Journal of Structural Geology - Special Issue - Fault Zones
Journal of Structural Geology - Special Issue - Fault Zones
Journal of Structural Geology - Special Issue - Fault Zones
V
V
0
b ln
V
0
q
D
c
(1)
where m and m
0
are the friction coefcient and the initial friction
coefcient respectively, a and b are experimentally derived
constants, V and V
0
are the new sliding velocity and the initial
sliding velocity respectively, q is the state variable and D
c
is the slip
weakening distance.
The evolution of the state variable is a function of time, normal
stress and displacement and has units of time. It has been explained
in terms of the age of the load-supporting contacts and the time
required for a new set of contacts to develop following a pertur-
bation of the system (e.g. displacement rate, change in normal
stress, etc.). Dieterich and Kilgore (1994) showed direct evidence
that the state variable is related to the evolution of the area of the
load supporting contacts over time and various normal stresses,
termed asperity creep. Two formulations are widely used to model
the state evolution, the aging (or slowness) law and the slip law
(Eqs. (2) and (3)).
_
q 1
Vq
D
c
Aging; Dieterich or Slowness law (2)
_
q
Vq
D
c
ln
Vq
D
c
Slip or Ruina law (3)
These two formulations produce quite different styles of
nucleation and rupture (e.g. Ampuero and Rubin, 2008).
The range of geological materials that have been characterized
in a rate-and-state framework are few. Most comprehensive studies
involve quartz and crushed granite powders. These granular
materials generally show velocity weakening behaviour (where
aeb is negative, see Fig. 8) at low slip rates, opening the possibility
for unstable slip (Green and Marone, 2002). Most natural fault
zones contain at least a proportion of phyllosilicate minerals but
experimental studies on these materials are even sparser (Morrow
et al., 1992; Scruggs and Tullis, 1998; Reinen, 2000; Saffer et al.,
2001; Moore and Lockner, 2008; Ikari et al., 2009; Smith and
Faulkner, 2010). These studies generally suggest that phyllosili-
cate-rich gouges exhibit velocity strengthening behaviour at low
slip rates (where aeb is positive) and thus may be associated with
fault creep (Faulkner et al., 2003). Talc, in particular, exhibits
inherently stable, velocity-strengthening behaviour under all
conditions tested (Moore and Lockner, 2008), although these
conditions do not include seismic slip speeds (w0.1 to 1 m/s). Note,
however, that montmorillonite and serpentinite gouge can both
exhibit velocity weakening behaviour (Reinen, 2000; Saffer et al.,
2001). Indeed, a recent compilation of experimental work
suggests that even materials which exhibit velocity-strengthening
behaviour at lower slip rates become velocity-weakening above
w0.1 m/s (Wibberley et al. 2008). Some experimental studies have
shown that phyllosilicates can exhibit negative b values (Saffer and
Marone, 2003; Ikari et al., 2009; Smith and Faulkner, 2010), which
are difcult to interpret physically as a negative b value is generally
assumed to indicate an increase in contact surface area with faster
slip. Karner et al. (1997) and Blanpied et al. (1998) also report
negative b values for granular quartz and granite.
The microphysical processes responsible for the observed rate-
and state-dependent behaviour are thermally activated and follow
Arrhenius-type behaviour (Chester, 1994; Blanpied et al., 1998;
Nakatani, 2001; Rice et al., 2001). They presumably include sub-
critical crack growth, crystal plasticity, diffusion and possibly
reaction at grain contacts. However, the rate and state formulation
does not include temperature. Chester (1994) showed that the
activation energy required for wet quartz gouge was consistent
with sub-critical crack growth at asperity contacts. This is sup-
ported by the conclusions of Frye and Marone (2002) that relative
humidity plays a role in the granular friction of quartz and alumina
as a result of chemically assisted mechanisms. This is clearly an
important area for future research, for if the physical mechanisms
responsible for frictional behaviour are known and characterized,
then the behaviour at a wider set of environmental conditions can
be better predicted. However, we note that future progress in this
area of research probably needs to combine experiments with
detailed microscopy (SEM and TEM) and theoretical and compu-
tational studies due to the complexity of the interactions that occur
at the nanoscale (Szlufarska et al., 2008; Mo et al., 2009).
3.3. Mechanics of dynamic rupture
The properties of faults during rupture are typically studied
using inversion of seismic data recorded during earthquakes to
compute the slip distribution of rupture events (kinematic model).
Various models allow computation of the stress eld, from which
the physical properties of the rupture are inferred (see review by
Kanamori and Brodsky, 2004). Recently, complementary laboratory
studies and eld measurements have led to a better understanding
of dynamic rupture, although this is currently a rapidly developing
area of research.
The slip history during large earthquakes appears to take the
form of a slip pulse rather than a self-similar crack-like rupture
(Heaton, 1990). In the slip-pulse model, only part of the fault plane
ruptures at any one time. Most seismological models now assume
source-time functions of small nite duration and thus implicitly
apply the slip-pulse model. However, this mode of slip still raises
Slow, V
1
Fast, V
2
Slip
F
r
i
c
t
i
o
n
c
o
e
f
f
i
c
i
e
n
t
a*ln(V /V )
2 1
b*ln(V /V )
2 1
D
c
Fig. 8. Idealized rate- and state-dependent frictional behaviour, in which the
mechanical response is dened by a stepwise change in velocity.
D.R. Faulkner et al. / Journal of Structural Geology 32 (2010) 1557e1575 1565
a number of unanswered questions. First, the spectral signature of
large earthquakes suggests that the corner frequency (lowest
frequency seismic waves) scales with rupture area which runs
contrary to the idea of a similar-size slip pulse regardless of the size
of the rupture. Additionally, in the early stages of failure the slip-
pulse model must accumulate approximately the correct amount of
slip appropriate to the overall rupture size. This leads to the ques-
tion of whether an earthquake event knows a priori how large it
will be (Marone and Richardson, 2006). Finally, modelling using
a rate-and-state framework has shown that slip pulses can only
develop under a restricted set of conditions where the initial shear
stress on the fault is low(Zheng and Rice, 1998). Lowstress (weak)
faults might well exist (Section 3.1), but a large proportion of faults
are inferred to have high stress levels (Townend and Zoback,
2000). Does this mean that slip pulse-type behaviour is not
possible for these faults?
Seismic records allow modelling of the mechanical properties of
rupture, but, in recent years, laboratory and eld measurements
have provided independent constraints. All these data yield an
understanding of the rupture process in terms of the energy
balance for an earthquake (Eq. (4); see Kanamori and Rivera, 2006).
The change in potential energy DU
e
(the sum of the elastic strain
energy and gravitational energy during earthquake slip) is the sum
of surface energy U
s
(to produce new crack surface area), kinetic
energy U
k
(radiated as seismic waves) and frictional energy U
f
(dissipated as heat):
DU
e
U
s
U
k
U
f
: (4)
Constraints on the kinetic energy are well known from model-
ling of seismic waves. Chester et al. (2005) estimated the fracture
energy for the Punchbowl fault in California by accounting for all
the fracture surface energy in both the damage zone and, more
importantly, the core zone, where nano-sized particles are present
and account for a signicant proportion of the fracture area. Hert-
zian fracture models (where fracturing occurs at grain contacts of
spherical grains) suggest that it is mechanically very difcult to
produce such small particle sizes. Sammis and Ben-Zion (2008)
suggested that shock loading and sub-critical crack growth under
compressive stress, or high strain rate tensile stress may be
responsible. The fracture energy expended during the formation of
pulverized rocks (Section 2.1) that are thought to develop during
seismic slip near the surface is not currently known. However,
Biegel et al. (2008) showed that off-fault damage will affect the
velocity of an earthquake slip pulse.
Di Toro et al. (2005) showed that a simple analysis of pseudo-
tachylyte veins provides broad constraints on the dynamic stress
during rupture. This analysis indicated very low(in terms of Byerlee
friction) shear stresses driving rupture. Di Toro et al. (2006)
corroborated these results with measurements, using the same
protolith, of dynamic friction at seismogenic slip velocity. While
frictional melting might result in low shear stresses driving slip
(after the effects of viscous braking have been overcome; Tsutsumi
and Shimamoto, 1997), the friction of other granular materials at
high velocity were unknown until recently.
Technical developments over the past 15 years (notably in the
laboratories of Shimamoto and co-workers) have allowed
measurement of the stresses during high-velocity frictional testing
in rotary shear apparatus. Fig. 9 shows typical results froma friction
experiment conducted at seismogenic slip velocity. These data
complement the data modelled from natural earthquake ruptures.
A key feature of the data in Fig. 9 is the dramatic weakening from
Byerlee levels of friction down to levels between 0.1 and 0.2. The
reasons for this weakening are many, and dependent on the
material tested. They include ash heating at asperity contacts
(Bowden and Tabor, 1950), silica gel formation (Goldsby and Tullis,
2002), thermal pressurization (Hirose and Bystricky, 2007), fric-
tional melt lubrication (Di Toro et al., 2006) and thermal decom-
position (Han et al., 2007). Recent modelling of the earthquake
process has started to combine and incorporate some of these
additional thermal factors into rate- and state-frictional frame-
works (Rempel and Rice, 2006; Rudnicki and Rice, 2006; Segall and
Rice, 2006; Noda, 2008; Noda et al., 2009).
The results in Fig. 9 have important implications. First, they
explain the long-standing debate on the development of frictional
melting in fault zones. Simple analyses show that extreme
temperatures are quickly reached due to frictional heating (see
Rice, 2006 for a summary). All these models assume friction coef-
cients commensurate with Byerlees law. If the shear traction
required for seismic slip reduces dramatically then the frictional
energy converted to heat is also dramatically reduced. It can also
explain the lack of any heat ow anomaly (Section 3.1) over the
seismogenic parts of the San Andreas fault.
Another feature of high velocity laboratory tests is the magnitude
of D
c
which is on the order of decimetres to metres. In slowfrictional
testing (Section 3.3) to determine rate- and state-friction parame-
ters, the slip-weakening distance is on the order of microns. A long-
standing debate focuses on the apparent discrepancy of D
c
values
derived from the laboratory versus values inferred from seismolog-
ical data, the latter being onthe order of a metre. This parameter may
scale with fault roughness or fault thickness (Scholz, 1988; Marone
and Kilgore, 1993). The emergence of models that account for the
different physical processes that occur during seismic slip as
opposed to slow laboratory frictional sliding seem to provide an
answer to the discrepancy. It appears the seismically derived D
c
is
a different parameter to that measured at slow slip rates in the
laboratory, and involves fundamentally different physics. This is
supported by the range of behaviour observed in high velocity
experiments, modelling and from other geological observations that
include thermal pressurization (Wibberley and Shimamoto, 2005;
Bizzarri and Cocco, 2006), ash heating, thermal decomposition
and frictional melting. The value of D
c
might be envisaged to increase
as thermally activated processes continue to produce weakening
with continued slip (in the same way that the boundary between
surface energy and heat varies during slip as suggested by Tinti et al.,
2005). It is clear that smaller critical slip distances must exist,
otherwise small earthquakes could never nucleate and that a sensi-
tive balance exists between the energy required for the work of
fracture and that converted to heat. Unfortunately, testing these
hypotheses with seismological data is hampered by the limited
frequency bandwidth from which kinematic models of rupture are
derived and also the constraints on the resolution and uniqueness of
Fig. 9. Results from a typical high-velocity friction experiment on kaolinite (unpub-
lished data from Faulkner, Mitchell, Hirose and Shimamoto). It shows a number of
characteristic features including the peak friction, slip weakening distance and steady-
state sliding. The strength recovery upon deceleration of slip is also shown.
D.R. Faulkner et al. / Journal of Structural Geology 32 (2010) 1557e1575 1566
parameters such as D
c
(Spudich and Guatteri, 2004; Tinti et al.,
2009). Field observations by Kirkpatrick and Shipton (in press)
conrm that slip weakening mechanisms are likely to be spatially
and temporally variable across an earthquake fault surface.
4. Fluid ow in fault zones
In recent years our understanding of uid ow through faults
has advanced greatly. The typical structure of fault zones (Section
2.1), with a core and damage zone, has provided the framework
within which to place laboratory measurements of the uid ow
properties of natural and synthetic fault products into context
(Fig. 10). Fault-related uid ow has also been investigated via
a number of indirect data sources such as migrating seismicity at
depth, shallow reservoir-induced seismicity, springs, geysers and
geothermal systems. These sources have provided some rst-order
constraints on the rates of uid ow in natural fault zones at depth,
and at length scales unavailable to lab experiments.
In the original Caine et al. (1996) fault core and damage zone
model of fault architecture the fault core was visualised as
providing an across-fault barrier to ow and the fractured damage
zone as an along/up-fault conduit. However, the varying fault
architectures outlined in Section 2.1 gives rise to a much more
complex set of fault zone hydraulic behaviours. The intricate
structure of lowand high permeability features within a fault zone
can lead to extreme permeability heterogeneity and anisotropy.
The permeability of a fault zone, both in-plane and perpendicular to
the plane (across-fault) is governed both by the permeability of the
individual fault rocks/fractures and, critically, by their geometric
architecture in three dimensions (e.g. Lunn et al., 2008). For
example, rocks from the fault core are commonly rich in phyllosi-
licates, which typically have low permeability, but only form
barriers to ow if they are continuous throughout the fault plane
(Faulkner and Rutter, 2001). Open fractures and slip surfaces (both
within the fault core and the surrounding damage zone) have
a permeability governed by their aperture distribution, which is in
turn inuenced by their orientation to the present day local stress
eld. Such fractures and slip surfaces may have a substantially
greater permeability than the host rock; however, their net effect
on bulk along-fault and across-fault ow, depends entirely on their
connectivity and ability to cross-cut other lower permeability units.
4.1. The hydraulic properties of the fault core and its inuence on
uid ow
In natural faults two distinct types of gouge are present. The rst
are granular materials composed of broken, irregular but roughly
equant clasts (in the sense that their long and short axes are
approximately equal), and the second are gouges that contain some
proportion of phyllosilicate material. Relatively few data on the
permeability of granular gouges are available but they tend to
develop a characteristic grain size distribution (Sammis et al., 1987;
Marone and Scholz, 1989) that may suggest a similar permeability
development for all these materials. Zhang and Tullis (1998)
measured the permeability development in synthetic quartz
gouge at a normal stress of 25 MPa. They found that at shear strains
up to w10 the permeability is reduced by two to three orders of
magnitude. This is in agreement with more recent ndings of
Crawford et al. (2008) and Main et al. (2000). Beyond this shear
strain (to a shear strain w200), Zhang and Tullis (1998) found the
permeability dropped by a further two to three orders of magnitude
and that a permeability anisotropy of one order of magnitude
developed. This was due to the formation of localized, ne-grained
Y shears. These laboratory data are in agreement with eld obser-
vations and permeability measurements from boreholes that
suggest a signicant drop in cross-fault permeability in deforma-
tion band-dominated faults as the fault core develops through-
going slip surfaces (Shipton et al., 2002, 2005).
Fault zones rich in phyllosilicate material tend to have lower
permeabilities than quartz and/or framework silicate-rich gouges.
Information on the uid ow properties of phyllosilicate-rich fault
zones is necessary to understand uid ow associated with fault
creep (e.g. Faulkner and Rutter, 2001) and earthquake slip (e.g.
Wibberley and Shimamoto, 2005; Yamashita and Suzuki, 2009), as
many large faults contain signicant proportions of clays. Where
the uid-ow properties of fault zones are needed to evaluate the
robustness of a fault-bounded hydrocarbon prospect or the eld
compartmentalizing effects of intra-reservoir faults, estimating the
possible phyllosilicate content of the fault zone is critical, along
with reservoir juxtaposition geometry. Based on eld-observations
of fault zones, the two main mechanisms that entrain phyllosili-
cates (typically shale and/or clays) into fault zones in layered
sandstone e shale sequences are shale or clay smearing (e.g. van
der Zee and Urai, 2005) and abrasional mixing (e.g. Yielding et al.,
fracture
density
permeability
a b
Fig. 10. Some physical properties of fault zones related to their structure (damage zone and fault core). (a) Single fault core and (b) multiple fault core, which illustrates the resulting
complexity in characterizing the resultant properties.
D.R. Faulkner et al. / Journal of Structural Geology 32 (2010) 1557e1575 1567
1997). The necessity of estimating fault properties from limited
datasets led to algorithms for estimating fault zone composition
which assume one or the other mechanism is operative. For shale/
clay smearing, the Clay Smear Potential (Weber et al., 1978) or the
Shale Smear Factor (Lindsay et al., 1993) rely on parameters such as
the thickness of the shale/clay source bed, distance of a point on the
fault from that source bed, and/or throw. These algorithms only
predict whether or not the smear along the fault is discontinuous
(likely leading to leakage) for a given fault throw and, if so, where.
On the other hand, the abrasional mixing mechanism led Yielding
et al. (1997) to propose the Shale Gouge Ratio (SGR) algorithm,
a ratio or percentage of shale in a silicilastic fault zone, which
simply assumes that any one point on the fault has a composition
identical to the average composition of the sequence past which
that point has slipped. In terms of sandstone eshale sequences, this
is extremely practical to implement, because the net volume of clay
(Vcl) logs from nearby wells can be extrapolated onto the fault (in
cases of simple stratigraphy), from which SGR is calculated for all
points on the fault for which wall-rock Vcl data exist. In reality, fault
zones are much more complex and local small-scale variations can
exist even in abrasive fault zones where the rule generally holds.
However, quantitatively constraining this variation may in future
help predict uncertainties in SGR-based evaluations.
In evaluating the sealing potential of a fault, analysis of juxta-
positions of reservoir/carrier beds against other reservoir beds is
critical. Basic geometric evaluation of such likely leak points is
easily done by fault-plane mapping of hanging-wall / footwall
juxtapositions, commonly called Allan diagrams (Allan, 1989).
However, eld-based studies show that faults are commonly both
segmented (in both dip and strike, e.g. Nemser and Cowan, 2009),
have multiple slip surfaces and involve varying degrees of ductile
deformation (fault-related folding; Section 2.3). Thus the net throw
observed on one fault at the scale of seismic resolution is in reality
often divided over two or more slip surfaces which share the
displacement at the sub-seismic scale. Such fault zone structure
therefore implies there might be reservoir-reservoir juxtapositions
across individual fault strands even where the entire fault
completely offsets the reservoir. Thus, better prediction of the likely
segmentation and slip zone bifurcation is needed, particularly the
incorporation of lenses of host-rock into the fault zone (e.g. Van der
Zee et al., 2008, Schmatz et al., 2010).
The recognition that fault zones in siliclastic sedimentary basins
are typically sand-shale gouges with uid barrier/transmission
behaviour governed by similar principles as shale-rich top-seals, led
to the application of capillary sealing theory. By using calibration
data for given fault-rock clay compositions and using laboratory
measurements (e.g. Sperrevik et al., 2002) or eld post-mortem
results (e.g. Bretan et al., 2003), the SGR method can be used to map
estimated capillary threshold pressures on the fault, yielding the
likely maximum hydrocarbon column height trapped against the
fault. The morerecent recognitionthat fault zone heterogeneitymay
lead to connected weak points which may provide a pathway for
hydrocarbon leakage has emphasized the need to predict better the
variability in fault zone structure and composition.
Similarly, fault zone permeability, particularly for incorporating
fault impact into reservoir simulators via the transmissibility
multiplier (Manzocchi et al., 1999), is commonly estimated from
SGR e permeability algorithms. Generally, there is a non-linear
dependence of the permeability on the fault zone clay content
under hydrostatic conditions due to the grain size difference and
compaction characteristics. For example, as the clay content
increases to between 25 and 40 volume % (the theoretical porosity
minimum, Revil et al., 2002) the clay particles sit in the pore space
between quartz, and the compaction characteristics are largely
controlled by the quartz framework. The permeability is strongly
controlled by the fraction of clay (Takahashi et al., 2007; Crawford
et al., 2008). At larger percentages of clay, the permeability is less
sensitive to the magnitude of the clay fraction. As the clay compacts
more readily than the quartz, the porosity minimum varies with
effective pressure (Crawford et al., 2008). Shear-enhanced
compaction is much less pronounced for the clay component than
for the quartz end-member. As a result, the clay-rich mixtures do
not reduce their permeability by much in comparison to the quartz-
rich gouges, which undergo a signicant permeability reduction
(Takahashi et al., 2007; Crawford et al., 2008).
Many measurements of natural clay-rich fault rocks are avail-
able (Faulkner and Rutter, 2000; Faulkner et al., 2003; Wibberley
and Shimamoto, 2003; Tsutsumi et al., 2004; Mizoguchi et al.,
2008). They all demonstrate the very low permeability of this
material. Natural gouge can have permeability anisotropy of up to
three orders of magnitude (Faulkner and Rutter, 2000). For
synthetic phyllosilicate gouges this value appears to be much lower
(Zhang et al., 1999), presumably due to the nature and distribution
of the authigenic clay phases that develop in natural gouges. Recent
work has measured the intensity of clay fabrics by, for example, X-
ray texture goniometry (Solum and van der Pluijm, 2009; Haines
et al., 2009). The gouges have a generally weak preferred orienta-
tion of the clays. Solum and van der Pluijm suggest that this indi-
cates that the clay fabrics are localized phenomena. Haines et al.
(2009) imply that the lack of clay fabric indicates the perme-
ability anisotropy must also be low. However, previous work on clay
fabrics shows that the permeability anisotropy is not due to clay
alignment; indeed this can only account for a permeability
anisotropy less than an order of magnitude (Faulkner and Rutter,
1998). Alternating microlayers of porous granular material and
ne grained clay-rich material are observed in the microstructure
and explain the anisotropy (Faulkner and Rutter, 1998; Faulkner,
2004). Furthermore, authigenic clay growth observed in TEM
images shows growth randomly in pore space in the stress shadows
of larger relict grains (Rutter et al., 1986).
The pressure dependence of the permeability is remarkably
similar for most clay-rich fault gouges, whether synthetic or natural
(see Faulkner, 2004). Faulkner and Rutter (2000, 2003) showed that
temperature and pore uid chemistry can strongly affect the
permeability of natural clay-rich rocks by altering physico-chemical
interactions between the rocks and aqueous pore uid.
The stress history of fault gouge, particularly clay-rich gouge,
has been shown to be an important control on the permeability.
Bolton et al. (1998) and Zhang et al. (1999) have shown that
permeability is reduced in sheared gouges that have undergone
normal consolidation, but may increase their permeability if they
have undergone overconsolidation and are subsequently sheared at
lower effective pressure conditions. Indeed the log-linear rela-
tionship of both porosity and permeability with mean effective
stress in clay-rich gouges fromthe Median Tectonic Line, Japan, and
dependence on the anisotropy of the stress regime suggest that
uid owmodelling in such fault zones may be based around a Cam
Clay -type soil mechanics framework (Wibberley et al., 2008).
The temporal evolution of the permeability of fault gouges,
particularly at hydrothermal conditions, has been a recent focus of
research. Pure quartz gouges (Nakatani and Scholz, 2004a;
Yasuhara et al., 2005; Giger et al., 2007) and quartzofeldspathic
mixtures (Olsen et al., 1998; Tenthorey et al., 1998; Morrow et al.,
2001) have been tested experimentally. Permeability reduction is
found to be initially rapid and progressively slows with time. The
rate of permeability reduction increases with increasing tempera-
ture (Olsen et al., 1998; Tenthorey et al., 2003; Nakatani and Scholz,
2004a; Yasuhara et al., 2005; Giger et al., 2007). The processes
responsible for the permeability reduction are interpreted to be
solution-precipitation processes (Nakatani and Scholz, 2004a;
D.R. Faulkner et al. / Journal of Structural Geology 32 (2010) 1557e1575 1568
Giger et al., 2007) and also precipitation of authigenic minerals
following the breakdown of quartz and feldspar (Tenthorey et al.,
1998). These processes have been successfully modelled
(Aharonov et al., 1998; Nakatani and Scholz, 2004b). These studies
have demonstrated that sealing will occur on time scales
commensurate with those of earthquake recurrence times (Morrow
et al., 2001).
4.2. Damage zone permeability
The permeability of fault damage zones is governed by the host-
rock permeability and the presence and geometric composition of
both macro-scale fracture networks (which will increase perme-
ability), and of low permeability deformation and compaction
bands (which will decrease permeability). Fault damage zone
permeability in low porosity rocks (Balsamo et al., 2010) is gener-
ally fracture-dominated and governed by the connectivity of the
macro-scale fracture network (Fig. 10). This contrasts with high
porosity rocks where the damage zone may be more complex and
permeability is governed by the frequency and connectivity of both
low permeability deformation bands and of high permeability slip
surfaces (Lunn et al., 2008). Both macro-scale fracture networks
and deformation bands have been shown to decrease in frequency
with increasing distance from the fault core (Rawling et al., 2001;
Shipton et al., 2002; Wilson et al., 2003; Mitchell and Faulkner,
2009). Wibberley and Shimamoto (2003) measured the perme-
ability of rocks collected from the surface trace of the Median
Tectonic line in Japan. Their measurements showed permeability
variations over several orders of magnitude, which can be
explained by the lithological variation of the fault zone and also the
structural complexity. Fig. 10 shows how this structural complexity
might result in permeability heterogeneity from the variation in
microfracture densities observed around multiple fault cores.
The permeability of the host rock within the damage zone is
controlled by the frequency and orientation of microfractures. One
potential problem with direct measurement of damage zone
permeability from rocks collected from the surface outcrop of fault
zones is that they may be subject to weathering or modication
since the fault related microfracture network was produced
(MorrowandLockner, 1994). Measurements of rocks recoveredfrom
depth fromfault-zone drilling projects may overcome this problem.
Experiments indicate permeability of initially lowporosity rocks
taken to failure increase by two to three orders of magnitude
(Simpson et al., 2001; Oda et al., 2002; Uehara and Shimamoto,
2004; Mitchell and Faulkner, 2008). For initially high porosity
rocks, the permeability may signicantly decrease with deforma-
tion in the damage zone (Main et al., 2000). Measurements of
permeability fromthe intact state of rocks to their failure stress can
be scaled to the levels and distribution of damage seen surrounding
fault zones (see Section 2.1), if a common factor between the two
can be found. For example one such common factor is the micro-
fracture density which can be measured in the eld and then
compared to that in experiments at various stages of damage,
where the permeability may be readily measured.
4.3. Estimating bulk fault zone permeability
The physical characteristics of fault damage zones are described
in Section 2.1. Estimates of the bulk fault zone permeability, for fault-
perpendicular and fault-parallel ow, are derived from numerical
models that simulate owthrough the fault zone (Brown and Bruhn,
1998; Jourde et al., 2002; Matthai and Belayneh, 2004; Odling et al.,
2004; Lunn et al., 2008) Such models showsignicant channelling of
ow within fault zones into a small number of focussed ow paths.
Similar channelling effects are observed in ow experiments within
individual fractures (Brown et al., 1998; Beeler and Hickman, 2004).
Observations from boreholes that penetrate faults, as well as along-
fault and across-fault pumping tests, are necessary to determine bulk
fault-zone permeability and to validate numerical models (Evans
et al., 2005; Medeiros et al., 2010).
Very few studies have measured bulk fault zone permeability
directly using boreholes. However, a number of secondary data
sources allow estimation of along-fault permeability. Talwani et al.
(1999) estimated the permeability of a shallow fault zone using
sinusoidal pressure oscillations in boreholes from lake level uctua-
tions. Their analysis shows fault zone permeability, in faults that are
subject to reservoir-induced seismicity, are between 1.1 10
15
and
1.78 10
15
m
2
. Tadokoro et al. (2000) estimated along-strike fault
(damage) zone permeabilities around 1e10 10
15
m
2
from the
migration rates of induced seismicity during borehole injection
experiments in the Nojima fault zone following the Kobe 1995
earthquake. At deeper levels, Shapiro et al. (1997) found crustal
permeability to be w10
16
m
2
in the KTB borehole in the depth
interval 7.5e9 km. Acompilation of fault zone permeabilities derived
from reservoir-induced seismicity data can be found in Talwani et al.
(2007). Usingthesametechniques of migratingpatterns of seismicity,
but at deeper levels in the brittle crust, Miller et al. (2004) estimated
fault zone permeability to be 4 10
11
m
2
immediately following the
M6 Colorito earthquake sequence of 1997 in central Italy. Noir et al.
(1997) inferred a higher fault zone permeability of 10
8
m
2
for the
Dobi earthquake sequence in Afar in 1989.
4.4. Spatial and temporal variability of fault zone permeability
A number of recent studies have shown that fault zone
permeability is highly heterogeneous both spatially and temporally.
Studies examining the distribution of geothermal spring tempera-
tures along the Borax Lake fault, Oregon, USA, show that neigh-
bouring springs a few metres apart can have widely different
temperatures (Fairley and Hinds, 2004). These springs suggest that
high permeability pathways exist as discrete structures in the Borax
fault damage zone, and that individual ow paths have the capa-
bility to transport uids rapidly from depth parallel to the fault
plane. Dockrill and Shipton (2010) use observations of natural
leakage of CO
2
along faults in Utah, USA, to show that along-fault
ow is occurring at a few discrete locations along strike, and that
these discrete along-fault ow channels have migrated over time.
The presence of a modern oil seep at one location on the fault also
indicates the existence of discrete unconnected along-fault pipes
that provide pathways for uid owfromlithologies at depth to the
surface and are unconnected to shallower horizons. Evans et al.
(2005), during an injection test in the Soultz borehole, observed
that some 95% of the ow entered the rock mass at just 10 major
owing fractures. Do Nascimento et al. (2005) show that pressure
changes as low as 0.5 kPa (equivalent to 5 cm of water head) are
enough to trigger transient changes in permeability that are
spatially correlated, and related to seismicity below a water
reservoir.
In the hydrocarbon exploration and production industry,
evidence for the spatial and temporal variation of fault perme-
ability exists but is usually limited to indirect inferences from
reservoir uid and pressure data either side of the faults, because
faults are generally avoided when drilling as they give rise to
a number of drilling problems (e.g. Grauls et al., 2002). Neverthe-
less, studies of charge timing in fault-bounded blocks (e.g. Residual
Salt Analyses of uid inclusions in reservoir pore cements) can
show an increase in the hydrocarbon buoyancy pressure differen-
tial across the fault through time, attributed to increasing sealing
properties as compaction affects the fault during progressive burial.
Transient increases in up-fault permeability during periodic
D.R. Faulkner et al. / Journal of Structural Geology 32 (2010) 1557e1575 1569
reactivation may lead to leakage of fault-bounded hydrocarbon
traps, as has been found for several cases in the northern North Sea
for example (Wiprut and Zoback, 2002). On production time scales,
faults which initially sealed signicant hydrocarbon-related pres-
sure differences may become leaking as one compartment is
depleted (e.g. Dincau, 1998) and stress changes related to depletion
may render the fault unstable in certain cases, leading to up-fault
leakage during slip (Cuisiat et al, 2010).
One exception to the general paucity of direct industry
measurements of fault permeability is the Pathnder well on
Eugene Island, Gulf of Mexico, which shows an active normal
growth fault in an overpressured siliclastic setting to have a rela-
tively high up-fault permeability of w1 mD which sharply
increases to around 1D as uid pressure is further increased
towards the minimum effective principal stress (Losh and Haney,
2006). A similar observation was made from data from an ODP
borehole through the basal thrust of the Barbados accretionary
prism (Screaton et al., 1990). Seismic data shot on Eugene Island at
two different times over a seven year interval image the same
overpressure pulse in two different places e around 1 km apart,
leading to large-scale permeability estimates of around 0.1 Darcy
and the suggestion that faults can burp such overpressure pulses
upwards (Haney et al., 2005).
Healing of macro and microfracture networks in damaged rocks,
in the same way as for healing of fault gouge, is an important
temporal process in fault zones. The lifetime of fracture networks is
needed to predict cyclic fault zone permeability at depth. Seismo-
logical evidence has shown that the recovery of P and S wave
velocities (presumably related to healing of fracture damage)
following the 1992 Landers earthquake is quite rapid (<10 years)
(Vidale and Li, 2003). This recovery may only partially heal fault
fracture damage as the low-velocity zone surrounding faults
appears to be long-lived (Cochran et al., 2009).
