
orogen, Georgian Bay, Ontario: polystage convergence In: L AW , R. D., S EARLE , M. P. & G ODIN , L. (eds)
and extension in the lower orogenic crust. Tectonics, Channel Flow, Ductile Extrusion and Exhumation in
16, 966–982. Continental Collision Zones. Geological Society,
D AIGNEAULT , R. & A LLARD , G. O. 1994. Transform- London, Special Publications, 268, 1– 23.
ation of Archaean structural inheritance at the Grenvil- G ODIN , L., G LEESON , T. P., S EARLE , M. P., U LLRICH ,
lian Foreland – Parautochthon transition zone, T. D. & P ARRISH , R. R. 2006b. Locking of southward
Chibougamau Quebec. Canadian Journal of Earth extrusion in favour of rapid crustal-scale buckling
Sciences, 31, 470–488. of the Greater Himalayan sequence, Nar Valley,
D ALLMEYER , R. D. & R IVERS , T. 1983. Recognition of central Nepal. In: L AW , R. D., S EARLE , M. P. &
extraneous argon components through incremental- G ODIN , L. (eds) Channel Flow, Ductile Extrusion
release 40Ar/39Ar analysis of biotite and hornblende and Exhumation in Continental Collision Zones.
across the Grenvillian metamorphic gradient in south- Geological Society, London, Special Publications,
western Labrador. Geochimica et Cosmochimica Acta, 268, 269 –292.
47, 413–428. G OSCOMBE , B., G RAY , D. & H AND , M. 2006. Crustal
D ARLING , R. S., F LORENCE , F. P., L ESTER , G. W. & architecture of the Himalayan metamorphic front in
W HITNEY , P. R. 2004. Petrogenesis of prismatine- eastern Nepal. Gondwana Research, 10, 232– 255,
bearing metapelitic gneiss along the Moose River, doi: 10.1016/j.gr2006.05.003.
west-central Adirondacks, New York. In: T OLLO , G OWER , C. F. & K ROGH , T. E. 2002. A U –Pb geochrono-
R. P., C ORRIVEAU , L., M C L ELLAND , J. & logical review of the Proterozoic history of the eastern
B ARTHOLOMEW , M. J. (eds) Proterozoic tectonic Grenville Province. Canadian Journal of Earth
evolution of the Grenville orogen in North America. Sciences, 39, 795– 829.
Geological Society of America Memoir, 197, G OWER , C. F., R YAN , A. B., B AILEY , D. G. & T HOMAS ,
325– 336. A. 1980. The position of the Grenville Front in eastern
D AVIDSON , A. 1990. Evidence for eclogite metamorph- and central Labrador. Canadian Journal of Earth
ism in the southwestern Grenville Province. Geologi- Sciences, 17, 784– 788.
cal Survey of Canada, 90– 1C, 113– 118. G OWER , C. F., H EAMAN , L. M., L OVERIDGE , W. D.,
D AVIDSON , A. 1998. An overview of the Grenville Pro- S CHÄRER , U. & T UCKER , R. D. 1991. Grenvillian
vince geology, Canadian Shield. In: L UCAS , S. B. & granitoid plutonism in the eastern Grenville Province.
S T -O NGE , M. R. (eds) Geology of the Precambrian Precambrian Research, 51, 315–336.
Superior and Grenville Provinces and Precambrian G REEN , A. G., M ILKREIT , B. ET AL . 1988. Crustal struc-
Fossils in North America. Geological Survey of ture of the Grenville Front and adjacent terranes.
Canada, Geology of Canada. Geology, 16, 788 –792.
D AVIDSON , A. & VAN B REEMEN , O. 1988. Baddeleyite- G RUJIC , D., H OLLISTER , L. & P ARRISH , R. R. 2002.
zircon relationships in coronitic metagabbro, Grenville Himalayan metamorphic sequence as an orogenic
Province, Ontario: implications for geochronology. channel: insight from Bhutan. Earth and Planetary
Contributions to Mineralogy and Petrology, 100, Science Letters, 198, 177– 191.
291– 299. G RUJIC , D., C OUTLAND , I., B OOKHAGEN , B., B LYTHE ,
D EWEY , J. F. 1988. Extensional collapse of orogens. A. & D UNCAN , C. 2006. Climatic forcing of erosion,
Tectonics, 7, 1123– 1139. landscape, and tectonics in the Bhutan Himalayas.
D ICKIN , A. P. & G UO , A. 2001. The location of the Geology, 34, 801 –804.
Allochthon Boundary Thrust and the Archean– H AGGART , M. J., J AMIESON , R. A., R EYNOLDS , P. H.,
Proterozoic suture in the Mattawa area of the Grenville K ROGH , T. E., B EAUMONT , C. & C ULSHAW , N. G.
Province: Nd isotopic evidence. Precambrian 1993. Last gasp of the Grenville Orogeny: thermo-
Research, 107, 31–43. chronology of the Grenville Front Tectonic Zone
near Killarney, Ontario. Journal of Geology, 101, in the Manicouagan area: new structural, petrologic
575–589. and U– Pb age constraints. Precambrian Research,
H EAMAN , L. M., E RDMER , P. & O WEN , J. V. 2002. U– Pb 130, 199–228.
geochronologic constraints on the crustal evolution of I NDARES , A. & M ARTIGNOLE , J. 1989. The Grenville
the Long Range Inlier, Newfoundland. Canadian Front south of Val d’Or. Tectonophysics, 157, 221–239.
Journal of Earth Sciences, 39, 845 –865. I NDARES , A. & M ARTIGNOLE , J. 1990a. Metamorphic
H ERD , R. K., A CKERMAN , D., W INDLEY , B. F. & constraints on the evolution of the gneisses from the
R ONDOT , J. 1986. Sapphirine– garnet rocks, parautochthonous and allochthonous polycyclic belts,
St. Maurice area, Québec: petrology and implications Grenville Province, western Quebec. Canadian
for tectonics and metamorphism. In: M OORE , J. M., Journal of Earth Sciences, 27, 357– 370.
D AVIDSON , A. & B AER , A. J. (eds) The Grenville I NDARES , A. & M ARTIGNOLE , J. 1990b. Metamorphic
Province. Geological Association of Canada, Special constraints on the tectonic evolution of the allochtho-
Paper, 31, 241 –253. nous monocyclic belt, Grenville Province, western
H ODGES , K. V. 1998. The thermodynamics of Himalayan Quebec. Canadian Journal of Earth Sciences, 27,
orogenesis. In: T RELOAR , P. J. & O’B RIEN , P. J. 371– 386.
(eds) What Drives Metamorphism And Metamorphic I NDARES , A. & M ARTIGNOLE , J. 1993. Étude régionale
Reactions? Geological Society, London, Special du Supergroupe de Wakeham, Province de Grenville.
Publications, 138, 7 –22. Ministère de l’Énergie et des Resources, Québec,
H ODGES , K. V. 2006. A synthesis of the channel flow – MB; 91– 21, 1– 73.
extrusion hypothesis as developed for the Himalayan I NDARES , A. & R IVERS , T. 1995. Textures, metamorphic
– Tibetan orogenic system. In: L AW , R. D., S EARLE , reactions and thermobarometry of eclogitized meta-
M. P. & G ODIN , L. (eds) Channel Flow, Ductile Extru- gabbros: a Proterozoic example. European Journal of
sion and Exhumation in Continental Collision Zones, Mineralogy, 7, 43– 56.
Geological Society, London, Special Publications, I NDARES , A., D UNNING , G., C OX , R., G ALE , D. &
268, 71–90. C ONNELLY , J. 1998. High-pressure, high-temperature
H ODGES , K. V., P ARRISH , R. R. & S EARLE , M. P. 1996. rocks from the base of thick continental crust: geology
Tectonic evolution of the central Annapurna Range, and age constraints from the Manicouagan Imbricate
Nepalese Himalaya. Tectonics, 15, 1264–1291. Zone, eastern Grenville Province. Tectonics, 17,
H OFFMAN , P. F. 1989. Precambrian geology and tectonic 426– 440.
history of North America. In: B ALLY , A. W. & I NDARES , A., D UNNING , G. & C OX , R. 2000.
P ALMER , A. R. (eds) The Geology of North America Tectono-thermal evolution of deep crust in a Mesopro-
– an overview. Geological Society of America, terozoic continental collision setting: the Manicouagan
Decade of North American Geology, A, 447–512. example. Canadian Journal of Earth Sciences, 37,
H YNES , A. 1994. Gravity, flexure, and the deep 325– 340.
structure of the Grenville front, eastern Quebec and J AMIESON , R. A., C ULSHAW , N. G. & C ORRIGAN , D.
Labrador. Canadian Journal of Earth Sciences, 31, 1995. North-west propagation of the Grenville
1002–1011. orogen: Grenville structure and metamorphism near
H YNES , A., I NDARES , A., R IVERS , T. & G OBEIL , A. Key Harbour, Georgian Bay, Ontario, Canada.
2000. Lithoprobe line 55: integration of out-of-plane Journal of Metamorphic Geology, 13, 185– 207.
seismic results with surface structure, metamorphism, J AMIESON , R. A., B EAUMONT , C., N GUYEN , M. H. &
and geochronology, and the tectonic evolution of the L EE , B. 2002. Interaction of metamorphism, defor-
eastern Grenville Province. Canadian Journal of mation, and exhumation in large convergent orogens.
Earth Sciences, 37, 341 –358. Journal of Metamorphic Geology, 20, 1 –16.
I NDARES , A. 1995. Metamorphic interpretation of high- J AMIESON , R. A., B EAUMONT , C. & M EDVEDEV , S.
pressure –temperature metapelites with preserved 2004. Crustal channel flows: 2. Numerical models
growth zoning in garnet, eastern Grenville Province, with implications for metamorphism in the Himalayan-
Canadian Shield. Journal of Metamorphic Geology, Tibetan orogen. Journal of Geophysical Research,
13, 475–486. 109, B06407, doi: 10.1029/2003JB002811.
I NDARES , A. 1997. Garnet-clinopyroxenites and garnet- J AMIESON , R. A., B EAUMONT , C., N GUYEN , M. H. &
kyanite restites from the Manicouagan Imbricate C ULSHAW , N. G. 2007. Synconvergent ductile
Zone: a case of high-P– high-T metamorphism in the flow in variable-strength continental crust: numerical
Grenville Province. Canadian Mineralogist, 35, models with application to the western Grenville
1161–1171. Province. Tectonics, 26, TC5005, doi: 10.1029/
I NDARES , A. 2003. Metamorphic textures and P– T evol- 2006TC002036.
ution of high-P granulites from the Lelukuau terrane, J OHNSON , E. L., G OERGEN , E. T. & F RUCHEY , B. L.
NE Grenville Province. Journal of Metamorphic 2004. Right lateral oblique slip movements
Geology, 21, 35–48. followed by post-Ottawan (1050–1020) Ma oroge-
I NDARES , A. & D UNNING , G. 1997. Coronitic metagabbro nic collapse along the Carthage-Colton shear
and eclogite from the Grenville Province of western zone: data from the Dana Hill metagabbro body,
Quebec: interpretation of U–Pb geochronology and Adirondack Mountains, New York. In: T OLLO , R. P.,
metamorphism. Canadian Journal of Earth Sciences, C ORRIVEAU , L., M C L ELLAND , J. & B ARTHOLO-
34, 891–901. MEW , M. J. (eds) Proterozoic tectonic evolution of
I NDARES , A. & D UNNING , G. 2004. Crustal architecture the Grenville orogen in North America. Geological
above the high-pressure belt of the Grenville Province Society of America Memoir, 197, 357–378.
442 T. RIVERS
J ORDAN , S. L., I NDARES , A. & D UNNING , G. 2006. Allard (SNRC 12L/11). Ministère des Richesses natur-
Partial melting of metapelites in the Gagnon terrane elles, Québec, RG 96– 05, 30.
below the high-pressure belt in the Manicouagan area M ARTIGNOLE , J. & C ALVERT , A. J. 1996. Crustal-scale
(Grenville Province): pressure– temperature (P– T) shortening and extension across the Grenville Province
and U– Pb age constraints and implications. Canadian of western Quebec. Tectonics, 15, 376– 386.
Journal of Earth Sciences, 43, 1309–1329. M ARTIGNOLE , J. & F RIEDMAN , R. 1998. Geochronologi-
K ARLSTROM , K. E., Å HÄLL , K.-I., H ARLAN , S. S., cal constraints on the last stages of terrane assembly in
W ILLIAMS , M. L., M C L ELLAND , J. & G EISSMAN , the central part of the Grenville Province (Quebec).
J. W. 2001. Long-lived (1.8– 1.0 Ga) convergent Precambrian Research, 92, 145–164.
orogen in southern Laurentia, its extensions to M ARTIGNOLE , J. & M ARTELAT , J.-E. 2005. Proterozoic
Australia and Baltica, and implications for refining mafic dykes as monitors of HP granulite facies meta-
Rodinia. Precambrian Research, 111, 5 –30. morphism in the Grenville Front Tectonic Zone
K ELLETT , R. L., B ARNES , A. E. & R IVE , M. 1994. The (western Quebec). Precambrian Research, 138,
deep structure of the Grenville Front: a new perspec- 183–207.
tive from western Quebec. Canadian Journal of M ARTIGNOLE , J. & R EYNOLDS , P. 1997. 40Ar/39Ar ther-
Earth Sciences, 31, 282– 292. mochronology along a western Québec transect of the
K ETCHUM , J. W. F. & D AVIDSON , A. 2000. Crustal archi- Grenville Province, Canada. Journal of Metamorphic
tecture and tectonic assembly of the Central Gneiss Geology, 15, 283 –296.
Belt, southwestern Grenville Province, Canada: a M ARTIGNOLE , J., C ALVERT , A. J., F RIEDMAN , R. &
new interpretation. Canadian Journal of Earth R EYNOLDS , P. 2000. Crustal evolution along a
Sciences, 37, 217–234. seismic section across the Grenville Province (western
K ETCHUM , J. W. F. & K ROGH , T. E. 1997. U– Pb con- Quebec). Canadian Journal of Earth Sciences, 37,
straints on high-pressure metamorphism in the Central 291–306.
Gneiss Belt, southwestern Grenville orogen. Geologi- M ARTIGNOLE , J., J I , S. & N ANTEL , S. 2006. A section
cal Association of Canada – Mineralogical Association through the Morin Terrane, Grenville Province
of Canada, Program with Abstracts, 22, A78. (Québec). Geological Association of Canada – Miner-
K ETCHUM , J. W. F. & K ROGH , T. E. 1998. U– Pb con- alogical Association of Canada, Joint Annual Meeting,
straints on high-pressure metamorphism in the south- Montréal 2006, Field Trip A2 Guidebook, 21p.
western Grenville orogen. Goldschmidt Conference M C L ELLAND , J., D ALY , J. S. & M C L ELLAND , J. M. 1996.
1998, Abstracts volume, Mineralogical Magazine, The Grenville orogenic cycle (c. 1350–1000 Ma): an
62A, 775–776. Adirondack perspective. Tectonophysics, 265, 1 –28.
K ETCHUM , J. W. F., H EAMAN , L. M., K ROGH , T. E., M C L ELLAND , J., H AMILTON , M., S ELLECK , B.,
C ULSHAW , N. G. & J AMIESON , R. A. 1998. Timing M C L ELLAND , J., W ALKER , D. & O RRELL , S. 2001.
and thermal influence of late orogenic extension in Zircon U–Pb geochronology of the Ottawan
the lower crust: a U–Pb geochronological study from Orogeny, Adirondack Highlands, New York: regional
the southwest Grenville orogen, Canada. Precambrian and tectonic implications. Precambrian Research,
Research, 89, 25–45. 109, 39– 72.
K RAUSS , J. B. & R IVERS , T. 2004. High-pressure M EZGER , K., R AWNSLEY , C. M., B OHLEN , S. R. &
granulites in the Grenvillian Grand Lake thrust H ANSON , G. N. 1991. U–Pb garnet, sphene, monazite,
system, Labrador: pressure-temperature conditions and rutile ages: implications for the duration of high-
and tectonic evolution. In: T OLLO , R. P., C ORRIVEAU , grade metamorphism and cooling histories, Adirondack
L., M C L ELLAND , J. & B ARTHOLOMEW , M. J. (eds) Mtns, New York. Journal of Geology, 99, 415–428.
Proterozoic tectonic evolution of the Grenville M ILLER , B. V., D UNNING , G. R., B ARR , S. M., R AESIDE ,
orogen in North America. Geological Society of R. P., J AMIESON , R. A. & R EYNOLDS , P. H. 1996.
America Memoir, 197, 105–134. Magmatism and metamorphism in a Grenvillian
K ROGH , T. E. 1994. Precise U–Pb ages for Grenvillian fragment: U– Pb and 40Ar/39Ar ages from the Blair
and pre-Grenvillian thrusting of Proterozoic and River Complex, northern Cape Breton Island, Nova
Archaean metamorphic assemblages in the Scotia. Geological Society of America Bulletin, 108,
Grenville Front tectonic zone, Canada. Tectonics, 13, 127–140.
963– 982. M OORE , J. M. & T HOMPSON , P. H. 1980. The Flinton
K ROGH , T. E., K AMO , S., G OWER , C. F. & O WEN , J. V. Group: a late Precambrian metasedimentary succes-
2002. Augmented and reassessed U– Pb geochronolo- sion in the Grenville Province. Canadian Journal of
gical data from the Labradorian– Grenvillian front in Earth Sciences, 17, 221–229.
the Smokey archipelago, eastern Labrador. Canadian O’B RIEN , P. J., C ARSWELL , D. A. & G EBAUER , D. 1990.
Journal of Earth Sciences, 39, 831– 843. Eclogite formation and distribution in the European
L I , Z. X., B OGDANOVA , S. V. ET AL . 2008. Assembly, Variscides. In: C ARSWELL , D. A. (ed.) Eclogite
configuration, and break-up of Rodinia: a synthesis. Facies Rocks. Blackie, Glasgow & London, 204– 224.
Precambrian Research, 160, 179–210. O NSTOTT , T. C. & P EACOCK , M. W. 1987. Argon reten-
L OEWY , S. L., C ONNELLY , J. N., D ALZIEL , I. W. D. & tivity of hornblendes: a field experiment in a slowly
G OWER , C. F. 2003. Eastern Laurentia in Rodinia: cooled metamorphic terrane. Geochimica et Cosmo-
constraints from whole-rock Pb and U/Pb geochronol- chimic Acta, 51, 2891–2903.
ogy. Tectonophysics, 375, 169– 197. O WEN , J. V., R IVERS , T. & G OWER , C. F. 1986. The
M ADORE , L., V ERPAELST , P., B RISEBOIS , D., H OCQ , M. Grenville Front on the Labrador coast. In: M OORE ,
& D ION , D. J. 1999. Géologie de la région du lac J. M., D AVIDSON , A. & B AER , A. J. (eds) The
Grenville Province. Geological Association of Canada, R IVERS , T., K ETCHUM , J., I NDARES , A. & H YNES , A.
Special Paper, 31, 95–106. 2002. The High Pressure Belt in the Grenville Pro-
O WEN , J. V., D ALLMEYER , R. D., G OWER , C. F. & vince: architecture, timing and exhumation. Canadian
R IVERS , T. 1988. Metamorphic conditions and Journal of Earth Sciences, 39, 867– 893.
40
Ar/39Ar contrasts across the Grenville Front zone, R IVERS , T., C ULSHAW , N., H YNES , A., I NDARES , A.,
coastal Labrador, Canada. Lithos, 21, 13– 35. J AMIESON , R. & M ARTIGNOLE , J. in review. The
P ECK , W. H., D E A NGELIS , M. T., M EREDITH , M. T. & Grenville Orogen. In: P ERCIVAL , J. & C OOK , F.
M ORIN , E. 2005. Polymetamorphism of marbles in (eds) Variations in Tectonic Styles Revisited: a Litho-
the Morin terrane, Grenville Province, Quebec. Cana- probe Perspective. Geological Association of Canada,
dian Journal of Earth Sciences, 42, 1949–1965. Special Paper.
Q UIRKE , T. T. & C OLLINS , W. H. 1930. The disappear- R OSENBERG , C. L. & H ANDY , M. R. 2005. Experimental
ance of the Huronian. Geological Survey of Canada, deformation of partially melted granite revisited:
Memoir, 160. implications for the continental crust. Journal of Meta-
R EY , P., V ANDERHAEGUE , O. & T EYSSIER , C. 2001. morphic Geology, 23, 19–28.
Gravitational collapse of the continental crust: defi- S EARLE , M. P., L AW , R. D. & J ESSUP , M. J. 2006.
nition, regimes and modes. Tectonophysics, 342, Crustal structure, restoration and evolution of the
435–449. Greater Himalaya in Nepal-South Tibet: implica-
R EYNOLDS , P. H., C ULSHAW , N. G., J AMIESON , R. A., tions for channel flow and ductile extrusion of the
G RANT , S. L. & M C K ENZIE , K. J. 1995. 40Ar/39Ar middle crust. In: L AW , R. D., S EARLE , M. P. &
traverse – Grenville Front Tectonic Zone to Britt G ODIN , L. (eds) Channel Flow, Ductile Extrusion
domain, Grenville Province, Ontario, Canada. and Exhumation in Continental Collision Zones.
Journal of Metamorphic Geology, 13, 209–221. Geological Society, London, Special Publications,
R IVERS , T. 1983a. The northern margin of the Grenville 268, 355–378.
Province in western Labrador – anatomy of an ancient S LAGSTAD , T., H AMILTON , M. A., J AMIESON , R. A. &
orogenic front. Precambrian Research, 22, 41– 73. C ULSHAW , N. G. 2004. Timing and duration of
R IVERS , T. 1983b. Progressive metamorphism of pelitic melting in the mid orogenic crust: constraints from
and quartzofeldspathic assemblages in the Grenville U– Pb (SHRIMP) data, Muskoka and Shawanaga
Province of western Labrador – tectonic implications domains, Grenville Province, Ontario. Canadian
of bathozone 6 assemblages. Canadian Journal of Journal of Earth Sciences, 41, 1339– 1365.
Earth Sciences, 20, 1791–1804. S PEAR , F. S. & M ARKUSSEN , J. C. 1997. Mineral zoning,
R IVERS , T. 1997. Lithotectonic elements of the Grenville P-T-X-M phase relations, and metamorphic evolution
Province: review and tectonic implications. Precam- of some Adirondack granulites. Journal of Petrology,
brian Research, 86, 117 –154. 38, 757 –783.
R IVERS , T. 2008. Assembly and preservation of lower, T IMMERMANN , H., P ARRISH , R. R., J AMIESON , R. A. &
mid and upper orogenic crust in the Grenville Province C ULSHAW , N. G. 1997. Time of metamorphism
– Implications for the evolution of large, hot long- beneath the Central Metasedimentary Belt Boundary
duration orogens. Precambrian Research, 167, Thrust Zone, Grenville Orogen, Ontario: accretion at
237–259. 1080 Ma? Canadian Journal of Earth Sciences, 34,
R IVERS , T. & C HOWN , E. H. 1986. The Grenville orogen 1023– 1029.
in eastern Quebec and western Labrador – definition, T IMMERMANN , H., J AMIESON , R. A., P ARRISH , R. R. &
identification and tectonometamorphic relationships C ULSHAW , N. G. 2002. Coeval migmatites and granu-
of autochthonous, parautochthonous and allochtho- lites, Muskoka domain, southwestern Grenville Pro-
nous terranes. In: M OORE , J. M., D AVIDSON , A. & vince, Ontario. Canadian Journal of Earth Sciences,
B AER , A. J. (eds) The Grenville Province. Geological 39, 239 –258.
Association of Canada, Special Paper, 31, 31– 50. T OFT , P. B., S COWEN , P. A. H., A RKANI -H AMED , J. &
R IVERS , T. & I NDARES , A. 2006. Assembly of a 60 km F RANCIS , D. 1993. Demagnetization by hydration
crustal section through the Grenville orogen from in deep-crustal rocks in the Grenville province of
eclogites through granulites to the non-metamorphic Quebec, Canada: implications for magnetic anoma-
lid – implications for Grenvillian tectonic evolution. lies of continental collision zones. Geology, 21,
Geological Association of Canada – Mineralogical 999– 1002.
Association of Canada, Abstracts, 31, 128. V AN B REEMEN , O. & C ORRIVEAU , L. 2005. U –Pb age
R IVERS , T. & M ENGEL , F. C. 1988. Contrasting assem- constraints on arenaceous and volcanic rocks of the
blages and petrogenetic evolution of corona and non- Wakeham Group, eastern Grenville Province. Cana-
coronaa gabbros in the Grenville province of western dian Journal of Earth Sciences, 42, 1677– 1697.
Labrador. Canadian Journal of Earth Sciences, 25, V AN G OOL , J., R IVERS , T. & C ALON , T. 2008. The
1629–1648. Grenville Front zone, Gagnon terrane, southwestern
R IVERS , T., M ARTIGNOLE , J., G OWER , C. F. & Labrador: configuration of a mid-crustal foreland
D AVIDSON , A. 1989. New tectonic divisions of the fold-thrust belt. Tectonics, 27, TC1004, doi:
Grenville Province, southeast Canadian shield. 10.1029/2006TC002095.
Tectonics, 8, 63–84. W ARDLE , R. J., R YAN , A. B., P HILIPPE , S. & S CHÄRER ,
R IVERS , T., VAN G OOL , J. A. M. & C ONNELLY , J. N. U. 1990. Proterozoic crustal development in the
1993. Contrasting tectonic styles in the northern Gren- Goose Bay region, Grenville Province, Labrador,
ville Province: implications for the dynamics of oro- Canada. In: G OWER , C. F., R IVERS , T. & R YAN ,
genic fronts. Geology, 21, 1127– 1130. A. B. (eds) Mid-Proterozoic Laurentia–Baltica.
444 T. RIVERS
Geological Association of Canada Special Paper, 38, Pb–Pb de la region de Sept-Îles–Natashquan, Pro-
197– 214. vince de Grenville, moyenne Côte-Nord. In: B RISE-
W ASTENEYS , H. A., K AMO , S. L., M OSER , D., K ROGH , BOIS , D. & C LARK , T. (eds) Géologie et resources
T. E., G OWER , C. F. & O WEN , J. V. 1997. U–Pb geo- minerales de la partie est de la Province de Grenville.
chronological constraints on the geological evolution Ministère des Richesses naturelles, Québec, DV2002–
of Pinware terrane and adjacent areas, Grenville Pro- 2003, 59–117.
vince, southeastern Labrador, Canada. Precambrian W YNNE -E DWARDS , H. R. 1972. The Grenville Province.
Research, 81, 101–128. In: P RICE , R. A. & D OUGLAS , R. J. W. (eds) Vari-
W ODICKA , N., P ARRISH , R. R. & J AMIESON , R. A. ations in Tectonic Styles in Canada. Geological
1996. The Parry Sound domain: a far-travelled alloch- Association of Canada Special Paper, 11, 263–334.
thon? New evidence from U–Pb zircon geochronol- Y ANG , P. & I NDARES , A. D. 2005. Mineral zoning, phase
ogy. Canadian Journal of Earth Sciences, 33, relations, and P– T evolution of high-pressure granu-
1087–1104. lites from the Lelukuau terrane, northeastern Grenville
W ODICKA , N., D AVID , J., P ARENT , M., G OBEIL , A. & Province, Quebec. Canadian Mineralogist, 43,
V ERPAELST , P. 2003. Géochronologie U– Pb et 443–462.
Neoproterozoic reworking of the Palaeoproterozoic Capricorn
Orogen of Western Australia and implications for the
amalgamation of Rodinia
SANDRA A. OCCHIPINTI1,2 & STEVEN M. REDDY1*
1
The Institute for Geoscience Research, Dept of Applied Geology, Curtin University of
Technology, GPO Box U1987, Perth, WA 6845, Australia
2
Now at: AngloGoldAshanti, Level 13, St. Martin’s Tower, 44 St. George’s Terrace,
WA 6000, Australia
*Corresponding author (e-mail: S.Reddy@curtin.edu.au)
Abstract: Argon isotopic data from mica from the southern Capricorn region of Western Australia
record complex intra- and inter-grain systematics that reflect modification due to a range of
processes. However, 40Ar/39Ar age distributions, though complex, generally show early Neopro-
terozoic ages in the west, increasing to Mesoproterozoic ages in the east. Palaeoproterozoic ages
associated with cooling after the c. 1.8 Ga Capricorn Orogen or c. 1.6 Ga Mangaroon Orogen
are not preserved. These data reflect cooling from a c. 300 8C thermal overprint that took place
prior to 960 Ma that is related to the enigmatic Edmundian Orogeny. These data, combined with
sediment provenance data from the Early Neoproterozoic Officer Basin and U– Pb age data
from the nearby Pinjarra Orogen, indicate that the late Mesoproterozoic– Neoproterozoic Pinjarra
and Edmundian events are dynamically linked and reflect tectonic activity on the western margin of
the amalgamated West Australian Craton. The temporal framework for this event suggest a link to
the evolving Rodinian supercontinent and reflect the oblique collision of either Greater India or
Kalahari cratons with the West Australian Craton. These results illustrate that the temporal
evolution of poorly preserved orogens can be constrained by low-temperature thermochronology
in the adjacent cratons.
40
Supplementary material: Summary of Ar/39Ar results reported in detail is available at
http://www.geolsoc.org.uk/SUP18357.
The cratonic cores of continental interiors are presenting mica 40Ar/39Ar data from the Palaeopro-
commonly typified by ancient high-grade meta- terozoic Capricorn Orogen of West Australia to
morphic rocks that have seen little tectonic acti- provide temporal constraints on tectonic activity
vity since the Archaean. In contrast, the margins along the western margin of the West Australian
of these cratons often record a complex geological Craton during the formation of Rodinia.
evolution involving cycles of rifting, accretion and
collision due to the global reorganization of con- Geological background
tinental fragments during the repeated dispersal
and amalgamation of supercontinents. Geological The West Australian Craton comprises the Archaean
analysis of craton margins therefore provides a Pilbara and Yilgarn cratonic blocks and a series of
valuable means of constraining ancient super- tectonically complex basement rocks and basins of
continent cycles. However, processes such as con- the Palaeoproterozoic Capricorn Orogen (Myers
tinental rifting, subduction erosion and crustal 1993) (Fig. 1). To the south, the West Australian
reworking that may take place at cratonic margins Craton passes into the Albany –Fraser Belt, a
can mask or destroy the evidence of earlier tectonic complex series of high-grade metamorphic rocks
activity, thereby limiting the ability of such areas to that were strongly deformed during the Mesoproter-
successfully assist in the reconstruction of tectonic ozoic collision of the West Australian Craton with
histories and palaeogeography. An alternative the South Australian –East Antarctic continent
approach is to attempt to identify and characterize (Clark et al. 2000). The eastern margin of the West
the far-field effects of craton margin tectonism Australian Craton is overlain and completely
within the craton, and use these data to constrain hidden by sediments of the Proterozoic Officer
the temporal evolution of processes taking place at Basin, the Phanerozoic Canning Basin and the
the margin. This approach is highlighted by Tertiary Eucla Basin (Trendall & Cockbain 1990).
446 S. A. OCCHIPINTI & S. M. REDDY
orogenesis (Occhipinti & Reddy 2004; Reddy &