5. Concluding remarks
Recent advances in the study of the structure, mechanics and
uid ow properties of faults and fault systems have been
reviewed. The importance of the interplay of these three properties
was emphasized. We conclude this work by highlighting some of
the key areas of ongoing research.
In terms of fault zone structure, the heterogeneity and along
fault variability are still poorly known. For example, seismological
methods are necessary to determine the structure of fault zones at
depth. Hence, if we are to reconcile the structure of fault zones at
depth with that observed at the surface, we must better understand
the dependence of seismological parameters on the physical
properties of the fault rocks. Basin-scale fault systems can benet
from revised models of fault growth related to geometric and
kinematic coherence. The impact of new 3D seismic datasets has
the potential to greatly enhance our understanding of fault growth
by providing a detailed view of growth strata that are intimately
associated with the growth of faults and development of fault
systems.
Our understanding of the mechanics of faulting and earthquakes
is still limited. The currently unresolved question of weak faulting
has helped to focus efforts, but direct observation via fault-drilling
projects is necessary to resolve 2010. Even then, the drill hole has to
penetrate a representative portion of the fault zone, with all the
inherent problems of fault-zone heterogeneity previously
mentioned. We can improve our knowledge of the mechanics of
earthquakes by integrating data from seismology, experiments and
eld geology. A key aspect is to improve our understanding of the
manifestations of seismic slip and its thermal effects in natural fault
rocks, if this is possible. This may be aided in the future with
comparison with microstructures developed in high velocity
experiments. Experiments (at both high and lowvelocity) must aim
to understand the underlying physics of slip.
Fluid ow around faults is dictated by 3D fault zone structure
and, as previously mentioned, this is likely to be heterogeneous.
Investigations of along-fault ow, in particular within crystalline
rocks at depth, showthat owcan be dominated by a small number
of fractures within the surrounding damage zone. Recent efforts
have helped to characterize the damage surrounding faults and we
can potentially reconcile this with laboratory measurements of
various fault-zone components (e.g. fault core or damage zone).
However, there remains a pressing need for a greater number of in
situ measurements over large areas, such as those exposed within
tunnels that characterize both the permeability of individual
features and whole fault zones at depth. Our understanding of the
temporal evolution of fault-zone permeability is still limited, but
we can address this by a combination of eld and laboratory
measurements.
Overall, as many aspects of faulting and fault systems are highly
interrelated, an integrated approach is necessary to make progress
in understanding their structure, mechanics and uid ow prop-
erties. This integrated approach will necessarily involve many
different disciplines, from eld geology, laboratory measurements,
geophysical measurements, modelling (numerical and experi-
mental), and direct observation of faults through drilling.
Acknowledgements
DRF thanks the Departamento de Geologia, Universidad de Chile
for their hospitality during the preparation of this manuscript. We
thank Bob Holdsworth for providing a thorough review that
improved the manuscript.
References
Aharonov, E., Tenthorey, E., Scholz, C.H., 1998. Precipitation sealing and diagenesis e
2. Theoretical analysis. Journal of Geophysical Research e Solid Earth 103 (B10),
23969e23981.
Allan, U.S., 1989. Model for hydrocarbon migration and entrapment within faulted
structures. American Association of Petroleum Geologists Bulletin 73, 803e811.
Ampuero, J.P., Rubin, A.M., 2008. Earthquake nucleation on rate and state faults -
aging and slip laws. Journal of Geophysical Research e Solid Earth 113 (B1).
Anastasio, D.J., Erslev, E.A., Fisher, D.M., 1997. Preface: fault-related folding. Journal
of Structural Geology 19, vevi.
Anders, M.H., Wiltschko, D.V., 1994. Microfracturing, paleostress and the growth of
faults. Journal of Structural Geology 16 (6), 795e815.
Axen, G.J., 2007. Research focus: signicance of large-displacement, low-angle
normal faults. Geology 35 (3), 287e288.
Balsamo, F., Storti, F., Salvini, F., Silva, A., Lima, C., 2010. Structural and petrophysical
evolution of extensional fault zones in low-porosity, poorly lithied sandstones
of the Barreiras Formation, NE Brazil. Journal of Structural Geology 32 (11),
1806e1826.
Baudon, C., Cartwright, J., 2008. Early stage evolution of growth faults: 3D seismic
insights from the Levant Basin, Eastern Mediterranean. Journal of Structural
Geology 30 (7), 888e898.
Beeler, N.M., Hickman, S.H., 2004. Stress-induced, time-dependent fracture closure
at hydrothermal conditions. Journal of Geophysical Research 109.
Berg, S.S., Skar, T., 2005. Controls on damage zone asymmetry of a normal fault
zone: outcrop analyses of a segment of the Moab fault, SE Utah. Journal Of
Structural Geology 27 (10), 1803e1822.
Beroza, G.C., Ellsworth, W.L., 1996. Properties of the seismic nucleation phase.
Tectonophysics 261 (1e3), 209e227.
Biegel, R.L., Sammis, C.G., 2004. Relating fault mechanics to fault zone structure.
Advances in Geophysics 47, 65e111.
Biegel, R.L., Sammis, C.G., Rosakis, A.J., 2008. An experimental study of the effect of
off-fault damage on the velocity of a slip pulse. Journal of Geophysical Research
e Solid Earth 113 (B4).
Bizzarri, A., Cocco, M., 2006. A thermal pressurization model for the spontaneous
dynamic rupture propagation on a three-dimensional fault: 2. Traction evolu-
tion and dynamic parameters. Journal of Geophysical Research e Solid Earth 111
(B5).
Blanpied, M.L., Lockner, D.A., Byerlee, J.D., 1992. An earthquake mechanism based on
rapid sealing of faults. Nature 358 (6387), 574e576.
D.R. Faulkner et al. / Journal of Structural Geology 32 (2010) 1557e1575 1570
Blanpied, M.L., Marone, C.J., Lockner, D.A., Byerlee, J.D., King, D.P., 1998. Quantitative
measure of the variation in fault rheology due to uiderock interactions.
Journal of Geophysical Research e Solid Earth 103 (B5), 9691e9712.
Blenkinsop, T.G., 2008. Relationships between faults, extension fractures and veins,
and stress. Journal of Structural Geology 30 (5), 622e632.
Bolton, A.J., Maltman, A.J., Clennell, M.B., 1998. The importance of overpressure
timing and permeability evolution in ne-grained sediments undergoing shear.
Journal of Structural Geology 20 (8), 1013e1022.
Boness, N.L., Zoback, M.D., 2006. A multiscale study of the mechanisms controlling
shear velocity anisotropy in the San Andreas Fault Observatory at depth.
Geophysics 71 (5), F131eF146.
Bonnet, E., Bour, O., Odling, N.E., Davy, P., Main, I., Cowie, P., Berkowitz, B., 2001.
Scaling of fracture systems in geological media. Reviews of Geophysics 39 (3),
347e383.
Bos, B., Peach, C.J., Spiers, C.J., 2000. Frictional-viscous ow of simulated fault gouge
caused by the combined effects of phyllosilicates and pressure solution. Tec-
tonophysics 327 (3-4), 173e194.
Bos, B., Spiers, C.J., 2001. Experimental investigation into the microstructural and
mechanical evolution of phyllosilicate-bearing fault rock under conditions
favouring pressure solution. Journal of Structural Geology 23 (8), 1187e1202.
Bowden, F.P., Tabor, D., 1950. The Friction and Lubrication of Solids. Oxford
University Press.
Bretan, P., Yielding, G., Jones, H., 2003. Using calibrated shale gouge ratio to estimate
hydrocarbon column heights. AAPG Bulletin 87 (3), 397e413.
Brister, B.S., Stephens, W.C., Norman, G.A., 2002. Structure, stratigraphy, and
hydrocarbon system of a Pennsylvanian pull-apart basin in north-central Texas.
AAPG Bulletin 86 (1), 1e20.
Brown, S., Caprihan, A., Hardy, R., 1998. Experimental observation of uid ow
channels in a single fracture. Journal of Geophysical Research e Solid Earth 103
(B3), 5125e5132.
Brown, S.R., Bruhn, R.L., 1998. Fluid permeability of deformable fracture networks.
Journal of Geophysical Research e Solid Earth 103 (B2), 2489e2500.
Brune, J.N., Henyey, T.L., Roy, R.F., 1969. Heat ow, stress, and rate of slip along the
San Andreas Fault, California. Journal of Geophysical Research 74, 3821e3827.
Bull, J.M., Barnes, P.M., Lamarche, G., Sanderson, D.J., Cowie, P.A., Taylor, S.K.,
Dix, J.K., 2006. High-resolution record of displacement accumulation on an
active normal fault: implications for models of slip accumulation during
repeated earthquakes. Journal of Structural Geology 28 (7), 1146e1166.
Byerlee, J., 1978. Friction of rocks. Pageoph 116, 615e626.
Byerlee, J., 1990. Friction, overpressure and fault normal compression. Geophysical
Research Letters 17 (12), 2109e2112.
Byerlee, J., 1993. Model for episodic ow of high-pressure water in fault zones
before earthquakes. Geology 21 (4), 303e306.
Caine, J.S., Evans, J.P., Forster, C.B., 1996. Fault zone architecture and permeability
structure. Geology 24 (11), 1025e1028.
Caine, J.S., Bruhn, R.L., Forster, C.B., 2010. Internal structure, fault rocks, and infer-
ences regarding deformation, uid ow, and mineralization in the seismogenic
stillwater normal fault, Dixie Valley, Nevada. Journal of Structural Geology 32
(11), 1576e1589.
Cartwright, J.A., Trudgill, B.D., Manseld, C.S., 1995. Fault growth by segment
linkage - an explanation for scatter in maximum displacement and trace length
data from the canyonlands grabens of SE Utah. Journal of Structural Geology 17
(9), 1319e1326.
Cembrano, J., Gonzalez, G., Arancibia, G., Ahumada, I., Olivares, V., Herrera, V., 2005.
Fault zone development and strain partitioning in an extensional strike-slip
duplex: A case study from the Mesozoic Atacama fault system, Northern Chile.
Tectonophysics 400 (1e4), 105e125.
Chery, J., Zoback, M.D., Hickman, S., 2004. A mechanical model of the San Andreas
fault and SAFOD pilot hole stress measurements. Geophysical Research Letters
31 (15).
Chester, F.M., 1994. Effects of temperature on friction: constitutive equations and
experiments with quartz gouge. Journal of Geophysical Research 99 (B4),
7247e7261.
Chester, F.M., Logan, J.M., 1986. Implications for mechanical-properties of brittle
faults from observations of the Punchbowl fault zone, California. Pure and
Applied Geophysics 124 (1-2), 79e106.
Chester, F.M., Friedman, M., Logan, J.M., 1985. Foliated cataclasites. Tectonophysics
111 (1e2), 139e146.
Chester, F.M., Evans, J.P., Biegel, R.L., 1993. Internal structure and weakening
mechanisms of the San-Andreas fault. Journal of Geophysical Research e Solid
Earth 98 (B1), 771e786.
Chester, J.S., Chester, F.M., Kronenberg, A.K., 2005. Fracture surface energy of the
Punchbowl fault, San Andreas system. Nature 437 (7055), 133e136.
Chiaraluce, L., Chiarabba, C., Collettini, C., Piccinini, D., Cocco, M., 2007. Architecture
and mechanics of an active low-angle normal fault: Alto Tiberina Fault,
northern Apennines, Italy. Journal of Geophysical Research e Solid Earth 112
(B10), B10310. doi:10.1029/2007JB005015.
Childs, C., Manzocchi, T., Walsh, J.J., Bonson, C.G., Nicol, A., Schopfer, M.P.J., 2009. A
geometric model of fault zone and fault rock thickness variations. Journal of
Structural Geology 31 (2), 117e127.
Christie-Blick, N., Biddle, K.T., 1985. Deformation and basin formation along strike-
slip faults. Strike-slip deformation, basin formation, and sedimentation. pp. 1e34.
Cochran, E.S., Li, Y.G., Shearer, P.M., Barbot, S., Fialko, Y., Vidale, J.E., 2009. Seismic
and geodetic evidence for extensive, long-lived fault damage zones. Geology 37
(4), 315e318.
Collettini, C., Holdsworth, R.E., 2004. Fault zone weakening and character of slip
along low-angle normal faults: insights from the Zuccale fault, Elba, Italy.
Journal of the Geological Society 161, 1039e1051.
Collettini, C., Sibson, R.H., 2001. Normal faults, normal friction? Geology 29 (10),
927e930.
Collettini, C., Niemeijer, A., Viti, C., Marone, C., 2009a. Fault zone fabric and fault
weakness. Nature.
Collettini, C., Viti, C., Smith, S.A.F., Holdsworth, R.E., 2009b. Development of inter-
connected talc networks and weakening of continental low-angle normal
faults. Geology 37 (6), 567e570.
Cook, J.E., Dunne, W.M., Onasch, C.A., 2006. Development of a dilatant damage zone
along a thrust relay in a low-porosity quartz arenite. Journal of Structural
Geology 28 (5), 776e792.
Cosgrove, J.W., Ameen, M.S., 2000. Forced folds and fractures. In: Geological Society
of London, Special Publications, vol. 169.
Cowie, P.A., Attal, M., Tucker, G.E., Whittaker, A.C., Naylor, M., Ganas, A.,
Roberts, G.P., 2006. Investigating the surface process response to fault inter-
action and linkage using a numerical modelling approach. Basin Research 18
(3), 231e266.
Cox, S.F., Knackstedt, M.A., Braun, J., 2001. Principles of structural control on
permeability and uid ow in hydrothermal systems. Reviews in Economic
Geology 14, 1e24.
Crawford, B.R., Faulkner, D.R., Rutter, E.H., 2008. Strength, porosity, and perme-
ability development during hydrostatic and shear loading of synthetic quartz-
clay fault gouge. Journal of Geophysical Research e Solid Earth 113 (B3).
Cuisiat, F., Jostad, H.P., Andresen, L., Skurtveit, E., Skomedal, E., Hettema, M.,
Lyslo, K., 2010. Geomechanical integrity of sealing faults during depressuriza-
tion of the Statfjord eld. Journal of Structural Geology 32 (11), 1754e1767.
Cristallini, E.O., Allmendinger, R.W., 2001. Pseudo 3-D modeling of trishear fault-
propagation folding. Journal of Structural Geology 23 (12), 1883e1899.
dAlessio, M., Martel, S.J., 2005. Development of strike-slip faults from dikes,
Sequoia National Park, California. Journal of Structural Geology 27 (1), 35e49.
Dawers, N.H., Anders, M.H., 1995. Displacement-length scaling and fault linkage.
Journal of Structural Geology 17 (5), 607.
de Joussineau, G., Aydin, A., 2007. The evolution of the damage zone with fault
growth in sandstone and its multiscale characteristics. Journal of Geophysical
Research 112.
De Paola, N., Collettini, C., Faulkner, D.R., Trippetta, F., 2008. Fault zone architecture
and deformation processes within evaporitic rocks in the upper crust. Tectonics
27 (4).
Di Toro, G., Pennacchioni, G., 2005. Fault plane processes and mesoscopic structure
of a strong-type seismogenic fault in tonalites (Adamello batholith, Southern
Alps). Tectonophysics 402 (1e4), 55e80.
Di Toro, G., Nielsen, S., Pennacchioni, G., 2005. Earthquake rupture dynamics frozen
in exhumed ancient faults. Nature 436, 1009e1012.
Di Toro, G., Hirose, T., Nielsen, S., Pennacchioni, G., Shimamoto, T., 2006. Natural and
experimental evidence of melt lubrication of faults during earthquakes. Science
311 (5761), 647e649.
Dieterich, J.H., Kilgore, B., 1996. Implications of fault constitutive properties for
earthquake prediction. Proceedings of the National Academy of Sciences of the
United States of America 93 (9), 3787e3794.
Dieterich, J.H., Kilgore, B.D., 1994. Direct observation of frictional contacts - new
insights for state-dependent properties. Pure and Applied Geophysics 143
(1e3), 283e302.
Dincau, A.R., 1998. Prediction and timing of production induced fault seal break-
down in the South Marsh Island 66 gas eld. Gulf Coast Association of
Geological Societies Transactions XLVIII, 21e32.
Do Nascimento, A.F., Lunn, R.J., Cowie, P.A., 2005. Modeling the heterogeneous
hydraulic properties of faults using constraints from reservoir-induced seis-
micity. Journal of Geophysical Research e Solid Earth 110 (B9).
Dockrill, B., Shipton, Z.K., 2010. Structural controls on leakage from a natural CO
2
geologic storage site: central Utah. U.S.A. Journal of Structural Geology 32 (11),
1768e1782.
Dor, O., Ben-Zion, Y., Rockwell, T.K., Brune, J., 2006. Pulverized rocks in the Mojave
section of the San Andreas Fault Zone. Earth and Planetary Science Letters 245
(3e4), 642e654.
Douglas, M., Clark, I.D., Raven, K., Bottomley, D., 2000. Groundwater mixing
dynamics at a Canadian Shield mine. Journal of Hydrology 235 (1e2), 88e103.
Eichhubl, P., DOnfro, P.S., Aydin, A., Waters, A., McCarty, D.K., 2005. Structure, petro-
physics, and diagenesis of shale entrained along a normal fault at Black Diamond
Mines, California e implications for fault seal. AAPG Bulletin 89 (9), 1113e1137.
Eichhubl, P., Davatzes, N.C., Becker, S.P., 2009. Structural and diagenetic control of
uid migration and cementation along the Moab fault, Utah. AAPG Bulletin 93
(5), 653e681.
Ellsworth, W.L., Beroza, G.C., 1995. Seismic evidence for a earthquake nucleation
phase. Science 268 (5212), 851e855.
Escartin, J., Andreani, M., Hirth, G., Evans, B., 2008. Relationships between the
microstructural evolution and the rheology of talc at elevated pressures and
temperatures. Earth and Planetary Science Letters 268 (3e4), 463e475.
Evans, J.P., 1990. Thickness displacement relationships for fault zones. Journal of
Structural Geology 12 (8), 1061e1065.
Evans, J.P., Chester, F.M., 1995. Fluiderock interaction in faults of the San-Andreas
system e inferences from San-Gabriel fault rock geochemistry and micro-
structures. Journal of Geophysical Research e Solid Earth 100 (B7),
13007e13020.
D.R. Faulkner et al. / Journal of Structural Geology 32 (2010) 1557e1575 1571
Evans, K.F., Genter, A., Sausse, J., 2005. Permeability creation and damage due to
massive uid injections into granite at 3.5 km at Soultz: 1. Borehole observa-
tions. Journal of Geophysical Research e Solid Earth 110 (B4).
Fairley, J.P., 2009. Modeling uid ow in a heterogeneous, fault-controlled hydro-
thermal system. Geouids 9 (2), 153e166.
Fairley, J.P., Hinds, J.J., 2004. Field observation of uid circulation patterns in
a normal fault system. Geophysical Research Letters 31, 1e4.
Faulkner, D.R., 2004. A model for the variation in permeability of clay-bearing fault
gouge with depth in the brittle crust. Geophysical Research Letters 31 (19), 5.
Faulkner, D.R., Lewis, A.C., Rutter, E.H., 2003. On the internal structure and
mechanics of large strike-slip fault zones: eld observations of the Carboneras
fault in southeastem Spain. Tectonophysics 367 (3e4), 235e251.
Faulkner, D.R., Mitchell, T.M., Healy, D., Heap, M.J., 2006. Slip on weak faults by the
rotation of regional stress in the fracture damage zone. Nature 444 (7121),
922e925.
Faulkner, D.R., Mitchell, T.M., Rutter, E.H., Cembrano, J., 2008. On the structure and
mechanical properties of large strike-slip faults. In: Wibberley, C.A.J., Kurz, W.,
Imber, J., Holdsworth., R.E., Collettini, C. (Eds.), Structure of Fault Zones:
Implications for Mechanical and Fluid-ow Properties. Geological Society of
London Special Publication, vol. 299, pp. 139e150.
Faulkner, D.R., Rutter, E.H., 1998. The gas permeability of clay-bearing fault gouge at
20
C. In: Jones, G., Fisher, Q., Knipe, R.J. (Eds.), Faults, Fault Sealing and Fluid
Flow in Hydrocarbon Reservoirs. Geological Society of London, Special Publi-
cation, vol. 147, pp. 147e156.
Faulkner, D.R., Rutter, E.H., 2000. Comparisons of water and argon permeability in
natural clay-bearing fault gouge under high pressure at 20 degrees C. Journal of
Geophysical Research e Solid Earth 105 (B7), 16415e16426.
Faulkner, D.R., Rutter, E.H., 2001. Can the maintenance of overpressured uids in
large strike-slip fault zones explain their apparent weakness? Geology 29 (6),
503e506.
Faulkner, D.R., Rutter, E.H., 2003. The effect of temperature, the nature of the pore
uid, and subyield differential stress on the permeability of phyllosilicate-rich
fault gouge. Journal of Geophysical Research e Solid Earth 108 (B5).
Ferrill, D.A., Winterle, J., Wittmeyer, G., Sims, D., Colton, S., Armstrong, A.,
Morris, A.P., 1999. Stressed rock strains groundwater at Yucca Mountain,
Nevada. GSA Today 9, 1e8.
Fossen, H., Schultz, R.A., Shipton, Z.K., Mair, K., 2007. Deformation bands in sand-
stone: a review. Journal of the Geological Society 164, 755e769.
Freeman, B., Boult, P., Yielding, G., Menpez, S., 2010. Using empirical geological rules
to reduce structural uncertainty in seismic interpretation of faults. Journal of
Structural Geology 32 (11), 1668e1676.
Frye, K.M., Marone, C., 2002. Effect of humidity on granular friction at room
temperature. Journal of Geophysical Research e Solid Earth 107 (B11).
Gawthorpe, R.L., Sharp, I., Underhill, J.R., Gupta, S., 1997. Linked sequence strati-
graphic and structural evolution of propagating normal faults. Geology 25 (9),
795e798.
Giger, S.B., Tenthorey, E., Cox, S.F., Gerald, J.D.F., 2007. Permeability evolution in
quartz fault gouges under hydrothermal conditions. Journal of Geophysical
Research e Solid Earth 112 (B7).
Goldsby, D.L., Tullis, T.E., 2002. Low frictional strength of quartz rocks at subseismic
slip rates. Geophysical Research Letters 29.
Gratier, J.P., Guiguet, R., Renard, F., Jenatton, L., Bernard, D., 2009. A pressure solu-
tion creep law for quartz from indentation experiments. Journal of Geophysical
Research e Solid Earth 114.
Gratier, J.P., Renard, F., Labaume, P., 1999. How pressure solution creep and frac-
turing processes interact in the upper crust to make it behave in both a brittle
and viscous manner. Journal of Structural Geology 21 (8e9), 1189e1197.
Grauls, D., Pascaud, F., Rives, T., 2002. Quantitative fault seal assessment in
hydrocarbon-compartmentalised structures using uid pressure data. In:
Koestler, A.G., Hunsdale, R. (Eds.), Hydrocarbon Seal Quantication. NPF Special
Publication, vol. 11, pp. 141e156.
Green, H.W., Marone, C., 2002. Instability of deformation. Reviews in Mineralogy
and Geochemistry 51, 181e199.
Gueydan, F., Leroy, Y.M., Jolivet, L., Agard, P., 2003. Analysis of continental mid-
crustal strain localization induced by microfracturing and reaction-softening.
Journal of Geophysical Research e Solid Earth 108 (B2).
Haines, S.H., van der Pluijm, B.A., Ikari, M.J., Saffer, D.M., Marone, C., 2009. Clay
fabric intensity in natural and articial fault gouges: implications for brittle
fault zone processes and sedimentary basin clay fabric evolution. Journal of
Geophysical Research e Solid Earth 114.
Han, R., Shimamoto, T., Hirose, T., Ree, J.H., Ando, J., 2007. Ultralow friction of
carbonate faults caused by thermal decomposition. Science 316 (5826),
878e881.
Harding, T.P., 1990. Identication of wrench faults using subsurface structural data:
criteria and pitfalls. American Association of Petroleum Geologists Bulletin 74
(10), 1590e1609.
Haney, M.M., Snieder, R., Sheiman, J., Losh, S., 2005. A moving uid pulse in a fault
zone. Nature 437, 46.
Henza, A.A., Withjack, M.O., Schlische, R.W., 2010. Normal-fault development
during two phases of non-coaxial extension: An experimental study. Journal of
Structural Geology 32 (11), 1656e1667.
Heaton, T.H., 1990. Evidence for and implications of self-healing pulses of slip in
earthquake rupture. Physics of the Earth and Planetary Interiors 64 (1), 1e20.
Hickman, S., Zoback, M., 2004. Stress orientations and magnitudes in the SAFOD
pilot hole. Geophysical Research Letters 31 (15).
Hickman, S.H., Evans, B., 1995. Kinetics of pressure solution at halite-silica interfaces
and intergranular clay lms. Journal of Geophysical Research e Solid Earth 100
(B7), 13113e13132.
Hickman, S., Sibson, R., Bruhn, R., 1995. Introduction to special section e mechanical
involvement of uids in faulting. Journal of Geophysical Research e Solid Earth
100 (B7), 12831e12840.
Hirose, T., Bystricky, M., 2007. Extreme dynamic weakening of faults during dehy-
dration by coseismic shear heating. Geophysical Research Letters 34 (14).
Holdsworth, R.E., 2004. Weak faults e rotten cores. Science 303 (5655), 181e182.
Holyoke, C.W., Tullis, J., 2006. The interaction between reaction and deformation:
an experimental study using a biotite plus plagioclase plus quartz gneiss.
Journal of Metamorphic Geology 24 (8), 743e762.
Hubbert, M.K., Rubey, W.W., 1959. Role of uid pressure in mechanics of overthrust
faulting 1. Mechanics of uid lled porous solids and its application of over-
thrust faulting. Geological Society of America Bulletin 70 (2), 115e166.
Ikari, M.J., Saffer, D.M., Marone, C., 2009. Frictional and hydrologic properties of
clay-rich fault gouge. Journal of Geophysical Research 114.
Imber, J., Holdsworth, R.E., Butler, C.A., Strachan, R.A., 2001. A reappraisal of the
Sibson-Scholz fault zone model: the nature of the frictional to viscous (brit-
tleeductile) transition along a long-lived, crustal-scale fault, Outer Hebrides,
Scotland. Tectonics 20 (5), 601e624.
Ishii, E., Funaki, H., Tokiwa, T., Ota, K. Relationship between growth mechanism of
faults and permeability variations with depth in siliceous mudstone. Journal of
Structural Geology, in press.
Janecke, S.U., Vandenburg, C.J., Blankenau, J.J., 1998. Geometry, mechanisms and
signicance of extensional folds from examples in the Rocky Mountain Basin
and Range province, USA. Journal of Structural Geology 20 (7), 841e856.
Janssen, C., Wagner, F.C., Zang, A., Dresen, G., 2001. Fracture process zone in granite:
a microstructural analysis. International Journal of Earth Sciences 90 (1), 46e59.
Jefferies, S.P., Holdsworth, R.E., Shimamoto, T., Takagi, H., Lloyd, G.E., Spiers, C.J.,
2006a. Origin and mechanical signicance of foliated cataclastic rocks in the
cores of crustal-scale faults: examples from the Median Tectonic Line, Japan.
Journal of Geophysical Research e Solid Earth 111 (B12).
Jefferies, S.P., Holdsworth, R.E., Wibberley, C.A.J., Shimamoto, T., Spiers, C.J.,
Niemeijer, A.R., Lloyd, G.E., 2006b. The nature and importance of phyllonite
development in crustal-scale fault cores: an example from the Median Tectonic
Line, Japan. Journal of Structural Geology 28 (2), 220e235.
Johansen, T.E.S., Fossen, H., Kluge, R., 2005. The impact of syn-faulting porosity
reduction on damage zone architecture in porous sandstone: an outcrop
example from the Moab Fault, Utah. Journal of Structural Geology 27 (8),
1469e1485.
Jourde, H., Flodin, E.A., Aydin, A., Durlofsky, L.J., Wen, X.H., 2002. Computing
permeability of fault zones in eolian sandstone from outcrop measurements.
AAPG Bulletin 86 (7), 1187e1200.
Kanamori, H., Brodsky, E.E., 2004. The physics of earthquakes. Reports on Progress in
Physics 67 (8). doi:10.1088/0034-4885/67/8/R03 pii: S0034-4885(04)25227-7.
Kanamori, H., Rivera, L., 2006. Energy partitioning during an earthquake. In:
Abercrombie, R., McGarr, A., Kanamori, H., Di Toro, G. (Eds.), Earthquakes:
Radiated Energy and the Physics of Faulting. Geophysical Monograph Series,
vol. 170, pp. 3e15.
Karner, S.L., Marone, C., Evans, B., 1997. Laboratory study of fault healing and lith-
ication in simulated fault gouge under hydrothermal conditions. Tectono-
physics 277 (1e3), 41e55.
Kim, Y.S., Peacock, D.C.P., Sanderson, D.J., 2004. Fault damage zones. Journal of
Structural Geology 26 (3), 503e517.
Kim, Y.S., Sanderson, D.J., 2005. The relationship between displacement and length
of faults: a review. Earth-Science Reviews 68 (3e4), 317e334.
Kirkpatrick, J.D., Shipton, Z.K., Evans, J.P., Micklethwaite, S., Lim, S.J., McKillop, P.,
2008. Strike-slip fault terminations at seismogenic depths: the structure and
kinematics of the Glacier Lakes fault, Sierra Nevada United States. Journal of
Geophysical Research e Solid Earth 113 (B4).
Kohlstedt, D.L., Evans, B., Mackwell, S.J., 1995. Strength of the lithosphere - con-
staints imposed by laboratory experiments. Journal of Geophysical Research e
Solid Earth 100 (B9), 17587e17602.
Kristensen, M.B., Childs, C.J., Korstgard, J.A., 2008. The 3D geometry of small-scale
relay zones between normal faults in soft sediments. Journal of Structural
Geology 30 (2), 257e272.
Lachenbruch, A.H., Sass, J.H., 1980. Heat-ow and energetics of the San-Andreas
fault zone. Journal of Geophysical Research 85 (Nb11), 6185e6222.
Lindsay, N.G., Murphy, F.C., Walsh, J.J., Watterson, J., 1993. Outcrop studies of shale
smear on fault surfaces. In: International Association of Sedimentology, Special
Publications, vol. 15 113e123.
Logan, J.M., Friedman, M., Higgs, N.G., Dengo, C.A. and Shimamoto, T. 1979. Exper-
imental studies of simulated gouge and their application to studies of natural
fault zones. Analysis of actual fault zones in bedrock. U.S. Geological Survey
Open File Report 79-1239, pp. 305e343.
Logan, J.M., Rauenzahn, K.A., 1987. Frictional dependence of gouge mixtures of
quartz and montmorillonite on velocity, composition and fabric. Tectonophysics
144 (1e3), 87e108.
Lohr, T., Krawczyk, C.M., Oncken, O., Tanner, D.C., 2008. Evolution of a fault surface
from 3D attribute analysis and displacement measurements. Journal of Struc-
tural Geology 30 (6), 690e700.
Losh, S., Haney, M., 2006. Episodic uid ow in an aseismic overpressured growth
fault, northern Gulf of Mexico. In: Abercrombie, R., McGarr, A., Di Toro, G.,
Kanamori, H. (Eds.), Earthquakes: Radiated Energy and the Physics of Faulting.
D.R. Faulkner et al. / Journal of Structural Geology 32 (2010) 1557e1575 1572
American Geophysical Union Geophysical Monograph Series, vol. 170, pp.
199e206.
Lunn, R.J., Willson, J.P., Shipton, Z.K., Moir, H., 2008. Simulating brittle fault growth
from linkage of preexisting structures. Journal of Geophysical Research e Solid
Earth 113 (B7).