Occhipinti 2004; Sheppard et al. 2004). Magma-
tism, metamorphism and deformation during
the 1680–1620 Ma Mangaroon Orogeny is also
heterogeneously distributed through the region,
becoming more significant to the north (Sheppard
et al. 2005).
In the central and eastern parts of the Capricorn
Orogen a regionally extensive series of sediments
and volcanics (the Bangemall Supergroup) uncon-
formably overlie basement rocks. Dolerite sills
that intruded the base of this stratigraphic sequence
yield dates of 1465 + 3 Ma (Wingate et al. 2002)
and 1070 + 6 Ma (Wingate et al. 2004). Some of
these dykes (of unknown age) were deformed and
metamorphosed with the sediments and the under-
lying basement at low metamorphic grades during
the enigmatic Edmundian Orogeny, before being
cut by northerly trending dolerite dykes of the
750 Ma Mundine Well dyke swarm (Wingate &
Giddings 2000; Martin & Thorne 2004). These
overprinting relationships loosely constrain the
Edmundian event at 1070–750 Ma.
Despite the relatively well-constrained Palaeo-
proterozoic evolution of the Capricorn Orogen,
Fig. 1. Map showing major geological features of the
West Australian Craton.
there are no published 40Ar/39Ar data from the
orogen. In this paper, the results of a regional
40
Ar/39Ar study document a widespread Late
The geological evolution of the western margin of the Mesoproterozoic–Neoproterozoic reworking of
West Australian Craton is also enigmatic because the Capricorn Orogen. The data, combined with
basin formation associated with the rifting and dis- previously published sedimentological data from
persal of Australia and Greater India during the Cre- the Officer Basin, and U – Pb zircon data from
taceous breakup of Gondwana (Song & Cawood basement rocks of the Pinjarra Orogen constrain
2000) masks the earlier history. Despite this, the tectonic activity on the western margin of the
presence of Mesoproterozoic and Neoproterozoic West Australian Craton.
rocks of the Pinjarra Orogen (Fig. 1) attests to an
eventful geological evolution of the western margin
of the craton (Myers 1990; Fitzsimons 2003) follow- Analytical procedure
ing its amalgamation in the Palaeoproterozoic era.
The Capricorn Orogen lies between the A regional suite of samples from different terranes
Archaean Pilbara and Yilgarn cratons and contains of the Capricorn Orogen and immediately adjacent
a series of terranes that comprise early to late Yilgarn Craton have been analysed by the
40
Archaean granite and granitic gneiss, Palaeopro- Ar/39Ar dating technique. Details of the analysed
terozoic metasedimentary and mafic meta-igneous samples, including sample localities and rock unit
rocks, granite and granitic gneiss (Fig. 2). These descriptions, are given as Supplementary Data. In
units are locally overlain by various sedimentary many cases the analysed samples have an igneous
units deposited in a range of settings between the origin and are granitic in composition, though
Palaeoproterozoic era to the Permian period a few samples are amphibolite –granulite facies
(Cawood & Tyler 2004, and references therein). metasedimentary rocks. In all samples the primary
The Capricorn Orogen comprises rocks mineralogical assemblage has been retrogressed
deformed and metamorphosed during the 2000– to greenschist facies metamorphic assemblage
1950 Ma Glenburgh Orogeny and the 1830– (see Supplementary Data). In many cases this
1780 Ma Capricorn Orogeny (Occhipinti et al. reflects the Capricorn Orogeny phase of the tectonic
1998; Occhipinti et al. 2004; Sheppard et al. evolution (c. 1800 Ma) (Occhipinti et al. 2004),
2004). Mineral assemblages throughout the range although recent studies have illustrated potentially
of basement rocks in the Capricorn Orogen indicate younger metamorphic overprints within the north
a regional greenschist facies metamorphic and of the Capricorn region (Sheppard et al. 2005).
deformation overprint associated with tectonic acti- Sample preparation and analytical procedure
vity associated with the latest stages of Capricorn have been described in detail elsewhere (Occhipinti
NEOPROTEROZOIC REWORKING OF THE PALAEOPROTEROZOIC CAPRICORN OROGEN 447
Fig. 2. Simplified geological map of the southern Capricorn Orogen showing major geological units. Sample loca-
tions for 40Ar/39Ar analyses from the Yilgarn craton (Narryer terrane), the Errabiddy Shear Zone and the Gascoyne
Complex are shown.
2004; Reddy et al. 2004) so only a brief summary is Proterozoic age and low Cl characteristics of the
presented here. 40Ar/39Ar analyses were conducted samples and the short amount of time between
on muscovite and biotite using two different analyti- irradiation and the time of analyses.
cal approaches: Infrared laser total grain fusion and Following irradiation, Ar was extracted using
infrared laser step-heating. In both cases, samples a CW– Nd–YAG laser, fired through a Merchantek
were crushed and inclusion-free mica grains were computer-controlled X –Y –Z sample chamber
selected after examination with a binocular micro- stage and microscope system. A defocused
scope. Depending on grain size, single or multiple 200 mm beam (9.7–11 Amps for 120 s) was used
grain aliquots were used to ensure sufficient Ar for for infrared laser analyses. Data were corrected for
measurement. All samples were cleaned in methanol, mass spectrometer discrimination and nuclear inter-
then de-ionized water in an ultrasonic bath. Once ference reactions. Errors quoted on the 40Ar/39Ar
dry, the samples were packed in aluminium foil ages are 1s, and ages were calculated using usual
and loaded into an aluminium package with other decay constants (Steiger & Jager 1977). J values
samples. The package was then Cd-shielded are noted on the supplementary data tables. Back-
(0.4 mm) and irradiated in the H5 position of the ground Ar levels were monitored prior to and after
McMaster University Reactor, Hamilton, Canada each analysis and the mean of the two blanks was
for 90 hours. Biotite age standard Tinto B, with a used to correct each sample analysis. Ar data were
K –Ar age of 409 Ma (Rex & Guise 1995) was corrected for mass spectrometer discrimination,
37
placed at 5 mm intervals throughout the aluminium Ar decay, and 38Ar decay.
package to monitor the neutron flux gradient. Tinto
B is a standard that has seen widespread use in the lit-
erature (Kelley et al. 1994; Sherlock & Kelley 2002; Results
Reddy et al. 2004; Downes et al. 2006). Correction
factors are as follows: (36Ar/37Ar)Ca ¼ 0.000255, Total-fusion data and step-heating age spectra are
(39Ar/37Ar)Ca ¼ 0.00065 and (40Ar/39Ar)K ¼ presented as accompanying supplementary data
0.0015. Corrections for (38Ar/39Ar)K and and are shown in Figures 4, 5 and 6. The 40Ar/39Ar
(38Cl/39Ar)K were not undertaken because of the data are also summarized in Table 1 and on the
448
Table 1. Summary of 40Ar/39Ar results reported in detail in the supplementary data. Bolded italicized ages are those interpreted as the best estimates of isotopic closure based on detailed
analysis of the composition, grains size, Ar relationships by Occhipinti (2004). WM, white mica (muscovite); B, biotite; IRF, Infrared fusion; IRSH, Infrared step heated; SGA, single grain
analyses; MGA, multiple grain analyses (number of grains indicated in parentheses). All ages are quoted at 1s. For ‘sample 142900’ mean ages were not calculated because the age range
was considerable
Area Sample Mineral Analytical % Atmospheric Weighted Unweighted Plateaux ages, Max Min Grain
name method Ar mean age mean age unweighted mean ages age age diameter
(range) (Ma) (Ma) using select data (Ma) (Ma) (Ma) (mm)
Narryer Terrane SO2-2a B GR1 IRSH 0– 1.60 1434 + 2 1440 + 38 1591 + 20 1389 + 10 313
B GR2 IRSH 0– 1.0 1249 + 2 1255 + 44 1341 + 6 1196 + 5 262
SO2-7a WM 834 + 2 834 + 21 874 + 5 773 + 15 332
SO2-08 B IRSH 0– 1.8 1532 + 2 1536 + 49 1552 + 30 (93.3% 39Ar) 1604 + 25 1452 + 6 578, 260
Errabiddy
Shear Zone
S. A. OCCHIPINTI & S. M. REDDY

East 142900 WM IRF 0.03 – 6.53 N/A N/A 1694 + 16 881 + 4 050– 200
142905 B IRF 0– 5.01 1623 + 2 1626 + 28 1690 + 13 1582 + 5 050– 150
West 142907 B IRF 1.47 – 14.01 927 + 1 927 + 23 994 + 6 893 + 10 050– 125
142907 WM IRF 0– 1.51 826 + 1 837 + 70 961 + 4 724 + 3 100– 200
142910 B IRF 1.78 – 9.11 961 + 1 962 + 10 984 + 6 973 + 5 075– 145
142910 WM IRF 0– 5.81 912 + 1 913 + 20 959 + 4 888 + 5 050– 200
142911 B IRF 0– 12.22 921 + 1 921 + 20 964 + 8 888 + 7 050– 115
168944 B IRF 1.72 – 7.09 1220 + 2 1224 + 26 1263 + 30 1178 + 5 050– 110
168944 WM IRF 0– 2.43 941 + 2 931 + 44 1021 + 4 896 + 9 050– 200
168946 B IRF 2.56 – 10.87 902 + 2 905 + 22 935 + 4 877 + 4 030– 150
SO2/10 B GR1 IRSH 0– 40.08 828 + 2 843 + 61 874 + 11 (65.6% 39Ar) 882 + 7 722 + 4 410
B GR2 2.90 – 23.9 789 + 1 794 + 46 820 + 14 (52.5% 39Ar) 701 + 3 844 + 4 650
SO2/14 WM GR1 IRSH 0– 13.56 884 + 2 866 + 43 901 + 24 (91.7% 39Ar) 942 + 4 801 + 7 260
WM GR2 0– 5.92 902 + 2 896 + 15 903 + 10 (91.59% 39Ar) 915 + 4 875 + 13 260
B IRSH 4.67 – 10.924 662 + 1 662 + 26 696 + 3 627 + 3 650
SAO-01/08 WM IRSH 0– 2.16 873 + 1 872 + 6 881 + 5 861 + 5 740
Gascoyne 142924 WM IRF 0.09 – 1.70 911 + 1 914 + 43 999 + 4 855 + 3 050– 350
Complex
142926 B IRF 3.56 – 12.04 1001 + 1 1001 + 16 1032 + 4 964 + 4 075– 175
142932 B IRF 0– 8.68 936 + 2 933 + 33 999 + 6 889 + 21 100– 200
142932 WM IRF 0– 2.61 898 + 2 906 + 30 963 + 7 880 + 4 150– 200
142933 B IRF 0.10 – 4.24 1102 + 1 1106 + 105 1176 + 5 967 + 6 60– 290
SO2-16A WM GR1 IRSH 0.13 – 13.29 886 + 2 940 + 137 895 + 7 (98.4% 39Ar) 1269 + 16 932 + 3 480
GR2 0– 1.04 950 + 2 966 + 93 925 + 10 (95.5% 39Ar) 1174 + 5 906 + 4 400
SO2-16C WM IRSH 0– 9.05 1045 + 1 1061 + 76 1182 + 17 914 + 4 890
620
CV-065 WM IRSH 0– 15.97 834 + 2 845 + 85 832 + 1 (plateau age) 1071 + 60 750 + 8 Not measured
B GR1 IRSH 0– 26.97 836 + 2 781 + 140 865 + 12 (61.9% 39Ar) 881 + 4 506 + 12 Not measured
B GR2 IRSH 0– 19.32 852 + 2 850 + 87 895 + 6 (61.1% 39Ar) 903 + 4 8+4 Not measured
regional geological map (Fig. 3). Overall the data with grain size but define several distinct age
show that apparent 40Ar/39Ar ages measured in peaks (Figs 4 & 5). By far the biggest peaks, and
biotite are often older than those measured in mus- therefore age distributions, for both muscovite and
covite from the same samples and the same range biotite, are of early Neoproterozoic age (Fig. 5).
of grain sizes (Figs 4 & 5). In addition, there is com- In detail age variations correspond to different
monly a wide range of 40Ar/39Ar ages within indi- regions. Biotite analysed by infrared total fusion
vidual samples. In some cases this is directly (Fig. 6a) show well-defined peaks between 1650
correlated to grain size and indicates the potential and 1580 Ma (Eastern Errabiddy Shear Zone,
presence of excess argon, particularly in some of n ¼ 10 of 11 analyses), 960 and 880 Ma (Western
the biotite samples, and heterogeneous Ar loss in Errabiddy Shear Zone, n ¼ 40 of 51 analyses) and
others. In all samples, measured 36Ar differs little 1020 and 930 Ma (Gascoyne complex, n ¼ 24 of
from background 36Ar levels. As a result, the use 38 analyses). Smaller peaks are present between
of isotope correlation diagrams (36Ar/40Ar vs 1270 and 1160 Ma (Western Errabiddy Shear
39
Ar/40Ar) is precluded as a means of recognizing Zone, n ¼ 11 of 51 analyses), and older ages up to
excess 40Ar. A detailed analysis of the complexity c. 1350 Ma are recorded from the Gascoyne
recorded at the individual sample level in the Complex (n ¼ 9). Muscovite total fusion ages
40
Ar/39Ar data has previously been described and between c. 1690– 880 Ma are recorded for the
analysed in considerable detail with respect to com- Eastern Errabiddy Shear Zone, but do not define
position (associated with mineral and fluid inclusion statistically valid peaks. For the Western Errabiddy
contamination), grain size and excess 40Ar and Ar Shear Zone, 15 out of 34 ages are between
loss (Occhipinti 2004). A summary of these data 920–880 Ma. This is broadly consistent with the
and age interpretations derived by Occhipinti greatest number of ages measured in the Gascoyne
(2004), taking these variables into account, are Complex between 900–870 Ma (n ¼ 7 of 11 analy-
given in Table 1. Readdressing the complexity is ses) (Fig. 5).
beyond the scope of this paper and we focus on Step-heating experiments on both biotite and
the broad patterns that emerge from the data and muscovite yield complex age spectra (Fig. 6) that
their tectonic significance. often correlate with compositional (Ca and Cl) vari-
Age data for the eastern Errabiddy Shear Zone ations (Occhipinti 2004). Generally the spectra do
(Fig. 3), the western Errabiddy Shear Zone, and not record statistically valid plateaux. However,
the Gascoyne Complex for single and multiple the distribution of Mesoproterozoic and Neoproter-
grain total-fusion show a wide range of apparent ozoic ages recorded by total fusion analyses are
40
Ar/39Ar ages that show complex relationships mimicked in the step-heating data and, despite the
Fig. 3. Simplified geological map of the Errabiddy Shear Zone and northern Yilgarn craton showing summary of
40
Ar/39Ar ages of biotite and muscovite. Note that age data from CV-065 and SO2_16A/C are not shown. Ages
represent interpretation after analysis of compositional, excess Ar and Ar loss in detailed by Occhipinti (2004).
complexity, attest to Neoproterozoic isotopic reset-