Main, I.G., Kwon, O., Ngwenya, B.T., Elphick, S.C., 2000. Fault sealing during defor-
mation-band growth in porous sandstone. Geology 28 (12), 1131e1134.
Manzocchi, T., Walsh, J.J., Nell, P., Yielding, G., 1999. Fault transmissibility multipliers
for ow simulation models. Petroleum Geoscience 5 (1), 53e63.
Mares, V.M., Kronenberg, A.K., 1993. Experimental deformation of muscovite.
Journal of Structural Geology 15 (9e10), 1061e1075.
Mariani, E., Brodie, K.H., Rutter, E.H., 2006. Experimental deformation of muscovite
shear zones at high temperatures under hydrothermal conditions and the
strength of phyllosilicate-bearing faults in nature. Journal of Structural Geology
28 (9), 1569e1587.
Marone, C., 1998. Laboratory-derived friction laws and their application to seismic
faulting. Annual Review of Earth and Planetary Sciences 26, 643e696.
Marone, C., Kilgore, B., 1993. Scaling of the critical slip distance for seismic faulting
with shear strain in fault zones. Nature 362 (6421), 618e621.
Marone, C., Richardson, E., 2006. Do earthquakes rupture piece by piece or all
together? Science 313 (5794), 1748e1749.
Marone, C., Scholz, C.H., 1989. Particle-size distribution and microstructures within
simulated fault Gouge. Journal of Structural Geology 11 (7), 799e814.
Martel, S.J., Pollard, D.D., Segall, P., 1988. Development of simple strike-slip fault
zones, Mount Abbot Quadrangle, Sierra Nevada, California. Geological Society of
America Bulletin 100 (9), 1451e1465.
Matthai, S.K., Belayneh, M., 2004. Fluid ow partitioning between fractures and
a permeable rock matrix. Geophysical Research Letters 31 (7).
McClay, K.R., 2004. Thrust tectonics and hydrocarbonsystems. AAPGMemoir 82, 667.
McGrath, A.G., Davison, I., 1995. Damage zone geometry around fault tips. Journal of
Structural Geology 17 (7), 1011e1024.
Medeiros W.E., do Nascimento, A.F., Alves da Silva, F.C., Destro, N., Demtrio, J.G.A.
Evidence of hydraulic connectivity across deformation bands from eld
pumping tests: Two examples from Tucano Basin, NE Brazil. Journal of Struc-
tural Geology, this issue.
Micarelli, L., Benedicto, A., Wibberley, C.A.J., 2006. Structural evolution and
permeability of normal fault zones in highly porous carbonate rocks. Journal of
Structural Geology 28 (7), 1214e1227.
Miller, S.A., Collettini, C., Chiaraluce, L., Cocco, M., Barchi, M., Kaus, B.J.P., 2004.
Aftershocks driven by a high-pressure CO
2
source at depth. Nature 427 (6976),
724e727.
Mitchell, T.M., Faulkner, D.R., 2008. Experimental measurements of permeability
evolution during triaxial compression of initially intact crystalline rocks and
implications for uid ow in fault zones. Journal of Geophysical Research e
Solid Earth 113 (B11).
Mitchell, T.M., Faulkner, D.R., 2009. The nature and origin of off-fault damage
surrounding strike-slip fault zones with a wide range of displacements: A eld
study from the Atacama fault zone, northern Chile. Journal of Structural
Geology 31, 802e816.
Mizoguchi, K., Hirose, T., Shimamoto, T., Fukuyama, E., 2008. Internal structure and
permeability of the Nojima fault, southwest Japan. Journal of Structural Geology
30 (4), 513e524.
Mo, Y.F., Turner, K.T., Szlufarska, I., 2009. Friction laws at the nanoscale. Nature 457
(7233), 1116e1119.
Moir, H., Lunn, R., Shipton, Z., Kirkpatrick, J. Simulating brittle fault evolution from
networks of pre-existing structures. Journal of Structural Geology, this issue.
Moore, D.E., Lockner, D.A., 2004. Crystallographic controls on the frictional behavior
of dry and water-saturated sheet structure minerals. Journal of Geophysical
Research e Solid Earth 109 (B3).
Moore, D.E., Lockner, D.A., 2008. Talc friction in the temperature range 25 degrees-
400 degrees C: relevance for fault-zone weakening. Tectonophysics 449 (1e4),
120e132.
Moore, D.E., Lockner, D.A., Tanaka, H., Iwata, K., 2004. The coefcient of friction of
Chrysotile gouge at seismogenic depths. International Geology Review 46 (5),
385e398.
Moore, D.E., Rymer, M.J., 2007. Talc-bearing serpentinite and the creeping section of
the San Andreas fault. Nature 448 (7155), 795e797.
Morrow, C.A., Lockner, D.A., 1994. Permeability differences between surface-derived
and deep drillhole core samples. Geophysical Research Letters 21 (19),
2151e2154.
Morrow, C.A., Moore, D.E., Lockner, D.A., 2001. Permeability reduction in granite
under hydrothermal conditions. Journal of Geophysical Research e Solid Earth
106 (B12), 30551e30560.
Morrow, C.A., Radney, B., Byerlee, J.D., 1992. Frictional strength and the effective
pressure law of montmorillonite and illite clays. In: Evans, B., Wong, T.-F. (Eds.),
Fault Mechanics and Transport Properties of Rocks. Academic Press, pp. 69e88.
Nakatani, M., 2001. Conceptual and physical clarication of rate and state friction:
frictional sliding as a thermally activated rheology. Journal of Geophysical
Research e Solid Earth 106 (B7), 13347e13380.
Nakatani, M., Scholz, C.H., 2004a. Frictional healing of quartz gouge under hydro-
thermal conditions: 1. Experimental evidence for solution transfer healing
mechanism. Journal of Geophysical Research e Solid Earth 109 (B7).
Nakatani, M., Scholz, C.H., 2004b. Frictional healing of quartz gouge under hydro-
thermal conditions: 2. Quantitative interpretation with a physical model.
Journal of Geophysical Research e Solid Earth 109 (B7).
Nemser, E.S., Cowan, D.S., 2009. Downdip segmentation of strike-slip fault zones in
the brittle crust. Geology 37 (5), 419e422.
Noda, H., 2008. Frictional constitutive law at intermediate slip rates accounting for
ash heating and thermally activated slip process. Journal of Geophysical
Research e Solid Earth 113 (B9).
Noda, H., Dunham, E.M., Rice, J.R., 2009. Earthquake ruptures with thermal weak-
ening and the operation of major faults at low overall stress levels. Journal of
Geophysical Research 114.
Noir, J., Jacques, E., Bekri, S., Adler, P.M., Tapponnier, P., King, G.C.P., 1997. Fluid ow
triggered migration of events in the 1989 Dobi earthquake sequence of Central
Afar. Geophysical Research Letters 24 (18), 2335e2338.
Numelin, T., Marone, C., Kirby, E., 2007. Frictional properties of natural fault gouge
from a low-angle normal fault, Panamint Valley, California. Tectonics 26 (2).
OHara, K., 2007. Reaction weakening and emplacement of crystalline thrusts:
diffusion control on reaction rate and strain rate. Journal of Structural Geology
29 (8), 1301e1314.
Oda, M., Takemura, T., Aoki, T., 2002. Damage growth and permeability change in
triaxial compressiontests of Inada granite. Mechanics of Materials 34 (6), 313e331.
Odling, N.E., Harris, S.D., Knipe, R., 2004. Permeability scaling properties of fault
damage zones in siliclastic rocks. Journal of Structural Geology 26 (9),
1727e1747.
Olsen, M.P., Scholz, C.H., Leger, A., 1998. Healing and sealing of a simulated fault
gouge under hydrothermal conditions: implications for fault healing. Journal of
Geophysical Research e Solid Earth 103 (B4), 7421e7430.
Peacock, D.C.P., Sanderson, D.J., 1991. Displacements, segment linkage and relay
ramps in normal fault zones. Journal of Structural Geology 13 (6), 721.
Price, R.A., 1988. The mechanical paradox of large overthrusts. Geological Society of
America Bulletin 100 (12), 1898e1908.
Rawling, G.C., Goodwin, L.B., Wilson, J.L., 2001. Internal architecture, permeability
structure, and hydrologic signicance of contrasting fault-zone types. Geology
29 (1), 43e46.
Reches, Z., Dewers, T.A., 2005. Gouge formation by dynamic pulverization during
earthquake rupture. Earth and Planetary Science Letters 235 (1e2), 361e374.
Reches, Z., Lockner, D.A., 1994. Nucleation and growth of faults in brittle rocks.
Journal of Geophysical Research 99 (B9), 18159e18173.
Reinen, L.A., 2000. Slip styles in a spring-slider model with a laboratory-derived
constitutive law for serpentinite. Geophysical Research Letters 27 (14),
2037e2040.
Rempel, A.W., Rice, J.R., 2006. Thermal pressurization and onset of melting in fault
zones. Journal of Geophysical Research e Solid Earth 111 (B9).
Renard, F., Gratier, J.P., Jamtveit, B., 2000. Kinetics of crack-sealing, intergranular
pressure solution, and compaction around active faults. Journal of Structural
Geology 22 (10), 1395e1407.
Revil, A., Grauls, D., Brevart, O., 2002. Mechanical compaction of sand/clay mixtures.
Journal of Geophysical Research e Solid Earth 107 (B11).
Rice, J.R., 1992. Fault stress states, pore pressure distributions, and the weakness of
the San Andreas fault. In: Evans, B., Wong, T.-F. (Eds.), Fault Mechanics and
Transport Properties in Rocks. Academic Press, p. 28.
Rice, J.R., 2006. Heating and weakening of faults during earthquake slip. Journal of
Geophysical Research e Solid Earth 111 (B5).
Rice, J.R., Lapusta, N., Ranjith, K., 2001. Rate and state dependent friction and the
stability of sliding between elastically deformable solids. Journal of the
Mechanics and Physics of Solids 49 (9), 1865e1898.
Roberts, G.P., Houghton, S.L., Underwood, C., Papanikolaou, I., Cowie, P.A., van
Calsteren, P., Wigley, T., Cooper, F.J., McArthur, J.M., 2009. Localization of
Quaternary slip rates in an active rift in 10(5) years: an example from central
Greece constrained by U-234eTh-230 coral dates from uplifted paleoshore-
lines. Journal of Geophysical Research e Solid Earth 114.
Rowland, J.V., Sibson, R.H., 2004. Structural controls on hydrothermal ow in
a segmented rift system, Taupo Volcanic Zone, New Zealand. Geouids 4 (4),
259e283.
Rudnicki, J.W., Rice, J.R., 2006. Effective normal stress alteration due to pore pres-
sure changes induced by dynamic slip propagation on a plane between
dissimilar materials. Journal of Geophysical Research e Solid Earth 111 (B10).
Rutter, E.H., Brodie, K.H., 1995. Mechanistic interactions between deformation and
metamorphism. Geological Journal 30 (3e4), 227e240.
Rutter, E.H., Maddock, R.H., Hall, S.H., White, S.H., 1986. Comparative microstruc-
tures of natural and experimentally produced clay-bearing fault gouges. Pure
and Applied Geophysics 124 (1e2), 3e30.
Rutter, E.H., Mainprice, D.H., 1979. On the possibility of slow fault slip controlled by
a diffusive mass transfer process. Gerlands Beitrage zur Geophysik 88, 154e162.
Saffer, D.M., Frye, K.M., Marone, C., Mair, K., 2001. Laboratory results indicating
complex and potentially unstable frictional behavior of smectite clay.
Geophysical Research Letters 28 (12), 2297e2300.
Saffer, D.M., Marone, C., 2003. Comparison of smectite- and illite-rich gouge frictional
properties: application to the updip limit of the seismogenic zone along
subduction megathrusts. Earth and Planetary Science Letters 215 (1-2), 219e235.
Saillet, E., Wibberley, C. Evolution of cataclastic faulting in high porosity sandstone,
Bassin du Sud-Est, Provence, France. Journal of Structural Geology, this issue.
Sammis, C., King, G., Biegel, R., 1987. The kinematics of gouge deformation. Pure and
Applied Geophysics 125 (5), 777e812.
Sammis, C.G., Ben-Zion, Y., 2008. Mechanics of grain-size reduction in fault zones.
Journal of Geophysical Research e Solid Earth 113 (B2).
Savage, H.M., Brodsky, E.E. Collateral damage: capturing slip delocalization in
fracture proles. Journal of Geophysical Research, submitted for publication.
D.R. Faulkner et al. / Journal of Structural Geology 32 (2010) 1557e1575 1573
Savage, H.M., Cooke, M.L., 2010. Unlocking the effects of friction on fault damage
zone models. Journal of Structural Geology 32 (11), 1732e1741.
Schlische, R.W., 1995. Geometry and origin of fault-related folds in extensional
settings. AAPG Bulletin e American Association of Petroleum Geologists 79 (11),
1661e1678.
Schlische, R.W., Withjack, M.O., Eisenstadt, G., 2002. An experimental study of the
secondary deformation produced by oblique-slip normal faulting. American
Association of Petroleum Geologists Bulletin 86 (5), 885e906.
Schmatz, J., Vrolijk, P., Urai, J., 2010. Clay smear in normal faults - the effect of
multilayers and clay cementation in water-saturated model experiments.
Journal of Structural Geology 32 (11), 1834e1849.
Scholz, C.H., 1987. Wear and Gouge formation in brittle faulting. Geology 15 (6),
493e495.
Scholz, C.H., 1988. The critical slip distance for seismic faulting. Nature 336 (6201),
761e763.
Scholz, C.H., 1998. Earthquakes and friction laws. Nature 391 (6662), 37e42.
Scholz, C.H., 2002. The Mechanics of Earthquakes and Faulting. Cambridge
University Press, Cambridge.
Scholz, C.H., 2006. The strength of the San Andreas fault: a critical analysis.
Earthquakes: Radiated Energy and the Physics of Faulting 170, 301e311.
Schulz, S.E., Evans, J.P., 2000. Mesoscopic structure of the Punchbowl Fault,
Southern California and the geologic and geophysical structure of active strike-
slip faults. Journal of Structural Geology 22 (7), 913e930.
Screaton, E.J., Wuthrich, D.R., Dreiss, S.J., 1990. Permeabilities, uid pressures, and
ow rates in the Barbados ridge complex. Journal of Geophysical Research e
Solid Earth and Planets 95 (B6), 8997e9007.
Scruggs, V.J., Tullis, T.E., 1998. Correlation between velocity dependence of friction
and strain localization in large displacement experiments on feldspar, musco-
vite and biotite gouge. Tectonophysics 295 (1e2), 15e40.
Segall, P., Rice, J.R., 2006. Does shear heating of pore uid contribute to earthquake
nucleation? Journal of Geophysical Research e Solid Earth 111 (B9).
Shapiro, S.A., Huenges, E., Borm, G., 1997. Estimating the crust permeability from
uid-injection-induced seismic emission at the KTB site. Geophysical Journal
International 131 (2), F15eF18.
Shaw, J.H., Connors, C., Suppe, J., 2005. Seismic interpretation of contractional fault-
related folds: an AAPG seismic atlas. American Association of Petroleum
Geologists Studies in Geology 53, 157.
Shipton, Z.K., Cowie, P.A., 2001. Damage zone and slip-surface evolution over mm to
km scales in high-porosity Navajo sandstone, Utah. Journal of Structural
Geology 23 (12), 1825e1844.
Shipton, Z.K., Cowie, P.A., 2003. A conceptual model for the origin of fault damage
zone structures in high-porosity sandstone. Journal of Structural Geology 25 (8),
1343e1345.
Shipton, Z.K., Evans, J.P., Robeson, K.R., Forster, C.B., Snelgrove, S., 2002. Structural
heterogeneity and permeability in faulted eolian sandstone: implications for
subsurface modeling of faults. AAPG Bulletin 86 (5), 863e883.
Shipton, Z.K., Evans, J.P., Thompson, L.B., 2005. The geometry and thickness of
deformation band fault core, and its inuence on sealing characteristics of
deformation band fault zones. American Association of Petroleum Geologists
Memoir 85, 181e195.
Shipton, Z.K., Soden, A.M., Kirkpatrick, J.D., Bright, A.M., Lunn, R.J., 2006. How thick
is a fault? Fault displacement-thickness scaling revisited. Earthquakes: Radiated
Energy and the Physics of Faulting 170, 193e198.
Sibson, R.H., 1990. Conditions for fault-valve behaviour. In: Geological Society of
London, Special Publication, vol. 54 15e28.
Sibson, R.H., 2001. Seismogenic framework for hydrothermal transport and ore
deposition. Reviews in Economic Geology 14, 25e50.
Sibson, R.H., 2009. Rupturing in overpressured crust during compressional inver-
sion-the case from NE Honshu, Japan. Tectonophysics 473 (3e4), 404e416.
Sibson, R.H., Xie, G.Y., 1998. Dip range for intracontinental reverse fault ruptures:
truth not stranger than friction? Bulletin of the Seismological Society of
America 88 (4), 1014e1022.
Simpson, G., Gueguen, Y., Schneider, F., 2001. Permeability enhancement due to
microcrack dilatancy in the damage regime. Journal of Geophysical Research e
Solid Earth 106 (B3), 3999e4016.
Smith, S.A.F., Faulkner, D.R., 2010. Laboratory measurements of the frictional
properties of a natural low-angle normal fault: The Zuccale fault, Elba Island,
Italy. Journal of Geophysical Research 115, B02407.
Smith, S.A.F., Holdsworth, R.E., Collettini, C., Imber, J., 2007. Using footwall struc-
tures to constrain the evolution of low-angle normal faults. Journal of the
Geological Society 164, 1187e1191.
Soliva, R., Benedicto, A., 2005. Geometry, scaling relations and spacing of vertically
restricted normal faults. Journal of Structural Geology 27 (2), 317e325.
Soliva, R., Maertan, F., Petit, J.-P., Auzias, V. Fault static friction and fracture orien-
tation in extensional relays; insight from eld data, photoelasticity and 3D
numerical modeling. Journal of Structural Geology, this issue.
Solum, J.G., van der Pluijm, B.A. Quantication of fabrics in clay gouge from the
Carboneras fault, Spain and implications for fault behavior. Tectonophysics 475
(3e4), 554e562.
Sperrevik, S., Gillespie, P.A., Fisher, Q., Halvorsen, T., Knipe, R.J., 2002. Empirical
estimation of fault rock properties. In: Koestler, A.G., Hunsdale, R. (Eds.),
Hydrocarbon Seal Quantication. NPF Special Publications, vol. 11, pp. 109e125.
Spudich, P., Guatteri, M., 2004. The effect of bandwidth limitations on the inference
of earthquake slip-weakening distance from seismograms. Bulletin of the
Seismological Society of America 94 (6), 2028e2036.
Streit, J.E., Hillis, R.R., 2004. Estimating fault stability and sustainable uid pressures
for underground storage of CO2
in porous rock. Energy 29 (9e10), 1445e1456.
Suppe, J., 1983. Geometry and kinematics of fault-bend folding. American Journal of
Science 283 (7), 684e721.
Swanson, M.T., 2005. Geometry and kinematics of adhesive wear in brittle strike-
slip fault zones. Journal of Structural Geology 27 (5), 871e887.
Szlufarska, I., Chandross, M., Carpick, R.W., 2008. Recent advances in single-asperity
nanotribology. Journal of Physics D e Applied Physics 41 (12).
Tadokoro, K., Ando, M., Nishigami, K., 2000. Induced earthquakes accompanying the
water injection experiment at the Nojima fault zone, Japan: seismicity and its
migration. Journal of Geophysical Research e Solid Earth 105 (B3), 6089e6104.
Takahashi, M., Mizoguchi, K., Kitamura, K., Masuda, K., 2007. Effects of clay content
on the frictional strength and uid transport property of faults. Journal of
Geophysical Research e Solid Earth 112 (B8).
Talwani, P., Cobb, J.S., Schaeffer, M.F., 1999. In situ measurements of hydraulic
properties of a shear zone in northwestern South Carolina. Journal of
Geophysical Research e Solid Earth 104 (B7), 14993e15003.
Talwani, P., Chen, L., Gahalaut, K., 2007. Seismogenic permeability, k(S). Journal of
Geophysical Research e Solid Earth 112 (B7).
Tembe, S., Lockner, D.A., Solum, J.G., Morrow, C.A., Wong, T.F., Moore, D.E., 2006.
Frictional strength of cuttings and core from SAFOD drillhole phases 1 and 2.
Geophysical Research Letters 33 (23).
Tenthorey, E., Scholz, C.H., Aharonov, E., Leger, A., 1998. Precipitation sealing and
diagenesis e 1. Experimental results. Journal of Geophysical Research e Solid
Earth 103 (B10), 23951e23967.
Tenthorey, E., Cox, S.F., Todd, H.F., 2003. Evolution of strength recovery and
permeability during uid-rock reaction in experimental fault zones. Earth and
Planetary Science Letters 206 (1e2), 161e172.
Tindall, S.E., Davis, G.H., 1999. Monocline development by oblique-slip fault-prop-
agation folding: the East Kaibab monocline, Colorado Plateau, Utah. Journal of
Structural Geology 21 (10), 1303e1320.
Tinti, E., Cocco, M., Fukuyama, E., Piatanesi, A., 2009. Dependence of slip weakening
distance (D-c) on nal slip during dynamic rupture of earthquakes. Geophysical
Journal International 177 (3), 1205e1220.
Tinti, E., Spudich, P., Cocco, M., 2005. Earthquake fracture energy inferred from
kinematic rupture models on extended faults. Journal of Geophysical Research
e Solid Earth 110 (B12).
Townend, J., Zoback, M.D., 2000. How faulting keeps the crust strong. Geology 28
(5), 399e402.
Tsutsumi, A., Shimamoto, T., 1997. High-velocity frictional properties of gabbro.
Geophysical Research Letters 24 (6), 699e702.
Tsutsumi, A., Nishino, S., Mizoguchi, K., Hirose, T., Uehara, S., Sato, K., Tanikawa, W.,
Shimamoto, T., 2004. Principal fault zone width and permeability of the active
Neodani fault, Nobi fault system, Southwest Japan. Tectonophysics 379 (1e4),
93e108.
Uehara, S., Shimamoto, T., 2004. Gas permeability evolution of cataclasite and fault
gouge in triaxial compression and implications for changes in fault-zone
permeability structure through the earthquake cycle. Tectonophysics 378 (3e4),
183e195.
van der Zee, W., Urai, J.L., 2005. Processes of normal fault evolution in a siliciclastic
sequence: a case study from Miri, Sarawak, Malaysia. Journal of Structural
Geology 27 (12), 2281e2300.
Van der Zee, W., Wibberley, C.A.J., Urai, J.L., 2008. The inuence of layering and pre-
existing joints on the development of internal structure in normal fault zones:
the Lodve basin, France. In: Wibberley, C.A.J., Kurz, W., Imber, J.,
Holdsworth, R.E., Collettini, C. (Eds.), The Internal Structure of Fault Zones:
Implications for Mechanical and Fluid Flow Properties. Geological Society of
London, Special Publication, vol. 299, pp. 57e74.
van Diggelen, E. W., De Bresser, J.H., Peach, C.J., Spiers, C.J. High shear strain
behaviour of synthetic muscovite fault gouges under hydrothermal conditions.
Journal of Structural Geology, this issue.
Vermilye, J.M., Scholz, C.H., 1998. The process zone: a microstructural view of
fault growth. Journal of Geophysical Research e Solid Earth 103 (B6),
12223e12237.
Vidale, J.E., Li, Y.G., 2003. Damage to the shallow Landers fault from the nearby
Hector Mine earthquake. Nature 421 (6922), 524e526.
Walker, J.P.F., Roberts, G.P., Cowie, P.A., Papanikolaou, I.D., Sammonds, P.R.,
Michetti, A.M., Phillips, R.J., 2009. Horizontal strain-rates and throw-rates
across breached relay zones, central Italy: implications for the preservation of
throw decits at points of normal fault linkage. Journal of Structural Geology 31
(10), 1145e1160.
Walsh, J.J., Watterson, J., 1988. Analysis of the relationship between displacements
and dimensions of faults. Journal of Structural Geology 10 (3), 239e247.
Walsh, J.J., Nicol, A., Childs, C., 2002. An alternative model for the growth of faults.
Journal of Structural Geology 24 (11), 1669e1675.
Weber, K.J., Mandl, G., Pilaar, W.F., Lehner, F., Precious, R.G. 1978. The role of faults in
hydrocarbon migration and trapping in Nigerian growth fault structures. In:
10th Annual Offshore Technology Conference Proceedings 4, pp. 2643e2653.
Whittaker, A.C., Attal, M., Cowie, P.A., Tucker, G.E., Roberts, G., 2008. Decoding
temporal and spatial patterns of fault uplift using transient river long proles.
Geomorphology 100 (3e4), 506e526.
Wibberley, C.A.J., 2005. Initiation of basement thrust detachments by fault-zone
reaction weakening. In: Bruhn, D., Burlini, L. (Eds.), High Strain Zones: Structure
and Physical Properties, vol. 245. Geological Society, London, Special Publica-
tions, pp. 347e372.
D.R. Faulkner et al. / Journal of Structural Geology 32 (2010) 1557e1575 1574
Wibberley, C.A.J., McCaig, A.M., 2000. Quantifying orthoclase and albite muscovi-
tisation sequences in fault zones. Chemical Geology 165 (3e4), 181e196.
Wibberley, C.A.J., Shimamoto, T., 2003. Internal structure and permeability of major
strike-slip fault zones: the Median Tectonic Line in Mie Prefecture, Southwest
Japan. Journal of Structural Geology 25 (1), 59e78.
Wibberley, C.A.J., Shimamoto, T., 2005. Earthquake slip weakening and asperities
explained by thermal pressurization. Nature 436 (7051), 689e692.
Wibberley, C.A.J., Yielding, G., Di Toro, G., 2008. Recent advances in the under-
standing of fault zone internal structure; a review. In: Wibberley, C.A.J.,
Kurz, W., Imber, J., Holdsworth, R.E., Collettini, C. (Eds.), Structure of Fault
Zones: Implications for Mechanical and Fluid-ow Properties. Geological
Society of London Special Publication, vol. 299, pp. 5e33.
Wilkerson, M.S., Fischer, M.P., Apotria, T., 2002. Fault-related folds: the transition
from 2-D to 3-D e preface. Journal of Structural Geology 24 (4), 591e592.
Wilkins, S.J., Gross, M.R., 2002. Normal fault growth in layered rocks at Split
Mountain, Utah: inuence of mechanical stratigraphy on dip linkage, fault
restriction and fault scaling. Journal of Structural Geology 24 (9), 1413e1429.
Wilkins, S.J., Gross, M.R., Wacker, M., Eyal, Y., Engelder, T., 2001. Faulted joints:
kinematics, displacement-length scaling relations and criteria for their identi-
cation. Journal of Structural Geology 23 (2e3), 315e327.
Wilson, J.E., Chester, J.S., Chester, F.M., 2003. Microfracture analysis of fault growth
and wear processes, Punchbowl Fault, San Andreas System, California. Journal
of Structural Geology 25 (11), 1855e1873.
Wintsch, R.P., Christoffersen, R., Kronenberg, A.K., 1995. Fluid-rock reaction weak-
ening of fault zones. Journal of Geophysical Research e Solid Earth 100 (B7),
13021e13032.
Wiprut, D., Zoback, M., 2002. Fault reactivation, leakage potential, and hydro-
carbon column heights in the northern North Sea. In: Koestler, A.G.,
Hunsdale, R. (Eds.), Hydrocarbon Seal Quantication. NPF Special Publication,
vol. 11, pp. 203e219.
Withjack, M.O., Schlische, R.W., Olsen, P.E., 2002. Rift-basin structure and its
inuence on sedimentary systems. In: Society for Sedimentary Geology Special
Publication, vol. 73 57e81.
Yamashita, T., Suzuki, T., 2009. Quasi-static fault slip on an interface between
poroelastic media with different hydraulic diffusivity: a generation mechanism
of afterslip. Journal of Geophysical Research e Solid Earth 114.
Yasuhara, H., Marone, C., Elsworth, D., 2005. Fault zone restrengthening and fric-
tional healing: the role of pressure solution. Journal of Geophysical Research e
Solid Earth 110 (B6).
Yielding, G., Freeman, B., Needham, D.T., 1997. Quantitative fault seal prediction.
AAPG Bulletin e American Association of Petroleum Geologists 81 (6),
897e917.
Zhang, S.Q., Tullis, T.E., 1998. The effect of fault slip on permeability and perme-
ability anisotropy in quartz gouge. Tectonophysics 295 (1e2), 41e52.
Zhang, S.Q., Tullis, T.E., Scruggs, V.J., 1999. Permeability anisotropy and pressure
dependency of permeability in experimentally sheared gouge materials. Journal
of Structural Geology 21 (7), 795e806.
Zheng, G., Rice, J.R., 1998. Conditions under which velocity-weakening friction
allows a self-healing versus a cracklike mode of rupture. Bulletin of the Seis-
mological Society of America 88 (6), 1466e1483.
Zoback, M., Hickman, S., Ellsworth, W., 2010. Scientic drilling into the San Andreas
fault zone. Eos. Transactions American Geophysical Union 91 (22), 197e199.
D.R. Faulkner et al. / Journal of Structural Geology 32 (2010) 1557e1575 1575
Internal structure, fault rocks, and inferences regarding deformation, uid ow,
and mineralization in the seismogenic Stillwater normal fault, Dixie
Valley, Nevada
Jonathan Saul Caine
a,
*
, Ronald L. Bruhn
b
, Craig B. Forster
b
a
U.S. Geological Survey, P.O. Box 25046, MS 964, Denver, CO 80225, USA
b
Department of Geology and Geophysics, University of Utah, 115 South 1460 East, Salt Lake City, UT 84112, USA
a r t i c l e i n f o
Article history:
Received 15 January 2009
Received in revised form
28 January 2010
Accepted 9 March 2010
Available online 17 March 2010
This work is dedicated to the memory of
Craig B. Forster who died in a tragic accident
on December 28, 2008. It is a reection of
his exceptional enthusiasm and dedication
to bringing students and colleagues
together in the pursuit of collaborative
scientic research.
Keywords:
Fault zone
Seismicity
Fluid ow
Hydrothermal
Breccia textures
Silicication
a b s t r a c t
Outcrop mapping and fault-rock characterization of the Stillwater normal fault zone in Dixie Valley,
Nevada are used to document and interpret ancient hydrothermal uid ow and its possible relationship
to seismic deformation. The fault zone is composed of distinct structural and hydrogeological compo-
nents. Previous work on the fault rocks is extended to the map scale where a distinctive fault core shows
a spectrum of different fault-related breccias. These include predominantly clast-supported breccias with
angular clasts that are cut by zones containing breccias with rounded clasts that are also clast supported.
These are further cut by breccias that are predominantly matrix supported with angular and rounded
clasts. The fault-core breccias are surrounded by a heterogeneously fractured damage zone. Breccias are
bounded between major, silicied slip surfaces, forming large pod-like structures, systematically
oriented with long axes parallel to slip. Matrix-supported breccias have multiply brecciated, angular and
rounded clasts revealing episodic deformation and uid ow. These breccias have a quartz-rich matrix
with microcrystalline anhedral, equant, and pervasively conformable mosaic texture. The breccia pods
are interpreted to have formed by decompression boiling and rapid precipitation of hydrothermal uids
whose ow was induced by coseismic, hybrid dilatant-shear deformation and hydraulic connection to
a geothermal reservoir. The addition of hydrothermal silica cement localized in the core at the map scale
causes fault-zone widening, local sealing, and mechanical heterogeneities that impact the evolution of
the fault zone throughout the seismic cycle.
Published by Elsevier Ltd.
1. Introduction
The presence and ow of uids in the upper crust has a major
impact on the mechanics of faulting (Hubbert and Rubey, 1959; Nur
and Booker, 1972; Sibson, 1977, 1981, 1990, 1996; Power and Tullis,
1989; Bruhn et al., 1990, 1994; Parry and Bruhn, 1990; Scholz, 2002;
Chester et al., 1993; Rice, 1992; Byerlee, 1993; Keller and Loaiciga,
1993; Evans and Chester, 1995; Caine et al., 1996; Miller et al.,
1996; Seront et al., 1998; Tanaka et al., 2001; Wibberley,
2002;Faulkner et al., 2006; Lockner et al., 2009). Fluid ow and
its interactions with heterogeneous permeability structures in
a fault zone can control the magnitude of local principal stresses
(Nemcok et al., 2002). This, in turn, affects local uid-pressure
gradients, mechanical failure, propagation of pressure transients,
uid inltration into and out of a fault zone via fault-valve mech-
anisms (e.g., Sibson, 1992), and fault-zone sealing and healing (e.g.,
Faulkner et al., 2008). Fluid ow in fault zones can control the
location, emplacement, and evolution of economic mineral
deposits and geothermal systems (e.g., Newhouse, 1942; Cox et al.,
2001; Sibson, 2001; Micklethwaite, 2009), and may also impact the
locations and magnitudes of foreshock, earthquake and aftershock
distributions (Miller et al., 2004). Yet fault zones are heterogeneous
geological and hydrological structures that commonly are not well
exposed. Even in well-exposed fault zones direct links between
internal structure, fault rocks, and mineral assemblages that
uniquely indicate a seismogenic origin are uncommon (cf. Sibson,
1986b; Cowan, 1999; Ujiie et al., 2007; Woodcock et al., 2007;
Smith et al., 2008). Thus, the study of exposed, seismogenic fault
zones that may record uid ow-related processes associated with
earthquakes remains important for understanding the mechanics
of faulting.
* Corresponding author.
E-mail address: jscaine@usgs.gov (J.S. Caine).
Contents lists available at ScienceDirect
Journal of Structural Geology
j ournal homepage: www. el sevi er. com/ l ocat e/ j sg
0191-8141/$ e see front matter Published by Elsevier Ltd.
doi:10.1016/j.jsg.2010.03.004
Journal of Structural Geology 32 (2010) 1576e1589
Fault zones are commonly composed of distinct, three-dimen-
sional, mappable components that include a fault core and damage
zone within relatively undeformed protolith (Chester and Logan,
1986; Smith et al., 1990; Forster et al., 1991; Caine et al., 1996).
Most of the strain is accommodated in a fault core indicated by
rocks such as fault-related breccias and clay-rich gouge. Fault zones
can also have multiple core zones interspersed with pods of
heterogeneously deformed host rock (cf. Faulkner et al., 2006). A
damage zone is the mappable network of subsidiary structures that
surrounds a fault core or fault-core zone and is related to the
nucleation, evolution, and growth of the fault zone (Chester and
Logan, 1986; Scholz, 2002; Caine et al., 1996; Knipe et al., 1998).
Damage-zone fracture networks commonly have orientations
mechanically related to the master fault and are of higher intensity
than found in the protolith (e.g., Caine and Forster, 1999). The fault
core and damage zone are surrounded by the protolith where fault-
related structures are generally absent.
The bulk permeability structure and strength of a fault zone are
controlled by preexisting and newly developed structures, the
regional and local stress state, fault-zone geometry, and changes
in lithology resulting from the coupling of mechanical, thermal,
uid ow, and reactive geochemical processes. For example, the
creation of new hydraulically contrasting lithologies and struc-
tures, such as clay-rich cataclasites and complex fracture
networks, has been documented to result from as well as impact
uid ow in diverse brittle fault-zone settings (Sibson, 1986a;
Chester and Logan, 1986; Scholz, 2002; Bruhn et al., 1994;
Antonellini and Aydin, 1994; Goddard and Evans, 1995; Caine
et al., 1996; Faulkner et al., 2006). These fault-related physical
attributes in the upper crust create hydraulic and mechanical
heterogeneity and anisotropy that have a signicant impact on
rupture and the arrest of failure (Parry et al., 1991; Byerlee, 1993;
Miller et al., 1996; Seront et al., 1998) as well as growth and
widening of a fault zone.
Previous theoretical research in earthquake mechanics has
focused on the role of uid circulation and hydrothermal alteration
associated with faulting processes (Sibson, 1981; Bruhn et al., 1994;
Parry et al., 1991; Rice, 1992; Byerlee, 1993; Scholz, 2002; Unsworth
et al., 1997). Although there have been studies of exhumed and
well-exposed seismogenic fault zones (e.g., Hancock and Barka,
1987; Ghisetti et al., 2001), details regarding the physical
pathways along which uid ow occurs, and the characteristics of
structures and rock types that result fromcoupled deformation and
uid ow remain sparsely documented.
This paper describes eld observations fromthe Mirrors locality
of the Stillwater Fault Zone (SFZ) in Dixie Valley, Nevada (Fig. 1).
This is an area of geological interest due to exposures of exhumed
portions of the footwall of this normal fault with a record of historic
earthquakes and surface ruptures associated with the fault. There
are also epithermal gold deposits, and a productive geothermal
reservoir hydraulically connected to the fault zone. Outcrop
mapping, hand-sample and thin-section fault-rock studies are used
to extend previous work and (1) document the internal structure
and geometry of part of the fault zone, (2) infer the paleo-perme-
ability structure, (3) document the textural attributes, composition,
and spatial and temporal distribution of fault rocks, and (4) infer
deformation-related uid owprocesses associated with seismicity
and growth of the fault zone.
2. Geologic setting and previous work
The SFZ is historically active and capable of generating magni-
tude (M) > 6 earthquakes. Ground surface rupture associated with
the 1954 M 6.8 earthquake was 30e40 km long (Caskey et al.,
1996). The SFZ, also called the Dixie Valley Fault, is the eastern
range-bounding fault between the Stillwater Mountains and Dixie
Valley graben (Fig. 1). Fault segments that range from several
kilometers to a few tens of kilometers in length form the SFZ. The
SFZ is one segment in a 300 km long belt of normal and normal-
oblique slip faults (Wallace and Whitney, 1984; Caskey et al., 1996).
The Stillwater Range is composed of Mesozoic metasedimentary,
plutonic, and volcanic rocks. Igneous rocks include the Jurassic
gabbroic Humboldt igneous complex, Cretaceous granites, a multi-
phase Oligocene graniteegranodioriteequartz monzonite complex,
and various extrusive rocks of Oligocene and Miocene age (Page,
1965; Speed and Armstrong, 1971; Wilden and Speed, 1974;
Speed, 1976).
The SFZ has been seismically active since Oligocene to early
Miocene times (Parry et al., 1991; Bruhn et al., 1994; Caskey et al.,
1996; Seront et al., 1998). The Holocene fault scarps that cut the
basin ll along the eastern base of the Stillwater Range are ground
surface ruptures probably formed during major earthquakes within
Surface trace of Stillwater
fault zone, bar and ball on
hanging wall.
Tertiary granitic rocks
Tertiary volcanic rocks
Jurassic granodiortic to
gabbroic rocks
Jr/Tr sedimentary rocks
Box
Canyons
Oxbow
Geothermal
Plant
The Mirrors
0 10 20 30 40
Km
N
Nevada
Stillwater
Range
40 N Latitude
D
i
x
i
e
V
a
l
l
e
y
118 W Longitude
Explanation
S
t
i
l
l
w
a
t
e
r
R
a
n
g
e
Fig. 1. Location map for the Stillwater Fault Zone in Dixie Valley, Nevada showing the primary study area of the Mirrors Locality. Modied from Parry and Bruhn (1990).
J.S. Caine et al. / Journal of Structural Geology 32 (2010) 1576e1589 1577
the last 12,000 years (Wallace and Whitney, 1984; Caskey et al.,
1996). The eastern slope of the Stillwater Range is the exhumed
footwall of the fault zone. An along-strike differential surface-uplift
history of the Stillwater Range is suggested by rocks exhumed from
a minimumdepth of 6 km(Parry et al., 1991) in the southern part of
the range and as little as 2 km at temperatures less than 270
C in
the central part of the range (Power and Tullis, 1989).
The rocks exposed along the SFZ and at the Mirrors locality in
the central part of the Stillwater Range have undergone varying
degrees of hydrothermal alteration that includes the following
mineral assemblages and paragenetic sequence: biotite K-feld-
spar, chlorite sphene epidote magnetite; mixed-layer
chloriteesmectite smectite goethite; quartz illite
calcite ferroan dolomite calcite barite; quartz kaolinite;
and quartz calcite (Parry et al., 1991; Lutz et al., 1997). The
assemblages in this paragenetic sequence were interpreted to
preserve a record of exhumation, with the earlier assemblages also
representing deeper crustal conditions. Power and Tullis (1989)
present a detailed analysis of slip surfaces at the Mirrors and
make estimates of deformation temperatures based on the ther-
modynamic stability of the quartz kaolinite mineral assemblage
in distinctive fault-related breccias and the tectonostratigraphic
position of this locality with respect to overlying volcanic rocks.
They observed crystallographic preferred orientation of quartz c-
axes in samples of the slickensided surfaces and their preferred
interpretation of this alignment involves pressure solution and
growth/dissolution rate anisotropy during continuous, low
temperature, low strain-rate deformation. Additionally, Power and
Tullis (1989) interpreted multiple generations of fault-related
brecciation and thus interpreted alternating periods of continuous
deformation during the interseismic phases of the earthquake cycle
followed by discontinuous deformation during coseismic phases of
the earthquake cycle.
Hydrothermal alteration in the Oligocene granitic complex
along the base of the Stillwater Range reects the interplay of
faulting, uid circulation, and exhumation associated with the SFZ
(Parry et al., 1991; Lutz et al., 1997). At some localities this alteration
may reect fault-related uid interactions with an ancient, and
presently producing, geothermal reservoir as well as a number of
surcial expressions of the geothermal system such as hot springs,
sinter mounds, and fumaroles (e.g., Lutz et al., 2002). Barton et al.
(1997) and Hickman et al. (1997, 1998, 2000) studied borehole log
and in situ stress data near the Oxbow geothermal plant, several
kilometers northeast of the Mirrors locality (Fig. 1). They found that
the hydraulically conductive fractures in the geothermal reservoir
were also critically stressed for shear failure in the regional stress
eld. Finally, structurally favorable sites and fault-related fracture
networks also create loci for hydrothermal uid ow that produced
economic epithermal gold deposits (e.g., Vikre, 1994).
3. The Mirrors map area exposure, overview, and methods
Each fault-zone component is well exposed in the bedrock at the
Mirrors locality (Figs. 2e4). This footwall remnant of the SFZ
extends approximately 250 m vertically upward from the base of
the Stillwater Range. Although the hanging wall is composed of
Quaternary basin ll in depositional contact with the crystalline
footwall and no recent ground surface rupturing fault scarps were
observed directly against it, we have mapped this as a fault contact
to portray the crystalline fault rocks in juxtaposition with the ll
(Figs. 2 and 3). It is likely that recent fault scarps exist in the ll but
mapping these was beyond the scope of our work. There are no
exposures of the crystalline hanging wall at the Mirrors. The
exposed fault core is complex but ranges in thickness from w1 to
5 m. However, partial erosion of the core makes it difcult to
estimate the in situ thickness or geometry of the entire fault core.
Although previously mapped as hornblende gabbros and anor-
thosites of the Jurassic Humboldt igneous complex (Speed, 1976),
the protolith at the Mirrors locality is a ne to medium-grained
granodiorite with chloritization of biotite and hornblende in the
protolith and variable chloritic to argillic to silicic alteration in the
damage zone.
Semiquantitative X-ray diffraction (XRD) patterns were
obtained from representative, whole-rock samples of each fault
zone component (Table 1). The samples were X-rayed with a quartz
standard and Cu-Ka radiation from 5
to 90
two-theta. Spectra
1
3
5
3
DZ
68
FC
DZ 68
DZ
PL
76
64
PL Protolith Geologic Contact
DZ Damage Zone
FC Fault Core Topographic Contour
Qal Quaternary Alluvium
Strike and Dip of Contact
64
70
N
~ Contour Interval 37 meters
0 ~ Scale (meters) 305
Qal
Qal
PL
PL
FC
72
69
60
FC
1
2
0
7
The Stillwater
Fault Zone
Dashed where gradational.
Bar and ball on hanging wall.
FI IG G. .
6AA
FI IG G. .
4
Fig. 2. Simplied geologic map of the Stillwater Fault Zone at the Mirrors Locality. Note that the protolith/damage-zone contact is gradational (dashed line) and the damage-zone/
fault-core contact is abrupt (solid line). The approximate locations of Figs. 4 and 6A are shown for reference.
J.S. Caine et al. / Journal of Structural Geology 32 (2010) 1576e1589 1578
were manually matched to library peak intensity data for individual
mineral determination and the estimated error for each reported
mineral is 10 weight percent.
The contacts between, and physical properties of, the fault core,
damage zone, and protolith were mapped using standard tape and
compass techniques using a 1:24,000 topographic base map blown
up to a scale of approximately 1:6,000 (Figs. 2 and 3). Detailed
fracture and vein data (e.g., orientation, intensity, and failure mode
interpretation, i.e., shear or extension fracture) were also collected
and used to help dene each architectural component. Fracture
data were collected using standard scanline and grid counting
methods (e.g., Priest, 1993) in three perpendicular directions
related to strike, dip, and slip direction of the fault zone (e.g., Seront
et al., 1998; Caine and Forster, 1999) at a variety of representative
outcrops. Data were integrated across a number of scales of
observation with centimeter-scale petrographic analyses, meter-
scale outcrop mapping, and at tens to hundreds of meters-scales
using low-elevation aerial photographs.
4. Fault zone structure and fault rocks
4.1. Fault zone orientation, component contact relationships,
and mineralogy
The SFZ strikes northeast to east-northeast and dips from 32
to
70
3
3
Damage Zone
n=594 1
2
4c
3
4a
4b
1
2
B
4c
N
Protolith
n=147
4b
4a
4c
4c
4b
A
N
N
Geothermal
Reservoir
Fractures
C
Fault Core
n
surfaces
=34
n
lineations
=21
SFZ
Fig. 5. Lower hemisphere, equal area projections showing orientations of structures in each fault-zone component at the Mirrors locality (A and B, Kamb contoured poles to
fractures, contour interval 2.0 sigma with the mean SFZ great circle shown in black, modied from Caine and Forster, 1999; C, great circles and slip lineations). A) Data from
protolith only. B) Data from damage zone only. In A and B, fractures broken into sets from raw data are assigned to a mode of formation based on interpretation of mechanical
compatibility with an Andersonian model of a normal fault. Relative to the master fault zone the following sets are dened: 1 extension; 2 shear; 3 step; 4a, 4b, and 4c cross
fractures. C) Gray great circles are slip surfaces and small gray dots are the corresponding slickenlines. The black great circle and large black dot show the mean slip surface and
slickenline orientations and the large black square is the pole to the mean SFZ slip surface. The gray star shows the approximate mean pole to hydraulically conductive and critically
stressed fractures derived from borehole logging in the Dixie Valley geothermal reservoir within several kilometers of the Mirrors locality (from Barton et al., 1997). D) Schematic,
eld-based cross-section of the traces of idealized fracture sets associated with an Andersonian normal fault (modied from Caine and Forster, 1999).
J.S. Caine et al. / Journal of Structural Geology 32 (2010) 1576e1589 1580
cut by fault-related shear fractures, the cross fractures may have
formed prior to faulting. Additionally, some cross fractures are lled
with a distinctive beige quartz kaolinite mineral assemblage that
emanates from the fault core, through the damage zone and into
the protolith suggesting they were possibly hydraulically
conductive during faulting. Extension, shear, and step fractures do
not generally occur in the protolith, suggesting that they are
mechanically related to faulting as their presence is largely
restricted to the damage zone (Fig. 5).
The damage zone consists of intensely fractured and veined
granodiorite (Figs. 2e7). The median fracture intensity inclusive of all
fracture types observed from 16 localities in the damage zone is 50
fractures per meter with a minimum of 32 and maximum of 99 per
meter. Althoughnosuitable outcrops were found where macroscopic
fracture intensity changes could be continuously measured from the
damage zone into the protolith, the change did not appear to be
systematic from one location to another where partial observation
could be made. Alteration and microfracture intensity also increase
towardthe fault core andsilicicationinthe damage zoneis generally
restricted to veins (Fig. 7; Table 1). Damage-zone veins at all sites are
lled with the same quartz kaolinite mineral assemblage found in
the fault core and many commonly have angular fragments of wall
rock suspended in the matrix (Fig. 7). The veins range in width,
perpendicular to their walls, from a few millimeters up to about
10 cmand have trace lengths of just under a meter up to nearly 10 m.
They occur at intensities of approximately 1 vein per meter to
stockworks with many tens of veins per meter. Veins are planar to
curviplanar but when they occur as stockworks they can have quite
irregular and jagged shapes.
4.3. Fault rocks
A representative sample from the fault core is described from
hand-sample to thin-section scales (Figs. 8 and 9). This sample
contains a major silicied slip surface similar to those found
bounding some of the breccia pods as well as the range of different
types of fault-related breccias with important cross-cutting rela-
tionships representative of the Mirrors locality. Sections cut parallel
and perpendicular to slip were cut to document variations in
textures, microstructures, and fault-rock compositions.
The fault core at the Mirrors locality comprises several struc-
turally and texturally distinct fault-rock types described using
elements of classications by Sibson (1977, 1986a), Sillitoe (1985),
Jbrak (1997), and Mort and Woodcock (2008) including: breccias
with rounded clasts; clast-supported breccias with highly angular
and interlocking clasts; matrix-supported breccias with angular
clasts that could be pieced back together if the matrix was
removed; breccias with open spaces between the clasts; and fault
gouge described in detail below.
4.3.1. Breccias with rounded clasts
Breccias with rounded clasts are predominantly matrix sup-
ported but are also locally clast supported. They are poorly sorted
with respect to size, composed of well-rounded clasts and sub-
angular clasts, and are typically associated with cross-cutting slip
surfaces. These textures are observed in the fault rocks at both the
hand-sample and thin-section scales (Figs. 8 and 9). The uppermost
millimeter of the hand sample shown in Fig. 8 is a polished and
striated, silicied slip surface representative of those described
above. In the rst centimeter of the sample, from the slip surface
toward the base of the polished face (Figs. 8 and 9), the sample
exhibits an ultra-ne-grained, highly indurated, whiteepink
colored breccia matrix. Most of the clasts in this matrix show
distinct clasts within clasts textures. This part of the sample is
nearly identical to samples studied by Power and Tullis (1989).
In thin section, quartz is the primary mineral with minor
amounts of kaolinite within about a centimeter of the slip surface
(Fig. 9A and D). In Fig. 9D the section is oriented perpendicular to
the strike of the fault and parallel to the slip direction. The silicied
hanging wall material shows crystallographic preferred orientation
Fig. 6. A) Exposure of the main Mirrors slip surfaces (SS) and large breccias pods (BP)
outlined in dashed yellow. Approximate contacts between Quaternary alluvium (Qal),
fault core (FC), damage zone (DZ), and protolith (PL) are shown with black lines. The
large corrugated pod is composed of all breccia types, bounded by polished and stri-
ated curviplanar slip surfaces, and is elongated parallel to the average slip vector
indicated by the red arrows. B) Outcrop photograph of a small (compass for scale)
breccia pod with matrix-supported breccias. The long axis of the pod is subparallel to
the average orientation of the master fault zone shown by the red arrow. A possible
feeder vein (FV) connected to a network of very long and wide veins can be seen at the
bottom central portion of the pod. C) Breccia pod veins with jigsaw puzzle breccia.
Note the en chelon clasts and closing taper of the veins away from a larger breccia
pod to the right of the photograph. Each photograph is a northward to northwest-
erward-looking view along the strike of the Stillwater Fault Zone.
J.S. Caine et al. / Journal of Structural Geology 32 (2010) 1576e1589 1581
(CPO) of quartz, indicated by linear domains of common extinction
in synthetic and antithetic microfaults as described in detail by
Power and Tullis (1989) and is typical of breccia pod bounding slip
surfaces. Scanning electron microscopy shows quartz occurs as
microcrystalline prisms on the order of 10 mm long with approxi-
mate alignment with the synthetic microfaults (Fig. 9E). Quartz
with CPO also occurs as micron-scale coatings surrounding many of
the clasts, with common extinction at average angles of about
60e120
from the slip surface (Fig. 9A). There are also numerous,
discrete rounded clasts and areas of matrix with internal domains
of CPO of quartz that do not make regular angles with the slip
surface consistent with observations by Power and Tullis (1989).
Many of the silicied clasts have varying degrees of optical contrast
with the matrix suggestive of varying degrees of silicication and
Fig. 8. A) Photograph of a polished hand-sample from the fault core. The top of the
sample is a slip surface and the polished face is parallel to the slip direction. The three
major breccia types are exhibited. Note that the breccia with rounded clasts grades into
the clast-supported breccia. In contrast, the contact between the clast-supported
breccia and the matrix-supported breccia is relatively abrupt indicating the cross-
cutting of the former by the latter. B) Example of breccia with vuggy open space. Each
iron oxide stained clast is coated with a rind of carbonate with euhedral crystals. Note
the 10 hand lens for scale.
Fig. 7. Photomicrographs of representative rock samples from the protolith and
damage zone. A) Polarized light image of protolith granodiorite. Equigranular
quartz plagioclase potassium feldspar biotite, which is altered to chlorite. Note
the absence of microfracturing. B) Polarized light image of damage-zone micro-
fractures in granodiorite w10 m from the fault core. C) Polarized light image of
damage-zone granodiorite w5 m from the fault core shows extensive microfracturing,
horsetail fractures, and silicied veins with cataclastic particles.
J.S. Caine et al. / Journal of Structural Geology 32 (2010) 1576e1589 1582
clay mineral content that result in translucent gray to brown,
layered to clotted, felt-like masses. Little to no carbonate is
observed in these breccias which grades into clast-supported
breccias as the polished face of the representative hand sample is
traversed further down, where the carbonate content increases
(Fig. 8). In other samples and exposures the breccias with rounded
clasts abruptly cut clast-supported breccias or are themselves cut
by discrete slip surfaces (e.g., Fig. 9D).
4.3.2. Clast-supported breccias
In outcrop and hand samples the most common fault-related
breccias are the predominantly clast-supported breccias with local
matrix-supported textures. These breccias have an average matrix-
to-clast ratio of w30:70. The clasts are poorly sorted, angular to
subangular, show interclast veining, and are predominantly
composed of silicied granodiorite protolith (Figs. 8 and 9B). These
breccias show little evidence for cataclastic transport of rock, such
as relative movement or rounding of clasts. This rock type is the
background fault rock that is cut by slip surfaces, breccias with
rounded clasts, matrix-supported breccias, and breccias with open
space. In some samples, clasts of brecciated protolith in the clast-
supported breccias are recognized in thin section and appear
brecciated multiple times. These types of textures are the exception
and are found where the breccias are locally more matrix sup-
ported. Some clasts are locally incorporated into a ne-grained,
quartz-rich breccia but none were observed with domains of quartz
CPO. The clasts and silicic matrix are both cut by carbonate veins,
which are also cut by ne-grained quartz kaolinite veins. The
latter assemblage is cut again by a set of young calcite veins. The
veins commonly include clasts of the protolith, preexisting vein
materials, breccias, and other fault-related rocks that show
multiple brecciation events.
4.3.3. Matrix-supported breccias
Matrix-supported breccias are found in distinctive pod-like
bodies and in veins that emanate from and (or) connect to other
breccia bodies. These breccias are also characterized by angular
breccia clasts of local wall rock that could be pieced back together
if the matrix that separates them was removed. Yet, locally they
have subrounded clasts of the other breccia types discussed
Fig. 9. Polarized light and scanning electron microscope (SEM) micrographs of the three breccias found in Fig. 8. A) Silicied breccia with rounded clasts that formed directly under
a polished and striated slip surface shown by a white arrow. Note the rounded and angular clasts, completely silicied relict clasts (RRC), partially clast-supported texture, and
crystallographic preferred orientation (CPO) of quartz. B) Moderately silicied predominantly clast-supported breccia showing subangular clasts and poor sorting. C) Highly silicied
matrix-supported breccia characterized by angular and rounded clasts sitting in an undeformed microcrystalline, quartz-rich matrix. D) Oriented section of the sharp juxtaposition
of two different breccias with rounded clasts (BRC) separated by a polished and striated silicied slip surface (the section separated along the thick dark irregular line with the white
arrow in it). E) SEM, secondary electron image of the predominantly quartz (Q) matrix with minor kaolinite (K) from the hanging wall of the sample shown in D (mineralogy
determined by energy dispersive spectroscopy). The image is roughly oriented as image D. F) Cathodoluminescence image of a clast within the microcrystalline quartz matrix (black
background). The triangular clast is interpreted as deformed quartz and the background quartz is interpreted as undeformed quartz. Note that the clast is laced with several
generations of healed microfractures (black and gray stringers of quartz).
J.S. Caine et al. / Journal of Structural Geology 32 (2010) 1576e1589 1583
above, some with exotic compositions from the wall rock. In
either case the clasts are generally widely separated from one
another (Figs. 6, 8 and 9C). Breccias with angular clasts that can be
pieced back together have been previously referred as jigsaw
puzzle breccias and implosion or explosion breccias (cf. Sibson,
1986a; Jbrak, 1997; Power and Tullis, 1989; Mort and
Woodcock, 2008). At the Mirrors, these breccias are matrix sup-
ported and the matrix has a distinctive beige color. They are
poorly sorted, highly silicied, lack veins or fractures, and exhibit
no open, macroscopic pore space. The breccia pods in which the
matrix-supported breccias occur abruptly and irregularly cut the
other breccia types (Fig. 8).
The clasts in the matrix-supported breccias range from 1 mm to
3 cm in size, contain multiple clasts of varying composition with
cross cutting internal brecciation but are compositionally distinct
from the beige matrix (Fig. 8). Clast size distribution is variable but
most clasts are orders of magnitude larger than the matrix (Figs. 8
and 9C). Clasts generally contain deformed angular quartz grains
visible in thin section. Cathodoluminescence analysis of the matrix-
supported breccia samples shows extensive fracturing and multiple
vein-lling events isolated within the clasts with no such defor-
mation or veining in the matrix (Fig. 9F). Clast lithologies include
adjacent wall rock breccias in the jigsawpuzzle breccias. Other clast
lithologies include protolith granodiorite, porphyritic volcanic
rocks, iron-stained and zoned dolomite vein crystals, relicts of
highly silicied breccia clasts in contact with their matrix material,
polished and striated fault surface materials, and possible meta-
sedimentary rock fragments. Calcite occurs as intraclast vein llings
in association with ferroan dolomite and as a late-stage precipitate
that lls some isolated void spaces, small stringer-like veins, and as
dismembered vein fragments.
The clasts are encased in a distinctive ultra-ne-grained
(w0.1e10 mm), beige colored, microcrystalline quartz-rich matrix
with minor amounts of intergrown kaolinite, trace amounts of
carbonate and amorphous material (Table 1). The matrix shows
rather uniformly sized anhedral to subhedral interlocking quartz
crystals with salt and pepper or mosaic-like microtexture (Fig. 9C).
Sub-equant, poorly developed quartz prisms with anhedral quartz
intergrowths are occasionally observed with somewhat reticulate
texture. All quartz grains are completely intergrown and conform
to every surface they are in contact with, showing no evidence for
preferred crystallographic orientation, spherulitic, or feathering
textures. Quartz crystals have uniform to domainal extinction
within individual grains and within masses of grains. The matrix
shows no evidence of brecciation, layering, overgrowths, cross-
cutting veins, or crack-seal textures. Carbonate appears to ll minor
void space as late-stage precipitate.
Maximum syndeformational porosity in the breccia pods was
estimated from representative hand samples of the matrix-sup-
ported breccias. Syndeformation porosity was estimated by digital
imaging of serial sections from several slabbed and polished
samples ranging in size from about 20 by 10 cm to about 3 by 5 cm
(e.g., the lower part of Fig. 8). The samples were scanned and a gray-
scale histogram was captured for each image where the dark
histogram signal was assumed to represent the clasts and the light
histogram signal assumed to represent the matrix. The measure-
ments also assume that the matrix of these breccias represent the
pore space formed during a single deformation event, consistent
with it cross-cutting other main breccias types. This is also
reasonable because the matrix itself is compositionally homoge-
nous, unlike the clasts, and shows no evidence for subsequent
deformation. The average, maximum syndeformation porosity
from this analysis is 54 percent and the analysis conrms the
largely matrix-supported texture using this pseudo three-dimen-
sional approach.
4.3.4. Breccias with open space
At several locations in exposures of the fault core there are
breccias that exhibit signicant open space and vugs (Fig. 8B).
These breccias include generally angular, iron oxide stained clasts
of other breccia types and are clast-supported fault rocks. The
volume of the interclast open space is on the order of microns to
several tens of cubic centimeters. Many of the spaces are coated
with banded and multilayered carbonate rinds that consist of
euhedral calcite crystals that conformably overlay previous layers
and the clasts. There also are euhedral carbonate veins that cut
through the clasts and in some cases connect to the euhedral rinds.
4.3.5. Fault gouge
The nal fault-rock category found at the Mirrors is unconsoli-
dated fault gouge (cf. Sibson, 1977). Only one thin seam, less than
10 cm in thickness, of this material was found. The gouge is ne to
medium grained and contains silicied clasts of the breccia types
described above. The gouge is not clay-rich as is commonly found in
some shallowly formed, brittle faults (cf. Morrow et al., 1984;
Foxford et al., 1998; Heynekamp et al., 1999; Vrolijk and van der
Pluijm, 1999; Aydin and Eyal, 2002). Although the exposure of
gouge cuts the clast-supported breccias its temporal relationships
to the other breccias types were not observed.
4.4. Internal structure of the fault core
A characteristic feature of the fault core is the occurrence of
matrix-supported breccias and other fault rocks that form pod-like
bodies (Figs. 3 and 6). The pods are generally lenticular with their
long axes subparallel to the average orientation of the slip vector for
the master fault zone. The short axes of the breccia pods are typi-
cally orthogonal to slip in the plane of the fault. Large slip surfaces
have elongated scoop-like depressions and associated ridges
parallel to the slip direction. The walls of the breccia pods that are
not bounded by slip surfaces are irregular and where exposed, they
are enveloped in fractured protolith. One pod-shaped structure
shows an along-strike dimension of several meters, a down-dip
dimension of w35 m, and a thickness of less than w5 m(Fig. 6). The
bulk of the pod is predominantly composed of silicied clast-sup-
ported breccias with interior meter-scale pods of the beige matrix-
supported breccias. This pod was one of the largest in outcrop; pods
on the order of w2 m or less in length are more typical.
The dimensions of exposed curviplanar or corrugated slip
surfaces that bound breccia bodies were characterized using
wavelength and amplitude measurements made with a tape
measure in the plane of the fault orthogonal to the slip direction.
The dimensions range from a maximum of approximately 35 m in
length to 3 m in amplitude at the map scale and 4 m in length to
0.5 m in amplitude at the outcrop scale. There is insufcient
exposure to make full measurements of anisotropy parallel to slip
at the map scale; however, at the outcrop scale on average the pods
vary in length from roughly 2 m to 15 m, with length to amplitude
ratios between 7:1 and 20:1. There is a difference in the wavelength
to amplitude ratio of breccia body corrugations in a direction
perpendicular to slip versus parallel to slip. At the outcrop scale,
corrugations parallel to slip may be up to three times as long as
those perpendicular to slip. In spite of the crude nature of these
estimates, they are consistent with other more sophisticated
measurements of fault surface roughness for the Mirrors and other
localities over a number of scales (cf. Lee and Bruhn, 1996; Sagy
et al., 2007).
En chelon, highly angular beige matrix-supported breccia
clasts in veinlets that extend from the breccia pods, indicate wall
rock fracture, shear, and subsequent uid ow within the breccia
pod interiors prior to mineral precipitation and sealing (Fig. 6).