ting. It is noticeable that biotite from the Narryer
Terrane yields Mesoproterozoic ages that are con-
siderably different to those measured in both the
Gascoyne Complex and Errabiddy Shear Zone
(Fig. 6). However a single muscovite spectrum
from the Narryer Terrane also yields Neoprotero-
zoic ages and is similar to ages further north.
Discussion
Despite complexities in grain size–age relation-
ships (Fig. 4) and Ar age–composition spectra
(Occhipinti 2004), general patterns in 40Ar/39Ar
age distributions indicate differences in the
thermal history of different parts of the orogen at a
time substantially postdating the last major orogenic
(Palaeoproterozoic) event in the region. The
apparent overlap of c. 960–820 Ma ages in the
single-grain fusion and step-heating data from
the western Errabiddy Shear Zone and Gascoyne
Complex indicate a previously unrecognized isoto-
pic resetting associated with regional heating of
the western Capricorn Orogen during the early
Neoproterozoic era. The extent of this resetting
event (17 000 km2) is significant and indicates a
regional, not local, thermal perturbation. Closure
temperature models for micas suggest temperatures
of 350–270 8C (calculated using a cooling history
assumption of 10 8C/Ma) are required to cause
this resetting. This temperature range is consistent
with the low metamorphic grades seen in Mesopro-
terozoic sedimentary rocks in the region. Older ages
from the Eastern Errabiddy Shear Zone indicate a
lesser degree of isotopic resetting but still indicate
some Ar isotopic disturbance significantly after the
c. 1800 Ma greenschist facies metamorphism
associated with Palaeoproterozoic Capricorn oro-
genesis. The pattern of isotopic resetting generally
decreases towards the east, indicating that the
cause of resetting was more proximal to the west.
Mesoproterozoic aged biotites in the Narryer
Terrane may also reflect incomplete resetting
after the Capricorn event. However, a single musco-
vite age of 834 Ma may indicate a component of
excess 40Ar in the biotite data. New Ar data from
the Archaean rocks of the Narryer Terrane to the
south of this study area indicates both a Capricorn
Fig. 4. Age– grainsize relationships for muscovite and (c. 1750 Ma) overprint and a younger, weaker and
biotite analysed by infrared laser total grain fusion. heterogeneously distributed overprint (Spaggiari
Samples are from the Gascoyne Complex (a– d), the et al. 2008).
western Errabiddy Shear Zone (e–l) and eastern The Edmundian Orogeny is associated with loca-
Errabiddy Shear Zone (m –n). n ¼ number of analyses
lized north and east trending folds and faults recor-
for each sample. Where analyses plot over each other the
number of analyses is noted in parentheses besides the ded throughout the Capricorn region, particularly
analysis symbol. Note that error bars are often small and in the Bangemall Supergroup. The age of this defor-
are hidden behind the symbol used to represent mation event is poorly constrained but has been
the analyses. bracketed by age data on deformed and younger
Fig. 5. Histograms summarizing the frequency of ages measured on muscovite and biotite by total grain total fusion of
single and occasional multiple grain fusion of small grains (usually ,200 mm) for different regions of the south
Capricorn region. Distributions were calculated using a bin size of 20 Ma.
undeformed dolerite dykes that yield ages of 1070 an Edmundian thermal perturbation associated with
and 750 Ma, respectively (Wingate & Giddings early Neoproterozoic tectonism.
2000; Martin & Thorne 2004; Wingate et al. 2004). The Neoproterozoic Officer Basin forms part of
The regional causes of Edmundian orogenesis has the Centralian Superbasin (Walter et al. 1995).
remained unclear. However, the 960– 820 Ma age Palaeocurrent data from the base of the northwes-
range for regional mica resetting is consistent with tern part of the Officer Basin, the Sunbeam Group,
Fig. 6. Single and multiple grain IR-laser step-heated data from samples from the Central and Southern Gascoyne
Complex, the Errabiddy Shear Zone and the Yilgarn Craton (Narryer Terrane).
indicates that during the early Neoproterozoic era, and deposition in non-marine to shallow-marine
sediments were derived from the west (Fig. 7). Sedi- conditions (Grey et al. 2005). Detrital zircon popu-
ment characteristics are consistent with sourcing lations from the Tarcunyah Group of the Officer
from the eroding Bangemall Supergroup and a Basin, immediately north of the Sunbeam Group,
rapid increase in sediment supply (Williams 1992) also record U –Pb ages consistent with derivation
Fig. 7. Summary of palaeocurrent data from the Sunbeam Group of the Neoproterozoic Officer Basin (modified after
Williams 1992). Arrows show the direction of flow inferred from palaeocurrent indicators and, along with
petrological evidence, indicate provenance of sediments from the eroding Edmundian Orogeny.
Fig. 8. Simplified model for development of the western margin of the West Australian Craton, during the
Mesoproterozoic and Neoproterozoic based on data presented here and by Fitzsimons (2003). 1150 Ma: deposition of
sediments now preserved in the proto-Mullingarra and Northampton Complexes. 1080 Ma: metamorphism of these
sediments and continuing dextral transpressional deformation along the western part of Western Australia lead to
northward migration of the Mullingarra and Northampton complexes to their current positions. 950 Ma: Continuing
deformation on the western margin of the West Australian Craton caused the initiation of the Edmundian Orogeny in the
Capricorn Orogen. Thermal effects associated with collisional activity give rise to resetting of Ar isotopes in mica.
Increasing uplift leads to increased erosion of the evolving orogen and led to deposition of sedimentary detritus into the
Neoproterozoic Officer Basin.
from the Gascoyne Complex (Bagas 2003). Asym- Craton during the early Neoproterozoic assembly
metric sediment dispersion within the Sunbeam of the supercontinent Rodinia (Fig. 8). Depending
Group, coarsening upward successions, and the on which Rodinian reconstruction is preferred,
lack of any active volcanism associated with its either Kalahari or Greater India are potential
development, together support the possibility that candidates for collision with the West Australian
the sediments were deposited in an intracontinental Craton during the early Neoproterozoic (Pisarevsky
foreland basin-like setting (Jordan 1995; Miall et al. 2003; Li et al. 2008). Irrespective of which of
1995) with an evolving orogen situated to the these may or may not be correct, the isotopic
west (Williams 1992). Although, alternative resetting associated with Pinjarra/Edmundian tec-
interpretations of deposition environment are poss- tonism within the Capricorn Orogen suggests that
ible from sedimentological observations (Bagas deformation and metamorphism of the region took
2003; Grey et al. 2005), the available data from place considerably earlier than the 750 Ma minimum
the age-equivalent parts of the Officer Basin are age of orogenesis deduced from the earlier studies
consistent with Neoproterozoic tectonism localized (Wingate & Giddings 2000; Fitzsimons 2003;
in the western Capricorn Orogen. Martin & Thorne 2004; Wingate et al. 2004). Con-
To the west of the amalgamated West Australian sequently, future testing of Rodinian reconstruc-
Craton, the Pinjarra Orogen (Fig. 1) comprises high tions requires that early Neoproterozoic tectonism
grade metamorphic rocks of the Northampton and be used for temporal correlation with potential
Mullingarra complexes in which granulite and collisional candidates.
amphibolite facies metamorphism has been dated
at 1079 and 1058 Ma, respectively (Bruguier et al.
1999; Fitzsimons 2003). Zircon provenance data Conclusions
from paragneiss units within these two complexes 40
indicate the Mesoproterozoic Albany –Fraser Ar/39Ar dating of micas from the Palaeoprotero-
Orogen as a likely sediment source and, combined zoic Capricorn Orogen yield cooling ages indicating
with the limited extent of high-grade metamorph- early Neoproterozoic lower greenschist facies
ism, has been used to suggest that the Northampton metamorphic conditions of regional extent. Vari-
and Mullingarra Complexes are allochtonous and ations in the degree of resetting suggest higher
are derived from much further south (Fitzsimons temperatures in the west of the region and point
2003). Mafic dykes that are emplaced both within towards the thermal event being associated with
the Northampton Complex and the Yilgarn Craton, tectonic activity on the western margin of the
indicate that tectonic juxtaposition of the two com- amalgamated West Australian Craton. This
plexes with the West Australian Craton must have interpretation is supported by evidence from the
taken place prior to c. 750 Ma (Wingate & Giddings Neoproterozoic Officer Basin and the Pinjarra
2000). Consequently, there is some evidence for Orogen. The results indicate the potential of low
tectonism within the Pinjarra Orogen between temperature thermochronology to provide within-
1080 and 750 Ma (Fitzsimons 2003). From the craton evidence of far-field tectonic activity that
geometry of the 750 Ma mafic dykes and brittle- may be poorly preserved at the craton margins.
ductile northerly trending dextral shear zones In this case the results suggest a dynamic link
that post-date their emplacement (Byrne & Harris between the Pinjarra and Edmundian orogenies
1993), Fitzsimons (2003) argues that dextral trans- associated with collisional orogenesis during the
pressional within the Pinjarra Orogen took place early Neoproterozoic amalgamation of Rodinia.
The authors thank the Australian Research Council
around 750 Ma. Although broadly consistent with
(A00106036) for funding this research. The Tectonics
this model, the Ar data presented here indicate Special Research Centre and the Geological Survey of
much earlier tectonic activity along this western Western Australia, particularly I. Tyler and S. Sheppard,
margin of the West Australian Craton. are thanked for their support. P. Betts, D. Corrigan,
The coincidence of thermal resetting of Capri- D. Foster and an anonymous reviewer are thanked for
corn micas, the provenance of sediments from the their comments on the manuscript. This paper is TIGeR
Officer Basin and zircon data from the high-grade publication 96.
metamorphic complexes of the Pinjarra Orogen
are consistent with tectonic activity along the
western margin of the West Australian Craton References
during the early Neoproterozoic. In this scenario
B AGAS , L. 2003. Zircon Provenance in the Basal Part of
the Pinjarra and Edmundian orogenies are linked the Northwestern Officer Basin, Western Australia.
in that they represent the local and far-field effects Western Australia Geological Survey, Annual
of this tectonism respectively. On a larger scale, Review, 2002–2003.
this tectonism is likely to represent convergence of B RUGUIER , O., B OSCH , D., P IDGEON , R. T., B YRNE , D. I.
continental material with the West Australian & H ARRIS , L. B. 1999. U– Pb chronology of the
Northampton Complex, Western Australia – evidence Zone, Western Australia. In: A LSOP , G. I., H OLDS-
for Grenvillian sedimentation, metamorphism and WORTH , R. E., M C C AFFREY , K. J. W. & H AND , M.
deformation and geodynamic implications. Contri- (eds) Flow Processes in Faults and Shear Zones.
butions to Mineralogy and Petrology, 136, 258– 272. Geological Society, London, Special Publications,
B YRNE , D. & H ARRIS , L. B. 1993. Structural controls on 224, 229–248.
the base-metal vein deposits of the Northampton O CCHIPINTI , S. A., S HEPPARD , S., N ELSON , D. R.,
Complex, Western Australia. Ore Geology Reviews, M YERS , J. S. & T YLER , I. M. 1998. Syntectonic
8, 89– 115. granite in the southern margin of the Palaeoproterozoic
C AWOOD , P. A. & T YLER , I. M. 2004. Assembling and Capricorn Orogen, Western Australia. Australian
reactivating the Proterozoic Capricorn Orogen: Journal of Earth Sciences, 45, 822– 822.
lithotectonic elements, orogenies, and significance. O CCHIPINTI , S. A., S HEPPARD , S., P ASSCHIER , C.,
Precambrian Research, 128, 201–218. T YLER , I. M. & N ELSON , D. R. 2004. Palaeoprotero-
C LARK , D. J., H ENSEN , B. J. & K INNY , P. D. 2000. Geo- zoic crustal accretion and collision in the southern
chronological constraints for a two-stage history of the Capricorn Orogen: the Glenburgh Orogeny. Precam-
Albany-Fraser Orogen, Western Australia. Precam- brian Research, 128, 237– 255.
brian Research, 102, 155–183. P ISAREVSKY , S. A., W INGATE , M. T. D., P OWELL , C. M.,
D OWNES , P. J., W ARTHO , J. A. & G RIFFIN , B. J. J OHNSON , S. & E VANS , D. A. D. 2003. Models of
2006. Magmatic evolution and ascent history of the Rodinia assembly and fragmentation. In: Y OSHIDA ,
Aries Micaceous Kimberlite, central Kimberley Basin, M., W INDLEY , B. E. & D ASGUPTA , S. (eds) Protero-
Western Australia: evidence from zoned phlogopite zoic East Gondwana: Supercontinent Assembly and
phenocrysts, and UV laser 40Ar/39Ar analysis of Breakup. Geological Society, London, Special Publi-
phlogopite-biotite. Journal of Petrology, 47, 1751. cations, 206, 35– 55.
F ITZSIMONS , I. C. W. 2003. Proterozoic basement pro- R EDDY , S. M. & O CCHIPINTI , S. A. 2004. High-strain
vinces of southern and southwestern Australia, and zone deformation in the southern Capricorn Orogen,
their correlation with Antarctica. In: Y OSHIDA , M., Western Australia: kinematics and age constraints.
W INDLEY , B. F. & D ASGUPTA , S. (eds) Proterozoic Precambrian Research, 128, 295– 314.
East Gondwana: Supercontinent Assembly and R EDDY , S. M., C OLLINS , A. S., B UCHAN , C. & M RUMA ,
Breakup. Geological Society, London, Special Publi- A. H. 2004. Heterogeneous excess argon and Neopro-
cations, 206, 93–130. terozoic heating in the Usagaran Orogen, Tanzania,
G REY , K., H OCKING , R. M. ET AL . 2005. Lithostrati- revealed by single grain 40Ar/39Ar thermochronology.
graphic nomenclature of the Officer Basin and cor- Journal of African Earth Sciences, 39, 165–176.
relative parts of the Paterson Orogen, Western R EX , D. C. & G UISE , P. G. 1995. Evaluation of argon
Australia. Western Australia Geological Survey standards with special emphasis on time scale
Report, 93. measurements. In: O DIN , G. S. (ed.) Phanerozoic
J ORDAN , T. E. 1995. Retroarc foreland and Related Time Scale. IUGS Subcommision on Geochronology,
Basins. In: B USBY , C. J. & I NGERSOLL , R. V. (eds) 21–23.
Tectonics of Sedimentary Basins. Blackwell Science, S HEPPARD , S., O CCHIPINTI , S. A. & T YLER , I. M. 2004.
Cambridge, MA, 331–362. A 2005–1970 Ma Andean-type batholith in the
K ELLEY , S. P., A RNAUD , N. O. & T URNER , S. P. 1994. southern Gascoyne Complex, Western Australia. Pre-
High-spatial-resolution 40Ar/39Ar investigations cambrian Research, 128, 257–277.
using an ultra-violet laser probe extraction technique. S HEPPARD , S., O CCHIPINTI , S. A. & N ELSON , D. R.
Geochimica et Cosmochimica Acta, 58, 3519– 3525. 2005. Intracontinental reworking in the Capricorn
L I , Z. X., B OGDANOVA , S. V. ET AL . 2008. Assembly, Orogen, Western Australia: the 1680– 1620 Ma Man-
configuration, and break-up history of Rodinia: a syn- garoon Orogeny. Australian Journal of Earth Sciences,
thesis. Precambrian Research, 160, 179 –210. 52, 443 –460.
M ARTIN , D. & T HORNE , A. M. 2008. Tectonic setting and S HERLOCK , S. & K ELLEY , S. P. 2002. Excess argon evol-
basin evolution of the Bangemall Supergroup in the ution in HP-LT rocks: a UVLAMP study of phengite
northwestern Capricorn Orogen. Precambrian and K-free minerals, NW Turkey. Chemical Geology,
Research, 128, 385 –409. 183, 619–636.
M IALL , A. D. 1995. Collision-Related Foreland Basins. S ONG , T. & C AWOOD , P. A. 2000. Structural styles in the
In: B USBY , C. J. & I NGERSOLL , R. V. (eds) Tectonics Perth Basin associated with the Mesozoic break-up of
of Sedimentary Basins. Blackwell Science, Cambridge, Greater India and Australia. Tectonophysics, 317,
MA, 393–424. 55–72.
M YERS , J. S. 1990. Pinjarra Orogen. Geology and Mineral S PAGGIARI , C. V., W ARTHO , J.-A. & W ILDE , S. A. 2008.
Resources of Western Australia. Geological Survey of Proterozoic Deformation in the Northwest of the
Western Australia, Perth, 265 –274. Archean Yilgarn Craton, Western Australia. Precam-
M YERS , J. S. 1993. Precambrian history of the West brian Research, 165, 354– 384.
Australian craton and adjacent orogens. Annual S TEIGER , R. H. & J AGER , E. 1977. Subcommision on
Review of Earth and Planetary Sciences, 21, 453– 485. Geochronology: convention on the use of decay
O CCHIPINTI , S. A. 2004. Tectonic Evolution of the constants in geo- and cosmochronology. Earth and
Southern Capricorn Orogen, Western Australia. Planetary Science Letters, 36, 359– 362.
Curtin University of Technology. T RENDALL , A. F. & C OCKBAIN , A. E. 1990. Introduction,
O CCHIPINTI , S. A. & R EDDY , S. M. 2004. Deformation in Chapter Four: basins. Memoir of the Geological Survey
a complex crustal-scale shear zone: Errabiddy Shear of Western Australia, 3, 291–293.
W ALTER , M. R., V EEVERS , J. J., C ALVER , C. R. & G REY , Australia-Laurentia connection at 755 Ma. Precam-
K. 1995. Neoproterozoic stratigraphy of the Centralian brian Research, 100, 335– 357.
Superbasin, Australia. Precambrian Research, 73, W INGATE , M. T. D., P ISAREVSKY , S. A. & E VANS ,
173– 195. D. A. D. 2002. Rodinia connections between Australia
W ILLIAMS , I. R. 1992. Geology of the Savory Basin, and Laurentia: no SWEAT, no AUSWUS? Terra
Western Australia, Western Australia Geological Nova, 14, 121– 128.
Survey, Bulletin, 141, 1–115. W INGATE , M. T. D., P IRAJNO , F. & M ORRIS , P. A. 2004.
W INGATE , M. T. D. & G IDDINGS , J. W. 2000. Age Warakurna large igneous province: a new Mesoproter-
and palaeomagnetism of the Mundine Well dyke ozoic large igneous province in west-central Australia.
swarm, Western Australia: implications for an Geology, 32, 105 –108.
The Palaeoproterozoic Trans-Hudson Orogen: a prototype of
modern accretionary processes
D. CORRIGAN*, S. PEHRSSON, N. WODICKA & E. DE KEMP
Geological Survey of Canada, 615 Booth Street, Ottawa, Ontario, Canada, K1A 0E8
*Corresponding author (e-mail: dcorriga@NRCan.gc.ca)
Abstract: The Trans-Hudson Orogen (THO) of North America is one of the earliest orogens in
Earth’s history that evolved through a complete Wilson Cycle. It represents c. 150 Ma of
opening of the Manikewan Ocean, from c. 2.07 –1.92 Ga, followed by its demise in the interval
1.92– 1.80 Ga, during the final phase of growth of the Supercontinent Columbia (Nuna). It
is characterized by three lithotectonic divisions: (i) Churchill margin (or peri-Churchill);
(ii) Reindeer Zone; and (iii) Superior margin (or peri-Superior). The peri-Churchill realm records
progressive outward continental growth by accretion of Archaean to Palaeoproterozoic micro-con-
tinents (Hearne, Meta Incognita/Core Zone, Sugluk) and eventually arc terranes (La Ronge– Lynn
Lake) to the Slave-Rae nuclei, with attendant development of orogenies and basin inversions related
to the specific accretion events (1.92–1.89 Ga Snowbird; 1.88– 1.865 Ga Foxe; 1.87–1.865 Ga
Reindeer orogenies). The Reindeer Zone is characterized by primitive to evolved oceanic arcs,
back-arc basins, oceanic crust and ocean plateaus that formed during closure of the Manikewan
Ocean, and accretion of a micro-continent (Sask Craton) and smaller Archaean crustal fragments.
The terminal phase of the Trans– Hudson orogeny represents collision between the Superior craton,
the Reindeer Zone and the composite western Churchill Province during the interval 1.83– 1.80 Ga.
Introduction of South Dakota and perhaps even further to the

Grand Canyon area (Bickford & Hill 2007), over a
The Palaeoproterozoic Era (2.5– 1.6 Ga) forms a distance of c. 3000 km across the Canadian Prairies,
unique period of Earth’s evolution, highlighted by Hudson Bay, Baffin Island, Greenland and Scandi-
profound changes in continental plate configuration navia (Hoffman 1988). It may originally have
and tectonic processes (e.g. Griffin et al. 2008; Hou been even more extensive, as both ends are now
et al. 2008), ocean and atmospheric compositions truncated by younger orogenic belts. In Canada,
(e.g. Anbar & Knoll 2002) and the biosphere (e.g. the THO is generally well exposed and remained
Konhauser et al. 2002). By 2.5 Ga, Archaean relatively intact after its formation, providing the
cratons had grown by tectonic and magmatic accre- opportunity to study its evolution both parallel to
tion into large, stable continental masses buoyed by and across orogenic strike.
thick, depleted lithospheric roots that could accom- The THO is the site of closure of the Manikewan
modate the deposition of laterally extensive passive Ocean (Stauffer 1984), the oceanic plate that once
margin sequences (e.g. Artemieva & Mooney 2001; existed between the Superior Craton and the Rae
Bédard 2006; Percival 2007; Francis 2003). Wilson- Craton (Fig. 2a), stretching from South Dakota to
Cycle tectonics also appeared in the Palaeoprotero- the Ungava Peninsula in northern Québec (present-
zoic, resulting in the oldest documented obducted day geometry). It is likely that this ocean once contin-
ophiolites and development of tectonostratigraphic ued eastwards past the Ungava Peninsula, but its
sequences akin to those observed in the Phanerozoic extent from there is unclear. Closure of the Manike-
(e.g. Hoffman 1988). Within the Canadian Shield, wan Ocean was traditionally associated with for-
the Trans-Hudson Orogen (THO, Fig. 1; Hoffman mation of juvenile crust and its eventual accretion
1988; Lewry & Collerson 1990) represents the to the Superior craton and composite western
largest and most completely preserved segment Churchill Province (Gibb & Walcott 1971; Gibb
amongst a number of Palaeoproterozoic collisional 1975). Recent studies have highlighted that it also
belts that formed during the amalgamation of super- involved accretion of a number of smaller interven-
continent Nuna (also known as Columbia) during ing Archaean to earliest-Palaeoproterozoic conti-
the interval 2.0–1.8 Ga (Rogers & Santosh 2002; nental fragments, including the Sask Craton (Lewry
Pesonen et al. 2003; Hou et al. 2008). Before the et al. 1994; Hajnal et al. 1996). From a tectonostrati-
opening of the North Atlantic Ocean, the THO graphic perspective, the THO preserves a relatively
formed part of a nearly continuous orogenic complete Wilson-Cycle, from early c. 2.45 –
system that extended from the present-day location 1.92 Ga rift to drift sedimentary assemblages
458 D. CORRIGAN ET AL.
Fig. 1. Simplified geological map of the Canadian Shield and west Greenland, with extension of western Precambrian
terranes shown beneath Phanerozoic cover (pale colours SW of thick hashed line). The area affected by Trans–Hudson
age tectonothermal reactivation (1.83– 1.80 Ga) is delimited by the thick red hashed line. Limit of the Reindeer
Zone is shown (thick stippled blue line). Minimum limit of Trans-Hudson age tectonothermal overprint (thick hashed
red line). Terrane names and major faults: BH, Buffalo Head terrane; BK, Black Hills; C, Chinchaga belt; CZ, Core
Zone; FS, Fort Simpson terrane; FFB, Foxe Fold Belt; GB, Great Bear; GSL, Great Slave Lake shear zone; H, Hottah
terrane; KS, Ksituan belt; L, Lacombe belt; N, Nahanni belt; NO, New Québec; NV, Nova terrane; RB, Rimbey Belt;
STZ, Snowbird Tectonic Zone; T, Thorsby basin; TA, Tasiuyak domain; W, Wabamun belt.
deposited along Archaean craton margins, to for- Hoffman (1981) to identify the Palaeoproterozoic
mation of c. 2.00– 1.88 Ga oceanic and pericratonic collision zone that formed between the Superior
arcs and back-arc basins, as well as younger (1.88– Province and the Archaean domains of the western
1.83 Ga) continental arcs, foredeep and collisional Churchill Province (Green et al. 1985). This usage
basins, and eventual terminal collision (Ansdell differs from the one proposed by Lewry et al.
2005). This paper provides a synopsis of the accre- (1985) and Lewry & Collerson (1990), who prefer-
tionary and collisional evolution of the THO in ably based their definition on timing and contiguous
Canada. The objective is to provide the reader with extent of peak ‘Hudsonian-age’ tectonothermal
an appreciation of the complexity of this composite overprint, defined at c. 1.83– 1.80 Ga. Within that
orogenic system, the scale and age of colliding framework, the latter authors refer to Hoffman’s
crustal fragments, and the type of geodynamic pro- THO as the Reindeer Zone, an area represented by
cesses that were in place during the Palaeoprotero- the Palaeoproterozoic juvenile internides (Fig. 1).
zoic. This paper considers the whole of the They include within the definition of the THO a
Canadian THO, in comparison to Ansdell (2005), vast region represented by reactivated Archaean
Corrigan et al. (2007) and St-Onge et al. (2006) rocks of the western Churchill Province encompass-
that provide overviews of selected parts of the ing nearly all of the Hearne Craton, as well as a large
orogen. The evolution of the eastern Churchill Pro- portion of the Rae Craton (Fig. 1).
vince is not discussed in detail in this paper, but uti- The THO as defined by Lewry & Collerson (1990)
lizes the overview published by Wardle et al. (2002). is more inclusive and helps understand the orogen in its
entirety in terms of ocean closure, terrane accretion,
Tectonic context continental margin reactivation and syn- to post-
orogenic tectonothermal overprint. The THO is
Historically, the usage of the term ‘Trans-Hudson herein more broadly defined as an area of contiguous
Orogen’ has varied widely. It was introduced by or nearly-contiguous structural and metamorphic
THE PALAEOPROTEROZOIC TRANS-HUDSON OROGEN: A PROTOTYPE OF MODERN 459
overprint caused by c. 1.83–1.80 Ga collision The Sugluk block [introduced by Hoffman

between the Superior Craton, the Reindeer Zone, (1985) as ‘Sugluk Terrane’] contains crystalline
and a previously amalgamated continental collage basement rocks that fall in a narrow Mesoarchaean
north of the Manikewan Ocean consisting of the age range of c. 2.8–3.2 Ga (whole rock Rb– Sr
Slave, Rae and Hearne cratons, with the latter two col- on sedimentary rocks, Doig 1987; whole rock
lectively referred to as the ‘western Churchill Sm –Nd, Dunphy & Ludden 1998; U –Pb on
Province’ (Fig. 1). This zone of tectonometamorphic zircon, Wodicka & Scott 1997; Rayner et al.
overprint reaches a maximum width of c. 1800 km 2008) and in contrast to the Meta Incognita micro-
across strike of the orogen, from the reactivated edge continent, has no known 2.4 –2.5 Ga magmatism.
of the Superior Craton to the northwestern limit of Its main exposure lies in the northwestern edge of
the orogenic front in the western Churchill Province Ungava Peninsula where it has been variably
(Lewry & Collerson 1990; Corrigan 2002; Pehrsson named ‘Sugluk Terrane’ (Hoffman 1985, 1990),
et al. 2004). The southeastern Churchill Province, ‘Narsajuaq Terrane’ (St-Onge & Lucas 1990) and
which separates the Superior Craton from the North ‘Narsajuaq arc’ (St-Onge et al. 2006). On Ungava
Atlantic Craton, is commonly interpreted as a branch Peninsula, the Sugluk block is intruded by juvenile
of the THO, although the actual correlation between to variably enriched granitoids ranging in age
distinct tectonostratigraphic units from the western from 1.86 –1.80 Ga, collectively referred to as the
and southeastern ‘arms’ is still a matter of debate Narsajuaq suite (Dunphy & Ludden 1998). Aero-
(e.g. Bourlon et al. 2002; St-Onge et al. 2002; magnetic lineaments (Fig. 3) and Bouger gravity
Wardle et al. 2002). anomalies characteristic to this crustal block
The lifetime of ‘Manikewan Ocean’ can be esti- suggest that it potentially extends northwards to
mated on the basis of the history of the bounding the southern edge of Baffin Island, where it is pre-
cratons involved in the THO. Rift-related mafic sently in structural contact with the Meta Incognita
sequences dated at c. 2.07 Ga along the margins micro-continent. It appears to extend westwards
of the Hearne and Superior cratons, including under Hudson Bay, corresponding in large part to
the 2075+ Ma Courtenay Lake porphyry (Ansdell the ‘Hudson protocontinent’ defined by Roksandic
et al. 2000) and 2072+3 Ma Cauchon dykes (1987). The Sugluk block described herein forms
(Heaman & Corkery 1996), respectively, provide a the continental crust onto which the Narsajuaq arc
maximum age for formation of this oceanic was built, as described by Dunphy & Ludden
crust. Mafic and ultramafic rocks interpreted as (1998), and is of different age and nature than the
remnants of obducted ophiolite from the Watts Meta Incognita continental crust which hosted
Group in the Cape Smith Belt are dated at Narsajuaq-age intrusions on Baffin Island (e.g.
c. 2.0 Ga (Parrish 1989), in agreement with the Rayner et al. 2007). We therefore consider it a
postulated presence of oceanic crust during that separate crustal entity.
period (Scott et al. 1992). The Meta Incognita micro-continent (e.g.
On the basis of geophysical, geochronological St-Onge et al. 2000, 2006) underlies much of
and tracer isotope data, at least three large continen- southern Baffin Island and comprises basement rocks
tal fragments or micro-continents are identified in consisting of Neoarchaean crust (c. 2.7–2.6 Ga) and
the Manikewan Ocean ‘realm’ and identified as early Palaeoproterozoic (c. 2.40–2.15 Ga) granitoid
the Sask, Sugluk and Meta Incognita–Core Zone intrusions (N. Wodicka, unpublished data) that are
blocks in Figure 2a. The predominantly buried structurally overlain by a clastic –carbonate shelf
Sask Craton (Lewry et al. 1994) has been identified sequence (Lake Harbour Group), which was depo-
in three relatively small basement windows in the sited during the 2.01 –1.90 Ga interval (Scott et al.
Flin Flon –Glennie complex and, on the basis of 2002; Wodicka et al. 2008). In contrast with the
seismic images, is interpreted to underlie a large Sugluk block, evidence for older Mesoarchaean
part of the Reindeer Zone at middle to lower crust is provided indirectly from U –Pb inheritance
crustal levels (Hajnal et al. 2005). Where exposed, in U –Pb zircon ages (e.g. Rayner et al. 2007) and
the Sask Craton footwall is separated from over- TDM model ages (Whalen et al. 2008). On the
lying juvenile Proterozoic rocks by high meta- basis of U –Pb age, magnetic and gravity anomalies,
morphic grade mylonite zones with top-to-the SW as well as lithological associations, the Meta Inco-
sense of shear (Ashton et al. 2005). Tracer isotope gnita micro-continent may extend southeastwards
and U –Pb geochronological data obtained mainly (present-day co-ordinates) and correlates, at least
from the limited surface exposures and drill cores in part, with the Core Zone of the New Québec
suggest that the Sask Craton is predominantly orogen (Scott & St-Onge 1998; Bourlon et al.
formed of 2.4–2.5 Ga felsic to mafic igneous 2002; Wardle & Hall 2002).
rocks intruded into 3.1–2.8 Ga calc-alkaline Closure of the Manikewan Ocean had initiated
orthogneiss (Chiarenzelli et al. 1998; Bickford by at least c. 1905 Ma, which is the oldest age
et al. 2005; Rayner et al. 2005b). obtained for oceanic arc rhyolite in the Reindeer
Fig. 2. Series of cartoons illustrating a map view of the geological evolution of the THO, with the main tectonic
elements shown: (a) main lithotectonic elements during the interval 1.98–1.92 Ga, prior to the onset of convergence;
(b) the interval 1.92–1.89 Ga; (c) the interval 1.88 –1.865 Ga; (d) the interval 1.865–1.85 Ga; (e) the interval 1.85–
1.83 Ga; and (f) the interval 1.83 –1.80 Ga. The blue coloured hashed lines represent possible extensions of the
Superior Craton at Moho. Although not discussed in text, evolution of North Atlantic Craton (shown here to include
Archaean basement rocks and their cover on Hall and Cumberland peninsulas on Baffin Island (e.g. Scott 1999;
Jackson & Berman 2000) and tectonostratigraphic elements of the New Quebec and Torngat orogens are after
Wardle et al. (2002). Abbreviations: Am, Amer group; B, Burwell arc; BeS, Bergeron suture; BS, Baffin suture; BiS,
Big Island suture; Bf, Bravo formation; Ch, Chukotat; FFG, Flin Flon–Glennie Complex; FR, Fox River belt; GFtz,
Great Falls tectonic zone; GLsz, Granville Lake structural zone; Hoare Bay group; KB, Kisseynew Basin; Ke,
Ketyet group; LL, La Ronge–Lynn Lake belts; LTSz, Lac Tudor shear zone; M, Molson dykes; MHB, Medicine Hat
block; M.I., Meta Incognita micro-continent; Pe, Penrhyn group; Pi, Piling group; PS, Parent arc and Spartan
forearc; PT, Pelican thrust; SL, Snow Lake belt; SS, Sugluk suture; STZ, Snowbird Tectonic Zone; TA, Tasiuyak
domain; TNB, Thompson Nickel Belt.
Fig. 3. Oblique view of greyscale, shaded relief total field aeromagnetic anomaly map of the Ungava Peninsula and
southern Baffin Island area (illumination from NW). Red line is contour of shoreline. Thick blue lines are main
tectonostratigraphic boundaries. Thin blue line is trace of the Kovik antiform hinge. The cross section on bottom (drawn
to scale) is after Hoffman (1985) and St-Onge et al. (2002). Abbreviations are: Ch, Chukotat group; K, Kovik antiform;
LT, Labrador Trough; NA, Narsajuaq continental arc (emplaced in Sugluk block); na, metamorphically retrograded
equivalent of Narsajuaq Arc rocks in Sugluk suture zone; Po, Povungnituk group; PS, Parent arc and Spartan fore-arc;
W, Watts group oceanic crust. See text for explanations.
Zone (Baldwin et al. 1987; Stern et al. 1995; dating back to c. 1.92 Ga. This 100 Ma time frame
Corrigan et al. 2001a). However, the more primitive of THO evolution is herein subdivided into a series
arcs likely formed earlier, perhaps c. 1.92 Ga, which of orogenies, in much the same manner as the
is the oldest age of detrital zircon found in arc- Appalachian orogen, for example, providing a fra-
derived clastic sediments within the La Ronge and mework in which to better characterize distinct
Lynn lake belts (Ansdell 2005; Corrigan et al. periods of intraoceanic or pericratonic crustal accre-
2005). Note that contrary to Tran et al. (2008), we tion, magmatic accretion in continental arc batho-
do not attribute the inclusion of c. 1.91 –1.92 Ga liths and terminal continent-continent collision.
detrital zircons in the Hearne margin cover sequence The early period, termed the Snowbird orogeny
(lower Wollaston Group) to a putative ‘Rottenstone (c. 1.92–1.89 Ga, Berman et al. 2007), is recognized
arc’, but rather to local derivation from c. 1.91– as a phase of amalgamation of the Hearne and Rae
1.92 Ga intrusions in the Peter Lake Domain cratons, concurrent with oceanic arc formation
(e.g. Porter Bay complex; Rayner et al. 2005a). within the Manikewan Ocean (Ansdell 2005;
Taken at face value, these geochronological Corrigan et al. 2005, 2007) and subduction beneath
constraints suggest that the Manikewan Ocean had the western Superior Craton margin (Percival et al.
formed over a time span of at least 150 Ma before 2005). The next major phase, c. 1.88– 1.865 Ga,
it started to close, which is comparable in duration involves formation of the Glennie –Flin Flon –
to most modern oceans. The c. 1.83– 1.80 Ga term- Snow Lake intraoceanic collage (Lucas et al. 1994)
inal collision with the Superior Province, which and accretion of the La Ronge– Lynn Lake arcs to
historically defines the THO sensu stricto, was thus the southeastern Hearne Craton margin (Bickford
preceded by episodes of accretion and convergence et al. 1990, 1994; Corrigan et al. 2005), with the
latter event referred to herein as the Reindeer block and in a klippe (Cape Smith Belt) thrust
Orogeny. Penecontemporaneous with the latter, but onto the Ungava Promontory of the Superior
structurally unrelated, was the accretion of the Craton in northern Québec (Hoffman 1985). Not
Meta Incognita micro-continent to the eastern Rae much is known about the nature of the Reindeer
Craton (St-Onge et al. 2006) with attendant tecto- Zone beneath the Hudson Bay basin, although a
nothermal reworking (Foxe Orogeny). potentially substantive magmatic arc (Severn arc)
The third phase, informally named herein the has been interpreted to exist on the basis of continu-
c. 1.865 –1.845 Ga Wathaman Orogeny, is marked ity of a large region of aeromagnetic high appearing
by voluminous magmatic accretion along the south- to correspond on land to the Wathaman– Chipewyan
eastern margin of the western Churchill Province batholith (Green et al. 1985; Hoffman 1990). In the
(Thériault et al. 2001; Corrigan et al. 2005) and western Reindeer Zone, juvenile crust occurs in two
accretion of the Sugluk block to the latter (this separate belts, which are: (i) an intra-oceanic assem-
paper, and Berman et al. 2007). A fourth phase of blage (Flin Flon– Glennie Complex; Ashton 1999)
micro-continent accretion and continental arc mag- composed of c. 1.91–1.88 Ga primitive to evolved
matism followed during the interval 1.84 –1.82 Ga, island arc, ocean floor, ocean plateau and associated
as the Sask Craton collided with the Flin Flon– sedimentary and plutonic rocks that developed
Glennie Complex and arc magmatism (Narsajuaq during closure of the Manikewan Ocean (e.g.
sensu stricto) continued in the eastern Manikewan Stauffer 1984; Syme & Bailes 1993; Lucas et al.
Ocean. Terminal collision with the Superior 1996; Ansdell 2005); (ii) laterally discontinuous
Craton eventually closed the Manikewan Ocean belts of ocean arc, back-arc, ocean crust and associated
and caused the widespread tectonothermal rework- sediments, and sub-arc plutonic rocks that evolved as
ing characterizing the 1.83–1.80 Ga ‘Hudsonian pericratonic arcs during the interval 1.91–1.88 Ga
orogeny’ (cf. Stockwell 1961), or Trans-Hudson before accretion to the southeastern Hearne Craton
Orogen sensu stricto. The following sections con- margin (northwestern Reindeer Zone; Maxeiner
sider each of these orogenies in turn, providing a et al. 2004; Corrigan et al. 2005; Zwanzig 2000);
brief description of accretionary or collisional and (iii) magmatic arcs that were emplaced in
effects both within the Manikewan Ocean and continental arc and oceanic successor arc settings
along the cratonic margins. (Fumerton et al. 1984; Whalen et al. 1999).
Initiation of Manikewan Ocean closure Flin Flon – Glennie Complex

and Snowbird orogeny: (1.92 – 1.89 Ga) From east to west the Flin Flon –Glennie Complex
(Fig. 4) comprises the Snow Lake arc assemblage,
The time period 1.92 –1.89 Ga marks the beginning the Amisk collage, Hanson Lake block and the
of Manikewan Ocean closure, with attendant effects Glennie Domain, interpreted as part of a single
both within the internal Reindeer Zone, Rae and crustal entity formed at c. 1.87 Ga as a result of
Hearne Provinces, and Superior Craton margin intraoceanic accretion (Lewry & Collerson 1990;
(Fig. 2a). We consider these effects in turn, focusing Lucas et al. 1996). The various tectonostratigraphic
on distinct intraoceanic, intracratonic or pericra- assemblages forming the complex are fold-repeated
tonic effects, respectively. The original extent of and thrust-stacked, and structurally overlie the
the Manikewan Ocean is equivocal, although some Archaean to earliest-Palaeoproterozoic Sask Craton
estimates based on palaeomagnetic data suggest an above a basal décollement (Pelican Thrust) that
original width of at least 5000 km (Symons & was active as early as c. 1.84 Ga (Ashton et al.
Harris 2005). 2005). The lithotectonic evolution of Flin Flon –
Glennie complex is generally regarded as having
Juvenile crust formation: the western involved five main development stages (Lucas
Reindeer Zone et al. 1996; Stern et al. 1999), consisting of: (i)
1.92 –1.88 Ga formation of juvenile arcs, back-arc
The Reindeer Zone of the THO consists mainly of basins, ocean plateaus (Fig. 2b); (ii) 1.88 –1.87 Ga
juvenile crust of Palaeoproterozoic age that was intraoceanic accretion (Fig. 2c); (iii) 1.87 –1.84 Ga
formed during the closure of the Manikewan post-accretion development of successor arc intru-
Ocean (Lewry & Collerson 1990). An overwhelm- sions and inter-arc basins (Fig. 2d); and (iv) 1.84–
ing proportion of this crust is preserved in the 1.83 Ga terminal collision stage (Fig. 2d, e), first
western part of the Reindeer Zone in the Canadian with the Sask Craton at c. 1.84–1.83 Ga and later,
Prairies, west and NW of the Superior Craton. By at 1.83 –1.80 Ga, with the Superior Craton (Bleeker
contrast, a relatively smaller fraction is identified 1990; Ellis & Beaumont 1999; Ashton et al. 2005).
in the eastern part of the Reindeer Zone, now pre- In general, the pre-accretionary period records an
served mainly as plutonic rocks in the Sugluk evolution from primitive arc tholeiites to evolved
Fig. 4. Simplified map of the THO showing main lithotectonic domains of the Reindeer Zone and bounding domains.
Cree Lake zone is the reactivated portion of the Hearne Craton. Abbreviations: A, Amisk Lake block; FRB, Fox River
belt; G, Glennie domain; H, Hanson Lake block; NFSZ, Needle Falls shear zone; SBZ, Superior Boundary Zone;
SL, Snow Lake belt; TNB, Thompson Nickel Belt.
calc-alkaline arc rocks (Maxeiner et al. 1999; Bailes accreted terrane c. 50 km wide by 750 km long
& Galley 1999). Although the arcs are mostly juven- that is exposed from Lac La Ronge in north-
ile, Nd isotopic data suggest variable input from central Saskatchewan to north-central Manitoba
Archaean crust, perhaps recycled in sediments or as (Fig. 4). The La Ronge Domain hosts the 1905–
rifted crustal fragments locally incorporated in the 1876 Ma (Van Schmus et al. 1987; Corrigan et al.
early arc assemblages (Stern et al. 1995). In the 2001a) Central Metavolcanic Belt, interpreted as a
Flin Flon Domain, Archaean crust is actually pre- composite volcanic, plutonic and sedimentary
served as thrust-imbricated slices interleaved with assemblage formed in an ensimatic island arc
juvenile arc and back-arc crust (David & Syme setting transitional into an ensialic setting (Lewry
1994; Lucas et al. 1996; Syme et al. 1999). 1981; Watters & Pearce 1987; Thomas 1993). Arc
The Snow Lake Belt, which is located between assemblages contain a lower succession of ultrama-
the Flin Flon– Glennie Complex and the Superior fic tectonite, mafic to minor intermediate volcanic
Craton, also comprises remnants of an oceanic arc rocks, pelite, and exhalative horizons, as well as
volcano-plutonic complex and has historically oxide- and silicate-facies iron formation (collec-
been linked to the latter (e.g. Galley et al. 1986; tively named the Lawrence Point Assemblage),
Lucas et al. 1994; Bailes & Galley 1999). How- overlain by an upper succession of felsic volcanic
ever, recent models suggest that the Snow Lake and pyroclastic rocks, interbedded with and laterally
Belt may have evolved as a pericratonic arc outboard grading into volcaniclastic rocks (Reed Lake
of the Superior Craton (Fig. 2b), implying a possibly Assemblage) (Maxeiner et al. 2004, 2005). The
independent evolution from the remainder of the Flin stratigraphic top of the Reed Lake Assemblage
Flon–Glennie Complex prior to terminal collision conformably grades into a thick, laterally continu-
(Percival et al. 2004; Corrigan et al. 2007). ous and homogeneous succession of graphite-
bearing, mostly semi-pelitic to pelitic gneiss
The La Ronge – Lynn Lake domains interpreted as a remnant of either a forearc basin
or accretionary prism formed on the northern flank
The La Ronge Domain (Saskatchewan) and the of the La Ronge and Lynn Lake arcs, prior to
Lynn Lake/Leaf Rapids domains (Manitoba) form c. 1.88 Ga collision with the Hearne Craton
the core of a more or less continuous curvilinear, margin (Corrigan et al. 2005).
The Lynn Lake Domain (Lewry et al. 1990) Assemblage and lateral equivalents in the Rotten-
broadly correlates with the La Ronge Domain in stone and Southern Indian domains; Corrigan
the sense that both contain tectonically imbricated et al. 1998; 2001b), interpreted as a foreland or
and complexly folded volcanic, plutonic and sedi- molasse basin deposited on the previously accreted
mentary rocks. However, there are some major arc assemblages (Corrigan et al. 2005). All the
differences as well. Zwanzig et al. (1999) subdi- above assemblages are intruded by the c. 1863–
vided supracrustal rocks of the Lynn Lake Domain 1850 Ma Wathaman –Chipewyan Batholith, imply-
(Wasekwan Group) into a northern belt, predomi- ing that they had been tectonically accreted to
nantly consisting of mafic volcanic and volcaniclas- the Hearne Craton margin prior to continental
tic rocks of tholeiitic affinity and a southern belt arc magmatism.
consisting of tholeiitic to calc-alkaline volcanic
rocks with minor amounts of mid-oceanic ridge
basalt (MORB). Rare felsic volcanic rocks from Pericratonic Accretion: 1.88– 1.865 Ga
the northern belt have yielded U –Pb zircon ages
ranging from 1915 þ 7/26 Ma to 1886 + 2 Ma
Reindeer and Foxe orogenies
(Baldwin et al. 1987; Beaumont-Smith & Böhm The period 1.88–1.865 Ga was characterized by
2002). These rocks are significantly older than increased accretionary activity in pericratonic and
felsic to intermediate volcanic rocks of the southern intra-oceanic settings (Fig. 2c), including accretion
belt, which have so far yielded U –Pb zircon ages of of the La Ronge –Lynn Lake –Rusty Lake arcs to
1881 þ 3/22 Ma, 1864 + 2 Ma, 1856 + 2 Ma and the southeastern Hearne Craton margin (Reindeer
c. 1842 Ma (Beaumont-Smith & Böhm 2002, 2003). Orogeny) and accretion of the Meta Incognita
Recent Nd isotope data suggest that igneous rocks of micro-continent to the eastern Rae Craton (Foxe
the northern belt, which yield 1Nd values of 22.3 to Orogeny). It was also characterized by the develop-
þ1.3 and TDM model ages of c. 2.5 Ga, have been ment of a juvenile oceanic arc in the eastern Mani-
contaminated by Archaean crust or sediments with kewan Ocean (Parent and Spartan Groups of the
a high proportion of Archaean detritus (Beaumont- Cape Smith Belt: Picard et al. 1990; St-Onge et al.
Smith & Böhm 2003). In contrast, igneous rocks 1992; Barette 1994) and by renewed extension
of the southern belt yield 1Nd values of þ4 to þ5 along the northern and eastern Superior Craton
and TDM model ages ,2.2 Ga, suggesting a juvenile margin, resulting in voluminous mafic to ultramafic
origin (Beaumont-Smith & Böhm 2003). Both the magmatism (Baragar & Scoates 1981; Hulbert et al.
northern belt and parts of the southern volcanic 2005). This time period also coincided with col-
belt are intruded by tonalitic plutons of the 1876 – lision between the Core Zone, Tasiuyak arc and
1871 Ma Pool Lake intrusive suite (Baldwin et al. North Atlantic Craton, forming the Burwell arc
1987; Turek et al. 2000). and Torngat Orogen (Wardle et al. 2002).
The Rusty Lake Belt (Baldwin 1987) forms part
of the Leaf Rapids Domain (Fig. 4). It comprises
remnants of an evolved arc with mafic flows and Pericratonic arc accretion on southeastern
sills, minor dacitic and rhyolitic flows, volcaniclas- Hearne Craton margin: the Reindeer Orogeny
tic rocks, polymictic conglomerate and meta-
turbidite. Precise U –Pb geochronology in the The northwestern Reindeer Zone (Fig. 4) consists of
Ruttan Block has yielded an 1883 + 2 Ma zircon an autochthonous cover sequence of volcanic and
age for a quartz-feldspar porphyry rhyolite, and a sedimentary rocks (Lower Wollaston Group), an
unimodal, 1880 –1920 Ma age-range detrital allochthonous sequence of oceanic arc fragments
source for meta-turbidites directly overlying the (La Ronge, Lynn Lake and Rusty Lake belts),
volcanic edifice, consistent with the age range of their plutonic root (Rottenstone and South Indian
arc volcanic rocks from the Reindeer Zone inter- domains), an accretionary prism (Crew Lake,
nides (Rayner & Corrigan 2004). Milton Island and Zed Lake belts) and flysch/
Mafic to felsic plutonic rocks intruding the meta- molasse deposits (Upper Wollaston Group, Park
volcanic rocks of the La Ronge, Lynn Lake and Island and Partridge Breast belts) that either
Rusty Lake belts range in age from c. 1894– formed on or accreted to the southeastern margin
1871 Ma and are interpreted as the arc plutonic of the Hearne Craton during the interval 1.88–
root (Baldwin et al. 1987; Turek et al. 2000; 1.865 Ga (Corrigan et al. 2005). The upper age
Corrigan et al. 2005). These essentially form the limit of this time bracket corresponds to the
core of the Rottenstone and Southern Indian waning stages of ocean arc-type magmatic activity
domains (Fig. 4). The supracrustal and plutonic in the arc assemblages (Corrigan et al. 2005), and
rocks of the northwestern Reindeer Zone are uncon- the lower age bracket as the earliest intrusion of
formably overlain by fluvial to littoral facies silici- continental arc type magmas (e.g. Van Schmus
clastic metasedimentary rocks (Park Island et al. 1987).
A direct consequence of tectonic accretion of the that were deposited in a continental margin prism
arc terranes on the Hearne margin along north- setting (e.g. Morgan et al. 1975, 1976) that formed
vergent thrust faults was the deposition of a foreland at the edge of the Rae Craton during the approxi-
or molasse basin whose relics now form the Upper mate interval 2.16 –1.88 Ga, possibly evolving in
Wollaston Belt (Yeo 1998), as well as inliers an incipient oceanic basin analogous to the Red
within the northwestern Reindeer Zone (e.g. the Sea (e.g. Bohannon & Eittreim 1991). This basin
Park Island and Partridge Breast assemblages was rapidly filled by a flysch-type, thick greywacke
Corrigan et al. 1998, 1999, 2001b). Both these sequence during the interval 1.92 –1.90 Ga and was
units contain abundant detrital zircons originating folded and metamorphosed at greenschist to low-
from the juvenile La Ronge –Lynn Lake –Rusty pressure granulite facies as a result of northerly-
Lake supracrustal belt, with a peak at c. 1.92– vergent thrusting between c. 1.88 and 1.865 Ga
1.88 Ga (Tran et al. 2008; D. Corrigan, unpublished (St-Onge et al. 2006; Gagné et al. 2009). It has
data) and are intruded by the c. 1.865 –1.85 Ga been postulated that this orogenic event was
Wathaman–Chipewyan Batholith, suggesting caused by docking and subsequent collision of the
deposition sometime during the interval 1.88– Meta Incognita micro-continent (see below) with
1.865 Ga (Corrigan et al. 2005). Thermal effects the southeastern margin of the Rae Craton (Corrigan
of this accretion event have not been directly 2002; St-Onge et al. 2006). There are no constraints
dated, and may have been reset by ubiquitous at present on the precise nature of the suture between
Hudsonian-age (c. 1.83 –1.80 Ga) peak metamorph- the Rae Craton and Meta Incognita micro-continent.
ism. However, a penetrative deformation event in It has been named the Baffin Suture (St-Onge et al.
the form of early tight folds and layer-parallel foli- 2002) and its present position is approximate, based
ation (D1) has been recorded in protoliths that are primarily on the distribution of Piling Group versus
younger than c. 1.88 Ga and that are cross-cut by Lake Harbour Group and Hoare Bay Group tecto-
Wathaman–Chipewyan age intrusions (Corrigan nostratigraphic marker units. The northward ver-
et al. 1998). The actual extent of the Reindeer gence of structures in the Piling Group and
Orogeny north and west of the Wollaston Belt is evidence for moderate tectonic burial of the latter
as yet unconstrained. suggests that the Meta Incognita micro-continent
may have formed the upper plate in this collisional
zone. This thrust vergence is concordant to that
The Foxe Orogeny: accretion of Meta inferred for the Snowbird Tectonic Zone (Ross
Incognita micro-continent 2002; Berman et al. 2007), although there is pre-
sently insufficient data to determine if the latter
Farther NE along the western Churchill Province and the Baffin Suture actually link. During this
margin, the Foxe Orogeny (Hoffman 1990) formed period, renewed deformation and magmatism was
as a result of collision between the eastern also recorded further in the hinterland region in
Rae Craton and the Meta Incognita micro- the Committee Bay region (Berman et al. 2005),
continent (Fig. 2c) during the time interval 1.88– potentially as a far-field effect of the Foxe and/or
1.865 Ga (St-Onge et al. 2006). Collision resulted Reindeer orogenies.
in closure of the Penrhyn and Piling basins,
which had formed earlier during the interval
2.16 1.88 Ga (Wodicka et al. 2007a, 2008). Intraoceanic Accretion: the Flin
The Penrhyn and Piling sequences comprise: Flon – Glennie Complex
(i) shelf-margin assemblages (e.g. Lower Piling
Group) deposited after 2.16 Ga but prior to 1.98– Intra-oceanic tectonic accretion is a well-
1.94 Ga; (ii) a continental to incipient oceanic rift documented evolutionary process in the develop-
sequence (e.g. Bravo Lake formation) formed ment of the Flin Flon–Glennie Complex. It was
between c. 1.98 and 1.88 Ga, consisting of mafic first recognized in the Flin Flon belt (Lewry et al.
and ultramafic flows and sills, iron formation, and 1990), which contains an imbricate stack of
volcaniclastic and siliciclastic rocks (Johns et al. oceanic arc rocks, ocean island basalt, ocean
2006; Modeland & Francis 2010); and (iii) deep- plateau and associated plutonic and sedimentary
water basinal strata including sulphidic black shale rocks that were juxtaposed along a series of thrust
and metapelite (e.g. Astarte River Formation) and faults soon after their formation on the Manikewan
a thick sequence of greywacke (e.g. Longstaff ocean floor. U –Pb ages on plutonic rocks that cut
Bluff Formation) that was rapidly deposited, the earliest faults suggest that these volcanic assem-
buried and folded before being intruded by hyper- blages were structurally juxtaposed at c. 1.87 Ga
solvus granites at c. 1.90 Ga (Wodicka et al. (Lucas et al. 1996). Although most of the arc
2002a, 2007a). We interpret the Penrhyn and crustal fragments are juvenile, some have negative
Piling groups as a sequence of supracrustal rocks 1Nd values, which suggest that they may have been
built on fragments of Archaean crust (David et al. Development of a large continental