J.S. Caine et al. / Journal of Structural Geology 32 (2010) 1576e1589 1584
Breccia-lled veins at the margins of the pods taper away from the
pods and terminate. Some of these veins, similar to that seen in
Fig. 6, are connected to an extensive network of veins that are
irregularly shaped, show variable thickness, and extend into the
damage zone.
5. Discussion
The SFZ at the Mirrors locality comprises a complex set of
structures and rock types that include multiple slip surfaces,
distinct fault rocks and breccia bodies, fault and non-fault-related
fracture networks. We infer that several processes acted together
during the formation of the fault zone and that some of the rocks
may record coseismic deformation and associated uid ow. In the
following discussion we present key observations and a conceptual
model that links deformation and uid ow to architecture,
permeability structure, and fault-zone growth.
5.1. Origin of the breccias and breccia pods
The key observations and interpretations that bear on the origin
of the different breccia types and breccia pods as recording
coseismic processes, and the growth of the SFZ include the
following. The breccia pods are discontinuous, elongate, and
tapered bodies aligned subparallel to the slip direction of the
master fault zone. These pods are primarily composed of clast-
supported breccias with local matrix-supported breccias. Silicied
slip surfaces and associated breccias with rounded clasts cut the
clast-supported breccias but are also locally brecciated. The matrix-
supported breccias are emplaced into the surrounding clast-sup-
ported breccias, slip surfaces, and breccias with rounded clasts
within the pods. Pod morphology, spatial restriction to the fault
core, kinematic compatibility with normal faulting, and large esti-
mated syndeformational porosity of the matrix-supported breccias
suggest that there was a component of dilatational strain during
deformation and pod formation. Slip surfaces and some matrix-
supported breccias with en chelon clasts also record a component
of shear deformation.
The topography of many of the large slip surfaces that bound the
pods includes elongated scoop-like depressions that may have
formed along local extensional fault jogs. These relatively small-
scale depressions may also record local dilatational strain and
permeability enhancement during coseismic deformation (e.g.,
Power and Tullis, 1989). Sibson (1986a) inferred that brecciation
could be caused by pore uid-pressure differentials between the
wall rock and an incipient, dilatant opening within a fault zone. If
the effective stress acting on a fault zone exceeds the tensile
strength of the wall rock, catastrophic failure can cause implosion
of the wall rock into openings. Many of the breccia pods, however,
do not occur near slip surfaces and are larger than openings due to
different topography along adjacent slip surfaces. This, and the
fault-slip parallel morphology of the larger breccia pods, suggests
that some pods may have originated by processes other than
dilatational strain in jogs caused by variations in slip-surface
topography. These larger breccia pods may have formed along
mechanically weak sites within the fault core such as preexisting
polished and striated slip surfaces, macroscopic fractures that were
partially mineral-lled, microfractured regions, or zones of
compositional variation.
The clast-within-clast textures found in the breccias provide
a record of multiple faulting or slip events localized in the core. The
matrix-supported jigsaw puzzle breccias have clasts that are
composed of the adjacent wall rock and these are generally angular
suggesting little transport. However, locally in the breccia pods,
where there are widely separated clasts in the same matrix as the
jigsaw puzzle breccias, the clasts are composed of exotic parent
lithologies and are subangular to subrounded indicating signicant
transport. The rounded clasts may also indicate mechanical attri-
tion fromgrain-size reduction given their proximity to slip surfaces.
Clast rounding may also have occurred during hydrothermal uid
owand associated abrasional wear of clasts against one another in
saturated, silica-rich rock our (cf. Sillitoe, 1985; Jbrak, 1997).
The matrix of breccia pods with matrix-supported clasts is
composed of distinctive beige colored, homogenous, ne-grained,
quartz with minor kaolinite and carbonate. The quartz has a mosaic
texture with uniformly sized (w0.1e10 mm) anhedral to subhedral
interlocking crystals that form domains of variable optical extinc-
tion and that completely inll and conform to all available space.
Lovering (1972), Fournier (1985), Saunders (1994), and Dong et al.
(1995) provide evidence that this mosaic texture forms by recrys-
tallization of amorphous silica that precipitated from the boiling of
hydrothermal uids and associated processes at temperatures
<180
C. At the temperatures of the modern geothermal reservoir,
a rapid pressure drop could result in such decompressional (iso-
enthalpic) boiling, phase separation, cooling, and rapid precipita-
tion of amorphous silica (Hedenquist and Henley, 1985; Saunders,
1994; Rimstidt, 1997). Such pressure drops could have occurred
in seismically induced dilatant openings localized in the fault zone
(e.g., Sibson, 1986a) and that could have also hydraulically con-
nected the reservoir to the surface. Decompressional boiling is
consistent with the minor kaolinite and carbonate minerals found
in small interstices of the mosaic quartz matrix. These phases are
also found in fault rocks elsewhere along the SFZ (Parry et al., 1991).
Loss of CO
2
to the vapor phase during boiling, with attendant
increase in pH, will cause clay and carbonate to precipitate (e.g.,
Fournier, 1985). Such decompression coupled with rapid amor-
phous silica deposition required rapid ow of uid and vapor up
through, and laterally within, the fault zone. Surcial expression of
this type of uid movement is evident as sinter mounds and
terraces, hot springs, and fumaroles such as those found at the
ground surface in association with the fault zone (cf. Caine and
Forster, 1999; Lutz et al., 2002).
Breccia-lled veins, with the same quartzekaolinite matrix as
the breccia pods emanate from and taper outward from the pod
margins and terminate. This may indicate that uid ow and the
pressure gradient were directed outward fromthe pod interior. The
pods then likely became permeability heterogeneities within the
fault core leading to a heterogeneous spatial distribution of
cemented or sealed portions of the fault zone. This, in turn, may
have lead to heterogeneities in fault-zone strength as the fault zone
evolved.
Observations of the juxtaposition of polished slip surfaces
against breccias with rounded clasts, cutting across clast-supported
breccias with angular clasts, then cut by matrix-supported breccias
with jigsaw puzzle and rounded clasts, open space breccias, the
irregular character of some of the larger breccia-lled veins, and
a seam of clay-rich gouge indicate that different mechanisms of
failure may have been operative in the fault zone over time. Some of
these mechanisms may be attributed to a non-static stress eld
changing dynamically during deformation or attributed to varia-
tions in uid pressures within the fault zone at different times (cf.
Caskey and Wesnousky, 1997; Micklethwaite, 2009). Multiple,
anastomosing, layered slip surfaces with curviplanar geometry and
different lineation directions indicate some portion of slip during
any deformation event was distributed along any number of pre-
existing surfaces from previous events. However, no evidence was
found to evaluate the nature of the actual slip distribution on an
event-by-event basis.
If breccia pod formation was seismogenic, it is unclear if they are
related to hypocentral earthquake nucleation sites or possibly
J.S. Caine et al. / Journal of Structural Geology 32 (2010) 1576e1589 1585
record rupturing in more distal and shallower parts of the fault
zone. The breccia pods and associated rocks may be unique
remnants of the deformation products linked to the distribution of
earthquake foreshocks, aftershocks, and microseismicity as
rupturing initiated and ceased, and pore pressure adjusted to
structural changes throughout the fault zone (cf. Nur and Booker,
1972; Byerlee, 1993; Miller et al., 1996, 2004). Episodic deforma-
tion and resulting fault rocks evolved into a hydraulically hetero-
geneous fault zone with local permeability contrasts great enough
to possibly have trapped uids under varying pressures. Byerlee
(1993) proposed a uid-pressure compartment model for fault
zones associated with earthquakes where uid sealed under high
pressure in hydraulically isolated compartments partially drains
into adjacent compartments of initially lower uid pressure during
a deformation event. Although not a direct comparison, Parry and
Bruhn (1990) and Parry et al. (1991) showed evidence for tran-
sient variations in uid pressure from uid inclusion data from the
SFZ about 40 km south of the Mirrors locality. The breccia pods are
not likely remnants of Byerlee-like pressure compartments per se,
however, the locations, sizes, shapes, and breaching of the breccia
pods were likely inuenced by variations in permeability, pore uid
pressure, and mechanical heterogeneities; features that may also
help to explain the variable uid pressures observed by Parry and
Bruhn (1990).
5.2. Heterogeneity and anisotropy within the fault zone
Several types of structural heterogeneities and anisotropies may
act together to impact uid ow, mechanics, and growth of the SFZ.
Fault-related permeability anisotropy in crystalline rock is
controlled by fracture networks and heterogeneously distributed
fault-rock types, such as the breccia bodies, with varying perme-
abilities (e.g., Seront et al., 1998; Caine and Forster, 1999). The
corrugated topography of the major slip surfaces that bound the
fault core and minor slip surfaces within the fault core may have
had a role in controlling the shapes and limited the growth of
breccia bodies. The elongation of these bodies parallel to the slip
direction should also contribute to permeability heterogeneity.
Anisotropic slip-surface topography may also be an important
cause of changes in local stress elds (Ferrill et al., 1999). The
distribution of slip surface and slip vector orientation data suggests
that most of the oblique slip is related to the accommodation of
deformation along the corrugated slip surfaces consistent with
attendant localized changes in the local stress eld (Fig. 5).
Outcrop-scale fracture network heterogeneities, such variations
in total fracture intensity between the damage zone and protolith,
the apparent unsystematic manner in which fracture intensity
changes with distance fromthe fault core, or internal heterogeneity
of the number of shear fractures and minor slip surfaces in any
given location, suggest that deformation and slip for any given
deformation event were heterogeneously distributed. Thus, fault
properties such as strain, strength, and permeability depend on the
range of scales over which fault-rock and damage-related hetero-
geneities exist. For example, the elongate nature of the bodies of
fault-core breccias, the contrast of fracture network intensity
between the fault core and damage zone, and the variable degree of
silica cementation in the fracture networks are evidence of signif-
icant hydrogeologic and mechanical heterogeneities and anisot-
ropies, implying there is signicant heterogeneity of fault-zone
strength and permeability from at least the outcrop scale to the
map scale.
Seront et al. (1998) measured permeability values on repre-
sentative core-plug samples that ranged from as high as 10
8
m
2
in
the damage zone to as low as 10
20
m
2
in the fault-core matrix.
Given the map-scale to micro-scale change in fracture intensity is
higher in the damage zone than in the fault core and that the core is
the locus of silicication, a heterogeneous combined conduit-
barrier permeability structure is suggested for the fault zone during
interseismic periods (cf. Caine and Forster, 1999). Although specu-
lative, this permeability heterogeneity could also lead to pore-
pressure heterogeneity and consequently strength heterogeneity
throughout the seismic cycle.
Seront et al. (1998) also found no appreciable difference in
mechanical strength of small laboratory samples of each fault-zone
component at the Mirrors. Their experimental results highlight the
importance of scale when considering the relative mechanical
strength within a fault zone composed of discrete components. For
example, uncemented, smooth-walled and discontinuous shear
fractures observed along the fault-core/damage-zone contact, and
within the damage zone, would provide favorably oriented failure
surfaces for later deformation and uid ow events (e.g., Barton
et al., 1997). The mechanical strength of these macro-scale frac-
tures would depend on frictional properties of the fractures rather
than grain-scale cohesion or microfracture properties that control
strength at the core-plug scale.
5.3. Model for deformation induced uid ow
We propose the following conceptual model that relates defor-
mation, growth, fault-related uid ow, and stress cycling to
architecture and permeability structure (Fig. 10). The proposed
model is built on the work of Power and Tullis (1989), Parry and
Bruhn (1990), Parry et al. (1991), Muir-Wood (1994), Seront et al.
(1998), and Sibson (2001) and assumes a constant maximum and
decreasing minimum principal stress in an Andersonian normal
fault regime.
During interseismic periods, uids were stored in an ancient
equivalent of the modern geothermal reservoir, in primary pore
space, and secondary macroscopic fracture networks in and
surrounding the fault zone (Fig. 10). The hydrothermal silica
cements in the fault core were likely precipitated from uids that
initially equilibrated with rocks in this reservoir. Aseismic defor-
mation along slip surfaces may have occurred as hypothesized by
Power and Tullis (1989). As differential stress increased, preexist-
ing fractures primarily in the damage zone, could have opened due
to decreasing minimum compressive stress. New fractures may
also have formed which mobilized uids and possibly dissolved
silica from comminuted by-products formed in preexisting frac-
tures in the surrounding host rock.
During coseismic brittle failure events new hydraulic connec-
tions to the geothermal reservoir were formed as ruptures propa-
gated likely releasing seismic energy. Quartz-supersaturated uid
and vapor inltration, ow, boiling and (or) phase separation
occurred in these new, highly permeable pathways (cf. Miller et al.,
2004). This resulted in rapid quartz matrix precipitation and
suspension of clasts forming the matrix-supported breccias in
hybrid dilatant and shear related openings localized in the fault
core. The effective minimum compressive stress also increased,
closing optimally oriented damage-zone fractures further mobi-
lizing uids into the fault core with accompanying shear along both
preexisting and newly formed discrete fractures and fracture
networks (cf. Muir-Wood and King, 1993; Seront et al., 1998;
Fig. 10). Comminution and the development of breccias with
rounded clasts were localized along constrictions between discrete
slip surfaces. Adjacent to slip surfaces and in the wall rock, in situ
distributed failure, possibly crushing, resulted in predominantly
clast-supported breccias. Some of the clast-supported breccias with
localized areas of matrix-supported textures may be older equiva-
lents of breccias similar to the beige matrix-supported breccias.
This process appears to have been repeated numerous times and
J.S. Caine et al. / Journal of Structural Geology 32 (2010) 1576e1589 1586
may have resulted in an increase in width of the fault zone as well
as heterogeneous sealing of the fault core. Dilatancy hardening may
also have resulted in associationwith uid-pressure drops (cf. Parry
et al., 1991).
Alternatively, the formation of the matrix-supported breccias
in discrete pods and some of the clast-supported breccias may
have formed by the fault-valve mechanism (Sibson, 1990) where
failure was driven by pore uid pressure rather than tectonic
stress. The present geothermal reservoir near the Mirrors locality
appears to be in a heterogeneous pressure regime largely at or
near hydrostatic pressure but with areas of enhanced pressure
(Hickman et al., 2000). Whereas, deep wells drilled outside of the
productive geothermal reservoir have artesian pressures (Dick
Benoit, personal communication, 2009). It may have been possible
that there were volumes of the ancient reservoir with localized
areas of overpressure connected to the fault zone or zones within
the fault with anomalous uid pressures possibly due to perme-
ability heterogeneity as suggested above. In either case, the
various breccia textures do not provide unique evidence for
a tectonic stress versus pore uid-pressure mechanism for their
origin.
Exhumation of the Mirrors portion of the fault zone brought the
fault zone into progressively different temperature, pressure, and
uid chemical conditions. The breccias with open space may have
originated from low pH ascending uids and yet clasts coated with
euhedral carbonate rinds are consistent with downwelling, silica
undersaturated uids both possible processes during interseismic
periods. Cross-cutting seams of fault gouge like represent yet
another phase of deformation largely lacking signicant localized
uid ow and associated hydraulic connection.
Fault
Core
Damage
Zone
Protolith
Damage Zone
Protolith
Interseismic Period
3
Decreases
Fluids Accumulate
Coseismic Period
3
Increases
Hybrid Failure
Fluids Mobilize
Fault Core
Damage Zone
Protolith
Damage Zone
Protolith
1
Constant
3
3
1
Induced
Hydraulic
Connections
to Geothermal
Reservoir
Matrix
Supported
& Jigsaw Clast
Supported
Rounded
Clasts
Major Breccia Types
Fig. 10. Schematic diagram and conceptual model of deformation, breccia formation, hybrid shear and dilatant fault growth, and uid ow inferred from the Mirrors locality
(hydrologic effects adapted from Muir-Wood, 1994 and specic structures from this work). The regional maximum principal stress (s
1
) is assumed to be constant and vertical
throughout the evolution of the fault zone. The minimum principal stress (s
3
) is assumed to be horizontal. At an early time (top diagram) in its evolution, the fault core is
predominantly composed of clast-supported breccias cut by breccias with rounded clasts (middle gray ll) and associated anastomosing slip surfaces (black lines that bound and
weave through the fault core). In the interseismic period when s
3
decreases as differential stress increases, optimally oriented fracture networks dilate (stubby arrows indicate local
stress), take on and store uids (curved arrows indicate uid transport direction). During a coseismic period, mixed-mode deformation (shear and dilatancy) is localized in the fault
core creating new hydraulic connections to an ancient geothermal reservoir accompanied by uid ow into breccia pods (light gray ll), development of breccias with rounded
clasts, and rapid precipitation of quartz due to decompressional boiling, locally sealing the fault zone. Because effective s
3
is increased, uids may also ow into the high
permeability fault core from optimally oriented damage-zone fracture networks.
J.S. Caine et al. / Journal of Structural Geology 32 (2010) 1576e1589 1587
6. Conclusions
Field mapping along the Stillwater Fault Zone provides
a detailed view of the internal structure of the footwall of a seis-
mogenic normal fault zone. Distinct mechanical and hydraulic
components include a fault core, damage zone, and protolith (cf.
Seront et al., 1998). The different components are distinguished by
distinctive variations in lithology, amount and type of hydro-
thermal alteration, fracture intensity, matrix-scale hydraulic prop-
erties, and structure. The fault core shows evidence for localized
deformation along silicied slip surfaces near large dilatant breccia
pods indicative of distributed deformation (cf. Power and Tullis,
1989). The core is surrounded by damage-zone fracture networks
also indicative of distributed deformation. The juxtaposition of
these distinct deformational styles leads us to characterize the
architecture of the fault zone at the Mirrors locality as a composite
deformation zone (Caine and Forster, 1999) where different fault-
related breccias and other fault rocks are indicative of multiple
deformation events possibly under different stress and (or) uid-
pressure conditions as the fault zone evolved.
Laboratory-derived permeability values indicate up to approxi-
mately four orders of magnitude difference in permeability
between the fault rock, damage zone, and protolith samples at the
core-plug scale. This and fracture network modeling (Caine and
Forster, 1999) in combination with the mapped fault-zone
components lead us to infer that the bulk permeability structure of
the present day fault zone is a heterogeneous and anisotropic
combined conduit-barrier ow system (e.g., Caine et al., 1996).
During seismic events, much of the fault zone behaves as a complex
and heterogeneous conduit where ow paths are likely controlled
to a large degree by mechanical heterogeneities, and the local
stresses they may induce (e.g., Faulkner et al., 2006). These
heterogeneous features are formed during progressive deformation
and associated sealing events. The source of uids, uid tempera-
tures, and component chemical constituents also play a large role in
the distribution of permeability and mechanical heterogeneity
from one seismic cycle to another.
The breccia pods contain remnants of likely coseismic, outcrop
to map-scale openings with components of hybrid dilatant and
shear deformation. Distinctive matrix-supported and jigsaw puzzle
breccias within these pods have microcrystalline, quartz-rich
mosaic textures that may record rapid, hydrothermal uid owand
mineral precipitation induced by the coseismic dilatancy and
hydraulic connection to an ancient geothermal reservoir. Alterna-
tively, the matrix-supported breccias and associated pods may have
formed from uid overpressures either in connection with the
geothermal reservoir or locally within the core of the fault. In either
case, the various breccia textures do not provide unique evidence
for a tectonic stress versus pore uid-pressure mechanism for their
origin. Yet, if either hypothesis is correct, these distinctive fault
rocks are formed during high enough strain rates to qualify as being
coseismic, and thus the breccias represent seismogenic rock
textures (cf. Cowan, 1999). Although, the fault rocks at the Mirrors
locality may be coincidental and unique to the Stillwater Fault Zone,
the textures and associated inferences regarding their formation
may be underappreciated where similar combinations of rocks,
structures, and geothermal uid reservoirs occur.
Acknowledgements
Funding for this work was provided by a U.S. Geological Survey,
National Earthquake Hazards Reduction Program Grant # 1434-93-
G-2280 to Forster and Bruhn who supported Caine with a graduate
research assistantship at the University of Utah, Department of
Geology and Geophysics from 1992 through 1995. We thank Bill
Parry, Laurel Goodwin, Darrel Cowan, George Breit, Don Sweetkind,
and Albert Hofstra, for helpful comments on an early version of this
manuscript. Lyndsay Ball, Dan Faulkner, David Lockner, Steven
Micklethwaite, and Rick Sibson provided critical reviews that
greatly improved the manuscript. We also thank Bennet Leeper,
Lori Chadwell, Kathleen Royster, and Bernard Seront who provided
eld assistance.
References
Antonellini, M., Aydin, A., 1994. Effect of faulting on uid ow in porous sandstones;
petrophysical properties. AAPG Bulletin 78, 355e377.
Aydin, A., Eyal, Y., 2002. Anatomy of a normal fault with shale smear; implications
for fault seal. AAPG Bulletin 86 (8), 1367e1381.
Barton, C.A., Hickman, S., Morin, R., Zoback, M.D., Finkbeiner, T., Sass, J., Benoit, W.R.,
1997. Fracture permeability and its relationship to in-situ stress in the Dixie
Valley, Nevada, geothermal reservoir. Report SGP-TR-155. In: Proceedings of the
22nd Workshop on Geothermal Reservoir Engineering. Stanford University,
Stanford, California, pp. 147e152.
Bruhn, R.L., Yonkee, W.A., Parry, W.T., 1990. Structural and uid-chemical properties
of seismogenic normal faults; earthquake source processes. In: 19th IUGG
General Assembly, Symposium on Earthquake Source Processes, Vancouver, BC,
Canada, Aug. 19e21, 1987, vol. 175, pp. 139e157.
Bruhn, R.L., Perry, W.T., Yonkee, W.A., Thompson, T., 1994. Fracturing and hydro-
thermal alteration in normal fault zones; faulting, friction, and earthquake
mechanics; part 1. Pure and Applied Geophysics 142, 609e644.
Byerlee, J.D., 1993. Model for episodic ow of high-pressure water in fault zones
before earthquakes. Geology 21, 303e306.
Caine, J.S., Evans, J.P., Forster, C.B., 1996. Fault zone architecture and permeability
structure. Geology 24, 1025e1028.
Caine, J.S., Forster, C.B., 1999. Fault zone architecture and uid ow; insights from
eld data and numerical modeling. In: Haneberg, W.C., Mozley, P.S., Moore, J.C.,
Goodwin, L.B. (Eds.), Faults and Sub-surface Fluid Flow in the Shallow Crust.
AGU Geophysical Monograph, vol. 113, pp. 101e127.
Caskey, S.J., Wesnousky, S.G., Zhang, P., Slemmons, D.B., 1996. Surface faulting of the
1954 Fairview Peak (M
s
7.2) and Dixie Valley (M
s
6.8) earthquakes, central
Nevada. Bulletin of the Seismological Society of America 86, 761e787.
Caskey, S.J., Wesnousky, S.G., 1997. Static stress changes and earthquake triggering
during the 1954 Fairview Peak and Dixie Valley earthquakes, central Nevada.
Bulletin of the Seismological Society of America 87 (3), 521e527.
Chester, F.M., Evans, J.P., Biegel, R.L., 1993. Internal structure and weakening
mechanisms of the San Andreas Fault. Journal of Geophysical Research 98 (B1),
771e786.
Chester, F.M., Logan, J.M., 1986. Implications for mechanical properties of brittle
faults from observations of the Punchbowl fault zone, California. Pure and
Applied Geophysics 124, 79e106.
Cowan, D.S., 1999. Do faults preserve a record of seismic slip? A eld geologists
opinion. Journal of Structural Geology 21, 995e1001.
Cox, S.F., Knackstedt, M.A., Braun, J., 2001. Principles of structural control on
permeability and uid ow in hydrothermal systems. Reviews in Economic
Geology 14, 1e24.
Dong, G., Morrison, G., Jaireth, S., 1995. Quartz textures in epithermal veins,
Queensland; classication, origin and implication. Economic Geology and the
Bulletin of the Society of Economic Geologists 90, 1841e1856.
Evans, J.P., Chester, F.M., 1995. Fluiderock interaction in faults of the San Andreas
system: inferences from San Gabriel fault rock geochemistry and microstruc-
tures. Journal of Geophysical Research 100 (B7), 13007e13020.
Faulkner, D.R., Mitchell, T.M., Healy, D., Heap, M.J., 2006. Slip on weak faults by
the rotation of regional stress in the fracture damage zone. Nature 444,
922e925.
Faulkner, D.R., Mitchell, T.M., Rutter, E.H., Cembrano, J., 2008. On the structure and
mechanical properties of large strikeeslip faults. In: Wibberley, C.A.J., Kurz, W.,
Imber, J., Holdsworth, R.E., Collettini, C. (Eds.), The Internal Structure of Fault
Zones: Implications for Mechanical and Fluid Flow Properties. Geological
Society, London, Special Publications, vol. 299, pp. 139e150.
Ferrill, D.A., Stamatakos, J.A., Sims, D., 1999. Normal fault corrugation; implications
for growth and seismicity of active normal faults. Journal of Structural Geology
21, 1027e1038.
Forster, C.B., Evans, J.P., Torgersen, T., 1991. Hydrogeology of thrust faults and
crystalline thrust sheets; results of combined eld and modeling studies.
Geophysical Research Letters 18, 979e982.
Fournier, R.O., 1985. The behavior of silica in hydrothermal solutions. In:
Berger, B.R., Bethke, P.M. (Eds.). Geology and Geochemistry of Epithermal
Systems. Reviews in Economic Geology 2, 45e61. Society of Economic
Geologists.
Foxford, K.A., Walsh, J.J., Watterson, J., Garden, I.R., Guscott, S.C., Burley, S.D., 1998.
Structure and content of the Moab Fault Zone, Utah, USA, and its implications
for fault seal prediction. In: Jones, G., Fisher, Q.J., Knipe, R.J. (Eds.), Faulting, Fault
Sealing, and Fluid Flow in Hydrocarbon Reservoirs. Geological Society, London,
Special Publications, vol. 147, pp. 87e103.
J.S. Caine et al. / Journal of Structural Geology 32 (2010) 1576e1589 1588
Ghisetti, F., Kirschner, D., Vezzani, L., Agosta, F., 2001. Stable isotope evidence for
contrasting paleouid circulation in thrust faults and normal faults of the
central Apennines, Italy. Journal of Geophysical Research 106, 8811e8825.
Goddard, J.V., Evans, J.P., 1995. Chemical changes and uiderock interaction in
faults of crystalline thrust sheets, northwestern Wyoming, U.S.A. Journal of
Structural Geology 17, 533e547.
Hancock, P.L., Barka, A.A., 1987. Kinematic indicators on active normal faults in
western Turkey. Journal of Structural Geology 9, 573e584.
Hedenquist, J.W., Henley, R.W., 1985. Hydrothermal eruptions in the Waiotapu
geothermal system, New Zealand; their origin, associated breccias, and relation
to precious metal mineralization. Economic Geology and the Bulletin of the
Society of Economic Geologists 80, 1640e1668.
Heynekamp, M.R., Goodwin, L.B., Mozley, P.S., Haneberg, W.C., 1999. Controls on
fault-zone architecture in poorly lithied sediments, Rio Grande Rift, New
Mexico; implications for fault-zone permeability and uid ow; faults and
subsurface uid ow in the shallow crust. Geophysical Monograph 113,
27e49.
Hickman, S.H., Barton, C.A., Zoback, M.D., Morin, R., Sass, J.H., Benoit, R., 1997. In situ
stress and fracture permeability along the Stillwater fault zone, Dixie Valley,
Nevada. International Journal of Rock Mechanics and Mining Sciences & Geo-
mechanics Abstracts 34, 414.
Hickman, S.H., Barton, C.A., Zoback, M.D., Morin, R.H., Benoit, R., et al., 1998. In-situ
stress and fracture permeability along the Stillwater fault zone, Dixie Valley,
Nevada. Eos, Transactions, American Geophysical Union 79, 814.
Hickman, S.H., Zoback, M.D., Barton, C.A., Benoit, R., Svitek, J., Summers, R., 2000.
Stress and permeability heterogeneity within the Dixie Valley geothermal
reservoir: recent results from well. Report SGP-TR-165. In: Proceedings of the
25th Workshop on Geothermal Reservoir Engineering. Stanford University,
Stanford, California, pp. 82e85.
Hubbert, M.K., Rubey, W.W., 1959. Mechanics of uid-lled porous solids and its
application to overthrust faulting, [part] 1 of role of uid pressure in mechanics
of overthrust faulting. Geological Society of America Bulletin 70, 115e166.
Jbrak, M., 1997. Hydrothermal breccias in vein-type ore deposits; a review of
mechanisms, morphology and size distribution. Ore Geology Reviews 12 (3),
111e134.
Keller, E.A., Loaiciga, H.A., 1993. Fluid-pressure induced seismicity at regional scales.
Geophysical Research Letters 20, 1683e1686.
Knipe, R.J., Jones, G., Fisher, Q.J., 1998. Faulting, fault sealing and uid ow in
hydrocarbon reservoirs: an introduction. In: Jones, G., Fisher, Q.J., Knipe, R.J.
(Eds.), Faulting, Fault Sealing, and Fluid Flow in Hydrocarbon Reservoirs.
Geological Society, London, Special Publications, vol. 147, pp. viiexxi.
Lee, J.-J., Bruhn, R.L., 1996. Structural anisotropy of normal fault surfaces. Journal of
Structural Geology 18, 1043e1059.
Lockner, D.A., Tanaka, H., Ito, H., Ikeda, R., Omura, K., Naka, H., 2009. Geometry of
the Nojima fault at Nojima-Hirabayashi, Japan e I. A simple damage structure
inferred from borehole core permeability. Pure and Applied Geophysics 30,
1649e1667.
Lovering, T.G., 1972. Jasperoid in the United States; Its Characteristics, Origin, and
Economic Signicance. U.S. Geological Survey Professional Paper, Report: P
0710, 164 pp.
Lutz, S.J., Moore, J.N., Benoit, D., 1997. Geologic framework of Jurassic reservoir rocks
in the Dixie Valley geothermal eld, Nevada: implications from hydrothermal
alteration and stratigraphy. Proceedings e Workshop on Geothermal Reservoir
Engineering 155, 131e139.
Lutz, S.J., Caskey, S.J., Mildenhall, D.C., Browne, P.R.L., Johnson, S.D., 2002. Dating
sinter deposits in northern Dixie Valley, Nevada; the paleoseismic record and
implications for the Dixie Valley geothermal system. Proceedings e Workshop
on Geothermal Reservoir Engineering 171, 284e290.
Martel, S.J., 1990. Formation of compound strikeeslip fault zones, Mount Abbot
Quadrangle, California. Journal of Structural Geology 12, 869e882.
Micklethwaite, S., 2009. Mechanisms of faulting and permeability enhancement
during epithermal mineralisation; Cracow goldeld, Australia. Journal of
Structural Geology 31 (3), 288e300.
Miller, S.A., Nur, A., Olgaard, D.L., 1996. Earthquakes as a coupled shear stress e high
pore pressure dynamical system. Geophysical Research Letters 23, 197e200.
Miller, S.A., Collettini, C., Chiaraluce, L., Cocco, M., Barchi, M., Kaus, B.J.P., 2004.
Aftershocks driven by a high-pressure CO
2
source at depth. Nature 427,
724e727.
Morrow, C.A., Shi, L.Q., Byerlee, J.D., 1984. Permeability of fault gouge under
conning pressure and shear stress. Journal of Geophysical Research 89 (B5),
3193e3200.
Mort, K., Woodcock, N.H., 2008. Quantifying fault breccia geometry: Dent Fault, NW
England. Journal of Structural Geology 30, 701e709.
Muir-Wood, R., King, G.C.P., 1993. Hydrological signatures of earthquake strain.
Journal of Geophysical Research 98 (B12), 22035e22068.
Muir-Wood, R., 1994. Earthquakes, strain-cycling and the mobilization of uids. In:
Geological Society, London, Special Publications, vol. 78 85e98.
Nemcok, M., Henk, A., Gayer, R.A., Vandycke, S., Hathaway, T.M., 2002. Strikeeslip
fault bridge uid pumping mechanism: insights from eld-based
palaeostress analysis and numerical modeling. Journal of Structural Geology 24,
1885e1901.