1996; Stern et al. 1999). Some of this Archaean magmatic arc on southeastern margin
crust is actually preserved as thin, fault-bounded
blocks imbricated between the arc slices (David & of western Churchill Province: the
Syme 1994). Formation of this early intra-oceanic 1.865 – 1.85 Ga Wathaman Orogeny
collage played a role in the further development of
the Reindeer Zone, as the Flin Flon– Glennie The period 1.865– 1.85 Ga records yet another fun-
complex behaved as a coherent crustal block when damental step in the evolution of the THO, with the
it eventually collided with the northern Reindeer emplacement of voluminous, mainly felsic plutonic
Zone and later rifted from the latter during rocks in an Andaean-type continental margin setting
the opening of the Kisseynew back-arc basin (Fig. 2d). The larger batholiths (Wathaman –
(see below). Chipewyan and Cumberland batholiths) were
emplaced in the previously accreted southeastern
margin of the western Churchill Province (Bickford
Magmatic accretion along Superior et al. 1994; Thériault et al. 2001; Corrigan et al.
2005). We speculate that the subduction zone
Craton margin related to this event may have been continuous for
During the interval 1.88 –1.865 Ga, as the Reindeer at least 2000 km, along the northern edge of the
Zone internides and western Churchill Province Manikewan oceanic plate (Fig. 2d). Contempora-
margin were in an overall contractional geodynamic neous subduction-related magmatism was also
setting, most of the Superior Craton margin facing occurring, albeit in smaller volumes, within the
the Manikewan Ocean was undergoing renewed Flin Flon –Glennie Complex and in the Sugluk
extension, perhaps as a result of pull from the sub- Block.
ducting Manikewan oceanic plate. Extension,
dated between c. 1.88 and 1.87 Ga, resulted in Wathaman – Chipewyan Batholith
the generation of voluminous mafic to ultramafic
bodies emplaced as dykes, sills and flows The Wathaman –Chipewyan Batholith is a major
(Fig. 2c), forming the extensive Circum-Superior intrusive complex that extends c. 700 km from
Belt (Baragar & Scoates 1981). Most of the mafic north-central Saskatchewan to northeastern Mani-
and ultramafic magma generated from this event toba (Lewry et al. 1981; Fumerton et al. 1984;
was emplaced through attenuated Superior Craton Meyer et al. 1992; MacHattie 2001). It was
crust and parautochthonous Palaeoproterozoic emplaced during a relatively narrow time interval
cover, as suggested by the slightly contaminated between c. 1862 and 1850 Ma, in a continental
1Nd values (Chauvel et al. 1987; Hegner & Bevier arc setting (Fumerton et al. 1984). The main body
1991). Dated members of this magmatic ‘event’ of the batholith is composed of K –feldspar
include the Molson dykes (Heaman et al. 1986; megacrystic, biotite + hornblende + titanite-bearing
Halls & Heaman 1997), the Fox River Sill granite, granodiorite and quartz monzonite, with
(Hulbert et al. 2005), the Chukotat Sills in the lesser amounts of diorite and layered gabbro
Cape Smith Belt (Wodicka et al. 2002b; Mungall (Corrigan et al. 1999). Along its northern flank,
2007) and mafic-ultramafic sills in the Labrador the batholith intrudes basement and cover sequences
trough (Wardle et al. 1990; Wodicka et al. 2002b). of the Hearne Craton margin. On its southern flank it
Historically, these mafic/ultramafic bodies have intrudes previously accreted juvenile terranes
been interpreted as the result of plume interaction including the La Ronge, Lynn Lake and Rusty
with the margin of the Superior craton (e.g. Ernst Lake arcs (Corrigan et al. 2005). These intrusive
& Buchan 2001). However, although a plume contacts were zones of focused strike-slip defor-
source may be possible, recent models point to the mation during terminal Trans-Hudson collision at
penecontemporaneous relationship between empla- c. 1.82 Ga (Lafrance & Varga 1996; Corrigan
cement of continental arc type calc-alkaline et al. 2005). Nd tracer isotope data across the bath-
plutons in the Thompson Nickel and Snow Lake olith show progressive contamination towards the
belts, and intrusion of the c. 1.88 Ga Molson Lake Hearne Craton, with 1Nd values ranging from þ2
dykes to infer an ensialic back-arc setting for near the southern margin of the batholith to 27
the generation of the mafic/ultramafic magmas along the northern margin (MacHattie 2001). This
(Percival et al. 2004, 2005; Corrigan et al. 2007). suggests that the batholith is a stitching intrusion
In the Cape Smith Belt, the Chukotat sills, dated at that masks the suture between accreted terranes
c. 1.89–1.87 Ga, overlap in age with the oldest mag- and Hearne Craton margin. Intrusive relationships
matic rocks of the Parent arc (1874 þ 4/23 Ma; with surrounding rocks also imply that the juvenile
Machado et al. 1993), suggesting an overall active arcs must have been already accreted to the
geodynamic setting during their emplacement. Hearne Margin prior to 1.865 Ga, and thus implies
that the batholith was produced as a result of sub- Plutonic rocks of similar age range (but as young
duction flip towards the N –NW, beneath the as 1844 Ma) and of more intermediate composition,
accreted margin (Corrigan et al. 2007). occur within the Sugluk block on Ungava Peninsula,
The waning stage of Wathaman –Chipewyan and have been named the ‘gneissic suite of the Nar-
Batholith magmatism at c. 1.85 Ga is interpreted sajuaq arc’ (Dunphy & Ludden 1998). They consist
to have occurred when the Flin Flon–Glennie of highly transposed, granulite to amphibolite facies
complex collided with the previously accreted La orthogneiss of tonalitic to dioritic composition,
Ronge–Lynn Lake –Rusty Lake arcs (northwestern which have 1Nd values ranging from þ4 to 210.7
Reindeer Zone), essentially consuming the interven- and have been interpreted as the magmatic product
ing oceanic lithosphere (Corrigan et al. 2005). The of subduction beneath a tectonically attenuated con-
suture produced by this collision is interpreted as tinental margin containing an Archaean crustal
the Granville Lake Structural Zone (Zwanzig component (Dunphy & Ludden 1998). We agree
2000; Corrigan et al. 2005, 2007), which currently with this model and postulate that this Archaean
correlates with a strong, crustal-scale seismic aniso- crust forms the core of the Sugluk block (Fig. 2d).
tropy that links at depth with a 5 km vertical offset in Genetically linking the Cumberland Batholith with
the Moho and corresponding gravity anomaly the gneissic suite would require that the Sugluk
(White et al. 2000). block and Meta Incognita micro-continent were
welded by 1.865 Ga, which contradicts current
models that bring these two crustal blocks together
Cumberland Batholith and ‘gneissic suite’ only at c. 1.82 Ga (e.g. St-Onge et al. 2000). There-
of the Narsajuaq Arc fore, we postulate that the penecontemporaneous
gneissic suite was produced by a different subduc-
The c. 1.865 –1.85 Ga Cumberland Batholith tion zone presently defined by the Sugluk suture
(Figs 1 & 2d) is contemporaneous with the (Hoffman 1985) (Figs 2d & 3). Cumberland Batho-
Wathaman–Chipewyan Batholith and has also been lith magmatism ended when the Sugluk block and
interpreted as the result of continental arc magma- Meta Incognita micro-continent collided, shortly
tism (Thériault et al. 2001; St-Onge et al. 2006). It after c. 1.85 Ga. This event may also have produced
forms a volumetrically large complex that was rapid uplift and cooling of granulite facies rocks
emplaced mainly within the Meta Incognita micro- dated at c. 1.84 Ga in the Meta Incognita micro-
continent, but extends northwards into the south- continent (M1 thermal event in Lake Harbour
eastern margin of the Rae Craton in central Baffin group; see St-Onge et al. 2007), as the Sugluk
Island (St-Onge et al. 2006). The Cumberland Bath- block began to underthrust. One possible expla-
olith is dominated by rocks of felsic composition nation that would reconcile the apparent dichotomy
and in contrast to the Wathaman– Chipewyan Bath- in timing of collision is that the Sugluk and Meta
olith, includes only minor mafic to intermediate Incognita blocks did begin to collide at
compositions. Moreover, the former yields c. 1.865 Ga, and that the bounding fault (Big
only negative 1Nd values ranging from 22 to 212 Island suture) was reactivated at c. 1.82 Ga.
(Thériault et al. 2001; Whalen et al. 2008), On Ungava Peninsula, magmatism related to the
suggesting either assimilation of a greater pro- 1863–1844 Ma gneissic suite is restricted to the
portion of Archaean crust, or a post-collisional Sugluk block (Dunphy & Ludden 1998). We postu-
setting (e.g. Whalen et al. 2008). The age and distri- late that this subduction-related magmatism stopped
bution of neodymium TDM model ages across the when the juvenile Parent–Spartan –Watts collage
Cumberland Batholith led Whalen et al. 2008 to collided with the Sugluk block at c. 1.84 Ga, reacti-
suggest that the latter images the Meta Incognita vating the Bergeron suture on the trailing edge of the
micro-continent and its bounding Archaean accreted collage and initiating the Narsajuaq arc
cratons (i.e. southern Rae margin in the north and sensu stricto (Fig. 2e).
the Sugluk block in the south). If correct, this
would suggest impingement of the Sugluk block
beneath Meta Incognita micro-continent as early Far-field c. 1.85 Ga reactivation in the
as c. 1.865 Ga. Whalen et al. (2008) postulate Snowbird Orogen
that the unusually large volume of magma forming
the Cumberland Batholith could have been the Berman et al. (2007) identified a large region within
product of slab breakoff. A similar origin could the northeastern part of the Snowbird Orogen that
potentially be inferred for the Wathaman- was subjected to NW-vergent thrusting and low
Chipewyan Batholith, as suggested by the unusually pressure, moderate temperature metamorphism at
large volume of magma produced and the above- c. 1.85 Ga, and speculated that it was possibly pro-
mentioned geophysical anomalies associated with duced by far-field reactivation during accretion of
the Granville Lake Structural Zone. the Hudson protocontinent (Sugluk block in this
paper). Farther SW along the Snowbird Orogen, 1997; Wodicka & Scott 1997). This new definition
exhumation of high-pressure rocks by east-verging restricts the age of arc magmatism to a duration of
transpression along the Legs Lake shear zone at c. 20 Ma and allows the interpretation of this suite
c. 1.85 Ga is documented by Mahan et al. (2006). as the product of continental arc magmatism in an
These events are contemporaneous with the upper plate formed of the previously amalgamated
waning stages of Wathaman– Chipewyan magma- Sugluk and Meta Incognita micro-continents, after
tism and beginning of infringement of the sub- the extinction of the 1.865–1.85 Ga Cumberland
duction zone that produced the Wathaman– Batholith, during the time interval 1.84 –1.82 Ga
Chipewyan arc (Granville Lake Structural Zone) (Fig. 2e).
by the Flin Flon –Glennie Complex (Fig. 2d). This This new definition also permits a potential link
specific docking event may have been the trigger between the Narsajuaq arc sensu stricto and the sub-
that produced a far-field reactivation within the ther- duction system (Sugluk suture and Lac Tudor shear
mally softened Snowbird Orogeny. We also agree zone) that produced the 1.84– 1.82 Ga De Pas Bath-
with the postulation of Berman et al. (2005, 2007) olith (Thomas & Kearey 1980; van der Leeden et al.
that some of that reactivation, particularly in the 1990; St-Onge et al. 2002). The De Pas batholith
northeastern part of the Snowbird Orogeny, may intrudes along the western margin of the Core
have been specifically produced by accretion of Zone, a potential extension of Meta Incognita
the Sugluk block. Moreover, we propose that this micro-continent south of Ungava Bay (e.g.
collision was predominantly of sinistral transpres- figure 10 in Wardle et al. 2002). Narsajuaq arc mag-
sional nature and that the resulting fault zone may matism waned at c. 1.82 Ga when the Manikewan
have been part of a continuous structure, linking Ocean eventually closed and the Superior Craton
the Granville Lake Structural Zone with the Big began its terminal collision with the accreted
Island suture to the NE (Fig. 2d). The oblique orogen internides.
nature of convergence may explain the absence of
continental arc magmatism in the area located
between the Wathaman– Chipewyan and Cumber- Back-arc opening and inversion in
land batholiths. Another possibility is that the Gran- western Manikewan Ocean: 1.85– 1.84 Ga
ville Lake Structural Zone connects with the Sugluk Kisseynew Domain
Suture, as postulated in Hoffman (1990), and that
the postulated Severn arc is part of that system. The Kisseynew Domain (Fig. 4) forms the youngest
tectonostratigraphic entity in the Reindeer Zone and
consists in large part of migmatitic meta-turbidite of
Final magmatic accretion: the Narsajuaq the Burntwood Group, interpreted to have been
arc (sensu stricto) and De Pas Batholith deposited in a c. 1.85 –1.84 Ga back-arc basin
(c. 1.84 – 1.82 Ga) formed between the La Ronge– Lynn Lake and
Flin Flon– Glennie arc complexes (Zwanzig 1990;
Published definitions of the Narsajuaq arc are overly Ansdell et al. 1995) (Fig. 2e). This model implies
inclusive, comprising all plutons ranging in age that c. 1.85–1.84 Ga magmatism in the Flin Flon –
from 1898–1800 Ma that occur within the Ungava Glennie Complex was the result of rifting of the
Peninsula in northern Québec (Dunphy & Ludden latter from the NW Reindeer Zone, perhaps as the
1998), as well as c. 1842–1821 Ma meta-plutonic result of hinge roll-back and opening of the Burnt-
rocks, mostly anatectic melts, that intrude the wood back-arc basin (Ansdell et al. 1995; Corrigan
Meta Incognita micro-continent on Baffin Island et al. 2007). The Flin Flon –Glennie complex was
(e.g. Scott 1997; Wodicka & Scott 1997). The defi- indeed quite magmatically active during that
nition has been expanded to include a c. 1.87– period, relative to the northwestern Reindeer
1.86 Ga island arc and associated sediments Zone. This phase of plutonism, however, is gener-
(Parent and Spartan Group) built on c. 2.0 Ga ally more mafic and includes large, Cu-Ni
oceanic crust (Watts Group) (St-Onge et al. 1992, +PGE-hosting mafic to ultramafic plutons, such
2006). We propose herein to restrict the definition as the 1847 + 6 Ma Namew Lake complex
of the Narsajuaq arc to specifically comprise (Cumming & Krstic 1991) and 1840 + 4 Ma
c. 1.84– 1.82 Ga continental arc-derived plutons gabbros that intrude the Bear Lake Block in the
that occur along the exposed southeastern margin Flin Flon belt (Bailes & Theyer 2006).
of the Sugluk block (i.e. the younger suite of The Burntwood Group is unconformably and/or
Dunphy & Ludden 1998), as well as the c. 1842 – disconformably overlain along most of its perimeter
1821 Ma plutonic suite identified along the southern by fluvial-alluvial sediments deposited between
margin of the Meta Incognita micro-continent c. 1845 and 1835 Ma, collectively interpreted as
and historically correlated with the Narsajuaq syn-collisional molasse basins (Stauffer 1990).
arc plutons on Ungava Peninsula (e.g. Scott Closure of the Burntwood back-arc basin may
have occurred at c. 1.84 Ga, shortly after its depo- between c. 1.83 and 1.76 Ga. Schneider et al.
sition, perhaps as a result of collision of the Sask (2007, and references therein) presented a compre-
craton with the Flin Flon –Glennie Complex hensive set of new and previously published ther-
(Fig. 2e, f ). This age coincides with the develop- mochronometric data in the western Reindeer
ment of the Pelican Thrust (Pelican suture in Zone in Saskatchewan and Manitoba, that show an
Fig. 2e) at the interface between the Sask Craton essentially homogeneous distribution of c. 1.83 –
and overlying Glennie Domain (Ashton et al. 1.76 Ga monazite and titanite U –Pb ages that
2005). Inversion of the Burntwood basin is also range geographically from the eastern margin of
reflected in the upward coarsening of sediments the Thompson Nickel Belt to the western edge of
from turbidite to intraformational conglomerate the Hearne Craton, near the Snowbird Tectonic
(Corrigan & Rayner 2002). Zone, a distance of c. 600 km (Fig. 2f). Similarly,
Wodicka et al. (2007b, c) published a compilation
of U –Pb monazite and titanite ages from the
Terminal Collision: docking of Superior Ungava –Baffin segment of the THO, showing a
Craton at 1.83 – 1.80 Ga broad distribution of c. 1.83 –1.78 Ga peak- to post-
peak metamorphic cooling ages spanning from the
Terminal collision in the THO (Fig. 2f) began with Cape Smith Belt, across Baffin Island, to the north-
early collision of the Superior Craton with the Rein- ernmost extents of reactivated Rae Craton north of
deer Zone, resulting in final consumption of the the Piling Group along Isortoq Fault (Jackson &
Manikewan Ocean and widespread tectonothermal Berman 2000; Bethune & Scammell 2003), a total
overprint across the Reindeer Zone and a large distance of c. 800 km (Fig. 2f). Although the ther-
part of the western Churchill Province (Lewry mochronological dataset from the core of the
et al. 1985; Lewry & Collerson 1990). Docking of western Churchill Province is sparser, there is suffi-
the Superior Craton appears to have been slightly cient information to infer a more or less uniform
diachronous, with earliest impingement occurring Hudsonian tectonothermal overprint, although
in the Thompson Nickel Belt at c. 1.83 Ga locally at low metamorphic grade, across most of
(Bleeker 1990) and subsequent collision with the its southwestern half (Berman et al. 2000) (Fig. 2f).
Cape Smith Belt at c. 1.82 Ga (St-Onge et al.
2002, 2006). Collision with the Core Zone was
even later, at c. 1.81 Ga (Wardle et al. 2002), and Summary and conclusions
likely the result of indentation of the Ungava
Peninsula and clockwise rotation of the Meta Incog- From its incipient intra-oceanic contractional phase
nita–Core Zone crustal block (e.g. Hoffman 1990). to terminal collision with the Superior Craton, the
Along most of its periphery, the Superior Craton composite THO represents c. 120 Ma of gradual
remained the lower crustal block during collision, ocean basin closure, tectonic and magmatic accre-
forming the footwall to accreted juvenile terranes, tion, intracratonic basin opening and closure, defor-
such as the Cape Smith klippe (Hoffman 1985; mation and metamorphism. Much in the same way
Lucas 1989). One exception to this general rule is as the Grenville or Appalachian orogens, for
the Thompson Nickel Belt, which began as a example, the THO can be defined as a series of
normal, SE-vergent Proterozoic-on-Archaean specific, short duration ‘orogenies’ that comprise
imbricate stack, but eventually developed retro- specific accretionary phases within a much larger
shears that thrust the Superior Craton basement orogenic system (e.g. Rivers 1997; Rivers &
onto the orogen internides during 1.82 –1.77 Ga Corrigan 2000; Hanmer et al. 2000; Murphy et al.
sinistral transpression (Ellis & Beaumont 1999; 1999; van Staal 2005; Hibbard et al. 2007).
White et al. 2002). The initial crustal architecture, The THO can be broadly subdivided into three
with the Reindeer Zone overthrust onto the Superior lithotectonic divisions, which are the Churchill
Craton, may explain why most of the eastern portion margin (or peri-Churchill), the Reindeer Zone and
of the Kisseynew Basin is currently underlain by the Superior margin (or peri-Superior) ‘realms’.
Archaean crust (Percival et al. 2007). The initial demise of the Manikewan Ocean and for-
As opposed to the Superior Craton, where mation of juvenile arcs, beginning at c. 1.92 Ga, was
Hudsonian tectonothermal overprint is limited to a penecontemporaneous with amalgamation of micro-
relatively narrow zone along its periphery, the continents (Hearne, Sugluk and Meta Incognita/
western Churchill Province is by comparison reacti- Core Zone) along the Rae margin, and accretion of
vated at a much greater scale. Two nearly complete juvenile arc complexes (La Ronge –Lynn Lake –
transects across the orogen, in the western and Rusty Lake arcs) outboard of that continental
eastern segments, show widely distributed U –Pb collage, by c. 1.88 Ga. The collisions between the
monazite and titanite ages representing peak- or Rae Craton and the micro-continents and arc com-
post-peak metamorphic temperatures, that range plexes resulted in basin inversions and attendant
tectonothermal ‘events’ that are divisible into the numerous structural breaks. The actual width of tec-
c. 1.92 –1.89 Ga Snowbird Orogeny and the tonothermal reactivation in the western Churchill
c. 1.88–1.865 Ga Foxe and Reindeer orogenies. Province, in the order of c. 900 km at its widest
Along the peri-Superior realm facing the Manike- extent, is similar in scale to the Parautochthonous
wan Ocean, early convergent activity appears to Belt of the Grenville Orogen, which extends for a
be limited to the Thompson Nickel Belt where maximum distance of up to 800 km across strike
c. 1.90–1.88 Ga calc-alkaline plutons intruding of the orogen (Figs 1 and 2f). It is noteworthy
the crystalline basement and Palaeoproterozoic that the Grenville Parautochthonous Belt, like the
cover sequence suggest continent margin magma- western Churchill Province, also evolved as a
tism, which in turn implies subduction beneath the long-lived accretionary margin prior to terminal
Superior Craton margin during that period. The Grenvillian collision between the Laurentian and
emplacement of voluminous, margin-parallel Amazonian cratons (Rivers & Corrigan 2000), and
mafic and ultramafic dykes and sills (Molson hence was also thermally softened prior to terminal
suite) at c. 1.88 Ga may signify development of an collision.
ensialic back-arc basin and establishment of a peri- An interesting observation from the THO is the
cratonic active arc (Snow Lake) outboard of that apparent absence of large extensional detachments,
margin (e.g. Percival et al. 2005). Within the core complexes, or other structural features indica-
Reindeer Zone, oceanic arcs formed first in tive of orogenic collapse (although, see Schneider
the western part of Manikewan Ocean, with the et al. 2007, for alternative view). This is consistent
c. 1.91–1.87 Ga amalgamation of the Flin Flon– with the rarity of ultrahigh- to high-pressure meta-
Glennie Complex, followed at c. 1.87– 1.86 Ga by morphic rocks of c. 1.83 Ga and younger age,
development of arcs (Parent and Spartan arc which suggests that the orogen, at least in its term-
complex) in the eastern part (or 1.90 –1.86 Ga if inal phase, may have been wide but of relatively
the Cape Smith suite ‘plutonic root’ is included). modest thickness. Most late Trans-Hudson age
An important, intermediate phase in the develop- (i.e. c. 1.85 –1.80 Ga) palaeo-pressures at the
ment of the THO was the establishment of a sub- present-day surface are in the range of 3– 7 kb
duction complex along the peri-Churchill margin, (e.g. Gordon 1989; Kraus & Menard 1997; Orrell
beginning at c. 1.865 Ga and leading to the gener- et al. 1999; Berman et al. 2007; Gagne et al. in
ation of voluminous continental arc batholiths. press), with higher pressures of c. 8– 9 kb found
This occurred in two distinct phases with the first uniquely in the Meta Incognita micro-continent
one at 1.865 –1.85 Ga producing the Wathaman– and Cape Smith Belt (Bégin 1992; St-Onge &
Chipewyan batholith in the Hearne Craton margin, Ijewliw 1996; St-Onge et al. 2007), and perhaps
as well as the penecontemporaneous Cumberland parts of the southwestern Hearne Craton (Annesley
batholith in the Meta Incognita microcontinent et al. 2005).
and 1.865– 1.844 Ga ‘older gneiss’ suite in the Assuming a present-day crustal thickness of
Sugluk block. A younger phase, at c. 1.84– c. 40 km for most of the THO (White et al. 2005),
1.82 Ga, was subsequent to accretion of the this suggests predominant crustal thicknesses of
Sugluk block to the peri-Churchill margin, leading c. 52–58 km during peak orogenesis, which is sub-
to back-stepping of the subduction zone and pro- stantially less than thicknesses of c. 70– 80 km esti-
duction of the Narsajuaq and De Pas continental mated for most of the Himalayan or Grenville
magmatic arcs and related anatectic melts. All orogens (Hirn et al. 1984; Indares & Dunning
subduction-related calc-alkaline magmatism ended 2001; Jamieson et al. 2002; Li & Mashele 2008).
at c. 1.82 Ga, with the final consumption of the The relatively broad and modestly thick nature of
Manikewan Ocean. the THO may have been caused by early thermal
Within this context, the THO sensu stricto rep- softening in the Reindeer Zone and western
resents the latest and most extensive of these Churchill Province, which would in turn have
orogenic events, resulting from terminal collision favored lateral growth rather than vertical thicken-
between the peri-Churchill collage, the Reindeer ing (e.g. Molnar & Lyon Caen 1988).
Zone internides and the Superior Craton. The An exception to the observed homogeneity of
broad asymmetry in extent of tectonothermal reacti- thermochronometric ages and palaeo-pressure and
vation observed between the Superior Craton and temperature conditions across the western THO is
the western Churchill Province (Fig. 1), may be found in the Flin Flon –Glennie Complex. At that
explained by the fact that the former represents a location, greenschist facies rocks of the Amisk
unique, previously cratonized, cold and rigid litho- block are juxtaposed against mid- to upper-
spheric block, whereas the latter comprises a amphibolite rocks of the Hanson Lake block and
crustal mosaic of smaller cratons that were Glennie Domain, along the north –south trending
assembled shortly before terminal collision and Tabbernor and Sturgeon –Weir fault system, with
hence represents hotter, weaker lithosphere with the Amisk block (eastern side) down-dropped with
respect to the western side (Ashton et al. 2005). We sequence in orogen-parallel folds (Fig. 3), as both
speculate that this faulting may have been caused by terranes were carried in a piggy-back and out-of-
southern extrusion of the underlying Sask Craton, sequence style onto the continental margin
caused by impingement of the Superior Craton, (Hoffman 1985; Lucas 1989).
which collided at c. 1.83 Ga, c. 10 Ma after the Another important point that is worth addressing
Sask Craton had collided with the Reindeer Zone is the pre-collisional evolution of the western
internides (e.g. Hajnal et al. 1996). margin of the Superior Craton, and in particular
One of the many features that have become the nature of the Snow Lake arc. Historically, the
apparent from the series of time-slice cartoons illus- latter has been associated with the intra-oceanic
trated in Figure 2 is the gradual southward growth of Flin Flon –Glennie Complex (e.g. Lewry et al.
the western Churchill Province, with consecutive 1990). However, given its relatively higher degree
accretion of: (i) the Hearne Craton; (ii) La of contamination by Archaean age crust, and the
Ronge–Lynn Lake Domains and Meta Incognita recent recognition of early arc magmatism on the
micro-continent; (iii) Flin Flon –Glennie Complex Superior Craton margin (Percival et al. 2004,
and Sugluk block; (iv) Sask Craton; and (v) Superior 2005), the question may be asked whether the
Craton; beginning at 1.92 Ga and ending at Snow Lake arc did actually evolve outboard, as
c. 1.80 Ga. Moreover, once the Superior Craton part of the Flin Flon –Glennie Complex, or as a peri-
stabilized at c. 1.78 Ga (Bleeker 1990), overall cratonic arc built on a rifted slice of the Superior
accretion and convergence at the North American Craton. If the latter, then one has to consider that
scale did not stop but simply migrated to the the Snow Lake –Amisk block boundary may be a
southern margin of the Wyoming–Trans-Hudson– fundamental, late-orogenic suture within the Flin
Superior collage, which then evolved in an open, Flon –Glennie Complex, and not part of the earlier
Andean-type setting for c. 600 Ma until terminal c. 1.87 Ga thrust stack of genetically related ocean
collision during the c. 1.2–1.0 Ga Grenville arc slices.
Orogeny (Hoffman 1988; Rivers & Corrigan 2000;
Whitmeyer & Karlstrom 2007). Another feature The authors owe a debt of gratitude to the influence of
which stands out from this orogen-scale compilation numerous geologists who have through the years unra-
is the apparent continuity of first-order structures, velled the story of different parts of this large and
complex orogen, and in particular to P. Hoffman who
such as subduction zones, between the western
began to painstakingly put it all together during the
and eastern regions of the Manikewan Ocean, not Decade in North-American Geology (DNAG) years.
only in timing of displacement but also inferred sub- We thank B. Murphy for the invitation to contribute to
duction polarities. One of the palaeo-subduction this volume, as well as R. Berman, P. Bickford and
zones or sutures that might have been thus linked K. Ansdell for providing insightful reviews. This paper is
is the Granville Lake Structural Zone –Big Island Geological Survey of Canada Contribution no. 20090130.
Suture (Fig. 2d). This subduction system is of
similar scale to most operating today in the SW
Pacific, for example. References
The model presented herein for the Palaeoproter- A NBAR , A. D. & K NOLL , A. H. 2002. Proterozoic ocean
ozoic evolution of the Ungava –Baffin segment of chemistry and evolution: a bioinorganic bridge?
the THO varies somewhat from previously pub- Science, 297, 1137–1142.
lished models (cf. Dunphy & Ludden 1998; A NNESLEY , I. R., M ADORE , C. & P ORTELLA , P. 2005.
St-Onge et al. 2006). For one part, we interpret the Geology and thermotectonic evolution of the western
Sugluk Suture (Hoffman 1990) as the northernmost margin of the Trans-Hudson Orogen: evidence from
extent of exposed Superior Craton (Fig. 3). This the eastern sub-Athabasca basement, Saskatchewan.
Canadian Journal of Earth Sciences, 42, 573–597.
suture connects in Ungava Bay with the Lac Tudor A NSDELL , K. M. 2005. Tectonic evolution of the
Shear Zone, which is interpreted as the western Manitoba-Saskatchewan segment of the Palaeoproter-
margin of the Core Zone (Wardle et al. 2002). The ozoic Trans-Hudson Orogen, Canada. Canadian
slightly older Big Island Suture (new name intro- Journal of Earth Sciences, 42, 741–759.
duced herein) separates the Sugluk block from the A NSDELL , K. M., L UCAS , S. B., C ONNORS , K. & S TERN ,
Meta Incognita micro-continent, and merges with R. A. 1995. Kisseynew metasedimentary gneiss belt,
the Sugluk–Lac Tudor ‘suture’ (Fig. 2e). Within Trans-Hudson Orogen (Canada): back-arc origin and
this framework, the Bergeron Suture forms the collisional inversion. Geology, 23, 1039–1043.
A NSDELL , K. M., M AC N EIL , A., D ELANEY , G. D. &
sole thrust that carried the juvenile Watts –Parent–
H AMILTON , M. A. 2000. Rifting and development of
Spartan oceanic crust and arc assemblage onto the the Hearne passive margin: age constraint from the
edge of the Superior Craton, where it is currently Cook Lake area, Saskatchewan. Geological Associ-
preserved as a structural klippe (cf. Hoffman ation of Canada – Mineralogical Association of
1985). The Bergeron suture likely merged with the Canada, GeoCanada 2000, Extended abstract #777
Sugluk suture prior to folding of the basement-cover (CD-ROM).
A RTEMIEVA , L. M. & M OONEY , W. D. 2001. Thermal Quebec, Canada: three new bathograds for mafic
thickness and evolution of Precambrian lithosphere: a rocks. Journal of Metamorphic Geology, 10, 685–704.
global study. Journal of Geophysical Research, 106, B ERMAN , R. G., R YAN , J. J. ET AL . 2000. The case of
16,387– 16,414. multiple metamorphic events in the western Churchill
A SHTON , K. E. 1999. A proposed lithotectonic domain Province: evidence from linked thermobarometric
reclassification of the southeastern Reindeer Zone in and in-situ SHRIMP data, and jury deliberations.
Saskatchewan. Summary of Investigations 1999, GAC– MAC 2000 Program with Abstracts, Volume
Volume 1, Saskatchewan Geological Survey, Sas- 25; GeoCanada 2000 Conference CD, Abstract 836.
katchewan Energy and Mines, Miscellaneous Report B ERMAN , R. G., S ANBORN -B ARRIE , M., S TERN , R. &
99-4, 92–100. C ARSON , C. 2005. Tectonometamorphism c. 2.35
A SHTON , K. E., L EWRY , J. F., H EAMAN , L. M., H AR- and 1.85 Ga in the Rae domain, western Churchill Pro-
TLAUB , R. P., S TAUFFER , M. R. & T RAN , H. T. vince Nunavut, Canada: insights from structural, meta-
2005. The Pelican Thrust Zone: basal detachment morphic and in-situ geochronological analysis of the
between the Archaean Sask Craton and Palaeoprotero- southwestern Committee Bay belt. Canadian Mineral-
zoic Flin Flon– Glennie Complex, western Trans- ogist, 43, 409– 442.
Hudson Orogen. Canadian Journal of Earth Sciences, B ERMAN , R. G., D AVIS , W. J. & P EHRSSON , S. 2007. Col-
42, 685–706. lisional Snowbird Tectonic Zone resurrected: growth of
B AILES , A. H. & G ALLEY , A. G. 1999. Evolution of the Laurentia during the 1.9 Ga accretionary phase of the
Palaeoproterozoic Snow Lake arc assemblage and geo- Hudsonian orogeny. Geology, 35, 911–914.
dynamic setting for associated volcanic-hosted B ETHUNE , K. M. & S CAMMELL , R. J. 2003. Distinguish-
massive sulphide deposits, Flin Flon Belt, Manitoba, ing between Archaean and Palaeoproterozoic tecton-
Canada. Canadian Journal of Earth Sciences, 36, ism, and evolution of the Isortoq fault zone, Eqe Bay
1789–1805. area, north-central Baffin Island. Canadian Journal
B AILES , A. H. & T HEYER , P. 2006. Wonderland gabbro: of Earth Sciences, 40, 1111–1135.
U–Pb age and geological interpretations, Manitoba B ICKFORD , M. E. & H ILL , B. M. 2007. Does the arc accre-
(NTS 63K 12NE). Report of Activities, Manitoba tion model adequately explain the Paleoproterozoic
Science, Technology, Energy and Mines, Manitoba evolution of southern Laurentia?: An expanded
Geological Survey, 35–41. interpretation. Geology, 35, 167– 170.
B ALDWIN , D. A. 1987. Proterozoic subaerial felsic vol- B ICKFORD , M. E., C OLLERSON , K. D., L EWRY , J. F.,
canism in the Rusty Lake greenstone belt. PhD thesis, V AN S CHMUS , W. R. & C HIARENZELLI , J. 1990.
University of Manitoba, Winnipeg, Manitoba. Proterozoic collisional tectonism in the Trans-Hudson
B ALDWIN , D. A., S YME , E. C., Z WANZIG , H. V., orogen, Saskatchewan. Geology, 18, 14–18.
G ORDON , T. M., H UNT , P. A. & S TEVENS , R. D. B ICKFORD , M. E., C OLLERSON , K. D. & L EWRY , J. F.
1987. U –Pb zircon ages from the Lynn Lake and 1994. Crustal histories of the Hearne and Rae pro-
Rusty Lake metavolcanic belts, Manitoba: two ages vinces, southwestern Canadian Shield, Saskatchewan:
of Proterozoic magmatism. Canadian Journal of constraints from geochronologic and isotopic data.
Earth Sciences, 24, 1053–1063. Precambrian Research, 68, 1– 21.
B ARAGAR , W. R. A. & S COATES , R. F. J. 1981. The B ICKFORD , M. E., M OCK , T. D., S TEINHART III, W. E.,
Circum-Superior belt; a Proterozoic plate margin? In: C OLLERSON , K. D. & L EWRY , J. F. 2005. Origin of
K RONER , A. (ed.) Precambrian Plate Tectonics. Else- the Archaean Sask craton and its extent within the
vier, Amsterdam, 297–330. Trans-Hudson orogen: evidence from Pb and Nd isoto-
B ARETTE , P. D. 1994. Lithostratigraphy and map-scale pic compositions of basement rocks and post-orogenic
structure in the western Cape Smith Belt, northern intrusions. Canadian Journal of Earth Sciences, 42,
Quebec: a tentative correlation between two tectonic 659– 684.
domains. Canadian Journal of Earth Sciences, 31, B LEEKER , W. 1990. New structural-metamorphic con-
986–994. straints on Early Proterozoic oblique collision along
B EAUMONT -S MITH , C. J. & B ÖHM , C. O. 2002. Struc- the Thompson Nickel Belt, Manitoba, Canada. In:
tural analysis and geochronological studies in the L EWRY , J. F. & S TAUFFER , M. R. (eds) The Early Pro-
Lynn Lake greenstone belt and its gold-bearing shear terozoic Trans-Hudson Orogen of North America.
zones (NTS 64C10, 11, 12, 14, 15 and 16). Report of Geological Association of Canada, Special Paper, 37,
Activities 2002, Manitoba Industry, Trade and Mines, 57–73.
Manitoba Geological Survey, 159– 170. B OHANNON , R. G. & E ITTREIM , S. L. 1991. Tectonic
B EAUMONT -S MITH , C. J. & B ÖHM , C. O. 2003. Tectonic development of passive continental margins of the
evolution and gold metallogeny of the Lynn lake green- southern and central Red Sea with a comparison to
stone belt, Manitoba (NTS 64C10, 11, 12, 14, 15 and Wilkes Land, Antarctica. Tectonophysics, 198,
16). Report of Activities 2003, Manitoba Industry, 129– 154.
Trade and Mines, Manitoba Geological Survey, B OURLON , E., M ARESCHAL , J. C., R OEST , W. &
39– 49. T ELMAT , H. 2002. Geophysical correlations in the
B ÉDARD , J. H. 2006. A catalytic delamination-driven Ungava Bay area. Canadian Journal of Earth Sciences,
model for coupled genesis of Archaean crust and sub- 39, 625 –637.
continental lithospheric mantle. Geochimica and Cos- C HAUVEL , C., A RNDT , N. T., K IELINZCUK , S. & T HOM ,
mochimica Acta, 70, 1188– 1214. A. 1987. Formation of Canadian 1.9 Ga old continental
B ÉGIN , N. J. 1992. Contrasting mineral isograd sequences crust. I: Nd isotopic data. Canadian Journal of Earth
in metabasites of the Cape Smith Belt, northern Sciences, 24, 396–406.
C HIARENZELLI , J. R., A SPLER , L. B., V ILLENEUVE , M. Palaeoproterozoic Flin Flon and Kisseynew belts,
& L EWRY , J. F. 1998. Early Proterozoic evolution of Trans-Hudson Orogen, Manitoba, Canada. Precam-
the Saskatchewan craton and its allochthonous cover, brian Research, 80, 107–124.
Trans-Hudson orogen. Journal of Geology, 106, D OIG , R. 1987. Rb– Sr geochronology and metamorphic
247– 267. history of Proterozoic to early Archaean rocks north
C ORRIGAN , D. 2002. The Trans-Hudson Orogen: old of the Cape Smith Fold Belt, Quebec. Canadian
paradigms and new concepts. Annual Meeting of the Journal of Earth Sciences, 24, 813–825.
Geological Association of Canada – Mineralogical D UNPHY , J. M. & L UDDEN , J. N. 1998. Petrological and
Association of Canada, Program with Abstracts, geochemical characteristics of a Palaeoproterozoic
Saskatoon 2002. magmatic arc (Narsajuaq terrane, Ungava Orogen,
C ORRIGAN , D. & R AYNER , N. 2002. Churchill River– Canada) and comparisons to Superior Province grani-
Southern Indian Lake Targeted Geoscience Initiative toids. Precambrian Research, 91, 109 –142.
(NTS 64B, 64C, 64G, 64H), Manitoba: Update and E LLIS , S. & B EAUMONT , C. 1999. Models of convergent
New Findings. Report of Activities 2002, Manitoba boundary tectonics: implications for the interpretation
Industry, Trade and Mines, Manitoba Geological of Lithoprobe data. Canadian Journal of Earth
Survey, 144–158. Sciences, 36, 1711–1741.
C ORRIGAN , D., M AXEINER , R., B ASHFORTH , A. & E RNST , R. & B UCHAN , K. 2001. Large mafic magmatic
L UCAS , S. 1998. Preliminary report on the geology events through time and links to mantle-plume heads.
and tectonic history of the Trans-Hudson Orogen in Geological Society of America, Special Paper, 352,
the northwestern Reindeer Zone, Saskatchewan. 483–575.
Current Research, Part C; Geological Survey of F RANCIS , D. 2003. Cratonic mantle roots, remnants of a
Canada, Paper, 98-1C, 95– 106. more chondritic Archaean mantle? Lithos, 71,
C ORRIGAN , D., M AC H ATTIE , T. G. & C HAKUNGAL , J. 135–152.
1999. The Wathaman Batholith and its relation to the F UMERTON , S. L., S TAUFFER , M. R. & L EWRY , J. F.
Peter Lake Domain: insights from recent mapping 1984. The Wathaman Batholith: largest known Pre-
along the Reindeer Lake transect, Trans-Hudson cambrian pluton. Canadian Journal of Earth Sciences,
Orogen. Saskatchewan Geological Survey, Miscella- 21, 1082–1097.
neous Report, Summary of Investigations, 1999, G AGNÉ , S., J AMIESON , R. A., M AC K AY , R., W ODICKA ,
Volume 2, 99-4.2, 132–142. N. & C ORRIGAN , D. 2009. Texture, composition, and
C ORRIGAN , D., M AXEINER , R. O. & H ARPER , C. 2001a. age variations in monazite from lower amphibolite to
Preliminary U– Pb results from the La Ronge-Lynn granulite facies, Longstaff Bluff Formation, Baffin
Lake Bridge project. In: Saskatchewan Geological Island, Canada. Canadian Mineralogist, 47, 847– 869.
Survey, Miscellaneous Report. Summary of Investi- G ALLEY , A. G., Z IEHLKE , D. V., F RANKLIN , J. M.,
gations 2001, Volume 2, 2001-4.2, 111–115. A MES , D. E. & G ORDON , T. M. 1986. Gold mineraliz-
C ORRIGAN , D., T HERRIAULT , A. & R AYNER , N. 2001b. ation in the Snow Lake–Wekusko Lake region, Mani-
Preliminary results from the Churchill River – toba. In: C LARK , A. L. (ed.) Gold in the Western
Southern Indian Lake Targeted Geoscience Initiative. Shield. Canadian Institute of Mining and Metallurgy,
Report of Activities, Manitoba Industry, Trade and Special Volume, 38, 379–398.
Mines. Manitoba Geological Survey, 94–107. G IBB , R. A. 1975. Collisional tectonics in the Canadian
C ORRIGAN , D., H AJNAL , Z., N ÉMETH , B. & L UCAS , S. B. Shield. Earth and Planetary Sciences Letters, 27,
2005. Tectonic framework of a Palaeoproterozoic 378–382.
arc-continent to continent-continent collisional zone, G IBB , R. A. & W ALCOTT , R. I. 1971. A Precambrian
Trans-Hudson Orogen, from geological and seismic suture in the Canadian Shield. Earth and Planetary
reflection studies. Canadian Journal of Earth Sciences, Science Letters, 10, 417–422.
42, 421–434. G ORDON , T. M. 1989. Thermal evolution of the
C ORRIGAN , D., G ALLEY , A. G. & P EHRSSON , S. 2007. Kisseynew sedimentary gneiss belt, Manitoba: meta-
Tectonic evolution and metallogeny of the southwes- morphism at an early Proterozoic accretionary
tern Trans-Hudson Orogen. In: G OODFELLOW , margin. In: D ALY , J. S., C LIFF , R. A. & Y ARDLEY ,
W. D. (ed.) Mineral Deposits of Canada: A Synthesis B. W. D. (eds) Evolution of Metamorphic Belts. Geo-
of Major Deposit-types, District Metallogeny, the logical Society, London, Special Publications, 43,
Evolution of Geological Provinces, and Exploration 233–244.
Methods. Geological Association of Canada, Mineral G REEN , A. G., H AJNAL , Z. & W EBER , W. 1985. An evol-
Deposits Division, Special Publication, 5, 881– 902. utionary model for the western Churchill Province and
C UMMING , G. L. & K RSTIC , D. 1991. Geochronology at western margin of the Superior Province in Canada and
the Namew Lake Ni–Cu deposit, Flin Flon area, Man- the north-central United States. Tectonophysics, 116,
itoba, Canada: a Pb/Pb study of whole rocks and ore 281–322.
minerals. Canadian Journal of Earth Sciences, 28, G RIFFIN , W. L., O’R EILLEY , S. Y., A LFONSO , J. C. &
1328–1339. B EGG , G. C. 2008. The composition and evolution of
D AVID , J. & S YME , E. C. 1994. U–Pb geochronology of lithospheric mantle: A re-evaluation and its tectonic
late Neoarchean tonalites in the Flin Flon Belt, Trans- implications. Journal of Petrology, doi: 10.1093/pet-
Hudson Orogen: surprise at surface. Canadian Journal rology/egn033.
of Earth Sciences, 31, 1785–1790. H AJNAL , Z., L UCAS , S., W HITE , D., L EWRY , J.,
D AVID , J., B AILES , A. H. & M ACHADO , N. 1996. B EZDAN , S., S TAUFFER , M. R. & T HOMAS , M. D.
Evolution of the Snow Lake portion of the 1996. Seismic reflection images of high-angle faults
and linked detachments in the Trans-Hudson Orogen. intrusions and associated nickel ores from the Thomp-
Tectonics, 15, 427– 439. son Nickel Belt, Manitoba, Canada. Economic
H AJNAL , Z., L EWRY , J. F. ET AL . 2005. The Sask Craton Geology, 100, 29–41.
and Hearne Province margin: seismic reflection I NDARES , A. & D UNNING , G. 2001. Partial melting of
studies in the western Trans-Hudson Orogen. Cana- high P-T metapelites from the Tshenukutish Terrane
dian Journal of Earth Sciences, 42, 403–419. (Grenville Province): petrography and U– Pb geochro-
H ALLS , H. C. & H EAMAN , L. M. 1997. New constraints on nology. Journal of Petrology, 42, 1547– 1565.
the Palaeoproterozoic segment of the Superior Pro- J ACKSON , G. D. & B ERMAN , R. G. 2000. Precambrian
vince apparent polar wander path from U –Pb dating metamorphic and tectonic evolution of northern
of Molson Dykes, Manitoba. Geological Association Baffin Island, Nunavut, Canada. Canadian Mineralo-
of Canada – Mineralogical Association of Canada, gist, 38, 399–421.
Joint Annual Meeting, Program with Abstracts, J AMIESON , R. A., B EAUMONT , C., N GUYEN , M. H. &
22, A61. L EE , B. 2002. Interaction of metamorphism, defor-
H ANMER , S., C ORRIGAN , D., P EHRSSON , S. & N ADEAU , mation and exhumation in large convergent orogens.
L. 2000. SW Grenville Province, Canada: the case Journal of Metamorphic Geology, 20, 9 –24.
against post-1.4 Ga accretionary tectonics. Tectono- J OHNS , S. M., H ELMSTAEDT , H. H. & K YSER , T. K. 2006.
physics, 319, 33– 51. Palaeoproterozoic submarine intrabasinal rifting,
H EAMAN , L. M. & C ORKERY , T. 1996. U –Pb geochro- Baffin Island, Nunavut, Canada: volcanic structure
nology of the Split Lake Block, Manitoba: preliminary and geochemistry of the Bravo Lake Formation. Cana-
results. In: H AJNAL , Z. & L EWRY , J. (eds) Trans- dian Journal of Earth Sciences, 43, 593– 616.
Hudson Lithoprobe Transect. Lithoprobe Report, 55, K ONHAUSER , K. O., H AMADE , T., R AISWELL , R.,
60– 67. M ORRIS , R. C., F ERRIS , F. G., S OUTHAM , G. &
H EAMAN , L. M., M ACHADO , N., K ROGH , T. E. & C ANFIELD , D. E. 2002. Could bacteria have formed
W EBER , W. 1986. Presise U–Pb zircon ages for the the Precambrian banded iron formations? Geology,
Molson dyke swarm and the Fox River sill: constraints 30, 1079–1082.
for Early Proterozoic crustal evolution in northern K RAUS , J. & M ENARD , T. 1997. A thermal gradient at
Manitoba, Canada. Contributions to Mineralogy and constant pressure: implications for low- to medium-
Petrology, 94, 82– 89. pressure metamorphism in a compressional tectonic
H EGNER , E. & B EVIER , M. L. 1991. Nd and Pb isotopic setting, Flin Flon and Kisseynew domains, Trans-
constraints on the origin of the Purtiniq ophiolites Hudson Orogen, Central Canada. Canadian Mineralo-
and Early Proterozoic Cape Smith Belt, northern gist, 35, 1117– 1136.
Quebec, Canada. Chemical Geology, 91, 357– 371. L AFRANCE , B. & V ARGA , M. 1996. Structural studies of
H IBBARD , J. P., V AN S TAAL , C. R. & R ANKIN , D. W. the Parker Lake Shear Zone and the Reilly Lake Shear
2007. A comparative analysis of pre-Silurian crustal Zone, Reindeer Lake. Summary of Investigations 1996,
building blocks of the northern and southern Appala- Saskatchewan Geological Survey, Miscellaneous
chians Orogen. American Journal of Science, 307, Report 96-4, 119– 124.
23– 45. L EWRY , J. F. 1981. Lower Proterozoic arc-microcontinent
H IRN , A., L ÉPINE , J.-C. ET AL . 1984. Crustal structure and collisional tectonics in the western Churchill province.
variability of the Himalayan border of Tibet. Nature, Nature, 294, 69– 72.
307, 23–25. L EWRY , J. F. & C OLLERSON , K. D. 1990. The Trans-
H OFFMAN , P. F. 1981. Autopsy of Athapuscow Aulacogen: Hudson Orogen: extent, subdivisions and problems.
a failed arm affected by three collisions. In: C AMP- In: L EWRY , J. F. & S TAUFFER , M. R. (eds) The
BELL , F. H. A. (ed.) Proterozoic Basins of Canada. Early Proterozoic Trans-Hudson Orogen of North
Geological Survey of Canada, Paper, 81-10, 97–101. America. Geological Association of Canada, Special
H OFFMAN , P. F. 1985. Is the Cape Smith Belt (northern Paper, 37, 1– 14.
Quebec) a klippe? Canadian Journal of Earth L EWRY , J. F., S TAUFFER , M. R. & F UMERTON , S. 1981.
Sciences, 22, 1361– 1369. A Cordilleran-type batholithic belt in the Churchill
H OFFMAN , P. F. 1988. United plates of America, the birth Province in northern Saskatchewan. Precambrian
of a craton: early Proterozoic assembly and growth of Research, 14, 277–313.
Laurentia. Annual Review of Earth and Planetary L EWRY , J. F., S IBBALD , T. I. I. & S CHLEDEWITZ , D. C. P.
Science Letters, 16, 543–603. 1985. Variation in character of Archaean rocks in the
H OFFMAN , P. F. 1990. Subdivision of the Churchill western Churchill province and its significance. In:
province and extent of the Trans-Hudson Orogen. In: A YRES , L. D., T HURSTON , P. C., C ARD , K. D. &
L EWRY , J. F. & S TAUFFER , M. R. (eds) The Early Pro- W EBER , W. (eds) Evolution of Archaean Supracrustal
terozoic Trans-Hudson Orogen of North America. Sequences. Geological Association of Canada, Special
Geological Association of Canada, Special Paper, 37, Paper, 28, 239– 261.
15– 39. L EWRY , J. F., T HOMAS , D. J., M AC D ONALD , R. &
H OU , G., S ANTOSH , M., Q IAN , X., L ISTER , G. S. & L I , J. C HIARENZELLI , J. 1990. Structural relations in
2008. Configuration of the late Palaeoproterozoic sup- accreted terranes of the Trans-Hudson Orogen, Sas-
ercontinent Columbia: insights from radiating mafic katchewan: telescoping in a collisional regime? In:
dyke swarms. Gondwana Research, 14, 395– 409. L EWRY , J. F. & S TAUFFER , M. R. (eds) The Early
H ULBERT , L. J., H AMILTON , M. A., H ORAN , M. F. & Proterozoic Trans-Hudson Orogen of North America.
S COATES , R. F. J. 2005. U– Pb zircon and Re–Os Geological Association of Canada, Special Paper,
isotope geochronology of mineralized ultramafic 37, 75– 94.
L EWRY , J. F., H AJNAL , Z. ET AL . 1994. Structure of a M ODELAND , S. & F RANCIS , D. 2010. Enriched mantle
Palaeoproterozoic continent-continent collision zone; components in Palaeoproterozoic alkaline magmas of
a Lithoprobe seismic reflection profile across the the Bravo Lakes formation, central Baffin Island,
Trans-Hudson Orogen, Canada. Tectonophysics, 232, Nunavut. Journal of Petrology, in press.
143– 160. M OLNAR , P. & L YON -C AEN , H. 1988. Some simple phys-
L I , A. & M ASHELE , B. 2008. Crustal structure in the Paki- ical aspects of the support, structure, and evolution of
stan Himalayas. Joint Meeting of The Geological mountain belts. In: C LARK , S. P. ET AL . (eds) Processes
Society of America, Soil Science Society of America, in continental lithospheric deformation. Geological
American Society of Agronomy, Crop Science Society Society of America Special Paper, 218, 179– 207.
of America, Gulf Coast Association of Geological M ORGAN , W. C., B OURNE , J., H ERD , R. K., P ICKETT ,
Societies with the Gulf Coast Section of SEPM. J. W. & T IPPETT , C. R. 1975. Geology of the Foxe
Houston, Texas, Paper 204-4. Fold Belt, Baffin Island, District of Franklin. Report
L UCAS , S. B. 1989. Structural evolution of the Cape Smith of Activities, Part A, April– October 1974, Geological
thrust belt, and the role of out-of-sequence faulting in Survey of Canada, Paper 75-01A, 343– 347.
the thickening of mountain belts. Tectonics, 8, M ORGAN , W. C., O KULITCH , A. V. & T HOMPSON , P. H.
655– 676. 1976. Stratigraphy, Structure and Metamorphism of
L UCAS , S. B., W HITE , D. ET AL . 1994. Three-dimensional the West Half of the Foxe Fold Belt, Baffin Island.
collisional structure of the Trans-Hudson Orogen, Report of Activities Part B/Report of Activities Part
Canada. Tectonophysics, 232, 161–177. A, Geological Survey of Canada, Paper 76-01B,
L UCAS , S. B., S TERN , R. A., S YME , E. C., R EILLY , B. A. 387–391.
& T HOMAS , D. J. 1996. Intraoceanic tectonics and the M UNGALL , J. E. 2007. Crustal contamination of picritic
development of continental crust: 1.92– 1.84 Ga evol- magmas during transport through dikes: the Expo
ution of the Flin Flon belt, Canada. Geological Intrusive Suite, Cape Smith Fold Belt, New Quebec.
Society of America Bulletin, 108, 602–629. Journal of Petrology, 48, 1021– 1039.
M ACHADO , N., D AVID , J., S COTT , D. J., L AMOTHE , D., M URPHY , J. B., K EPPIE , D., D OSTAL , J. & N ANCE , D.
P HILIPPE , S. & G ARIÉPY , C. 1993. U– Pb geochronol- 1999. Neoproterozoic– early Paleozoic evolution of
ogy of the western Cape Smith Belt, Canada: new Avalonia. In: R AMOS , V. & K EPPIE , D. (eds)
insights on the age of initial rifting and arc magmatism. Laurentia– Gondwana Connections before Pangea.
Precambrian Research, 63, 211– 223. Geological Society of America Special Paper, 336,
M AC H ATTIE , T. G. 2001. Petrogenesis of the Wathaman 253–266.
Batholith and La Ronge Domain plutons in the Rein- O RRELL , S. E., B ICKFORD , M. E. & L EWRY , J. F. 1999.
deer Lake area, Trans-Hudson Orogen, Saskatchewan. Crustal evolution and age of thermotectonic reworking
MSc thesis, Memorial University, St. John’s, in the western hinterland of the Trans-Hudson Orogen,
Newfoundland. northern Saskatchewan. Precambrian Research, 95,
M AHAN , K., G ONCALVES , P., W ILLIAMS , M. & J ERCI- 187–223.
NOVIC , M. 2006. Dating metamorphic reactions and P ARRISH , R. R. 1989. U– Pb geochronology of the Cape
fluid flow: application to exhumation of high-P granu- Smith Belt and Sugluk Block, northern Quebec. In:
lites in a crustal-scale shear zone, western Canadian L UCAS , S. B., P ICARD , C. & S T -O NGE , M. R. (eds)
Shield. Journal of Metamorphic Geology, 24, Tectonic, Magmatic and Metallogenic Evolution of
193– 217. the Early Proterozoic Cape Smith Thrust Belt.
M AXEINER , R. O., S IBBALD , T. I. I., S LIMMON , W. L., Geoscience Canada, 16, 126– 130.
H EAMAN , L. M. & W ATTERS , B. R. 1999. Lithogeo- P EHRSSON , S. J. AND THE W ESTERN C HURCHILL
chemistry of volcano-plutonic assemblages of the M ETALLOGENY P ROJECT W ORKING G ROUP , 2004.
southern Hanson Lake Block and Southeastern Evolution of thought on the evolution of a craton:
Glennie Domain, Trans-Hudson Orogen: evidence for new perspectives on the origin and reworking of the
a single island arc complex. Canadian Journal of western Churchill Province. PanLithoprobe IV,
Earth Sciences, 36, 209– 225. Lithoprobe celebratory conference, Toronto, Program
M AXEINER , R. O., H ARPER , C., C ORRIGAN , D. & M AC- with abstracts, poster, CD-ROM.
D OUGALL , D. G. 2004. La Ronge-Lynn Lake Bridge P ERCIVAL , J. A. 2007. Geology and metallogeny of the
Project: geology of the Southern Reindeer Lake area. Superior Province. In: G OODFELLOW , W. D. (ed.)
Saskatchewan Industry and Resources. Open File Mineral Deposits of Canada: A Synthesis of Major
Report 2003-1, 2 CD-ROM set and 3 maps. Deposit-types, District Metallogeny, the Evolution of
M AXEINER , R. O., C ORRIGAN , D., H ARPER , C. T., Geological Provinces, and Exploration Methods. Geo-
M AC D OUGALL , D. G. & A NSDELL , K. 2005. Palaeo- logical Association of Canada, Mineral Deposits Div-
proterozoic arc and ophiolitic rocks on the northwest- ision, Special Publication, 5, 309– 328.
margin of the Trans-Hudson Orogen, Saskatchewan, P ERCIVAL , J. A., W HALEN , J. B. & R AYNER , N. 2004.
Canada; their contribution to a revised tectonic frame- Pikwitonei –Snow Lake, Manitoba transect, Trans-
work for the orogen. Precambrian Research, 136, Hudson Orogen – Superior Margin Metallotect
67–106. Project: initial geological, isotopic, and U –Pb
M EYER , M. T., B ICKFORD , M. E. & L EWRY , J. F. 1992. SHRIMP results. Report of Activities 2004, Manitoba
The Wathaman Batholith: an Early Proterozoic conti- Industry, Economic Development and Mines, Mani-
nental arc in the Trans-Hudson orogenic belt, toba Geological Survey, 120– 134.
Canada. Geological Society of America Bulletin, 104, P ERCIVAL , J. A., W HALEN , J. B. & R AYNER , N. 2005.
1073–1085. Pikwitonei – Snow Lake, Manitoba transect (parts of
NTS 63J, 63O and 63P), Trans-Hudson Orogen – R OSS , G. M. 2002. Evolution of Precambrian continental
Superior Margin Metallotect Project: new results lithosphere in Western Canada: results from Lithop-
and tectonic interpretation. Report of Activities robe studies in Alberta and beyond. Canadian
2005, Manitoba Industry, Economic Development Journal of Earth Sciences, 39, 413– 437.
and Mines, Manitoba Geological Survey, 69–91. S CHNEIDER , D. A., H EIZLER , M. T., B ICKFORD , M. E.,
P ERCIVAL , J. A., R AYNER , N., G ROWDON , M. L., W ORTMAN , G. L., C ONDIE , K. C. & P ERILLI , S.
W HALEN , J. B. & Z WANZIG , H. V. 2007. New field 2007. Timing constraints of orogeny to cratonization:
and geochronological results for the Osik– Atik– thermochronology of the Palaeoproterozoic Trans-
Footprint lakes area, Manitoba (NTS 63O13, 14, 15, Hudson orogen, Manitoba and Saskatchewan, Canada.
64B2, 3). Report of Activities 2007, Manitoba Precambrian Research, 153, 65–95.
Science, Technology, Energy and Mines, Manitoba S COTT , D. J. 1997. Geology, U–Pb and Pb– Pb geochro-
Geological Survey, 71–81. nology of the Lake Harbour area, southern Baffin
P ESONEN , L. J., E LMING , S.-Å. ET AL . 2003. Paleomag- Island: implications for the Palaeoproterozoic tectonic
netic configuration of continents during the Protero- evolution of north-eastern Laurentia. Canadian
zoic. Tectonophysics, 375, 289– 324. Journal of Earth Sciences, 34, 140– 155.
P ICARD , C., L AMOTHE , D., P IBOULE , M. & O LIVIER , R. S COTT , D. J. 1999. U –Pb geochronology of the
1990. Magmatic and geotectonic evolution of a Proter- eastern Hall Peninsula, southern Baffin Island,
ozoic oceanic basin system: the Cape Smith Thrust- Canada: a northern link between the Archaean of
Fold Belt (New Quebec). Precambrian Research, 47, West Greenland and the Palaeoproterozoic Torngat
223–249. Orogen of northern Labrador. Precambrian Research,
R AYNER , N. & C ORRIGAN , D. 2004. Uranium-lead geo- 93, 5 –26.
chronological results from the Churchill River – S COTT , D. J. & S T -O NGE , M. R. 1998. Proterozoic assem-
Southern Indian Lake transect, northern Manitoba. bly of northeast Laurentia revisited: a model based on
Geological Survey of Canada, Current Research, southward extrapolation of Ungava – Baffin crustal
2004-F1, 1 –14. architecture. In: W ARDLE , R. J. & H ALL , J. (eds)
R AYNER , N. M., M AXEINER , R. O. & C ORRIGAN , D. Eastern Canadian Shield Onshore-Offshore Transect
2005a. Progress report on U–Pb geochronology (ECSOOT). Transect Meeting (May 4– 5, 1998), The
results for the Peter Lake Domain Project. In: University of British Columbia, Lithoprobe Secreter-
Summary of Investigations 2005, Saskatchewan Geo- iat, Report No. 68, 134–147.
logical Survey. Miscellaneous Report, 2005-4.2, S COTT , D. J., H ELMSTAEDT , H. & B ICKLE , M. J. 1992.
CD-ROM, paper A-4, 12. Purtuniq ophiolite, Cape Smith belt, northern
R AYNER , N., S TERN , R. A. & B ICKFORD , M. E. 2005b. Québec, Canada: a reconstructed section of Early Pro-
Tectonic implications of new SHRIMP and TIMS terozoic oceanic crust. Geology, 20, 173 –176.
U–Pb geochronology of rocks from the Sask Craton, S COTT , D. J., S TERN , R. A., S T -O NGE , M. R. &
Peter Lake Domain, and Hearne margin, Trans-Hudson M C M ULLEN , S. M. 2002. U– Pb geochronology of
Orogen, Saskatchewan. Canadian Journal of Earth detrital zircons in metasedimentary rocks from
Sciences, 42, 635– 657. southern Baffin Island: implications for the Palaeo-
R AYNER , N., S ANBORN -B ARRIE , M., W ODICKA , N. & proterozoic tectonic evolution of Northeastern
S T -O NGE , M. 2007. New U– Pb geochronological Laurentia. Canadian Journal of Earth Sciences, 39,
constraints on the timing of deformation and the 611– 623.
nature of basement of SW Baffin Island, Nunavut. S TAUFFER , M. R. 1984. Manikewan and early Proterozoic
In: C AIRNS , S. & F ALCK , H. (compilers), 35th ocean in central Canada, its igneous history and
Annual Yellowknife Geoscience Forum Abstracts. orogenic closure. Precambrian Research, 25,
Northwest Territories Geoscience Office, Yellowknife, 257– 281.
NT. YKGSF Abstracts 49, 2007. S TAUFFER , M. R. 1990. The Missi Formation: an Aphe-
R AYNER , N. M., S COTT , D. J., W ODICKA , N. & K ASSAM , bian molasse deposit in the Reindeer Lake Zone of
A. 2008. New Geochronological Constraints from Mill, the Trans-Hudson Orogen, Canada. In: L EWRY , J. F.
Salisbury, and Nottingham Islands, Nunavut. Geologi- & S TAUFFER , M. R. (eds) The Early Proterozoic
cal Survey of Canada, Current Research 2008-22, 1– 14. Trans-Hudson Orogen of North America. Geological
R IVERS , T. 1997. Lithotectonic elements of the Grenville Association of Canada, Special Paper, 37, 75–94.
Province: review and tectonic implications. Precam- S TERN , R. A., S YME , E. C. & L UCAS , S. B. 1995. Geo-
brian Research, 86, 117 –154. chemistry of 1.9 Ga MORB- and OIB-like basalts
R IVERS , T. & C ORRIGAN , D. 2000. Convergent margin on from the Amisk collage, Flin Flon belt, Canada: evi-
southeastern Laurentia during the Mesoproterozoic: dence for an intra-oceanic origin. Geochimica et Cos-
tectonic implications. Canadian Journal of Earth mochimica Acta, 59, 3131–3154.
Sciences, 37, 359– 383. S TERN , R. A., M ACHADO , N., S YME , E. C., L UCAS , S. B.
R OGERS , J. J. W. & S ANTOSH , M. 2002. Configuration of & D AVID , J. 1999. Chronology of crustal growth and
Columbia, a Mesoproterozoic Supercontinent. Gond- recycling in the Palaeoproterozoic Amisk collage
wana Research, 5, 5 –22. (Flin Flon Belt), Trans-Hudson Orogen, Canada.
R OKSANDIC , M. M. 1987. The tectonics and evolution of Canadian Journal of Earth Sciences, 36, 1807– 1827.
the Hudson Bay region. In: B EAUMONT , C. & S T -O NGE , M. R. & I JLEWLIW , O. J. 1996. Mineral corona
T ANKARD , A. J. (eds) Sedimentary Basins and Basin- formation during high-P retrogression of granulitic
forming Mechanisms. Canadian Society of Petroleum rocks, Ungava Orogen, Canada. Journal of Petrology,
Geologists, Memoir, 12, 507–518. 37, 553 –582.
S T -O NGE , M. R. & L UCAS , S. B. 1990. Evolution of the T HOMAS , M. D. & K EAREY , P. 1980. Gravity anomalies,
Cape Smith Belt: early Proterozoic continent under- block-faulting and Andean-type tectonism in the
thrusting, ophiolites obduction and thick-skinned Eastern Churchill province. Nature, 283, 61– 63.
folding. In: L EWRY , J. F. & S TAUFFER , M. R. (eds) T RAN , H. T., A NSDELL , K., B ETHUNE , K. M., A SHTON ,
The Early Proterozoic Trans-Hudson Orogen of K. & H AMILTON , M. A. 2008. Provenance and tec-
North America. Geological Association of Canada, tonic setting of Palaeoproterozoic metasedimentary
Special Paper, 37, 313–351. rocks along the eastern margin of the Hearne craton:
S T -O NGE , M. R., L UCAS , S. B. & P ARRISH , R. R. 1992. constraints from SHRIMP geochronology, Wollaston
Terrane accretion in the internal zone of the Ungava Group, Saskatchewan, Canada. Precambrian
orogen, northern Quebec. Part 1: tectonostratigraphic Research, 167, 171–185.
assemblages and their tectonic implications. Canadian T UREK , A., W OODHEAD , J. & Z WANZIG , H. V. 2000.
Journal of Earth Sciences, 29, 746– 764. U–Pb age of the gabbro and other plutons at Lynn
S T -O NGE , M. R., S COTT , D. J. & L UCAS , S. B. 2000. Lake (parts of NTS 64C). Report of Activities 2000,
Early partitioning of Quebec: microcontinent for- Manitoba Industry, Trade and Mines, Manitoba
mation in the Palaeoproterozoic. Geology, 28, Geological Survey, 97– 104.
323– 326. V AN S CHMUS , W. R., B ICKFORD , M. E., L EWRY , J. F. &
S T -O NGE , M. R., S COTT , D. J. & W ODICKA , N. 2002. M ACDONALD , R. 1987. U–Pb geochronology in the
Review of crustal architecture and evolution in the Trans-Hudson Orogen, northern Saskatchewan,
Ungava Paninsula – Baffin Island area: connection to Canada. Canadian Journal of Earth Sciences, 24,
the Lithoprobe ECSOOT transect. Canadian Journal 407–424.
of Earth Sciences, 39, 589–610. V AN DER L EEDEN , J., B ÉLANGER , M., D ANIS , D.,
S T -O NGE , M. R., S EARLE , M. P. & W ODICKA , N. 2006. G ITRARD , R. & M ARTELAIN , J. 1990. Lithotectonic
Trans-Hudson Orogen of North America and domains in the high-grade terrain east of the Labrador
Himalaya-Karakoram-Tibetan Orogen of Asia: Struc- Trough (Québec). In: L EWRY , J. F. & S TAUFFER ,
tural and thermal characteristics of the lower and M. R. (eds) The Early Proterozoic Trans-Hudson
upper plates. Tectonics, 25, TC4006, doi: 10.1029/ Orogen of North America. Geological Association of
2005TC001907. Canada, Special Paper, 37, 371 –386.
S T -O NGE , M. R., W ODICKA , N. & I JEWLIW , O. 2007. V AN S TAAL , C. R. 2005. Northern Appalachians. In:
Polymetamorphic evolution of the Trans-Hudson S ELLEY , R. C., C OCKS , R. M. & P ILMER , I. R. (eds)
Orogen, Baffin Island, Canada: integration of petrolo- Encyclopedia of Geology. Oxford, Elsevier, 4, 81–91.
gical, structural, and geochronological data. Journal W ARDLE , R. J. & H ALL , J. 2002. Proterozoic evolution of
of Petrology, 48, 271 –302, doi: 10.1093/petrology/ the northeastern Canadian Shield: Lithoprobe Eastern
egl060. Canadian Shield Onshore-Offshore Transect
S TOCKWELL , C. H. 1961. Structural provinces, orogenies (ECSOOT), introduction and summary. Canadian
and time classification of rocks of the Canadian Journal of Earth Sciences, 39, 563–567.
Precambrian Shield. In: L OWDON , J. A. (ed.) Age W ARDLE , R. J., R YAN , B., N UNN , G. A. C. & M ENGEL , F.
Determinations by the Geological Survey of Canada, C. 1990. Labrador segment of the Trans-Hudson
Report 2: Isotopic Ages. Geological Survey of orogen: crustal development through oblique conver-
Canada, Paper 61–17, 108– 118. gence and collision. Geological Association of
S YME , E. C. & B AILES , A. H. 1993. Stratigraphic and Canada Special Paper, 37, 353–369.
tectonic setting of volcanogenic massive sulphide W ARDLE , R. J., J AMES , D. T., S COTT , D. J. & H ALL , J.
deposits, Flin Flon, Manitoba. Economic Geology, 2002. The southeastern Churchill Province: synthesis
88, 566–589. of a Palaeoproterozoic transpressional orogen. Cana-
S YME , E. C., L UCAS , S. B., B AILES , A. H. & S TERN , dian Journal of Earth Sciences, 39, 639–663.
R. A. 1999. Contrasting arc and MORB-like asse- W ATTERS , B. R. & P EARCE , J. A. 1987. Metavolcanic
mblages in the Palaeoproterozoic Flin Flon Belt, rocks of the La Ronge Domain in the Churchill Pro-
Manitoba, and the role of intra-arc extension in vince, Saskatchewan geochemical evidence for a vol-
localizing volcanic-hosted massive sulphide canic arc origin. In: P HARAOH , T. C., B ECKINSALE ,
deposits. Canadian Journal of Earth Sciences, 36, R. D. & R ICKARD , D. (eds) Geochemistry and Miner-
1767–1788. alization of Proterozoic Volcanic Suites. Geological
S YMONS , D. T. A. & H ARRIS , M. J. 2005. Accretion Society, London, Special Publications, 33, 167–182.
history of the Trans-Hudson Orogen in Manitoba and W HALEN , J. B., S YME , E. C. & S TERN , R. A. 1999. Geo-
Saskatchewan from paleomagnetism. Canadian chemical and Nd isotopic evolution of Palaeoprotero-
Journal of Earth Sciences, 42, 723– 740. zoic arc-type granitoids magmatism in the Flin Flon
T HÉRIAULT , R. J., S T -O NGE , M. R. & S COTT , D. J. 2001. Belt, Trans-Hudson Orogen, Canada. Canadian
Nd isotopic and geochemical signature of the Palaeo- Journal of Earth Sciences, 36, 227–250.
proterozoic Trans-Hudson Orogen, southern Baffin W HALEN , J. B., W ODICKA , N. & T AYLOR , B. E. 2008.
Island, Canada: implications for the evolution of Cumberland Batholith petrogenesis: implications for
eastern Laurentia. Precambrian Research, 108, Palaeoproterozoic crustal and orogenic processes.
113– 138. Geochimica Cosmochimica et Acta, Goldschmidt Con-
T HOMAS , D. J. 1993. Geology of the Star Lake-Otter Lake ference Abstracts 2008, 72, A1016.
Portion of the Central Metavolcanic Belt, La Ronge W HITE , D. J., Z WANZIG , H. V. & H AJNAL , Z. 2000. Crustal
Domain. Saskatchewan Geological Survey, Miscella- suture preserved in the Palaeoproterozoic Trans-
neous Report, 236, 1 –133. Hudson Orogen, Canada. Geology, 29, 527– 530.
W HITE , D. J., L UCAS , S. B., B LEEKER , W., H AJNAL , Z., Trans-Hudson Orogen in Ungava Peninsula. In:
L EWRY , J. F. & Z WANZIG , H. V. 2002. Suture-zone S T -O NGE , M. R., L AMOTHE , D., H ENDERSON , I. &
geometry along an irregular Palaeoproterozoic F ORD , A. (eds) Atlas géoscientifique numérique, cein-
margin: the Superior Boundary Zone, Manitoba, ture de Cape Smith et environs, péninsule d’Ungava,
Canada. Geology, 30, 735– 738. Québec-Nunavut/Digital geoscience atlas of the
W HITE , D. J., T HOMAS , M. D., J ONES , A. G., H OPE , J., Cape Smith Belt and adjacent domains, Ungava Penin-
N ÉMETH , B. & H AJNAL , Z. 2005. Geophysical trans- sula, Quebec-Nunavut. Geological Survey of Canada,
ect across a Palaeoproterozoic continent-continent col- Open File, 5117.
lision zone: the Trans-Hudson Orogen. Canadian W ODICKA , N., W HALEN , J. B., J ACKSON , G. D. &
Journal of Earth Sciences, 42, 385 –402. H EGNER , E. 2007c. Geochronological compilation of
W HITMEYER , S. J. & K ARLSTROM , K. E. 2007. Tectonic Trans-Hudson Orogen and environs, Baffin Island.
model for the Proterozoic growth of North America. In: S T -O NGE , M. R., F ORD , A. & H ENDERSON , I.
Geosphere, 3, 220–259. (eds) Digital geoscience atlas of Baffin Island (south
W ODICKA , N. & S COTT , D. J. 1997. A preliminary report of 708N and east of 808W), Nunavut. Geological
on the U–Pb geochronology of the Meta Incognita Survey of Canada, Open File, 5116.
Peninsula, southern Baffin Island, Northwest Terri- W ODICKA , N., S T -O NGE , M. R. & W HALEN , J. B. 2008.
tories. Geological Survey of Canada, Paper, 1997-C, Characteristics of two opposing continental margin
167–178. successions in northeast Laurentia. Geochimica Cos-
W ODICKA , N., S T -O NGE , M. R., S COTT , D. J. & C ORRI- mochimica et Acta, Goldschmidt Conference Abstracts
GAN , D. 2002a. Preliminary report on the U–Pb geo- 2008, 72, A1030.
chronology of the northern margin of the Y EO , G. 1998. A systems tract approach to the stratigra-
Trans-Hudson Orogen, central Baffin Island, phy of paragneisses in the southeastern Wollaston
Nunavut. Radiogenic Age and Isotopic Studies, Domain. Summary of Investigations 1998. Saskatche-
Report 15, Geological Survey of Canada, Current wan Geological Survey, Saskatchewan Energy and
Research, 2002-F7, 12. Mines, Miscellaneous Report, 98-4, 36– 47.
W ODICKA , N., M ADORE , L., L ARBI , Y. & V ICKER , P. Z WANZIG , H. 1990. Kisseynew gneiss belt in Manitoba:
2002b. Géochronologie U–Pb de filons-couches mafi- stratigraphy, structure, and tectonic evolution. In:
ques de la Ceinture de Cape Smith et de la Fosse du L EWRY , J. F. & S TAUFFER , M. R. (eds) The Early Pro-
Labrador: Dans: L’exploration minérale au Québec: terozoic Trans-Hudson Orogen of North America.
notre savoir, vos découvertes. Séminaire d’information Geological Association of Canada, Special Paper, 37,
sur la recherche géologique, Programme et résumés 95–120.
2002. Ministère des Ressources naturelles, Québec, Z WANZIG , H. V. 2000. Geochemistry and tectonic frame-
DV 2002-10, 48. work of the Kisseynew Domain – Lynn Lake belt
W ODICKA , N., S T -O NGE , M. R., C ORRIGAN , D. & boundary (part of NTS 64P/13). Report of Activities
S COTT , D. J. 2007a. Depositional age and provenance 2000, Manitoba Industry, Trades and Mines, Manitoba
of the Piling Group, central Baffin Island, Nunavut: Geological Survey, 102–114.
implications for the Palaeoproterozoic tectonic Z WANZIG , H. V., S YME , E. C. & G ILBERT , H. P. 1999.
development of the southern Rae margin. Geo- Updated trace element geochemistry of the c. 1.9 Ga
logical Association of Canada – Mineralogical Asso- metavolcanic rocks in the Palaeoproterozoic Lynn
ciation of Canada Annual Meeting, Abstracts, 32, 88. Lake Belt. Manitoba Industry, Trade and Mines,
W ODICKA , N., B REITSPRECHER , K. & W HALEN , Geological Services, Open File Report, 99-13, 1 –46,
J. B. 2007b. Geochronological compilation of plus accompanying map and disk.
Index
Page numbers in italic denote figures. Page numbers in bold denote tables.
Acadian orogenesis 274, 286, 291– 302, 305 Appalachians

Acadian slab 299, 300 northern microcontinents 276– 277
Acatlán Complex 239, 240, 256, 347 oceanic terranes 277
mafic rocks 346–350 apparent polar wander path, Laurentia 375, 376– 378,
palaeogeography 261– 265, 263, 264 379, 380, 383
U–Pb zircon analysis 241, 243– 256, Apulia, plate tectonics 90, 91, 92, 93
257–260, 348 Apulian–Hyblean foreland 115, 115, 117
40
Ackley Granite 272 Ar/39Ar dating, Capricorn Orogen 446– 451
Aconcagua fold and thrust belt 35, 37, 46 Arabian Platform 147, 149
Adria 90, 91, 92, 93, 117 Aracena massif 217, 219–220, 227
Aeolian Islands 114, 118 Aracena Metamorphic Belt 354
African–Anatolian plate boundary 127 –152 Arauco Basin 40
Akdag Massif 139, 140, 141 Arisaig Group 272, 275, 291
Albany–Fraser Belt 379– 380, 445 Armorican microplate 359
Alleghanian orogeny 164, 271, 273, 276, 346 Ası́s Lithodeme 243, 246, 257, 349
Allochthon Boundary Thrust 406, 407– 408, 411, asthenosphere, upwelling 216
419– 421, 421– 423, 425 Anatolia 133, 137– 138, 144, 145, 150, 151
hanging wall 423, 425, 429, 434 Appalachians 284, 286, 291
Orogenic Lid 432 Australia
Allochthonous High Pressure Belt 425, 426, 433, 436 Mawsonland 379– 380
Allochthonous Medium-Low Pressure Belt 425, Rodinia reconstruction 382, 384
427– 428, 429, 436 Austroalpine domain 93, 100, 103
Alpine cycle 104– 106 Austroalpine–Carpathian orogen 97–98
Alpine realm, plate tectonics 89– 107 Avalonia 276– 277, 318
Alpine Tethys 90, 95, 96, 97, 98, 103 accretion of Meguma 302– 303
subduction 105– 106 accretion to Laurentia 291–293, 300
Altaides 162, 163 Lizard Complex 355
Altiplano flat-slab segment 41–42, 47, 48
Amate Unit 244, 247, 249, 256, 258 Badger Basin 272, 285
Amazonia Baffin Suture 460–461, 468
collision with Laurentia 406 Baie Verte oceanic tract 272, 275, 277
Rodinia reconstruction 382, 389– 391 Baie Verte–Brompton Line 272, 281, 283
ANACONDA reconstruction 381, 383–384 Baltica 161, 163
Anatolia magnetic anomalies 167, 169
central orogenic belt passive margin 171 –173, 180
geology 139, 140, 141–143 rifting 167, 384
tectonic evolution 143, 144, 145 Rodinia reconstruction 377– 378, 381
eastern, geology 145–146, 147, 148– 151 Bamford Brook fault 272, 275, 289, 290
western Bangemall Supergroup 446, 447, 451, 452
geology 134, 135, 136 Basal Unit 352
tectonic evolution 136–139 Lizard Complex 355
Anatolian plate basalt, K Trig 15
boundary with African plate 127–152 basalt flooding 48
Anatolide block 134, 135, 138 Baskil arc 147, 148
Andes Bazar ophiolite 319, 320, 324– 325, 353
flat-slab subduction 31–48 origin and history 334– 335
arc magmatism 33, 35, 35 Benedict Fault 415, 417
sedimentation 35, 36 Benioff zone
tectonics 32, 35– 36 Chiapas fold-and-thrust belt 66, 67
uplift 46 see also Wadati– Benioff zone
Annieopsquotch accretionary tract 277, 279, 281, 283, Bergeron Suture 462 –463, 474
284, 305 Bermejo foreland basin 46
Annieopsquotch ophiolite belt 275 Big Island Suture 462, 471, 474
Apennine–Maghrebide peninsula 123 Bitlis massif 146, 146, 147, 148–149, 150, 151
Apennine–Sicilian fold-and-thrust belt 113–117 Bohemian Massif 197, 198, 318
buckling 119, 121– 123 Rheic Ocean mafic rocks 356, 357
palaeogeographic evolution 116, 120 U– Pb zircon analysis 204, 205–206, 208, 209, 210
Appalachian orogenesis 4, 164, 271–308, 345 Variscan orogeny 197 –212, 211, 356
482 INDEX
Bolkar Mountains 142 Cookson Group 272, 290