Newhouse, W.H., 1942. Ore Deposits as Related to Structural Features. Hafner
Publishing Co., New York, London, 280 pp.
Nur, A., Booker, J.R., 1972. Aftershocks caused by pore uid ow? Science 175,
885e887.
Page, B.M., 1965. Preliminary geologic map of a part of the Stillwater Range,
Churchill County, Nevada. In: Map 28-Nevada Bureau of Mines and Geology.
Parry, W.T., Bruhn, R.L., 1990. Fluid pressure transients on seismogenic normal
faults. Tectonophysics 179, 335e344.
Parry, W.T., Hedderly-Smith, D., Bruhn, R.L., 1991. Fluid inclusions and hydrothermal
alteration on the Dixie Valley fault, Nevada. Journal of Geophysical Research 96
(B12), 19733e19748.
Power, W.L., Tullis, T.E., 1989. The relationship between slickenside surfaces in ne-
grained quartz and the seismic cycle. Journal of Structural Geology 11, 879e893.
Priest, S., 1993. Discontinuity Analysis for Rock Engineering. Chapman and Hall.
Rice, J.R., 1992. Fault stress states, pore pressure distributions, and the weakness of
the San Andreas Fault. In: Evans, B., Wong, T.-F. (Eds.), Fault Mechanics and
Transport Properties of Rocks; a Festschrift in Honor of W.F. Brace. Academy
Press, pp. 475e503.
Rimstidt, J.D., 1997. Gangue mineral transport and deposition. In: Barnes, H.L. (Ed.),
Geochemistry of Hydrothermal Ore Deposits. John Wiley & Sons, New York, NY,
pp. 487e516.
Sagy, A., Brodsky, E.E., Axen, G.J., 2007. Evolution of fault-surface roughness with
slip. Geology 35, 283e286.
Saunders, J.A., 1994. Silica and gold textures in bonanza ores of the Sleeper Deposit,
Humboldt County, Nevada; evidence for colloids and implications for epi-
thermal ore-forming processes. Economic Geology and the Bulletin of the
Society of Economic Geologists 89, 628e638.
Scholz, C.H., 2002. The Mechanics of Earthquakes and Faulting, second ed. Cam-
bridge University Press, Cambridge.
Seront, B., Wong, T.-F., Caine, J.S., Forster, C.B., Bruhn, R.L., Fredrich, J.T., 1998.
Laboratory characterization of hydromechanical properties of a seismogenic
normal fault system. Journal of Structural Geology 20, 865e881.
Sibson, R.H., 1977. Fault rocks and fault mechanisms. Journal of the Geological
Society of London 133, 191e213.
Sibson, R.H., 1981. Fluid ow accompanying faulting; eld evidence and models. In:
Simpson, D.W., Richards, P.G. (Eds.), Maurice Ewing Series, no.4, pp. 593e603.
Sibson, R.H., 1986a. Brecciation processes in fault zones; inferences from earth-
quake rupturing. Pure and Applied Geophysics 124, 159e175.
Sibson, R.H., 1986b. Earthquakes and rock deformation in crustal fault zones.
Annual Review of Earth and Planetary Sciences 14, 149e175.
Sibson, R.H., 1990. Conditions for fault-valve behaviour. In: Knipe, R.J., Rutter, E.H.
(Eds.), Geological Society, London, Special Publications, vol. 54, pp. 15e28.
Sibson, R.H., 1992. Implications of fault-valve behaviour for rupture nucleation and
recurrence. In: Mikumo, T., Aki, K., Ohnaka, M., Ruff, L.J., Spudich, P.K.P. (Eds.),
Earthquake Source Physics and Earthquake Precursors. Tectonophysics 211,
283e293.
Sibson, R.H., 1996. Structural permeability of uid-driven fault-fracture meshes.
Journal of Structural Geology 18, 1031e1042.
Sibson, R.H., 2001. Seismogenic framework for hydrothermal transport and ore
deposition. Reviews in Economic Geology 14, 25e50.
Sillitoe, R.H., 1985. Ore-related breccias in volcanoplutonic arcs. In: Sawkins, F.J.,
Sillitoe, R.H. (Eds.), Economic Geology and the Bulletin of the Society of Economic
Geologists, 80. Economic Geology Publishing Company, Lancaster, pp. 1467e1514.
Smith, L., Forster, C., Evans, J., 1990. Interaction of fault zones, uid ow, and heat
transfer at the basin scale. In: Neuman, S.P., Neretnieks, I. (Eds.), Hydrogeology,
vol. 2, pp. 41e67.
Smith, S.A.F., Collettini, C., Holdsworth, R.E., 2008. Recognizing the seismic cycle
along ancient faults: CO
2
-induced uidization of breccias in the footwall of
a sealing low-angle normal fault. Journal of Structural Geology 30, 1034e1046.
Speed, R.C., Armstrong, R.L., 1971. Potassiumeargon ages of some minerals from
igneous rocks of western Nevada. Isochron/West 1, 1e8.
Speed, R.C., 1976. Geologic Map of the Humboldt Lopolith and Surrounding Terrane,
Nevada.
Tanaka, H., Hinoki, S.-I., Kosaka, K., Lin, A., Takemura, K., Murata, A., Miyata, T.,
Shimamoto, T., Fujimoto, K., Wibberley, C.A.J., 2001. Deformation mechanics and
uid behavior in a shallow, brittle fault zone during coseismic and interseismic
periods; results from drill core penetrating the Nojima Fault, Japan. Island Arc
10, 381e391.
Ujiie, K., Yamaguchi, A., Kimura, G., Toh, S., 2007. Fluidization of granular material in
a subduction thrust at seismogenic depths. Earth and Planetary Science Letters
259 (3e4), 307e318.
Unsworth, M.J., Malin, P.E., Egbert, G.D., Booker, J.R., 1997. Internal structure of the
San Andreas Fault at Parkeld, California. Geology 25, 359e362.
Vikre, P.G., 1994. Gold mineralization and fault evolution at the Dixie Comstock
Mine, Churchill County, Nevada. Economic Geology and the Bulletin of the
Society of Economic Geologists 89, 707e719.
Vrolijk, P., van der Pluijm, B.A., 1999. Clay gouge. Journal of Structural Geology 21
(8e9), 1039e1048.
Wallace, R.E., Whitney, R.A., 1984. Late Quaternary History of the Stillwater Seismic
Gap, Nevada.
Wibberley, C.A.J., 2002. Hydraulic diffusivity of fault gouge zones and implications
for thermal pressurization during seismic slip. Earth. Planets and Space 54 (11),
1153e1171.
Wilden, R., Speed, R.C., 1974. Geology and mineral resources of Churchill County,
Nevada. Nevada Bureau of Mines and Geology Bulletin 83, 95.
Woodcock, N.H., Dickson, J.A.D., Tarasewicz, J.P.T., 2007. Transient permeability and
reseal hardening in fault zones; evidence from dilation breccia textures. In:
Lonergan, L., Jolly, R.J.H., Rawnsley, K., Sanderson, D.J. (Eds.), Fractured reser-
voirs, 270. Geological Society Special Publications, London, pp. 43e53.
J.S. Caine et al. / Journal of Structural Geology 32 (2010) 1576e1589 1589
Evolution of cataclastic faulting in high-porosity sandstone, Bassin du Sud-Est,
Provence, France
Elodie Saillet
a,
*
, Christopher A.J. Wibberley
b
a
Geosciences Azur, UMR 6526, 250 Av. Albert Einstein, 06560 Valbonne, France
b
TOTAL EP, CSTJF, Av. Larribau, 64018 Pau, France
a r t i c l e i n f o
Article history:
Received 14 April 2009
Received in revised form
9 February 2010
Accepted 17 February 2010
Available online 6 March 2010
Keywords:
Reservoirs
Porous sandstones
Cataclastic Deformation Bands (CDBs)
Bassin du Sud-Est
a b s t r a c t
Cataclastic deformation structures in Cretaceous high-porosity sands in the Bassin du Sud-Est, SE France
were surveyed by scan-lines to examine: (i) the role of tectonic loading path on cataclastic deformation
band (CDB) network development, and (ii) the development of larger ultracataclastic faults as strain
increases. Deformation during PyreneaneProvenal shortening resulted in a persistent high density
(w10/m
2
) of conjugate reverse-sense CDB zones (displacements up to w30 cm), with no generation of
larger faults. Highelow-density undulations occur for each pair of the conjugate set in an alternating
manner, suggestive of network hardening, with a wavelength of several tens of metres being in the order
of mechanical bed thickness. For two study areas which experienced signicant OligoceneeMiocene
extension, a moderate, undulating background density (w4/m
2
) of normal-offset CDBs was recorded,
which became focussed in places into clusters (w50/m
2
) a few metres wide. Thus tectonic loading path
may strongly inuence strain distribution. CDB zones develop by the addition of successive bands at the
edges until, at a thickness of around 5 cm, new bands tend to stray further away from the zone edges.
Coarser sands have thicker CDB zones, suggesting that host grain size, along with mechanical bed
thickness, could be an important contributor to the scale limit in CDB zone growth. Larger ultracataclastic
faults and discrete slip zones localised within or at the edges of some clusters of CDB zones, post-date
cluster development rather than inducing it. This stage of deformation evolution is only reached in
extension, not in shortening, suggesting the infeasibility of achieving the critical state line during
horizontal compression.
2010 Elsevier Ltd. All rights reserved.
1. Introduction
Fluid circulation in the crust, and in particular hydrocarbon
migration in reservoirs, is highly dependant on fault geometrical
and hydromechanical properties (e.g. Manzocchi et al., 1998;
Matthi et al., 1998; Wibberley et al., 2008). Faulting in porous
sandstone often produces zones of deformation bands rather than
planar fracture surfaces (Aydin, 1978; Aydin and Johnson, 1978,
1983; Underhill and Woodcock, 1987; Antonellini et al., 1994;
Davis, 1999; Fossen et al., 2007). Cataclastic deformation bands
(CDBs) are brittle shear zones that form through a combination of
compaction and cataclasis. Porosity and grain size reduction asso-
ciated with CDB formation are thought to cause strain hardening,
further deformation then being accommodated by deformation of
the wall rock, adjacent to the initial band (Aydin, 1978; Aydin and
Johnson, 1978, 1983; Underhill and Woodcock, 1987). Continued
deformation may possibly result in the development of localised
slip surfaces at the edge of deformation band zones (Aydin and
Johnson, 1983; Antonelini and Aydin, 1995; Shipton and Cowie,
2001). Some, but not all of these different eld observations have
been understood through laboratory experiments (Wong et al.,
1997; Mair et al., 2000; Torabi et al., 2007).
Descriptions of such deformation distributions in high-porosity
sandstones are quite varied, ranging from examples of deformation
band localisation as damage zones around larger faults, in relay
zones (often expressed as ladder zones cf. Schultz and Balasko,
2003) or in fault-tip folds, to zones of deformation distributed
over distances much greater than typical damage zone widths (e.g.
w100 m) with no obvious relationship to larger structures (Aydin,
1978; Underhill and Woodcock, 1987; Jamison and Stearns, 1982;
Shipton and Cowie, 2001; Du Bernard et al., 2002a,b; Wibberley
et al., 2007). Yet although recent advances have been made in
understanding the mechanics of yielding to generate a single
deformation band (e.g. Schultz and Siddharthan, 2005; Aydin et al.,
2006; Wibberley et al., 2007), no unifying mechanical model exists
for explaining and predicting distributions of deformation bands in
* Corresponding author. Tel.: 33 492942682.
E-mail address: saillet@geoazur.unice.fr (E. Saillet).
Contents lists available at ScienceDirect
Journal of Structural Geology
j ournal homepage: www. el sevi er. com/ l ocat e/ j sg
0191-8141/$ e see front matter 2010 Elsevier Ltd. All rights reserved.
doi:10.1016/j.jsg.2010.02.007
Journal of Structural Geology 32 (2010) 1590e1608
terms of regional controls such as tectonic loading paths. Further-
more, the evolution of a network, and mechanical controls on this
evolution, are still unclear, particularly with respect to localisation
processes (a) at the scale of a single deformation band e what
controls deformation localisation into a band, and deformation
band growth? (b) at the scale of a cluster of deformation bands edo
they become clustered around previously formed larger faults as
fault damage zones, or do the clusters of deformation bands form
rst, by early deformation localisation, with continued localisation
of deformation generating the larger faults within or at the edges of
these clusters? Some example of these behaviours exist (Fossen
et al., 2000, 2007; Shipton and Cowie, 2001), yet understanding
the mechanical controls on fault distribution is fundamental in
order to provide better fault distribution prediction in reservoir
ow simulations from limited structural input data (Saillet, 2009).
This paper presents a statistical and quantitative eld study
aimed at providing a model of deformation patterns and fault
growth mechanisms in high-porosity sandstones. This study pres-
ents three different cases of Cretaceous high-porosity sands and
sandstones in the Bassin du Sud-Est, Provence, France (Fig. 1). The
eld data were recorded along scan-lines from excellent quarry
exposures from the Orange, Massif d'Uchaux and Bdoin areas
(Fig. 1). These eld data allow us to provide detailed descriptions of
the distribution of deformation and its evolution, leading to inter-
pretations in terms of fault growth mechanisms and fault network
development in high-porosity sandstones. This part of the study
concerns only the geometrical evolution of the deformation. The
impact of deformed structures in sandstones on uid ow, evalu-
ated from permeability and geometric properties of the structures,
will be addressed in a separate publication.
2. Geological setting and data acquisition
2.1. Regional context of the Bassin du Sud-Est
The Bassin du Sud-Est is a triangular region between the Massif
Central to the North West, the Alps to the East, and the Mediter-
ranean Sea to the South. It is a Mesozoic cratonic basin on the edge
of the Alpine orogen, w200 km long and 100e150 km wide. The
total thickness of the sedimentary units is up to 10,000 m in the
central area, but this thickness decreases to 2000e3000 m towards
the edge of the basin (Delfaud and Dubois, 1984). From the Triassic
to the Cretaceous, sedimentary deposits are essentially marine,
corresponding to basin rifting related to Tethys opening. In the
middle Cretaceous, the sedimentary units correspond to detrital
sands deposited during the beginning of basin inversion. From late
to end Cretaceous, the sands were deposited only in a continental
environment, with locally high sedimentary rates (Debrand-
Passard et al., 1984). In the West of the Bassin du Sud-Est, much of
the resulting CenomanianeTuronian deposits are high-porosity
sands and poorly to moderately consolidated sandstones.
This paper presents a combined study of three areas in the
Bassin du Sud-Est, the Bdoin, Massif d'Uchaux and Orange areas
(Fig. 1). These studies were carried out on high-porosity sand and
sandstone outcrops in active and abandoned quarries (Figs. 2e4),
which provide excellent 2-D and 3-D exposures of deformation
band networks and larger faults. These sands/sandstones are
composed of a large range of quartz grain sizes which vary between
the study areas. They also contain a fewclay lamellae and, in places,
limestone beds containing a few shallow marine fossils. These
sands generally have a marine origin from deltaic to beach sands,
Fig. 1. Simplied geological and structural map of the study areas in the Bassin du Sud-Est, Provence, Southeast France. Locations of the three study areas are denoted by boxes.
Modied from Wibberley et al. (2007).
E. Saillet, C.A.J. Wibberley / Journal of Structural Geology 32 (2010) 1590e1608 1591
with some beds of aeolian origin. The sand outcrops show a low to
moderate cohesion between the quartz grains due to a lack of
cement, but can formvertical quarry faces for some months to years
before erosion and outcrop collapse sets in. These sand units have
undergone a low depth of burial, possibly less than 800 m (Delfaud
and Dubois, 1984). Deformation is expressed in these sands and
sandstone units as cataclastic deformation bands (CDBs), larger,
ultracataclastic (mature) faults and discrete slip surfaces
(Wibberley et al., 2007). The orientation and kinematics of these
structures are highly variable, but, fromregional knowledge, can be
attributed to the three different regional tectonic events in Pro-
vence: (i) NortheSouth PyreneaneProvenal shortening accom-
modated by regional foreland thrusting, giving the Cretaceous
outcrop an EasteWest structural trend; (ii) NWeSE Oligocene-
eearly Miocene rifting which caused normal faulting in parts of the
Upper Cretaceous strata; (iii) Miocene left-lateral strike-slip reac-
tivation of some of the major pre-existing NEeSW faults in the
region.
2.2. Deformation features
The deformation features in the Cretaceous high-porosity sands
and sandstones of the Bassin du Sud-Est consist of cataclastic
deformation bands (CDBs) and larger ultracataclastic faults or
discrete slip surfaces. At the outcrop scale the CDBs appear as single
bands with millimetric offsets, usually grouped into zones of
narrowly spaced bands with offsets from a few millimetres to tens
of centimetres (Fig. 5). In most cases, the ne-scale sedimentary
lamellae make it easy to determine apparent offsets. At the
microscopic scale these CDBs are characterized by grain crushing
and compaction (Wibberley et al., 2007). Such grain crushing and
compaction resulted in a signicant decrease in the porosity and
permeability between the host sand and the deformed sand as
previously described elsewhere (e.g. Underhill and Woodcock,
1987; Antonelini and Aydin, 1994; Fowles and Burley, 1994). The
observations of these CDBs in the Bassin du Sud-Est suggest an
evolution with displacement increase and thickness growth of the
structures, similar to previous observations (Aydin and Johnson,
1978; Underhill and Woodcock, 1987; Antonelini and Aydin,
1995). The thinner features are deformation bands with very low
displacements (e.g. <10 mm). These bands correspond to a single
cataclastic fault strand with a thickness of one to a few millimetres.
The SEM observations of these CDBs (Wibberley et al., 2007) show
a preferential occurrence of fractures at grain contact points, sug-
gesting a process of Hertzian fracturing (Gallagher et al., 1974). This
observation also shows that, for a single CDB, this Hertzian frac-
turing process involves a zone typically not greater than ve quartz
grains wide. Scanning electron microscope (SEM) observations
show that the processes of fracture and cataclasis produce a new
generation of very ne grained fragments inlling the porosity
space.
Thicker features are multiple strand zones of cataclastic
deformation bands (e.g. 10e300 mm in displacement, 10e100 mm
wide) formed by generation of successive adjacent single bands,
referred to hereafter as CDB zones. The adjacent bands within
these CDB zones can be generated with a variable spacing. At the
outcrop scale it is not easy to distinguish the different bands within
a CDB zone, except where there are lenses of host sand between the
different bands. Field observations show that the number of indi-
vidual bands in a CDB zone correlates with offset, up to the scale
limit of the deformation bands, around 30 cm offset (Wibberley
et al., 2000). This nding contrasts with studies of the number of
discrete slip surfaces in fault core zones in high-porosity sand-
stones which show no correlation with offset (Shipton et al., 2005).
Larger faults, with offsets greater than around 1 m, are also
present in the high-porosity sands and sandstones of the Bassin du
Sud-Est. These faults appear to have been generated separately from
the smaller-offset CDBs, in some cases during a later tectonic event
Fig. 2. (a) Location of the study area at Orange. Modied from the 1/50,000 geological map of Orange, BRGM. (b) Location of the scan-line within the sandstones from the Quartier
de l'tang quarry.
E. Saillet, C.A.J. Wibberley / Journal of Structural Geology 32 (2010) 1590e1608 1592
than those forming the cataclastic deformation bands. These larger
faults present drastic differences in their properties and charac-
teristics compared to the CDBs. On the eld scale, these faults show
a fairly homogeneous deformation zone of ultracataclastic fault
rock, with sometimes one or more discrete slip surfaces in the fault
core, and a cluster of anastomosing cataclastic strands around the
edges of the fault zone (Fig. 6). At the microscopic scale there are
also large differences between these faults and the CBD micro-
textures. SEM photomicrographs (Wibberley et al., 2007) show
some rounded quartz grains of original sedimentary size and
a higher proportion of ultrane material than in CDBs. These
differences seem to reect a transition in fault growth mechanisms
fromCDBs to larger faults, yet there are a relatively small number of
these ultracataclastic fault zones in the study areas. In the Orange
and Bdoin areas, these mature faults are normal faults, post-dating
the earlier reverse-sense CDBs. In the Massif d'Uchaux area they are
large strike-slip faults, cross-cutting the earlier normal-sense CDBs.
Thus, in some cases at least, deformation localisation into
ultracataclastic fault zones is encouraged by the presence of an
earlier set of deformation bands of high density (Wibberley et al.,
2007).
2.3. Field data acquisition
In order to characterise the deformation distribution in each
area, the CDBs were sampled by scan-lines made on each outcrop.
The position of all visible CDBs or larger faults was recorded along
one metre-wide scan-lines (thus structural density, rather than
frequency, is recorded), perpendicular to the strikes of the main
structures. This methodology is similar to at least some of the
previous studies of this type elsewhere (e.g. Du Bernard et al.,
2002b). Because of our rectilinear quarry topography, it is not
necessary to make any statistical correction in the CDB positions on
the outcrop, although density corrections were performed for dip
obliquity between (horizontal) scan-line and non-vertical struc-
tures (Priest and Hudson, 1981). For each CDB, the orientation, the
Fig. 3. (a) Location of the study area in the Massif d'Uchaux. Modied from the 1/50,000 geological map of Orange, BRGM. (b) Location of the different scan-lines within the
Boncava quarry.
E. Saillet, C.A.J. Wibberley / Journal of Structural Geology 32 (2010) 1590e1608 1593
thickness, and the number of CBDs in a CDB zone and their offset
were recorded in order to statistically analyse each outcrop. CDB
zones are grouped into a single structure provided that the outer-
most strands are no further from the remaining structure than the
thickness of the remaining structure, and that they converge to
touch the structure. A total scan-line length of 717 m was recorded
for the 3 study areas combined.
3. The distribution of deformation
3.1. Outcrop description
3.1.1. Orange sandstones
The Orange study area is located in the abandoned Quartier de
l'Etang quarry, on the south side of the town (Fig. 2). The outcrop is
composed of different sedimentary units of Cenomanian age. The
majority of the quarry is composed of a cohesive (not disintegrating
to the touch) quartz sandstone unit, mineralogically mature, but
texturally immature, with a high porosity partly due to the small
proportion of quartz cement. This sandstone unit has a marine to
beach origin, and is composed of medium-size quartz grains of
average size of 550 mm. At outcrop, it is possible to observe some
sedimentary channel structures formed in a marine environment,
and in other parts of the outcrop some benthic shallow marine
fossils are present. The sandstone unit is covered by a calcarenite
unit, the contact between these two units dipping around 8
to the
South-West. The deformation of the sandstone unit of the Quartier
de l'Etang quarry consists of conjugate reverse-sense CBDs with
a low angle of dip (34
to 37
(1)
On the density proles, the D
corr
and the D
eld
are given. The
density proles obtained illustrate three important features
(Fig. 8):
1. The CDB density scan-line corresponding to the total data set of
the conjugate reverse fault population shows a persistently
high density (generally in the range 7e15/m
2
), and a more-or-
less continuous distribution of the deformation along the
outcrop, with a clustering in the South (Fig. 8a). The clusters of
CDBs on the South side of the outcrop correspond to the
presence of ladder zones of very thin CDBs and do not impact
signicantly on the overall distribution of strain.
2. The CDB density variations corresponding to the two different
(opposing dips) sets of the conjugate fault population show
a more heterogeneous distribution of deformation (Fig. 8b, c).
The density prole corresponding to the SW-dipping conjugate
population shows a high CDB concentration in the south,
middle and north of the outcrop but low-density voids
around 100 and 160 m (Fig. 8b). The density prole corre-
sponding to the opposite set (North-dipping) of the conjugate
population shows a high CDB concentration around 100 and
160 m but important voids in the South, in the centre at
around 120 m, and in the North (Fig. 8c). As illustrated by the
vertical dashed arrows in Fig. 8b and c, these voids or troughs
in the NE-dipping set density prole coincide closely with the
Fig. 5. Field appearance of conjugate arrays of cataclastic deformation bands in the Upper Cretaceous sandstones at the Quartier de l'Etang quarry, Orange. a) Detailed eld view of
the alternating chronology between structures of two conjugate sets. (i) Field photo; (ii) interpretative eld sketch. b) Broader eld view of the generally synchronous, but locally
alternating, relationships between two conjugate sets of reverse CDBs and CDB zones. (i) Field photo; (ii) interpretative eld sketch.
E. Saillet, C.A.J. Wibberley / Journal of Structural Geology 32 (2010) 1590e1608 1595
positions of high-density zones of the SW-dipping set, and vice
versa, i.e. that there is a spatial correlation between the high
density of one of the conjugate sets and a low density of the
other. A statistical analysis of deformation distribution, based
on the clustering tendency as used in Wibberley et al. (2007),
improved to analyse a range of data bin sizes, shows that each
of the two conjugate sets is much more clustered than the
overall population treated as a single ensemble (Fig. 9).
Furthermore, the cluster analysis is strongest for bin sizes in the
order of w60 m along the scan-line (Fig. 9), converting to
a wavelength of around 30 m when the dip correction is
applied to these low-angle structures, i.e. on the order of bed
thickness.
3. The larger normal fault, on the North side of the outcrop
(Fig. 2), corresponding to a later tectonic event, has no spatial
correlation with any of the reverse structures. It has a throw of
w10 m and is an ultracataclastic fault zone around 75 cmwide
where measured. The microstructures show an intense
decrease of grains and pores size, which can signicantly retard
uid ow (Saillet, 2009). Additionally, Wibberley et al. (2007)
show that larger faults form preferentially in regions which
have already recorded an earlier deformation event.
3.2.2. Massif d'Uchaux sands
The deformation of the Massif d'Uchaux is essentially expressed
by a population of normal-sense conjugate CDBs and later strike-
slip faults. A lownumber of small low-angle reverse-sense CDBs are
present, probably generated by an earlier tectonic event. The
density proles (Fig. 10) corresponding to the different outcrops
show essentially the conjugate and normal CDBs but some strike-
slip faults are also present including a large mature ultra-
cataclastic strike-slip fault which crosses the entire quarry outcrop
(Fig. 3). On each prole the position of this large strike-slip fault is
indicated. The density proles (Fig. 10) illustrate three important
features:
1. The different density proles show similar patterns of spatial
deformation distribution. The most striking feature of this
pattern is the relatively persistent, in relation to precedent
publications (Du Bernard et al., 2002b), yet undulating back-
ground distribution of the deformation over the length of the
45e60 m scan-lines.
2. For all the outcrops there are a few zones which present an
anomalous CDB density increase in addition to the moderate
overall density (Fig. 10). These zones of high-density concen-
tration are ladder zone features such as in the middle of the
outcrop South-1. The anomalous weighting effect of small
ladder zones on density proles has been checked by also
plotting proles of cumulative thickness per square metre (not
presented in this paper for simplicity). In such a case the
outcrop shows a more even deformation distribution because
each ladder zone is a high density of very thin features,
impacting much less on cumulative thickness proles. In other
Fig. 6. Field appearance of relationships between thin CDBs, large clusters and a slip zone in the Upper Cretaceous sands of the Le Cros quarry at Bdoin. a) Field view of the
relationships between earlier thin structures and two larger CDB clusters with strike-slip kinematics. (i) Field photo; (ii) interpretative eld sketch. b) Field view of the relationships
between a cemented slip zone within a large CDB cluster of normal fault kinematics. The eld observations show that earlier CDBs cluster is cut by the slip zone. (i) Field photo;
(ii) interpretative eld sketch.
E. Saillet, C.A.J. Wibberley / Journal of Structural Geology 32 (2010) 1590e1608 1596
parts of the outcrop there are increases of CDB density unre-
lated to specic structures such as a larger fault.
3. All the outcrops studied are cut by the same large ultra-
cataclastic strike-slip fault zone (with bedding-parallel stria-
tions) which is systematically correlated with a large CDB
density increase. Nevertheless, it is difcult to determine the
kinematics of these CDBs which are present around the larger
fault and hence infer the tectonic event which generated them.
The CDBs clustered around the main strike-slip fault zone show
normal apparent offset in the vertical outcrop views, but due to
the geometry of the dip of the CDBs and the dip of the stra-
tigraphy they may in reality be normal or strike-slip structures.
Hence the relative timing of background CDBs with respect to
clusters is not inferred by these observations.
All the Massif d'Uchaux scan-lines show relatively similar
deformation distribution patterns and CDB densities. These density
distribution diagrams illustrate three important features: (i) The
overall CDB density is moderate (mean values are given in Fig. 10),
with the predominance of single CDBs and thin CDB zones; (ii) A
few patches present anomalous CDB density increases, where CDB
zones are organized in clusters; (iii) These clusters are not
systematically associated with main slip surfaces or larger ultra-
cataclastic fault zones, but those slip surfaces or fault zones that are
present do exist within or at the edges of clusters of CDB zones.
3.2.3. Bdoin sands
Three different events have been recorded by the outcrops of
Bdoin and are evidenced along each of the four scan-lines recor-
ded in the study area (Fig. 4): the PyreneaneProvenal shortening,
the OligoceneeMiocene extension and the strike-slip faulting
during the Miocene. However, offset along the CDBs and CDB zones
is of a similar magnitude to the CDB zone thickness, and displace-
ment is only visible to the naked eye for the larger features. Hence it
Fig. 7. Lower-hemisphere equal-area stereographic projections (stereograms) of poles to planes for cataclastic deformation bands and larger faults from the studied areas. (a) Data
from the Orange area; (b) data from the Massif d'Uchaux area; (c) data from the Bdoin area.
E. Saillet, C.A.J. Wibberley / Journal of Structural Geology 32 (2010) 1590e1608 1597
is very often not possible to determine the kinematics of any one
CDB structure. The density proles obtained on the scan-lines
(totalling 250 m in length) illustrate three important features
similar to those of the Massif d'Uchaux area (Fig. 11):
1. There is generally an overall deformation pattern of undulating
moderate CDB density which constitutes a background
deformation along the outcrops.
2. All the different outcrops from the Bdoin study area show
some patches with an increase in the density of structures
corresponding to a localisation of the deformation within a few
clusters of CDB zones. These clusters are 5e10 m wide with
a CDB density comprised between 30 and 50 structures per
square metre, for a general density average comprised between
6 and 10 structures. These clusters do not generally coincide
with the position of any larger fault, except in one case (scan-
line 3, Fig. 11).
3. Although there are not many larger faults visible at the
outcrops in the Bdoin study area, it is possible to identify two
larger structures. The rst one is situated on scan-line number
3 (Figs. 4 and 11), where it is possible to observe a progressive
increase of the CDB density to a maximum in the centre of the
cluster, beyond which the density drops sharply to dene the
other edge of the cluster. Localisation of a discrete slip surface
occurred at this sharp transition region at the edge of the
cluster. On this rst structure it is difcult to determine the
relative timing of cluster and slip surface because of the poor
condition of the outcrop. The second structure is situated on
the south side of the quarry and was not sampled by any scan-
line because of the orientation of the fault with respect to the
outcrop. Nevertheless, it is also possible in this case to observe
a cluster of CDB zones around a central slip surface with the
same orientation and dip. Given that many clusters exist with no
localised slip planes or larger fault zones, it seems logical to
suggest that the clusters evolved rst, only some of which then
localised deformation into more evolved fault and/or slip zones.
The Bdoin scan-lines show relatively similar deformation
distribution patterns to the Massif d'Uchaux outcrops: (i) The
overall CDB density is undulating but generally moderate, with
Fig. 8. Graphs of CDB densities for the conjugate, reverse structures recorded along the 258 m long outcrop of the Quartier de l'Etang quarry in the Orange area. Density is dened
as the number of CDBs per square metre, directly measured on the eld or with angle correction (see text for more details). (a) The overall sum of the CDBs present in the scan-line
of the outcrop. (b) Density distribution of the SW-dipping CDB set. (c) Density distribution of the N-dipping CDB set. For each graph n is the total number of deformed structures and
D
av
is the mean average density directly measured on the eld (without angle correction). Cf is the cluster factor in each case (see text for details).