Bouguer gravity anomaly, Lake Taupo 24 Cordillera Blanca 37, 46
Bragança Complex 353 Cordillera de Marañón 37, 46
Briançonnais terrane 89, 91, 92, 99, 102, 105, 106 Coromandel Volcanic Zone 9, 12
Bridge River terrane 81, 83 Cosoltepec Formation 244, 250, 251, 256, 259, 264
Bucaramanga segment 32, 37, 39– 40, 46, 47 mafic rocks 349–350
buckling, Apennine–Sicilian fold-and-thrust belt 119, cratons, reconstruction 375–376
121– 123 Cretaceous
Burgeo batholith 272 Alpine tectonics 97–104
Burntwood back-arc basin 469–470 oceans 99–102
crystallization, assimilation fractional 142, 143, 144
Cabo Otegal Complex 320, 321 Cumberland Batholith 468, 469, 471
Cabot fault 272 Cut Throat Island Fault 415, 417
Cache Creek terrane 71, 72 Cyprean Arc 128, 132
geological setting 73, 75, 76 subduction zone 131 –132, 133
high-pressure rocks 81–82
extrusion model 82– 84 Dashwoods microcontinent 276, 277, 278, 279,
stratigraphy 74 281, 283, 303 –304, 306
Cadomian orogeny 199, 209 De Pas Batholith 469, 471
Calabrian nappes 115 deformation
Calabrian orocline 3, 113–117, 118 Acadian orogeny 295– 296
tectonic model 119, 118– 120 Andes 48
Calabrian subduction zone 117– 118 Ossa Morena Zone 216–234
Calabrian– Peloritanian nappes 114 Salinic orogeny 288, 289– 291
caldera volcanoes, Taupo Volcanic Zone 12, 16–25 Saxo-Thuringian Zone 201, 202
Caledonia fault 272, 291 Taconic orogeny 281, 306
Canadian Cordillera 71– 84, 72 Urals 182
Canadian Shield 457, 458 Dewar Formation 75
Canoas Unit 244, 247, 249, 256, 258, 350 Dog Bay Line 272, 289–290, 293, 297
Cantabrian zone 350, 351 Döhlen Basin 202, 203, 204, 207, 209, 211–212
Cape Ray fault–Victoria Lake shear zone 296 Dover fault 272, 291
Cappadocian volcanic province 140, 143 Dunn Point volcanics 275
Capricorn Orogen 5 –6, 445– 454, 446 Dunnage mélange 284, 285
Carballo– Bazar slice 324 dykes, mafic
Careón ophiolite 319, 320, 327– 329, 353 Appalachians 276–277
origin and history 333, 336–338 Grenville Front 413, 415– 416, 418, 423
Carnegie Ridge 32, 40 Trans-Hudson Orogen 467, 471
Carreiro shear zone 331, 333, 338 West Australian Craton 452
Central Anatolian Crystalline Complex 139, 140, see also Sudbury dykes
141– 142
late-Cretaceous plutons 142, 143, 144 East Anatolian High Plateau 145, 146, 148
Central Asian ocean 161 deformation and volcanism 149, 151
Central Atlantic opening 95, 96 tectonic evolution 150, 151
Central Uralian megazone 166 East Anatolian subduction-accretion complex
tectonostratigraphy 168 148 –149, 150, 151
Chain Lakes Massif 272 East Magnitogorsk Fault 165, 166, 185
channel flow, Grenville Province 406, 431, 432, tectonostratigraphy 168
434, 435, 436, 437 East Magnitogorskian mélange zone 166
Chazumba Formation 244, 252, 254, 256, 260 East Pontide arc 145– 146, 146
Chiapas fold-and-thrust belt 55–67, 56 East Uralian zone 165, 166, 185, 188
China tectonostratigraphy 168
North, Rodinia reconstruction 382, 388 –389, 392 Eastern Cordillera 39–40, 41, 42, 46
South, Rodinia reconstruction 382, 386 –388 Eastern Turkey Seismic Experiment 151
Choiyoi Group 44– 45 Edgecumbe volcanoes 13
Cilo ophiolite 147, 149 Edmundian Orogenesis 446, 451–452, 454
Cimmerian orogeny 182 El Epazote Unit 244, 247, 249, 256, 258, 261
Coaker porphyry 275 El Rodeo Formation 244, 247, 249, 256, 258, 262
Coastal arc 292, 294 Elazig– Palu nappe 147, 148
Coatlaco Unit 244, 250, 251, 256, 259, 350 Elbe Zone 198, 201–204, 203
Cobequid –Chedabucto fault 272, 303 exhumation 204, 212
COBRA reconstruction 390– 391 geochronology 209–212
Cold Spring Pond mélange 275 strike-slip processes 201–202, 202, 211
Congo craton, Rodinia reconstruction 380–381, U–Pb zircon analysis 204, 205– 206, 208, 209, 210
382, 383 –384 Eocene, Alpine tectonics 103, 106
INDEX 483
erosion, subduction Gondwana

Cache Creek HP rocks 82– 83 Appalachian microcontinents 276– 277
Chiapas fold-and-thrust belt 64 collision with Avalonia 216
Errabiddy Shear Zone 449– 450, 451, 452 collision with Laurussia 198, 319, 321
Erzgebirge nappe complex 201, 202 Mexican terranes 264 –265
Esperanza granitoids 243, 244, 245, 247 palaeogeography 345
ESRU–SB profile 185, 186, 187 Gondwanaland 391, 393– 394
Europe, Eastern, Rheic Ocean mafic rocks graben systems, western Anatolia 135, 138
357– 358 granite, biotite, Evora massif, U–Th– Pb analysis
Evora Massif 217, 218, 219, 227 222, 226, 229, 230, 232
U–Th– Pb analysis 222 –230, 232 granite magmatism, Urals 181
Évora–Aracena–Lora del Rı́o metamorphic belt Granite Supersuite plutons 140, 142, 144
216– 217, 232 granitoids
exhumation Acatlán Complex 244, 245, 247, 248, 256, 257,
Anatolia 138 258, 259, 261
Elbe Zone 204, 212 Appalachians 273, 284
Grenville Front 417, 420, 433– 434, 437 Elbe Zone 203– 204, 209
Ossa– Morena Zone 216, 232 Lausitz Block 202, 209
Piaxtla Suite 350 Ossa Morena Zone 217
Salinic 290 Granjeno Schist 240, 252, 255, 256, 257, 260, 261– 265
San Rafael block 43 Granville Lake Structural Zone 468–469, 472
Taconic 284 Green Bay fault 272
extension greenstone belt, Grenville Front 414, 416, 418
Andes 48 Grenville APW loop 376 –377, 385
eastern Mediterranean 137, 138– 139, 145 Grenville foreland 413, 415– 417, 418
Variscan, SW Iberian Massif 231–234 Grenville Front 406, 407, 409– 411, 413– 414,
414–417, 418, 420
faulting exhumation 417, 420, 433– 434, 437
Central Anatolian Crystalline Complex 145 Lithoprobe seismic study 409– 410, 411
East Anatolian High Plateau 149 Grenville Front Tectonic Zone 409, 410–411, 413,
Sonda de Campeche 63 414–416, 417, 418
STEP 133 Grenville Orogen 406 –407, 407, 420
Taupo Volcanic zone 13– 16, 14 Grenville Province
Fitzcarrald arch 37, 38 channel flow 406, 431, 432, 434, 435, 436, 439
Flin Flon– Glennie Complex 460, 463– 464, 466– 467, general features 406–409, 407
469– 470, 471, 472 interior, metamorphism and deformation 421 –429
flysch large hot long-duration orogen 5, 405– 438
Alpine 104 metamorphism 406– 409
Anatolian orogenic belt 136 Ottawan event 421– 429, 430, 432–433, 435
Uralides 164–165, 176, 181 Rigolet event 409–420, 412, 429, 432–433, 435
Variscides 164 northwest, metamorphism and deformation
Fogo Island pluton 272, 291, 297 409 –420
fold-and-thrust belts 2 Orogenic Lid 423, 429, 432, 434
Apennine–Sicilian 113–117 rebound 437
Chiapas 55–67, 56 Guañacos segement 40
Fournier Group 272 Guleman ophiolite 147, 148, 150, 151
Foxe Orogeny 466, 471
Fredericton trough 290 hanging wall, Allochthon Boundary Thrust 423,
Frontal Cordillera 33, 35, 43, 44, 45 425, 429, 434
Hearne Craton 460, 465–466, 467, 470, 472
Gagnon terrane 415, 417, 421 see also Western Churchill Province
Gander Group 290 Hellenic Arc 128, 132
Gander margin 276, 285, 289, 294–295 subduction zone 129, 130– 134
Gander River ultrabasic belt 272 Hermitage Bay fault 291
Ganderia 276 –277, 285, 286, 291, 300 Hikurangi Subduction Deformation Front 10
Ganderia– Avalonia suture 291, 294 Hikurangi Trough 10, 11
Gascoyne Complex 449– 450, 451, 452 Himalaya–Tibet Orogen 440
geochemistry, Takla Group 77–81, 82 Hodges Hill pluton 272, 275, 291
geochronology Huerta Unit 244, 247, 249, 256, 258, 261, 264
Elbe Zone 209–212 Huiznopala Gneiss 240, 241
Evora massif 222, 223, 224, 225, 226, 227–228, Humber margin 283, 291
229, 230 Humber seaway 278, 279
Ossa Morena Zone 222, 223, 224 Hungry Mountain Thrust 280, 290
Gevas ophiolite 147, 148 Huron Supergroup 413, 414– 415
484 INDEX
Iapetus Ocean 240, 343–344, 345 collision with Amazonia 406