E. Saillet, C.A.J. Wibberley / Journal of Structural Geology 32 (2010) 1590e1608 1598
mean scan-line averages between 6 and 10 structures per m; (ii) A
few zones of anomalously high CDB density exist (with density of
30e50 structures per m), where CDB zones are organized into
clusters; (iii) These clusters are not systematically associated with
discrete slip surfaces or ultracataclastic fault zones, but ultra-
cataclastic fault zones are sometimes present within the maximum
density of CDB zones. In the Bdoin area the deformation features
correspond to three different tectonic events and in some cases it is
not possible to determine the kinematics or displacement with the
naked eye.
4. Characterisation of individual CDB properties based
on eld data
4.1. Fault thickness data
Previous studies of the relationship between fault zone thick-
ness and fault displacement (Robertson, 1983; Scholz, 1987; Hull,
1988; Evans, 1990) have generally shown that although there is
an overall linear relationship between fault zone thickness and
displacement, it is necessary to take account of various problems
like the denition of the edges of the fault zone or the thickness
variation along a single fault when interpreting data. Fault zone
thicknessedisplacement data are presented in Fig. 12, for single
CDBs, CDB zones and larger ultracataclastic fault zones in the study
area of the sandstones from Orange and the incohesive sands from
Massif d'Uchaux and Bdoin. The data are presented with two
different thickness data measurements. Thicknesses using method
1 are given as the total distance between the outermost edges of the
outermost strands constituting the multiple-band CDB zones. With
this rst measurement method the host sand contained in between
the different strands of a multiple-band CDB zone is included in the
measurement. Thicknesses using method 2 are given as the sum of
the thicknesses of each individual band constituting a multiple-
band CDB zone. This second measurement method denes an
effective thickness of deformed material, taking into account that
material between the bands is undeformed host sand.
There are sufcient similarities between the data fromthe three
study areas to be able to generalize the observations in the
following way:
1. Fault (or CBD) thickness is often equal to (with a thick-
nessedisplacement ratio approximately equal to 1) or smaller
than displacement (with a thicknessedisplacement ratio
between 1 and 0.1), with a general linear relationship despite
this variation.
2. For the larger structures (d >w100e500 mm), a proportional
increase between thickness and displacement does not clearly
exist.
3. Where thickness data were also collected using method 2 or
effective thickness data, these thicknesses recalculated as the
sumof the individual strands showa steady general decrease in
ratio of effective thickness of deformed material to
displacement as displacement increases. In other words, as
displacement increases, additional increments of displacement
result in decreasing amounts of added volume of deformed
material.
4. Larger normal faults (throw>1 m) and CDBs/CDB zones are
plotted on the same graph for comparison. These faults fall in
a different class of structures and they show different micro-
structure and deformation mechanisms. These ultracataclastic
fault zones have much lower thicknessedisplacement ratios
than CDBs and CDB zones.
5. Where clay layers are intersected, the structures have thick-
nessedisplacement ratios at the lower end of the range of data.
6. Where structures were formed in different tectonic events at
the same outcrop, no signicant differences are detectable
between data corresponding to the different events.
4.2. Fault thickness distribution data
Fig. 13 shows the fault thickness distribution data for all of the
outcrops studied in the Bassin du Sud-Est. For each different study
area, all of the structures recorded along the scan-lines are pre-
sented. Data are not divided into kinematic groups, because kine-
matics was not discernable in many cases, and so leaving such data
out would generate articial trends, particularly in the exclusion of
the single CDBs and thinner CDB zones for which offsets were most
difcult to discern with the naked eye. The three different study
areas exhibit the same general pattern, that there is a predomi-
nance of thin CDB zones of small displacement. For example, only
20% of the structures have thickness greater than around 1 cm
(0.6e2 cm depending on the study area). There is a much lower
proportion of larger (wider) CDB zones and ultracataclastic fault
zones than small CDB zones. Differences in the fault thickness
distribution data between the three study areas relates closely to
the initial host sand grain size: the larger host sand grain size has
thicker structures, and the lower host sand grain size has thinner
ones (Fig. 13).
4.3. CDB, cluster and fault zone internal structure
Fig. 14 presents information on the internal structure of CDB
zones in terms of the number of individual cataclastic bands
identiable to the naked eye within a CDB zone in relation to its
thickness. These data are presented separately for the three
different study areas. The data corresponding to the Orange area
have been recorded only using thickness measurement method 1
(Fig. 14a), i.e. the overall width of the CDB zone fromone edge to the
other. Method 2 was not used because it was not possible to
precisely distinguish the limits of the zones of cemented host grains
from the white strands of the deformation band structure with the
naked eye. It was, however, possible to detect strands as whiter
zones and tentatively count them. Data from the Massif d'Uchaux
and Bdoin areas have been presented using both method 1
Fig. 9. Graphs of cluster factor versus bin size using the denition of Wibberley et al.
(2007) adapted for analysing the size of the most likely periodic clustering by
repeating the cluster factor analysis for a range of different bin sizes, displacing the bin
along the scan-line at 1 m intervals. Results are theoretically independent of sample
size, hence the result that each conjugate set has a cluster factor twice as high as the
ensemble of the conjugate fault population (peaking at around 50e60 m bin size)
suggests that the more even distribution of CDBs in the overall population is the result
of summing the more clustered distribution of peaks and troughs of the two conjugate
sets together.
E. Saillet, C.A.J. Wibberley / Journal of Structural Geology 32 (2010) 1590e1608 1599
(Fig. 14b(i) and c(i)) and method 2 (Fig. 14b(ii) and c(ii)) thickness
measurements. Thicknesses using method 2 are given as the sumof
the thicknesses of each of the individual bands constituting
a multiple-band CDB zone or individual high-strain zones in an
ultracataclastic fault zone.
The data on the cataclastic bands in the CDB zones corre-
sponding to the three different study areas show the same overall
patterns of increase in the number of strands with increase in
thickness. In general, the CDB zones show a linear dependence of
the number of bands with increase in thickness, suggesting that
CDB zones develop by formation of additional bands successively at
the edges of the structure as deformation continues. Multiple-band
CDB zones with identiable undeformed sand in between the
bands are typically recorded as thicker structures for a given
displacement than those CDB zones with no host sand within the
structure (Fig. 14b(i), c(i)) but when the effective thickness of fault
rock (thickness method 2) is plotted, it seems clearer that the
thickness of fault rock increases as a function of displacement in the
same way for the two types of structures: the best-t relation for
the effective thickness for both types of structure together (Fig. 14b
(ii), c(ii)), being very close to the best-t relation of overall
thickness (method 1) for the CDB zones with no host sand within
them. For the data from Bdoin (Fig. 14c), the CDB zones thicker
than around four centimetres a different trend is apparent:
a disproportionate thickness increase with addition of new bands
(Fig. 14b(i), c(i)) collapses onto a single linear trend when the data
are re-plotted in terms of effective thickness of deformed fault rock
(Fig. 14b(ii), c(ii)). This indicates that CDB zones growing wider than
around 50 mm do so by addition of successive bands further and
further away into the host sand until they are considered as sepa-
rate structures entirely.
The different eld observations show an evidence for a chro-
nology in the development of isolated CDBs, CDB clusters and slip
surfaces or narrow slip zones. Generally, the observations show
that clusters are cut by the slip surfaces and narrow slip zones
(Fig. 6b). The different scan-lines (Figs. 10 and 11) also show that
slip surfaces and narrow ultracataclastic fault zones can exist in the
middle or at the edge of clusters of CDB zones, but they never exist
without clusters. On the other hand, many clusters exist with no
associated slip surface or narrow ultracataclastic zone. These
observations suggest that the clusters of CDB zones formed rst,
some of which then went on to localise the deformation by
Fig. 10. Graphs of CDB densities for the normal and strike-slip structures recorded along scan-lines with a total length of 209 m in the Boncavai quarry in the Massif d'Uchaux area.
Density is dened as the number of CDBs per square metre. The CDB density overall is moderate, with localisation of deformation within clusters of deformation bands. For each
outcrop, n is the total number of deformed structures and D
av
is the mean average density.
E. Saillet, C.A.J. Wibberley / Journal of Structural Geology 32 (2010) 1590e1608 1600
generation of discrete slip planes or narrow slip zones of highly
deformed ultracataclasite.
5. Discussion
Spatial distributions and geometric properties of CDBs and
larger faults were measured in three different areas of the Bassin du
Sud-Est along scan-lines of totalling 717 m in length. Systematic
recording of the deformation structures has allowed us to better
understand the mechanisms which control the distribution of
deformation and fault growth in high-porosity sandstones.
5.1. Development of individual CDBs
The systematic recording of the deformation features along the
different outcrops allow us to characterise the deformation in the
high-porosity sands and sandstones and to propose a conceptual
model for CDB growth (Fig. 15). First, our results showa predominance
Fig. 11. Graph of CDB densities for the reverse, normal and strike-slip structures recorded along scan-lines with a total length of 250 m in the Le Cros quarry in the Bdoin area.
Density is dened as the number of CDBs per square metre. As for the Massif d'Uchaux case, the CDB density overall is moderate but variable, with localisation of deformation within
clusters of deformation bands. For each outcrop, n is the total number of deformed structures and D
av
is the mean average density.
E. Saillet, C.A.J. Wibberley / Journal of Structural Geology 32 (2010) 1590e1608 1601
of thinner CDB zones, with 80% of all CDB zones having a thickness
less than around a centimetre, as discussed in the preceding section
(Fig. 13). The frequency distribution of thickness/displacement,
skewed towards the narrower structures, is signicant in consid-
ering fault growth. The thinner features grow during continued
displacement with a proportional increase in thickness up to
around 5 cm, by an increase of the number of individual bands
which constitute a CDB zone (Stages 1 and 2 in Fig. 15a and b).
Additional bands become spaced further and further away fromthe
CDB zone, so that beyond around 100 mm thickness, later strands
are far enough away to be counted as separate structures (see
denition in the section on data acquisition). Nevertheless, the eld
data presented (Figs. 12 and 14) suggest an upper limit of CDB zone
thickness of around 0.1e0.2 m, beyond which deformation jumps
to a new band entirely.
There is, on average, a difference in the thicknessedisplacement
ratio between the smaller CDB zones and the larger faults, for all
those structures with displacements above around 0.5 m. This
difference is clearest for the data from Orange (Fig. 12a). For the
smaller features (d <200 mm), there is a proportional increase in
thickness with displacement, albeit with a scatter in thick-
nessedisplacement ratios between 0.1 and 2. The larger structures
(d >0.5 m) showthickness-to-displacement ratios less than 0.2, i.e.
at or below the thicknessedisplacement ratio range of the smaller-
displacement structures (Fig. 12), suggesting that the growth
process does not operate as efciently at larger displacements.
Indeed, the development of ultracataclastic fault rock textures in
the larger-displacement structures gives the impression of even
more localised deformation. Studies showthat ultracataclastic fault
core zones in high-porosity sandstones elsewhere also have much
narrower thickness-to-displacement ratios than associated CDBs
(e.g. Shipton et al., 2005). Where clay is involved, thick-
nessedisplacement faulted clay values have, on average, lower
ratios than the structures without clay, demonstrating that the
Fig. 12. Thicknessedisplacement relationships for CDBs, CDB zones and larger faults in the high-porosity Cretaceous sands and sandstones of Provence. (a) Data from the Orange
area; (b) data from the Massif d'Uchaux area; (c) data from the Bdoin area. Data from the Massif d'Uchaux and Bdoin areas have been recorded based on two different
measurement methods: (i) thickness dened as the distance between the outside edges of the outermost strands (method 1); (ii) thickness dened as the sum of individual strand
thickness (method 2). The cohesion of the host rock seems to have a very important role in the thicknessedisplacement relationships (difference of one order of magnitude between
sands and sandstones), and hence on the micro-mechanisms of deformation.
E. Saillet, C.A.J. Wibberley / Journal of Structural Geology 32 (2010) 1590e1608 1602
incorporation of even small amounts of clay into the deformation
bands may preferentially localise deformation or at least partially
inhibit wall rock wear and deformation band widening.
The observation that CDB zones grow by the addition of new
peripheral bands instead of continued deformation on the existing
ones may, at rst sight, be taken as evidence of work hardening (e.g.
Underhill and Woodcock, 1987). Nevertheless, stress perturbations
around the CDB zone which encourage localisation of the new
bands adjacent to the CDB zone, coupled with intragranular
microcracking around the CDB zone as it develops (Wibberley et al.,
2007), suggest that other factors may also play a role, particularly as
laboratory experiments do not showa strength increase of samples
as individual strands form in a developing CDB zone (Mair et al.,
2000). The data corresponding to the larger CDB zones (with
a thickness more than w10 cm) show a diminishing thickness
growth as displacement progressively increases, by a decreasing
role played by the addition of new bands to the CDB zone structure.
Thus the data fromthe different sand and sandstone outcrops show
two different growth mechanisms, which are scale dependent. The
thinner CDB zones grow by systematic addition of new bands,
according to a work-hardening process and/or by stress perturba-
tions encouraging adjacent intragranular fracturing. For CDB zones
wider than around 10 cm however, the data suggest that these
work hardening and/or other processes cease to dominate. The
reason for such a transition may be one or more of the following
possibilities:
1. The larger CDB zones are typically longer e their development
may be limited by bed thickness and mechanical properties of
adjacent beds (e.g. Schultz and Fossen, 2002) as predicted by
numerical modelling of instability development during multi-
layer deformation (e.g. Schueller et al., 2005). However, beds
are often thicker than the height of the outcrop (e.g. several
metres to several tens of metres), making this difcult to
assess.
2. Stress perturbations may vary with CDB zone thickness, length
and displacement gradient. It is possible that wider, longer CDB
zones will have wider zones of stress perturbation, so addi-
tional bands could be generated further and further away from
the original CDB zone to the point where they are no longer
recorded as part of the same single (multistrand) CDB zone.
3. Grain size and grain point contact geometry could be control-
ling factors (e.g. Cundall, 1989), particularly given that smaller
average host sand grain size seems to cause narrower CDBs
(Fig. 13).
5.2. Development of CDB networks
The three study areas show deformation patterns which can be
related to one or more different tectonic events. For the sandstones
from the Orange area, the structures are mostly conjugate reverse
CDB faults corresponding to PyreneaneProvenal shortening. Here,
only a few later, normal faults are present, related to Oligoce-
neeMiocene extension. Evidence of deformation structures corre-
sponding essentially to one single tectonic event at this outcrop is
highly signicant in terms of allowing us to make interpretations of
how this network of structures developed during the single
tectonic event. In terms of the spatial distribution of the CDBs
documented from the scan-line data (Fig. 8), two main interpre-
tations can be made:
1. Firstly, the distribution of the deformation along the outcrop is
relatively homogeneous with a moderately high density of
CDBs and an addition of high density of deformation which
corresponds to a localisation of the deformation in a few
clusters of CDBs such as ladder zones. Such a persistent
density distribution over a transect of c. 250 m is important
because it suggests an overall relatively homogeneous bulk
strain (shortening) accommodated by these CDBs over a rela-
tively large distance, in comparison to previous studies which
mostly suggest CDBs are formed as clusters of CDBs around
larger faults, such as damage zones developed by progressive
displacement along the fault (e.g. Shipton and Cowie, 2001; Du
Bernard et al., 2002b; Johansen and Fossen, 2008).
2. Secondly, there is a general inverse relationship between the
undulating density of one of the conjugate sets and the density
of the other (Fig. 8), with alternating density peaks and troughs
of one set mirroring the troughs and peaks of the conjugate set.
The sum of the two conjugate distributions is a much more
continuous deformation distribution than the clustered
tendency of each of the two conjugate sets individually (Fig. 9).
These eld observations allow us to suggest that deformation
proceeded by development of two conjugate sets of CDBs, of
opposing dips, during the same tectonic event. At the begin-
ning of the tectonic event the deformation is accommodated by
a random distribution of each conjugate set. As deformation
proceeds, further generation of the CDBs of one set is inhibited
in regions where earlier CDBs of the other set are already
present (Fig. 15a). The areas not deformed early on in the
process may experience a higher number of CDBs generated
later. The result of this growth process is that regions with
a high density of faults fromone set have a lower density in the
other, and vice-versa. This CDB growth process by competition
between two opposing conjugate sets during a single tectonic
event results in a broadly homogeneous total (albeit undu-
lating) deformation distribution along the outcrop.
Previous studies have suggested that deformation bands form
by work-hardening processes (e.g. Aydin and Johnson, 1983;
Underhill and Woodcock, 1987). Furthermore, work-hardening
processes at fault tips and relay zones have been evidenced by the
generation of ladder zones between segment tips (e.g. Schultz and
Balasko, 2003; Okubo and Schultz, 2005, 2006). However, our eld
observations show that the work-hardening nature inferred for
CDB generation may inuence the population at the network scale
as well as at the scale of the fault tip, thus causing an increase in the
mechanical strength of the rock mass as well as the individual
deformation band structure. Hence, it is easier for a conjugate set to
develop in a region with low density of earlier deformation struc-
tures. Although overall the two sets are generally synchronous, as
shown by the mutual cross-cutting relationships in different places
Fig. 13. Frequency distribution of CDB zone and fault zone thicknesses for cataclastic
deformation bands and larger faults in the three different study areas of Provence. The
thickness is dened as the distance between the outside edges of the outermost bands
(method 1). The average grain size of the host rock is given for each outcrop by the
symbol f.
E. Saillet, C.A.J. Wibberley / Journal of Structural Geology 32 (2010) 1590e1608 1603
(Fig. 5), the preferential development of one set in a certain region
probably depends on it arriving in that region by chance before the
other set. Despite this seeming randomness, it is nevertheless
possible that the spacing of peaks and troughs in the densities of
each conjugate set, in this case around 60 m along the scan-line, so
around 30 m perpendicular to any one set, is related to mechanical
unit thickness, although the base of the unit is below ground level
so we can not assess this.
It was not possible to test this model for conjugate CDB devel-
opment on the other study areas because they have recorded
deformation from two or three superimposed tectonic events, and
in many cases it is not possible to attribute an individual CDB to
a particular tectonic event. Because of this, the statistical clustering
analysis performed on the Orange data (Section 3.2.1) was not
performed on the other data sets as it would not have been
meaningful. However, all the Massif d'Uchaux outcrops expressed
similar patterns. For each outcrop there is a generally moderate
density of deformation structures with the addition of clusters of
CDB zones which localised the deformation (Fig. 10). The same
deformation patterns exist also in the outcrops fromthe incohesive
sands of the Bdoin area (Fig. 11). In these cases, the majority of the
CDBs are normal-sense structures generated during extension. It
therefore seems likely that the difference in clustering tendency
between these outcrops and the distributed nature of the defor-
mation at Orange is simply due to kinematic type of deformation,
with shortening encouraging further generation of CDBs
throughout the volume by system hardening during compression
(Fig. 15a). However, other factors such as differences in burial depth
Fig. 14. Relationship between number of bands and thickness for CDB zones and larger faults in the high-porosity Cretaceous sands and sandstones of Provence. (a) Data from the
Orange area; (b) data from the Massif d'Uchaux area; (c) data from the Bdoin area. Data from the Massif d'Uchaux and Bdoin areas have been recorded based on two different
measurement methods: (i) thickness dened as the distance between the outside edges of the outermost bands (method 1); (ii) thickness dened as the sum of individual band
thickness (method 2). The linear regressions presented on these graphs are the lines of best-t, most of which are constrained to pass through the origin; the equations are given for
each one. For the graphs (c), the second linear regression corresponds to the lines of best-t of the largest structures, which are not constrained to pass through the origin. The
regression in (c ii) is given as a dotted line because of the small number of these structures.
E. Saillet, C.A.J. Wibberley / Journal of Structural Geology 32 (2010) 1590e1608 1604
at the time of deformation, which is difcult to constrain here, may
also play a role. It seems likely that differences in tectonic loading
path have had an effect on the spatial distribution of the resulting
deformation.
5.3. Development of higher-displacement mature faults
and slip planes
Previous qualitative and semi-quantitative work on some of the
outcrops presented here found that larger ultracataclastic fault
zones and localised slip surfaces often formed where there was
a high density of CDBs from a previous tectonic event (Wibberley
et al., 2007). The new work expands on the previous study by the
thorough collection of statistical data from scan-lines, such as
spatial variations in density, as well as studying additional
outcrops. It has been found that larger-displacement fault zones
and localised slip surfaces can also form in the same tectonic event
as the CDBs, but in this case they exist within, or at the edge of
clusters of CDB zones. The precise relative chronology of CDBs and
larger faults is difcult to discern at outcrop in most cases, and it is
possible to suggest two different hypotheses for their genetic
relationship: either (a) all the background CDBs formed rst, with
a relatively constant density across the outcrop, followed by
localisation into the clusters; or (b) the CDBs initiated in a few
places, and continued deformation allowed both further develop-
ment into clusters and spreading of the deformation elsewhere to
form the other individual CDBs further away from the clusters.
The fact that all larger-displacement faults occur within regions
of high CDB density, but that many high-density clusters exist
without larger faults, suggests that the clusters formed rst, only
some of which suffered continued deformation localisation and
generation of ultracataclastic faults and discrete slip surfaces
(Fig. 15b, stage 4). In other words, a larger fault is not a prerequisite
for generating peripheral clusters of CDBs (as in a damage zone)
in the study areas examined, but these larger faults localise on pre-
existing cluster zones to accommodate further strain. In fact, in
extensional contexts such as those which generated the structures
studies in the Urchaux and Bdoin areas, it is much easier to form
a slip zone. In the case of reverse deformation, it seems that it is
much more difcult to obtain a localised slip plane or narrow slip
zone e the example of deformation bands accommodating short-
ening at Orange never achieved a state of localisation to evolve into
a large structure; deformation is likely to be transferred into
a different lithological layer before such a high deviatoric stress is
achieved.
The progressive localisation of deformation through time during
extension, from an undulating background CDB density pattern
and the clustering of CDBs through to formation of larger-
displacement ultracataclastic fault zones and initiation of discrete
slip surfaces may therefore be at least partly a function of the
extensional tectonic regime and associated loading path. It is often
tempting to associate such deformation localisation with material
softening of the slip zones, yet in reality localisation is also
controlled by other factors such as system geometric constraints
Fig. 15. Series of block diagrams showing the sequential evolution of deformation from a single band to different sets in a network, with possible localised faulting, during a single
tectonic event. (a) A compressive tectonic event. (Stage 1) Random distribution of the deformation at the beginning of the tectonic event. (Stage 2) Evolution of single CDBs to larger
CDB zones. (Stage 3) Further deformation results in cross-cutting of CDB zones by conjugate structures, but the previous structures inhibit further development of the conjugate set
in places. (Stage 4) As deformation continues, the structural network evolves towards a broadly constant density of conjugate structures. (b) An extension tectonic event. (Stage 1)
Initiation of single CDBs at the beginning of the tectonic event. (Stage 2) Evolution of single CDBs to larger CDB zones. (Stage 3) Localisation of the deformation within some clusters
of CDB zones. (Stage 4) As deformation continues, a slip surface or narrow slip zone grows within a zone of deformation band clustering.
E. Saillet, C.A.J. Wibberley / Journal of Structural Geology 32 (2010) 1590e1608 1605
and is not simply a function of material softening (e.g. Hobbs et al.,
1990). In the case of extension, it has been suggested that elastic
unloading of the wall during normal faulting may inhibit the
continued generation of new deformation bands around the fault,
whilst the main fault zone continues to deform. As the faults grow
up-dip and down-dip they will be controlled by other system
parameters such as bed thickness and the mechanical contrast with
adjacent beds, so that eventually the mechanical properties of the
fault will be controlled by system constraints at a larger scale,
possibly with damage zone accumulation contributing to further
fault zone development.
The importance of loading path on deformation structural
evolution can be illustrated by QeP
0
maps (Fig. 16). These plot the
stress paths in terms of differential stress (Q) against effective mean
stress (P
0
) as commonly done in soil mechanics. On such plots, the
range of stress states at the onset of plastic deformation can be
plotted (a yield envelope), as can the elds in which deformation
occurs by compaction, dilation or the critical state line of constant-
volume deformation between the two. Plastic yield envelopes
continuously evolve as the sand is subjected to compaction and/or
tectonic deformation until they reach a clasticeplastic yield
envelope at which point cataclastic deformation bands are gener-
ated (Bolton and McDowell, 1997; Wibberley et al., 2007). For this
reason, several authors have recently used a framework of soil
mechanics to better understand the initiation of cataclastic defor-
mation bands, seen as a phenomenon of plastic deformation which
may be either compactant, dilatant or constant-volume (e.g.
Schultz and Siddharthan, 2005; Aydin et al., 2006; Wibberley et al.,
2007). Deformation accommodating much larger slip than typical
deformation bands, usually by discrete slip surfaces and/or ultra-
cataclastic fault zones, is considered to occur on the critical state
line of isovolumetric deformation. Although there is no a priori
connection between the critical state line (as a range of stress
states) and localisation (as a description of deformation distribu-
tion), microstructural studies of the ultracataclastic fault zones
suggest constant-volume granular ow operated during shearing,
supported by classical soil mechanics experiments after a given
shear strain (e.g. Mandl et al., 1977).
Fig. 16 illustrates four hypothetical loading paths on a QeP
0
map of differential stress (Q) against effective mean stress (P
0
), the
rst path (A) being tectonic compression, the other ones being
different cases of extension (BeD). In all cases, cataclastic defor-
mation bands initiate at the clasticeplastic yield envelope (stage
i), in most cases on the right hand side (cap side) of the Q peak
in the yield envelope, suggesting compaction. The strength of the
CDB itself denes a new yield envelope, being stronger than the
host sand. In the case of shortening, CDBs continue to be gener-
ated until they start to cross each other, at which point network
hardening occurs and the stress path moves up towards the CDB
yield envelope (stage ii on stress path A). A critical density is
achieved at which the bulk stress state lies in fact on the CDB yield
envelope (stage iii). Further increase in stress is required to
continue deformation, but it is uncertain whether this can
continue to the critical state line (Fig. 16) e it is more likely that
failure of an adjacent bed will lead to a thrust fault propagating
into the porous sand bed.
In extension, on the other hand, the exact sequence and type of
structures generated is suggested to depend on the starting
position of the effective tension loading with respect to the clas-
ticeplastic and CDB yield envelopes. In loading path B (Fig. 16), the
path continues to increase differential stress and hence generate
CDBs until, like path A, a critical density is reached (network
hardening). Beyond this, differential stress continues to increase
until the critical state line of the (bulk) deformation band material
is reached at which point faulting occurs by ultracataclasic in
a zone with granular ow, and/or by forming discrete slip surfaces.
Field evidence indicates that slip surfaces in this context localise
within clusters of CDB zones or at the competence contrast
between CDB clusters and host sand. In loading path C (Fig. 16),
the critical state line for the sand is reached as the stress path
Fig. 16. A QeP
0
map of generation of deformation structures in high-porosity sand. CSL indicates Critical State Line. Four different hypothetical stress paths are indicated (labelled
AeD), each producing a different range of structures. See text for details.
E. Saillet, C.A.J. Wibberley / Journal of Structural Geology 32 (2010) 1590e1608 1606
continues to move through the compactant regime during
network hardening before the CDB yield envelope is reached,
suggesting that faulting will occur in the sand by localised ultra-
cataclasis. Finally, in loading path D (Fig. 16), the CDB initiation
occurred at greater differential stress than the critical state line
(and on the dilatant side to the left of the yield envelope peak),
suggesting that further deformation will quickly result in defor-
mation localisation in the sand on discrete slip surfaces along with
unloading. This softening effect may be enhanced by non-associ-
ated Coulomb plasticity softening (Mandl, 1988) as described by
Wibberley et al. (2007).
6. Conclusions
The statistical analyses of the spatial distributions and
geometrical properties of cataclastic deformation bands and larger
ultracataclastic fault zones in Cretaceous high-porosity sands and
sandstones from Provence, SouthEast France, allow us to evaluate
the controls on deformation distribution and fault growth mecha-
nisms. The main results can be summarized as follows:
1. CDB density distributions along the Orange outcrop suggest
that deformation was accommodated by two conjugate sets,
each of heterogeneous distribution, during the same tectonic
event. In any one place, where there is a higher density of faults
from the rst set, there is a lower density in the second, and
where there is a lower density of faults from the rst set, there
is a higher density in the second. The result of this fault growth
process is a broadly homogeneous total deformation along the
outcrops which developed by a system-hardening mechanism
during shortening. The frequency of the density undulations of
each of the two conjugate sets may be related to mechanical
bed thickness.
2. The Massif d'Uchaux and Bdoin outcrops have moderate,
undulating densities of mainly normal-sense CDBs associated
with extension. Some zones are present with anomalously
high CDB densities, where CDB zones are organized into
clusters, some of which contain larger-displacement ultra-
cataclastic fault zones or discrete slip surfaces. This distribu-
tion differs from the reverse-sense CDBs at Orange. Thus for
the rst deformation event to strongly affect the outcrop, the
clustering tendency is dependant on kinematics and therefore
tectonic loading path. Thickness growth of CDB zones,
however, is independent of these, but dependent on host grain
size with thicker CDB zones forming in coarser sand/
sandstone.
3. CDB zones grow by a proportional increase of thickness with
displacement due to the addition of new individual bands
within the zone, corresponding to a work-hardening process
at the scale of the zone. As thickness achieves a certain
value on the order of w10 cm, the generation of additional
bands does not continue in proportion to thickness growth.
In these larger CDB zones, the new outer bands are gener-
ated further and further into the host sand away from the
zone. Hardening processes cease to dominate for these
larger CDB zones, and further deformation occurs with
clustering of CDB zones followed in some cases by genera-
tion of the larger-displacement ultracataclastic faults and
discrete slip surfaces which grow with increasing local-
isation of deformation.
Acknowledgements
The rst author was supported by a doctorate student MRT
stipend from the French Ministry for Higher Education and
Research. Access to the active quarries was granted by the oper-
ating company SIFRACO who is gratefully acknowledged. Thorough
reviews by Geoffrey Rawling and Richard Schultz helped clarify
much of the text, and we are also grateful to Zoe Shipton for
editorial guidance.
References
Antonelini, M., Aydin, A., 1994. Effect of faulting on uid ow in porous sandstones:
petrophysical properties. American Association of Petroleum Geologists
Bulletin 78, 335e377.
Antonelini, M., Aydin, A., 1995. The effect of faulting on uid ow in porous
sandstones: geometry and spatial distribution. American Association of Petro-
leum Geologists Bulletin 79, 642e671.
Antonellini, M.A.A., Aydin, A., Pollard, D.D., 1994. Microstructure of deformation
bands in porous sandstones at Arches National Park, Utah. Journal of Structural
Geology 16, 941e959.
Aydin, A., 1978. Small faults formed as deformation bands in sandstones. Pure and
Applied Geophysics 116, 913e930.
Aydin, A., Johnson, D., 1978. Development of faults as zones of deformation bands
and as slip surfaces in sandstone. Pure and Applied Geophysics 116, 931e942.
Aydin, A., Johnson, D., 1983. Analysis of faulting in porous sandstones. Journal of
Structural Geology 5, 19e31.
Aydin, A., Borja, R.I., Eichhubl, P., 2006. Geological and mathematical framework for
failure modes in granular rock. Journal of Structural Geology 28, 83e98.
Bolton, M.D., McDowell, G.R., 1997. Clastic mechanics. In: Fleck, N.A., Cocks, A.C.F.
(Eds.), IUTAM Symposium on Mechanics of Granular and Porous Materials.
Kluwer, Dordrecht, pp. 35e46.
Cundall, P.A., 1989. Numerical experiments on localization in frictional materials.
Ingenieur-Archiv 59, 148e159.
Davis, G.H., 1999. Structural geology of the Colorado Plateau regional of southern
Utah, with special emphasis on deformation bands. Geological Society of
America Special Paper 342.
Debrand-Passard, S., Courbouleix, S., Lienhardt, M.J., 1984. Synthse gologique du
Sud-Est de la France. Mmoire du Bureau de Recherches Geologiques et Mini-
res, (BRGM), 125, Orleans, France.