Iberia Rodinia reconstruction 374, 381
geology 350–351 Laurussia 163, 171
NW collision with Kazakhstania 181, 188
ophiolites 317– 338, 320, 352–353, 359 passive margin 171, 172
Variscan extensional tectonics 233–234 collision with Magnitogorsk arc 176, 177,
Rheic Ocean mafic complexes 351– 355 178–179, 180
SW Lausitz Block 201, 204, 209
mafic complexes 354–355 granitoids 202, 209
Variscan extensional tectonics 216–234 U–Pb zircon analysis 204, 205– 206, 208, 209, 210
Iberian Pyrite Belt 354 Lausitz Thrust 203
India, Rodinia reconstruction 382, 386, 387– 388, 392 Letovice ophiolite 357
Inner-Tauride Ocean 143, 144 Liberty– Orrington thrust 275, 289, 290
Inner-Tauride suture zone 140, 141 Ligurides 114, 115
Interior Magmatic Belt 407, 433 Lithoprobe seismic study 273, 409–410, 411, 430, 431
Intermontane terranes 71, 72 Lizard Complex, mafic rocks 355 –357
Internal Platform 114–115 Lloyd’s River fault 290, 306
Ionian oceanic slab, subduction 117– 118 Lobster Cove fault 288, 290, 306
Isparta Angle 128, 129, 132–133 Long Lake volcanics 275
Ispendere ophiolite 147, 148, 151 Long Range ultramafic complex 278, 280
Italy Loon Bay pluton 275
Calabrian orocline 113 –123 Lora del Rı́o massif 217, 220– 221, 227
plate tectonics 90–93 Lushs Bight oceanic tract 272, 275, 277–278
tectonic elements 114–118 Lycian ophiolite nappes 135, 136
Ixcamilpa blueschists 243, 246, 258, 261, 262
Ixtaltepec Formation 241, 242, 256, 257 Maden Complex 147, 149, 151
Izmir–Ankara suture zone 90, 134, 135, 136, 141 mafic complexes, Rheic Ocean 346– 360
Magdalena Migmatite 244, 252, 254, 256, 260, 261
Juan Fernández ridge 32, 35 Maghrebides 120, 120, 122
Jurassic, Alpine realm, plate tectonics 94, 95, 96, 97 magmatism
arc
K Trig Basalts 15 Appalachian 273
Kaimai caldera 9, 12 Acadian 291–302
Kaingaroa Ignimbrite 20 Salinic 286, 289, 291, 292
Kalahari craton, Rodinia reconstruction 382, 384 –386 Taconic 281, 283–284
Kapenga Caldera 12, 16 Calabrian subduction zone 118
Kartaly (Troitsk) Fault 165, 166, 185 Chiapas fold-and-thrust belt 57
Kazakhstania 161, 163 Pampean flat-slab segment 33, 35, 35, 46
collision with Laurussia 181, 188 Trans-Hudson Orogen 467–468
Kazakhstanides 162, 163 Bohemian Massif 199, 200, 211
Kazdag Massif 134, 135, 138 Central Anatolian Crystalline Complex 143, 144
Keban metamorphics 147, 148 Neoarchian 303
Kibaran Belt 380, 383 Ossa Morena Zone 221–222
Killan Imbricate Unit 147, 149 slab-breakoff, Anatolia 137–139, 144, 145, 151
Kingston arc terrane 292 Trans-Hudson Orogen 469
Kirsehir Massif 139, 140, 141 Ural mountains 165, 174, 181–182
Kisseynew Domain 468 –470 Magnitogorsk arc 174, 175– 176, 185
Kömürhan ophiolite 15, 147, 148 collision with Laurussia 176, 177, 178–179, 180
Kungurian evaporites 165 Magnitogorsk zone, tectonostratigraphy 168
Küre basin 94, 96 Main Granitic Axis 165, 166, 180–181
Main Uralian Fault 165, 176, 179, 185, 188
La Poile basin 293 suture zone 166, 167, 180
La Ronge– Lake Lynn domains 460, 464– 465, 472 tectonostratigraphy 168
Lagonegro –Imerese– Sicanian basin 115– 116, 120, 121 Maksutovo Complex 177, 178, 179
Lake Lynn domain 462, 466–467 Malpica– Tui Complex 233–234
Lake Taupo, Bouguer gravity anomaly 24 Mamaku Ignimbrite 17–18, 19
large igneous province, Uralo– Siberian 181–182 Manawahe volcano 13
Las Calaveras Unit 244, 247, 249, 256, 258 Mangakino caldera 9, 12, 16
Laurentia 71, 72 Manikewan Ocean 457, 458, 460 –461, 462, 463, 467,
accretion of Avalonia 277, 291– 293 468 –470
accretion of Meguma 302– 303 marine basins, Ossa Morena Zone 222
Appalachian margin 273, 276 Markerbach Granite 203, 209, 210, 211, 212
apparent polar wander path reconstruction 375, Maroa Dome Complex 16, 17, 19
376–378, 379, 380, 383 Mascarene basin 293, 297
INDEX 485
Massif Central, Rheic Ocean mafic rocks 358 Namaqua– Natal belt 385
Matahina Ignimbrite 21 Narryer terrane 450–451, 452
Matapedia fore-arc basin 287, 289 Narsajuaq Arc 459, 462, 469, 471
Mawsonland 379– 380, 384, 392 gneissic suite 459, 468, 471
Mediterranean Nazca Ridge 32, 36, 37, 38
eastern Neoacadian orogeny 286, 302–303, 305
seismic structure 129– 133 Neotethys 89, 90– 91, 94
tectonic evolution 136–139 Ngongotaha rhyolite dome complex 18, 19
western, palaeogeographic evolution 122, 123 Nicola Group 73, 81, 82
Meguma 277, 286, 302– 303, 304 Niederbobritzsch granite 203– 204
Meissen Massif 203, 209, 210, 211, 212 Nigde Massif 139, 140, 141
mélange North Island Dextral Fault Belt 10, 13
Appalachians 278, 280, 284, 287, 289, 290 North Island Shear Belt 13
Somozas ophiolite 319, 320, 321– 323 Northampton Complex 452, 453
Meliata– Maliac slab 95–96, 97 Notre Dame arc 278, 281, 282, 284, 285, 292
Menderes Massif 135, 136, 138 Novillo Gneiss 240, 241, 256, 261
Meta Incognita micro-continent 459, 462, 466, Nuna 374, 391
468, 469, 472
metamorphism Oaxacan Complex 240, 245, 256, 262, 346, 349
Acatlán Complex 348– 349, 359 U– Pb zircon analysis 241, 242, 257, 261
Anatolian orogenic belt 136 Oaxaquia terrane 239, 240, 256, 346, 349
Appalachians palaeogeography 261–265, 263
Acadian 295, 296 U– Pb zircon analysis 241, 242
Salinic 290– 291 Officer Basin 446, 451– 452, 453
Taconic 278– 279, 281, 283 Ohakuri Caldera 19, 24
Bohemian Massif 199, 211, 359 Okataina Caldera Complex 12, 13, 15, 21– 22,
Cache Creek 75, 83 24–25
Central Anatolian Crystalline Complex 139, Olinalá Formation 244, 252, 253, 256, 259, 263
141, 143 ophiolites
Grenville Province 406–432 Appalachians 281, 283
Magnitogorsk arc 176, 177, 178, 179 East Anatolia Plateau 147, 148– 149
Menderes Massif 136 Eastern Europe 357, 359
NW Iberia 321, 353 Lizard Complex 355– 357, 359
Oaxacan complex 241 NW Iberia 317–338, 320, 353, 359
Ossa– Morena Zone 216–221, 230–234 Palaeouralian ocean 170–171, 188
metasedimentary rocks, Mexico 239– 265 Tethyan
Mexico IASZ 135, 136, 141
Acatlán Complex 346– 350 ITSZ 140, 141– 142, 143
Chiapas fold and thrust belt 55– 67, 56 Órdenes Complex 320, 324, 325, 327, 352, 353
palaeogeography 261– 265 oroclines, Calabrian 113–123
U–Pb zircon analysis 240–257 Orogenic Lid, Grenville Province 423, 429,
Mid-German Crystalline Zone 199, 356, 357 432, 434
Mid-Saxon Fault 202, 203 orogens
Middle America Trench 56, 57–58, 66, 67 ancient 3 –6
Miocene, Alpine tectonics 105, 106 collisional 405 –406
Mirovian Ocean 392–393 Grenville Province 406–438
Mischief mélange 278 large hot long-duration 5, 405– 406, 430, 433
Mistassini Group 414, 416 modern 1 –2
Mixteca terrane 239, 240, 255, 256, 346 Ortegal complex 352, 353
palaeogeography 261– 265 orthogneiss
Moeche ophiolite 319, 320, 325– 327, 352 Evora massif, U–Th– Pb analysis 222, 224, 225,
origin and history 336 226, 227–231, 232
Moho, Ural mountains 185, 187, 188 Oaxaquia terrane 241
Moldanubian Zone, Rheic Ocean mafic rocks 356, Oruanui eruption 22, 23–24
357– 358 Ossa–Morena Zone
Monzonite Supersuite plutons 142, 144, 145 deformation 216–222
Moosevale Formation 75 geochronology 222–231
Morais complex 352, 353 tectonic models 231 –234
Morais–Talhinas Unit 353 Ottawan Event
Motuhora volcano 13 Grenville Province 408– 409
Mount Peyton pluton 272, 275, 291, 297 deformation and metamorphism 421 –430,
Mullingarra Complex 452, 453 432– 433, 434, 435
Munzur platform 147, 148 orogenic collapse 432, 437
Murzinka–Adui zone 187 Ovacik Mélange 147, 148
486 INDEX
Paeroa Fault 13 quartzite, Acatlán Complex 243, 246

Paganzo Basin 45 Quesnellia terrane 71–73
Palaeo-Asian ocean 161 geological setting 73, 76
palaeogeography island arc 81– 84
Mexico 261 –265 stratigraphy 74
Pannotia 332 Takla Group 75–81, 82
Rheic Ocean 332
Rodinia 330, 371– 395 Rae Craton 466, 468, 470
palaeomagnetic data, Rodinia reconstruction 371– 395 see also Western Churchill Province
Palaeotethys Ocean 89, 94, 346, 359 rebound, Grenville Province 437
Palaeouralian ocean 163, 166, 167 reconstruction, Rodinia 371– 395
ophiolites 170 –171, 188 Red Indian Lake arc 275, 279, 305
Palaeozoic, flat-slab subduction 43, 44–46 Red Indian Line 272, 279, 284, 291, 293
magmatic evidence 45 Rehberg ophiolite 357
sedimentary evolution 44– 45 Reindeer Orogeny 465– 466, 471
tectonics 45–46 Reindeer Zone 458, 459, 463, 464, 465, 470
Pampean flat-slab segment 32, 33, 35, 35–36, 46, 47 Reporoa Caldera 19–21, 24
Pangaea Restigouche fault 272
amalgamation 163, 165, 216, 239–240, 265, 345 Rheic Ocean 5, 197, 199, 200, 212, 240, 263– 265,
break-up 93 319, 331, 332
Pannotia 167, 331, 332 closure 336, 338
paragneiss evolution 345 –360
Evora massif, U– Th– Pb analysis 222, 223, 225, mafic complexes 344, 346– 360
228, 232 opening 332, 333–336, 337
Oaxaquia terrane 241 Rheic suture 197, 198, 199, 263, 344, 356, 359
Parautochthonous Belt 407, 408, 410, 412, 413, 416, Rheno–Hercynian zone 356, 357
417, 418, 429, 433 rhyolite, Andes 35, 40, 41
Patlanoaya Formation 244, 259, 263 rifting
Pay–Khoy– Novozemelian foldbelt 163, 182 Alpine Realm 93, 97
Payenia flat-slab segment 42–44, 46, 47, 48 Uralides 167, 170, 188
peneplains, Urals 182 Rigolet event
Penobscot arc 276, 297 Grenville Province 408–409
peri-Gondwanan terranes 345 deformation and metamorphism 409– 420, 412,
peridotite, Lizard Complex 355 –356 424, 429, 432–434, 435
Peruvian flat-slab segment 32, 36– 37, 38, 46, 47 Rigolet Thrust 423
Petlalcingo Suite 244, 252, 254, 263 Rocky Brook–Millstream fault system 272
petrography, Takla Group 75– 77 Rodinia 5
Piaxtla Suite 243, 244, 245, 246, 257– 258, 262, apparent polar wander paths 375
346, 349 reconstruction 329, 330, 331, 371– 395,
Piemont – Penninic ocean 102, 104 372– 373, 455
Pilbara Craton 445, 446 Rotoiti eruption 21, 22
Pinjarra Orogen 446, 454, 455 Rotorua Caldera 17– 19, 24
Plata Craton, Rodinia reconstruction 390– 391, 393 Ruamoko Rift zone 13
Platinum-bearing Belt 165, 166, 173 Ruapehu volcanic massif 14, 15
Plughat Mountain Succession 75 Rusty Lake Belt 465
plume magmatism, Uralo-Siberian 165, 188
plutonism St George batholith 272, 293
Appalachian 273, 274, 279, 282, 292, 293, 295, Sakarya continent 134, 135, 136, 138
297–298, 300, 301, 304 Salda metamorphic complex 187
Bohemian Massif 199, 200 Salida Unit 244, 250, 251, 256, 259, 264
Pontide arc 150, 151 Salinic orogenesis 274, 276, 284–291, 305
Popelogan –Victoria arc 276, 279, 284, 285, 297, Salinic slab 291, 292, 297
304– 305 salt diapirs, Sonda de Campeche 63
Precordillera 33, 35, 43, 44–45 San Rafael block 42, 43–44, 45, 46
Preuralian foredeep 165 Santiago Formation 241, 242, 256, 257
Preuralian zone, tectonostratigraphy 168 São Francisco craton, Rodinia reconstruction 380– 381,
Principal Cordillera 33, 35, 44 382, 383– 384
Pukehangi rhyolite dome complex 18 Sask Craton 459, 463, 470, 472
Puna flat-slab segment 41, 42, 47, 48 Savage Mountain Formation 75
Purrido ophiolite 319, 320, 321, 322, 353 Saxo–Thuringian Zone
origin and history 329, 330, 331, 333, 334 geology 198– 199, 200, 202
Pütürge massif 146, 146, 147, 150, 151 Rheic Ocean mafic rocks 356, 357– 358
Pyrenean cycle 102–104 U–Pb zircon analysis 204, 205-206, 208, 209, 210
Pyrenean– Biscay ocean 100– 102, 104 Saxonian Granulitgebirge, metamorphism 211
INDEX 487
Scandian orogeny 345 subduction channel, Elbe zone 201

schist, Elbe Zone 201 –203, 211 subduction erosion
Série Negra paragneiss, U– Th–Pb analysis 222, 223, 225, Acadian 295
228, 231, 232 Cache Creek HP rocks 82–83
Serov– Mauk Fault 165, 166, 187 Chiapas fold-and-thrust belt 64
Serov– Mauk suture zone 181 Sudbury dykes 410–411, 413, 414, 419
Siberia 161, 181–182 Sugluk block 459, 462, 468, 472
Rodinia reconstruction 382, 388– 389, 391, 392 Sunbeam Group 454
Sicilides 114, 115 Superior Craton 459, 460 –461, 462, 467, 470, 471, 472
Sierra de Chiapas Syenite Supersuite plutons 142, 144, 145
seismic data 60, 63
stratigraphy 58, 59, 63 Taconic orogenesis 274, 277 –284, 303–305
Sierra Madre terrane 239, 240, 244, 252, 255, 256, 257 Taconic seaway 279
palaeogeography 261– 265, 264 Tagil arc 173, 174, 175, 180, 187
Sierras Pampeanas 31, 33, 35, 46 Tagilo–Magnitogorskian megazone 166, 173
slab breakoff 137– 139, 144, 145, 150, 151, 306 –307 Takla Group 72– 73, 75–77
Acadian 286, 300–301 geochemistry 77–81, 82
Salicic 291, 292 petrography 75–77
Taconic 279, 284 Tally Pond Group 272, 275
Slate Creek Succession 75 Taranaki Fault 10
Slide Mountain terrane 72, 73, 76 Tarawera volcano 15
stratigraphy 74 Tarim, Rodinia reconstruction 382, 386– 388
Smokey Archipelago, Grenville Front 415, 417, 418 Taupo Caldera Complex 22– 24, 25
Snow Lake Belt 460, 464, 472 Taupo Fault Belt 13
Snowbird orogen 460– 461, 462, 463, 468 –469, 471 Taupo Ignimbrite 23
Soeira Unit 353 Taupo Volcanic Zone 9– 25, 10, 12, 24
Somozas ophiolitic mélange 319, 320, 321–323 central 15–16
origin and history 334 caldera volcanoes 16– 24
Sonda de Campeche 57, 58 faulting and volcanism 13– 16
seismic data 62, 63 structure 13
stratigraphy 61, 63 subsurface 10– 11, 13
Sops Head/Boones Point Complex 275, 284 Tauranga caldera 9, 12
South Mountain batholith 272, 275 Tauride block 135, 136, 138, 140, 142– 143
STEP faults 133 Tecomate Formation 244, 250, 252, 253, 256, 259, 263
Stikinia terrane 71–73 Tehuantepec Transform/Ridge 55, 56, 57–58, 66, 67
geological setting 73, 76 Tehuitzingo serpentinite 348
island arc 81–84 Tetagouche–Exploits back-arc basin 276, 277, 279,
stratigraphy 74 285–286, 291, 297, 305
Takla Group 75, 77, 81, 82 Tethys Ocean 89– 90
strike-slip tectonics 215– 216 subduction 141, 143, 144, 145, 151
Elbe Zone 201–202, 202, 211 see also Alpine Tethys; Neotethys; Palaeotethys
Stuhini Group 73, 81, 82 Ocean
Subandean fold and thrust belt 42, 48 Tharandt Volcanic Complex 203
subduction Timanian orogeny 161, 163, 166, 167
Acatlán Complex 348– 349 Tiñu Formation 241, 242, 256, 257, 340, 346
African– Anatolian plate boundary 128– 129, 131, Tongariro Volcanic Centre 14, 24
134, 136 –138, 141, 143, 144, 145, 150, 151 Trans– Hudson orogen 6, 460
Alpine Tethys 105– 106 evolution 457–472, 460–461
Appalachians Transuralian zone 166, 185, 187, 188
Acadian 291–295, 305 tectonostratigraphy 168
Salinic 286, 289, 291 Triassic, Alpine realm, plate tectonics 94
Taconic 279, 281, 283 Turkestanian ocean 163
Calabria 117– 118, 121 Tyan-Shan mountains 162, 163, 182
ensialic 180 Tyrrhenian Sea 113–114, 114, 117, 118, 123
flat-slab 2
Andes 31– 48 U –Pb zircon analysis
transition to normal 46, 48 Elbe Zone 204, 205–206, 208, 209, 210
Gondwana– Laurussia collision 198 Mexico 240– 257
Neoarchian 303 U –Th–Pb zircon analysis, Evora massif 222– 230
normal, transition to flat-slab 46 uplift
Rheic Ocean 353– 354 Andes 46, 48
Taupo Volcanic zone 11 East Anatolian High Plateau 149, 151
Trans–Hudson Orogen 467– 468, 471 Salinic 290
Urals 173–181 thermal 48
488 INDEX
Ural mountains 161, 162 Quaternary, eastern Mediterranean 130, 131, 133
Cimmerian orogeny 182 rhyolitic
deep structure, seismic profiles 183–187 flat to normal slab transition 46
formation 181 –183 Taupo Volcanic Zone 9, 10, 15, 18, 20–22
neo-orogeny 182–183 Tagil arc 173, 174, 175
structural development 166– 167 Taupo Volcanic Zone 9 –25
subduction 173– 181 and faulting 13– 16
tectonic zones 165– 166 Uralo–Siberian LIP 181– 182
Ural– Tau antiform 165, 166, 176, 179, 185 Valerianovka subduction zone 181
tectonostratigraphy 168 volcanoes, caldera, Taupo Volcanic Zone 16– 24
Uralian orogen 3, 161–189
Uralides 162, 163 Wadati–Benioff zone
comparison with Variscides 163–164 Peru 38
structural development 167– 173 Taupo volcanic zone 10, 11
rifting 167, 170 see also Benioff zone
tectonostratigraphy 168 Wathaman orogeny 467–468
Uralo–Mongolian belt 162, 163 Wathaman–Chipewyan Batholith 467–468, 471
Uralo–Siberian large igneous province 181 –182 Weir Formation 286, 287
URSEIS profile 183, 184, 185, 187 West Africa, Rodinia reconstruction 382, 389– 391
Utopia granite 272 West Australian Craton 445, 446, 453
West Uralian megazone 165–166, 188
Valais ocean 102 tectonostratigraphy 168
Valais Trough 102, 104 Western Churchill Province 458, 460, 463, 470,
Valerianovka subduction zone 174, 179–181 471, 472
Vardar ocean 95–96, 102 see also Hearne Craton; Rae Craton
Variscan orogeny 3– 4, 318, 345, 346, 348 Western Dome Belt 17
Bohemian Massif 197–212, 211 Westlausitz Fault 203
Iberian Massif 216 Whakamaru Caldera 12, 16– 17, 19
NW Iberia 319, 321 Whakatane graben 13, 15
SW Iberia, extensional tectonics 231– 234 Whakatane seamount 13
Variscan suture, NW Iberia 319, 321 White Island volcano 13
Variscides 163–164, 318 White Rock Formation 275
ophiolites 317 –338 Wilson cycle
Victoria arc 272 Trans–Hudson Orogen 457
Vila de Cruces ophiolite 319, 320, 325, 326, 352 Uralides 187–188, 189
origin and history 335– 336 window, asthenospheric 137, 144, 145, 150, 151
Viséan, Variscan extensional tectonics, SW Iberia Wollaston Belt 468
231– 234
volcanism Xayacatlán Formation 243, 246, 257, 262
Andes 35
andesitic, Taupo Volcanic Zone 9, 10, 11, 13, 14 Yilgarn Craton 447, 446, 449, 452, 452
Cappadocian volcanic province 140, 143 Yukon–Tanana terrane 71, 72, 73
East Anatolian High Plateau 146, 149, 150, 151
Magnitogorsk arc 174, 175– 176 Zilair flysch formation 176, 177, 179
Ossa Morena Zone 222 zircon, metamorphic, Ossa Morena Zone 230–231

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Ancient Orogens and Modern Analogues

The Geological Society of London

Society Books Editors

Society Books Advisors

IUGS/GSL publishing agreement

Ancient Orogens and Modern Analogues

Published by The Geological Society from:

British Library Cataloguing in Publication Data

Mesozoic –Cenozoic orogens

RAMOS , V. A. & FOLGUERA , A. Andean flat-slab subduction through time 31

STAMPFLI , G. M. & HOCHARD , C. Plate tectonics of the Alpine realm 89

OCCHIPINTI , S. A. & REDDY , S. M. Neoproterozoic reworking of the Palaeoproterozoic 445

Tongariro Volcanic Centre of Tongariro Volcanic Centre. Alternatively, in

Central TVZ axis. Near Taupo, for example, a series of basalt

collapse structures, with rectangular geometries, References

Introduction The objective of the present study is to charac-

Introduction better understanding of modern analogs, such as the

composition of QTG clinopyroxene is comparable Newfoundland. The concentrations of Rb and Sr

EXTRUSION OF HIGH-PRESSURE CACHE CREEK ROCKS

Th/La, Ba/La, Tb/Lu (Fig. 7) and Nb/La ratios.

desite; Bas, basanite; Trach, trachyte; Neph, nephelinite.

argue against a significant role for garnet in their

Fig. 6. Chondrite-normalized REE patterns of the

Fig. 8. Primitive mantle-normalized trace element

Introduction For a long time, the Tethys was considered as a

G. M. STAMPFLI & C. HOCHARD

Introduction respectively) of the Apennine –Sicilian peninsula

Tectonic elements of southern Italy Three main tectono-stratigraphic units comprise

In contrast to the Apennines and the Calabrian Discussion

somewhat wedge-shaped (narrowing at depth) in initiating subduction remains unknown. The

Abstract: The modern Anatolian–African plate boundary is represented by a north-dipping

Introduction suture zone has been estimated to be c. 16 mm a21

zone (Le Pichon & Angelier 1979; Jolivet 2001;

Eastern Pontide arc Sanandaj–Sirjan continental block to the SE in

Introduction Early Palaeozoic, these continents were separated

be attached to Gondwana (e.g. Linneman et al.

arc-continent collision was accompanied by HP –

Tectonic zones of the Urals (Fig. 4)

Palaeozoic– Lower Jurassic stage as an example of a truncated by the N –NE-trending magnetic

The origin of the Palaeouralian

Fig. 8. The position of the Early Palaeozoic rift/plume

The younger ages of the metamorphic rocks

With the demise of the Magnitogorskian arc, suprasubductional tonalite–granodiorite massifs by

Introduction suture, see Fig. 1) (Scotese & Barret 1990).

SW. The geodynamic relationships and the appro-

TIMING OF DEXTRAL STRIKE-SLIP PROCESSES AND BASEMENT EXHUMATION

Fig. 8. Cathodoluminescence–SEM–Images of representative zircon grains of different samples, which have

Introduction of intra-continental extensional tectonics during

micaschists (greenschist-upper amphibolite facies); high-grade unit suggests a prograde metamorphic

Apparent age (Ma)

Apparent age (Ma)

Biotite granite ARL-6

Table 4. Monazite SHRIMP data from the orthogneiss VLV-3

Apparent age (Ma)

A more complex structural scenario was proposed

Introduction dominated by Palaeozoic siliciclastic and oceanic