Delfaud, J., Dubois, P., 1984. Le basin du Sud-Est. In: Gubler, Y. (Ed.), Dynamique des
bassins sdimentaires. Livre jubilaire BRGM, Orleans, France.
Du Bernard, X., Eichhbul, P., Aydin, A., 2002a. Dilation bands: a new form of
localized failure in granular media. Geophysical Research Letters 29, 2176.
doi:10.1029/2002GL015966.
Du Bernard, X., Labaume, P., Darcel, C., Davy, P., Bour, O., 2002b. Cataclastic slip band
distribution in normal fault damage zones, Nubian sandstones, Suez rift. Journal
of Geophysical Research 107, 103e114.
Evans, J.P., 1990. Thicknessedisplacement relationships for fault zones. Journal of
Structural Geology 12, 1061e1065.
Fossen, H., Odinsen, T., Frseth, R.B., Gabrielsen, R.H., 2000. Detachments and low-
angle faults in the northern North Sea rift system. In: Nttvedt, A. (Ed.),
Dynamics of the Norwegian margins. Special Publications, Geological Society,
London, vol. 167, pp. 105e131.
Fossen, H., Schultz, R.A., Shipton, Z.K., Mair, K., 2007. Deformation bands in sand-
stone: a review. Journal of the Geological Society of London 164, 755e769.
Fowles, J., Burley, S., 1994. Textural and permeability characteristics of faulted, high
porosity sandstones. Marine and Petroleum Geology 11, 608e623.
Gallagher Jr., J.J., Friedman, M., Handin, J., Sowers, G.M., 1974. Experimental studies
relating to microfracture in sandstone. Tectonophysics 21, 203e247.
Hobbs, B.E., Mhlhaus, H.-B., Ord, A., 1990. Instability, softening and localization of
deformation. In: Knipe, R.J., Rutter, E.H. (Eds.), DeformationMechanisms, Rheology
and Tectonics. Geological Society Special Publications, vol. 54, pp. 143e165.
Hull, J., 1988. Thicknessedisplacement relationships for fault zones. Journal of
Structural Geology 10, 431e435.
Jamison, W.R., Stearns, D.W., 1982. Tectonic deformation of Wingate Sandstone,
Colorado National Monument. American Association of Petroleum Geologists
Bulletin 66, 2584e2608.
Johansen, T.E.S., Fossen, H., 2008. Internal geometry of fault damage zones in
interbedded siliclastic sediments. In: Wibberley, C.A.J., Kurz, W., Imber, J.,
Holdsworth, R.E., Collettini, C. (Eds.), The Internal Structure of Fault Zones:
Implications for Mechanical and Fluid-Flow Properties. Geological Society
Special Publication, vol. 299, pp. 34e55.
Mair, K., Main, I., Elphick, S., 2000. Sequential growth of deformation bands in the
laboratory. Journal of Structural Geology 22, 25e42.
Mandl, G., 1988. Mechanics and Tectonics of Faulting: Models and Basic Concepts.
Elsevier, Amsterdam.
Mandl, G., de Jong, L.N.J., Maltha, A., 1977. Shear zones in granular material. Rock
Mechanics 9, 95e144.
Manzocchi, T., Ringrose, P.S., Underhill, J.R., 1998. Flow through fault systems in
high-porosity sandstones. In: Coward, M.P., Daltaban, T.S., Johnson, H. (Eds.),
Structural Geology in Reservoir Characterisation. Geological Society, London,
Special Publications, vol. 127, pp. 65e82.
Matthi, S.K., Aydin, A., Pollard, D.D., Roberts, S.G., 1998. Numerical simulation of
departures from radial drawn in faulted sandstone reservoirs with joints and
deformation bands. In: Jones, G., Fisher, Q., Knipe, R.J. (Eds.), Faulting, Fault
E. Saillet, C.A.J. Wibberley / Journal of Structural Geology 32 (2010) 1590e1608 1607
Sealing and Fluid Flow in Hydrocarbon Reservoirs. Geological Society, London,
Special Publications, vol. 147, pp. 157e192.
Okubo, C.H., Schultz, R.A., 2005. Evolution of damage zone geometry and intensity
in porous sandstone: insight gained from strain energy density. Journal of the
Geological Society of London 162, 939e950.
Okubo, C.H., Schultz, R.A., 2006. Near-tip stress rotation and the development of
deformation band stepover geometries in mode II. Geological Society of
America Bulletin 118, 343e348.
Priest, S.D., Hudson, J.A., 1981. Estimation of discontinuity spacing and tracing
length using scanline survey. International Journal of Rock Mechanics and
Mining Sciences and Geomechanics Abstracts 18, 183e197.
Robertson, E.C., 1983. Relationship of fault displacement to gouge and breccia
thickness. Mining Engineering 35, 1426e1432.
Saillet, E., 2009. La localisation de la dformation dans les grs poreux: car-
actrisation d'un analogue de rservoir faill dans le Bassin du Sud-Est,
Provence, France. Doctorate thesis, University of Nice-Sophia Antipolis,
France, 273 p.
Shipton, Z.K., Cowie, P.A., 2001. Damage zone and slip-surface evolution over mm to
km scales in high-porosity Navajo sandstone, Utah. Journal of Structural
Geology 23, 1825e1844.
Shipton, Z.K., Evans, J.P., Thompson, L.B., 2005. The geometry and thickness of
deformation-band fault core and its inuence on sealing characteristics of
deformation-band fault zones. In: Sorkhabi, R., Tsuji, Y. (Eds.), Faults, Fluid Flow
and Petroleum Traps. American Association of Petroleum Geologists Memoir,
vol. 85, pp. 181e195.
Scholz, C.H., 1987. Wear and gouge formation in brittle faulting. Geology 15,
493e495.
Schultz, R.A., Fossen, H., 2002. Displacementelength scaling in three dimensions:
the importance of aspect ratio and application to deformation bands. Journal of
Structural Geology 24, 1389e1411.
Schultz, R.A., Balasko, C.M., 2003. Growth of deformation bands into echelon and
ladder geometries. Geophysical Research Letters 30, 2033. doi:10.1029/
2003GL018449.
Schultz, R.A., Siddharthan, R., 2005. A general framework for the occurrence and
faulting of deformationbands inporous granular rocks. Tectonophysics 411, 1e18.
Schueller, S., Gueydan, F., Davy, P., 2005. Brittleeductile coupling: role of ductile
viscosity on brittle fracturing. Geophysical Research Letters 32, L10308.
doi:10.1029/2004GL022272.
Torabi, A., Braathen, A., Cuisiat, F., Fossen, H., 2007. Shear zones in porous sand:
insights from ring-shear experiments and naturally deformed sandstones.
Tectonophysics 437, 37e50.
Underhill, J.R., Woodcock, N.H., 1987. Faulting mechanisms in high porosity sand-
stones; New Red Sandstone, Arran, Scotland. In: Jones, M.E., Preston, R.M.F.
(Eds.), Deformation of Sediments and Sedimentary Rocks. Geological Society,
London, Special Publications, vol. 29, pp. 91e105.
Wibberley, C.A.J., Petit, J.P., Rives, T., 2000. Mechanics of cataclastic deformation
band faulting in high-porosity sandstone, Provence. Comptes Rendus de
l'Academie des Sciences, Paris 331, 419e425.
Wibberley, C.A.J., Petit, J.P., Rives, T., 2007. The mechanics of fault distribution and
localization in high-porosity sands, Provence, France. In: Lewis, H., Couples, G.D.
(Eds.), The Relationship between Damage and Localisation. Geological Society,
London, Special Publication, vol. 289, pp. 19e46.
Wibberley, C.A.J., Yielding, G., DiToro, G., 2008. Recent advances in the under-
standing of fault zone internal structure: a review. In: Wibberley, C.A.J.,
Kurz, W., Imber, J., Holdsworth, R.E., Collettini, C. (Eds.), The Internal Structure
of Fault Zones: Implications for Mechanical and Fluid-Flow Properties.
Geological Society Special Publication, vol. 299, pp. 5e33.
Wong, T.-F., David, C., Zhu, W., 1997. The transition frombrittle faulting to cataclastic
ow in porous sandstones: mechanical deformation. Journal of Geophysical
Research 102, 3009e3025.
E. Saillet, C.A.J. Wibberley / Journal of Structural Geology 32 (2010) 1590e1608 1608
Extensional faults in ne grained carbonates e analysis of fault core lithology
and thicknessedisplacement relationships
Eivind Bastesen
a, b,
*
, Alvar Braathen
b, c
a
Centre for Integrated Petroleum Research, University of Bergen, 5020 Bergen, Norway
b
Department of Earth Science, University of Bergen, 5020 Bergen, Norway
c
University Centre in Svalbard, 9171 Longyearbyen, Norway
a r t i c l e i n f o
Article history:
Received 4 July 2009
Received in revised form
26 August 2010
Accepted 18 September 2010
Available online 25 September 2010
Keywords:
Thicknessedisplacement
Extensional faults
Carbonates
Fault core
Fault facies
a b s t r a c t
A study of 103 extensional faults hosted by ne grained carbonates in western Sinai, Svalbard and Oman
reveals that faults vary geometrically between simple cores and cores comprising fault splays, lenses,
segment linkages and overlap structures. Fault core rocks are typically carbonate breccias, carbonate and
shale gouge, shale smear, secondary calcite cement and veins, and host rock lenses.
There is a signicant scatter in the core thickness for any given displacement, but the overall pattern is
that the thickness increases with displacement. This increase best ts a power law function (0.29D
0.56
)
that describes a gradual decrease in the thickness/displacement relationship for increasing slip along
faults. In more detail, the general function can be seen as the sum of two (power law) trend lines; the
rst representing thin localized fault cores with generally simple and planar geometry, the second
representing thicker fault cores with complex geometry of lenses and overlap structures and with fault
rock membranes.
Thestudiedfaults showasignicant changeincompositionandgeometryfromsmall (0e1m), tomoderate
(1e10m) andtolarge offset faults (10e400m). Theoverall patternis that fault initiates as fractures lledwith
calcite veins and thin shear fractures that hosts gouge membranes. With increased fault offset, complexity
increases with breakdownof veins, more extensive fault rock membranes, anda trend towards development
of lenses. Whenoffset exceeds 100 m, cores become complex, withmultiple slipzones, cementedbreccia and
shale smear membranes, and various types of lenses. We envision that the fault development as reected by
offset is dominated by forces (extension, compression) acting in the fault, mechanical heterogeneity of wall
rocks, the core lithologies and their developing rheology, and especially geometric effects arising from fault
irregularities.
2010 Elsevier Ltd. All rights reserved.
1. Introduction
Characterization and quantication of fault zones in outcrops is
a fundamental requirement for modelling and forecasting structural
reservoir heterogeneity in carbonate reservoirs. Key parameters for
fault characterization include fault thickness, composition, geom-
etry and displacement (e.g. Yielding et al., 1997; Manzocchi et al.,
1999; Braathen et al., 2009). A fault can be dened as a zone of
focused deformation that can be subdivided into domains/sub-
zones termed core and damage zone(s) (e.g. Chester and Logan,
1986). Alternatively, a fault can be considered an array of hard-
linked and soft-linked fault segments of various scales that affect
a restricted rock volume or fault envelope (e.g. Peacock, 2002;
Childs et al., 2009; Braathen et al., 2009). Descriptions of fault cores
(e.g. Caine et al., 1996; Childs et al., 1996; Lindanger et al., 2007;
Bonson et al., 2007; Wibberley et al., 2008; Bastesen et al., 2009;
Braathen et al., 2009) show a number of recurring elements such
as slip surfaces, fracture/deformation band sets, fault rocks (gouge,
breccias and cataclasites), shale smears, and lenses of protolith or
fault rock. Bulk strain of the core is semi-penetrative to penetrative,
and core elements in most cases exhibit signicantly altered uid
conductivity compared to the host rock fromwhich they are derived.
In contrast, bulk strain in the damage zones anking the core is non-
penetrative and hosts discrete structures including minor faults,
fractures and/or deformation band sets. Studies addressing the
width of the fault zone envelope vs. fault displacement have
revealed a substantial degree of variation and uncertainty (e.g. Hull,
1988; Knott, 1994; Shipton et al., 2006; Childs et al., 2009). In most
cases the thickness/displacement ratio (T/D) shows that thickness
* Corresponding author. Centre for Integrated Petroleum Research, Uni Research,
5020 Bergen, Norway. Tel.: 47 99230248.
E-mail address: Eivind.bastesen@Uni.no (E. Bastesen).
Contents lists available at ScienceDirect
Journal of Structural Geology
j ournal homepage: www. el sevi er. com/ l ocat e/ j sg
0191-8141/$ e see front matter 2010 Elsevier Ltd. All rights reserved.
doi:10.1016/j.jsg.2010.09.008
Journal of Structural Geology 32 (2010) 1609e1628
varies by up to three orders of magnitude for a given displacement,
reecting the geometric complexity arising from the presence of
linked and unlinked segments (Childs et al., 2009).
In this paper we present characteristics and scaling laws for
faults in ne grained carbonates. The dataset includes 103 faults
described using the fault facies characterization concept (Braathen
et al., 2009). Fault facies refers to any feature or rock body
deriving its properties from tectonic deformation (Tveranger et al.,
2005), and includes the main elements such as lenses, membranes
and fractures. Fault facies can be characterized in terms of dimen-
sions, geometry, internal structure and petrophysical properties,
thus facilitating quantication, pattern recognition and statistical
handling of structural elements in fault envelopes (Tveranger et al.,
2005; Braathen et al., 2009). In the present study we have dened
fault facies according to composition and geometry. Fault core
geometry and distribution of fault facies are analysed in relation to
the T/D ratio as established for each studied fault.
Fault architecture and related uid ow properties in carbonate
rocks have in the last years received increased attention (e.g.
Agosta and Kirschner, 2003; Cello et al., 2003; Micarelli et al., 2003;
Storti et al., 2003; Labaume et al., 2004; Agosta and Aydin, 2006;
Graham Wall et al., 2006; Bonson et al., 2007; Benedicto et al.,
2008; Bastesen et al., 2009; Putz-Perrier and Sanderson, 2010).
However, studies addressing scaling relationships of faults in such
rocks are scarce, and restricted to faults with small displacements
(Billi et al., 2003; Micarelli et al., 2006; Soliva and Benedicto, 2005),
or case studies (Micarelli et al., 2003; Agosta and Aydin, 2006;
Bonson et al., 2007; Bastesen et al., 2009). In this paper we
present a database of extensional faults from shallow buried
(<2 km) carbonates from three different regions: western Sinai
(Egypt), Central Spitsbergen, Svalbard (Arctic Norway), and Central
Oman (Adams Foothills). The bulk of the data were collected in
western Sinai, whereas data from Spitsbergen and Oman were
collected for comparison purposes (i.e. different tectonic regimes
and/or protoliths). All three areas exhibit thick successions of
sedimentary carbonates which are truncated by well exposed
extensional faults at different scales, with fault displacements
ranging from a fewcentimetres to several hundred metres. In Sinai
(e.g. Moustafa, 2004) and Spitsbergen (e.g. Steel and Worsley,
1984; Maher and Braathen, in press), faulting is related to
regional rifting events, whereas in Oman, extensional faults are
found along the crest of regional anticlines and domes, of which
the folding is controlled by deep-seated thrusting and salt move-
ments (e.g. Hanna, 1990). Protoliths range from massive homoge-
nous limestone to layered heterogeneous shale-rich carbonates
and marls.
2. Terminology
The studied fault cores exhibit several fault core facies associa-
tions (Braathen et al., 2009), which allow an identication and
a description of fault facies using lithology and fault core geometry
as descriptive parameters. Lithologically, fault core facies associa-
tions can be subdivided into shale smear (SS) (Lindsay et al., 1993;
Yielding et al., 1997), carbonate breccia (CB) (Billi, 2005; Micarelli
et al., 2003, 2006), secondary calcite (SCa) (Benedicto et al.,
2008), or composite cores; the latter displaying two or all three
fault core facies associations (Fig. 1).
The fault rocks observed in the present study formed at shallow
depths and at low temperature, and in most cases bear resem-
blance to the primary non-cohesive breccia series described by
Sibson (1977) and Braathen et al. (2004). Clast materials are frag-
ments of carbonate formed by brittle failure. The breccia matrix
consists of either nely crushed limestone fragments in a gouge
(mud fraction fault rock), ne breccia (Billi, 2005) exhibiting
varying degrees of cementation (CB), or shale; the latter giving rise
to shale supported breccias (SCB) (Fig. 1b,c). In this study gouge
was observed as very ne, crushed carbonate material and thin
(<1 mm) membranes of sheared calcareous clay along faults. Shale
smears are shale layers dragged into the fault, aligned parallel to
fault dip and connected to a source shale layer (Lindsay et al.,
1993). Secondary mineral precipitation was observed in the
shape of calcite veins and void llings and as pore space ll in
breccias (described above). In many places precipitated calcite
forms fault-parallel membranes and lenses, displaying the typical
crack-seal vein appearance advocated by Petit et al. (1999). In the
following descriptions the compositional elements are arranged
into fault facies such as lenses (Childs et al., 1997; Gabrielsen and
Clausen, 2001) and membranes (Braathen et al., 2009). Fault len-
ses are elongate pods of host rocks, fault rocks and/or calcite veins
which are completely separated from the surrounding fault
elements by slip surfaces with wall rocks (slip zones) (Fig. 1a).
Membranes are continuous, semi-continuous or patchy fault-
parallel sheets consisting of carbonate breccias, gouge, shale smear
or veins.
Overall fault geometries are classied in Fig. 2. The Type 1
geometry corresponds to faults with a simple geometry, which
further divides into planar geometry (1a) or fault cores that exhibit
jogs or bends of either releasing (1b) or restraining (1c) character.
Secondary shear fractures are classied according to their slip
direction relative to the main fault orientation (Petit, 1987; Bastesen
et al., 2009). These are termed R, P or R
0
shears and correspond to
Type 2 a, b and c, respectively. Type 3 geometries are soft-linked
releasing (3a) or restraining (3b) overlap structures (Rykkelid and
Fossen, 2002; Ferrill and Morris, 2003), while Type 4 geometries
represent breached relays. Fault lenses are assigned to the Type 5.
The Type 6 geometry includes faults with multiple slip surface
zones, several lenses and complex intrinsic composition.
3. Methods
3.1. Database
The database analysed in this study includes a total of 103
extensional faults: 68 from western Sinai, 22 from Oman and 13
from Central Spitsbergen. Fault displacements range from2.6 cmto
400 m, and fault core thicknesses from 1 mm to 11 m. For the
purpose of the present study, faults are termed small when the
displacement is smaller than 1 m, moderate when the displace-
ment is in the range of 1e10 m, and large for the range of
10e400 m. The faults are mainly exposed in 2D cliff sections. For
each fault core thickness, composition and the overall geometry
were recorded (Figs. 1 and 2). The thickness was measured normal
to the dip of the fault. Due to along-fault variations, the thickness
was measured in several places along individual fault outcrops,
covering the maximum and minimum thicknesses of the fault. This
yielded a database of 423 thickness points for a given displacement,
with assigned composition and geometry.
3.2. Fault core thickness
The precise boundaries of the fault core are in some cases
difcult to establish accurately (Childs et al., 2009). In such cases we
dene the core as the part of the fault envelope accommodating the
bulk of the displacement and delimits it by identifying intervals
where sedimentary structures are signicantly displaced (lenses)
or pervasively deformed (brecciated). The core can also consist
entirely of precipitated calcite along one or more fractures (Fig. 1a).
In some cases, the fault core and host rock are separated by a slip
surface, in other cases this transition may be gradual, forming
E. Bastesen, A. Braathen / Journal of Structural Geology 32 (2010) 1609e1628 1610
a fault core-damage zone transition (Billi et al., 2003). Relay
structures and fault bends require additional criteria, in that they
offer increased complexity. Breached relays were measured as
single fault strand; in un-breached relays the two fault strands
were measured individually.
3.3. Fault displacement
Measuring displacement is fairly straightforward, in cases
where the fault displacement does not exceed the height of the
outcrop. For larger faults, with displacements exceeding the height
Fig. 1. a) Schematic fault core with lithologies and geometries typically encountered in small to moderate offset extensional faults in ne grained carbonates. The fault core area is
dened by the bold lines, whereas the surrounding rock is deformed in the damage zone. The upper part of the fault core consists mainly of calcite veins and shale gouge (see inset),
where the calcite veins are truncated by slip surfaces coated with gouge to form sheared calcite veins and lenses. The middle part of the core is dominated by membranes of
carbonate breccias, while the lower part hosts shale smear and associated breccias that are arranged in composite lenses. b) Ternary compositional facies diagram based on three
end members; shale smear (SS), secondary calcite (SCa) and carbonate breccia (CB). Fault cores with combinations of these elements are represented within the diagram. In cases
where all three elements are observed, the fault core is described as a composite facies core that plots in the middle of the diagram. c) Table explaining the abbreviations of the fault
core composition.
E. Bastesen, A. Braathen / Journal of Structural Geology 32 (2010) 1609e1628 1611
of the exposure, local and/or regional stratigraphy and formation
thickness data must be used. Thus stratigraphic offset can be esti-
mated by applying the overall geometry of the fault, such as dip,
position of relay structures, and fault drag. The accuracy of this
method is dependent on the level of stratigraphic details available
and, consequently, the precision of the estimated displacement
commonly decreases with increasing displacement.
4. Geological setting
4.1. Western Sinai
The studied faults are located in carbonates exposed in the
eastern, exhumed ank of the Suez Rift (Figs. 3a and 4). The region
is characterized by large fault blocks (e.g. Hammam Faraoun and El
Qaa) bounded by basement-involved, west-facing extensional
master faults (Coastal fault belt and Eastern boundary fault belt)
with kilometre-scale displacement (Moustafa and Abdeen, 1992;
Sharp et al., 2000; Jackson et al., 2006). These large fault blocks
are broken up by subsidiary faults with maximum displacement of
a few hundred metres. The main period of fault movement, as
recorded by syn-rift deposits, occurred during the Oligocene to
mid-Miocene (e.g. Robson, 1971; Patton et al., 1994; Bosworth et al.,
2005). Faults included in the present database were found in the
Hammam Faraoun and El Qaa fault blocks (Fig. 4a).
Carbonates form a substantial part of the w500 m thick late
Cretaceous to early Tertiary El Egma Group (Moustafa and Abdeen,
1992; Bosworth et al., 2005), which can be further subdivided into
Fig. 2. Geometric classication scheme, showing possible geometries encountered in extensional faults. The diagram spans from planar faults (1a) to fault with complexities; such
as fault bends (1b,c), complex faults with splay structures (2), overlap structures (3), breached overlap (4), lenses (5), and multiple fault strands (6).
E. Bastesen, A. Braathen / Journal of Structural Geology 32 (2010) 1609e1628 1612
the Sudr, Esna, Thebes, Darat and Tanka formations (Fig. 4b). Most
data points for the present study were collected from faults in the
Sudr, Thebes, Darat and Tanka formations. The Thebes Formation
(Said, 1960; Kuss et al., 2000) is a massive, deep water, fossiliferous
limestone, characterized by abundant bands and concretions of
chert and layers of marl. The formation exhibits signicant lateral
variation, both in composition and thickness (Moustafa and
Abdeen, 1992), changing from deep marine, massive micrite
horizons in the north to cherty micritic wackestone in the south
(Moustafa, 2004). The unit exhibits successions of massive lime-
stone and heterogeneous intervals of bedded chert, marl, lime-
stones and 10e50 cm calcareous clay beds. The Thebes Formation
forms a minor reservoir, cap rock and potential source to the
hydrocarbon elds of the central Suez Rift (Alsharhan, 2003). The
Darat Formation consists of a succession of massive chalky lime-
stones and 0.5e2 mthick limestone beds intercalated with 1e5 cm
muddy to clayey limestone and up to 0.5 m thick calcareous clay
units. The Tanka Formation is a brilliantly white, bedded, chalky
limestone with thin layers of aky marl, deposited in a shallow
intertidal environment. This unit is only exposed in the northern
and central part of the Hammam Faraoun block, and is missing in
southern parts of the region due to uplift and rift-shoulder erosion
(Moustafa, 2004).
4.2. Central Oman
The study site is located in the carbonate platform of the
Adams Foothills. The investigated faults are situated in the two
mountain areas separated by a distance of 100 km; the Jebel
Quasaybah and the Jebel Madar (Fig. 3b) (Grlaud et al., 2006).
Jebel Quasaybah is located at the western end of a 70 km long,
EeW oriented anticline (Fig. 3b), whereas Jebel Madar is formed
by a local salt dome (Immenhauser et al., 2007). Outcrops are
found in the hillsides and along valleys cutting into folded and
faulted carbonates of Jurassic to Cretaceous age (Scott, 1990;
Wagner, 1990). The structural conguration reects the early
Tertiary plate-scale closure of the southeastern Persian Gulf
(Hanna, 1990), with major detachment folds located above blind
thrusts located in Paleozoic salt layers (Al-Kindi et al., 2006).
Within the major folds, there are smaller faults, distinguishable
as steeply dipping normal and strike-slip faults and thrust faults.
Faults are also located along the hinges of regional folds associ-
ated with salt diapir-driven exuring (Montenat et al., 2000;
Immenhauser et al., 2007).
Faults at the studied sites are hosted by early to middle Creta-
ceous carbonates (Grlaud et al., 2006). These can be divided into
three units, from base to top: the Shuaiba, Nahr Umr, and Natih
formations (Alsharhan and Nairn, 1997; Grlaud et al., 2006). The
Shuaiba Formation (40e125 m thick) is characterized by massive
limestones consisting of packstones and wackestone beds, whereas
the Nahr Umr Formation (150 mthick) is a thick green shale section
with some thin beds of micritic and marly limestones. The upper-
most Natih Formation (350 m) consists of layered mudstone and
wackstones/packstones (Alsharhan and Nairn, 1997).
4.3. Central Spitsbergen (Svalbard)
The dataset from Spitsbergen was collected from a rift-basin
found in the inner parts of the Billefjorden area (Fig. 3c). Several
phases of tectonic activity have been reported from the basin
bounding master fault system, the NeS trending Billefjorden fault
zone (e.g. McCann and Dallmann, 1996). Of special interest to this
study is mid- to late-Carboniferous to Permian (?) extension
causing the formation of the Billefjorden Trough; a more than
2000 m deep and 30e40 km wide, asymmetric rift-basin lled by
mixed clastic, carbonates and evaporites (Johannessen and Steel,
1992; Maher and Braathen, in press). Syn-rift deposits are found
in the Ebbadalen and Minkinfjellet formations, whereas the Wor-
diekammen Formation constitutes the late-rift succession. The
latter unit reects a transition to a regional, stable platform setting
that prevails in Spitsbergen and the Barents Shelf (Pickard et al.,
1996; Samuelsberg et al., 2003). The faults studied are formed in
the late Carboniferous to Permian Wordiekammen Formation. This
unit consists of 1e10 m thick micrite layers that may be subdivided
into the so-called black crags (Pickard et al., 1996). The crags are
separated by m-thick calcareous shales and wacke/packstones. The
base of the Wordiekammen Formation is in many places charac-
terized by breccia pipes known as the Fortet Breccia Formation.
These pipes represent paleo-karst breccias formed due to collapse
into cavities which were formed due to extensive karstication of
Fig. 3. a). Location of the Sinai study area (star), showing the current plate tectonic setting of the region. Arrows indicate relative plate motion. b) Tectonic map of the eastern part of
the Arabian Peninsula, showing the Hajjar mountain chain and the anticline structures of the study areas of the Adams Foothills in central Oman. The Jebel Madar (east) and Jebel
Quasaybah (west) study areas are denoted with stars. c) Tectonic map of the Spitsbergen island, showing large faults, including the Billefjorden fault zone (BFZ) and the map
distribution of pre-Carboniferous, late Palaeozoic and Mesozoic to Tertiary bedrock units. The star locates the study area.
E. Bastesen, A. Braathen / Journal of Structural Geology 32 (2010) 1609e1628 1613
underlying gypsum in the late Carboniferous Minkinfjellet Forma-
tion (Eliassen and Talbot, 2003).
5. Field data
5.1. Faults in Western Sinai
The majority of the studied extensional faults in western Sinai
are oriented NWeSE, i.e. parallel to the regional structural grain
(Fig. 4c). Some faults oriented NeS and NEeSWwere also observed.
Most faults juxtapose pre-rift carbonates, with a few exceptional
faults that juxtapose the Tanka Formation with syn-rift sedimen-
tary rocks. The majority of fault outcrops were observed in the
Thebes and Darat formations, but some faults juxtapose the Sudr
and/or Esna formations (shale) with the Thebes Formation. The
limestone beds fall into two categories; 1) massive- to bedded-
limestone (Thebes, Sudr and Tanka formations) and 2) interbedded
limestone shale (Darat Formation). The displacement on the
studied faults is listed in Table 1.
5.1.1. Small offset faults
Small faults in the massive limestone beds of Darat, Thebes, Sudr
and Tanka formations have a geometry characterized by straight
and slightly curved fault segments (Fig. 5a,b). Lenses and fault
splays appear in breached fault overlaps and near fault bends. In
most cases the fault core in these carbonates is composed of
membranes of calcite veins and thin (mm) clay gouge along slip
surfaces. Fault cores dominated by veins of crystalline calcite were
identied as a fault core facies in approximately 70% of all studied
small faults in Sinai. Most of these veins are cut by multiple fault-
parallel shear fractures, and are therefore classied as sheared
calcite veins. In many cases, these sheared calcite veins are asso-
ciated with thin clay gouge membranes and calcite gouge
membranes associated with striated slip surfaces. The latter are
commonly positioned along the fault core to host rock boundary.
Fig. 4. a) Geological map of the Sinai eld area, adopted from Moustafa (2004). Red squares indicate areas where faults have been studied. CBF e coastal fault belt, EBFB e eastern
boundary fault belt. b) Simplied stratigraphical column of western Sinai, highlighting the carbonate succession of the El Egma Group. c) Stereographic plot (lower hemisphere, equal
area stereo net) showing the orientation of the studied faults. (For interpretation of the references to colour in this gure legend, the reader is referred to the webversion of this article).
E. Bastesen, A. Braathen / Journal of Structural Geology 32 (2010) 1609e1628 1614
Table 1
Summary of fault core data from the Sinai study area. For each studied fault the displacement, maximum and minimum thicknesses, fault core composition and geometry are
listed. Abbreviations are given in Fig. 1.
Fault
loc.
Displacement
(cm)
Thickness
(cm)
Protolith Fm. Fault
geometry
Fault core
composition
Small offset faults
(>1 m)
Si1 4 0.3e0.7 Calcareous clay-limestone Darat 1a SCa G
Si2 11 0.7e3.3 Calcareous clay-limestone Darat 1a SCa SS
Si3 11 0.7e1.9 Calcareous clay-limestone Darat 5 SCa SS
Si4 2.6e12 0.2e2.5 Bedded limestone Tanka 1c&5 SCa SS
Si5 13 0.5e1.3 Calcareous clay-limestone Darat 1b SCa SS
Si6 14 1.5e2.8 Bedded limestone Tanka 3a, 1c SCa SS
Si7 17 2e2.4 Calcareous clay-limestone Darat 1a SCa SS
Si8 21.5 0.4e5 Calcareous clay-limestone Darat 1a&5 SCa SS
Si9 4e23 1e22 Limestone-clay-chert Darat 1b&3a CB G
Si10 25 0.5e2 Bedded limestone Tanka 1a SCa G
Si11 2.6e27.3 1.3e3.25 Bedded limestone Tanka 1c SCa
Si12 10e29 0.7e1.2 Calcareous clay-limestone Darat 1a SCa SS
Si13 32 1.8e3 Bedded limestone Tanka 1a SCa
Si14 25e33 0.2e2.4 Calcareous clay-limestone Tanka 4 SCa SS
Si15 35 0.3e1.5 Bedded limestone Tanka 2c SCa G
Si16 26e37 2.6e9.5 Calcareous clay-limestone Darat 1b CB SS
Si17 37 9e23 Bedded limestone Tanka 3a CB G SCa
Si18 40 2.8e7.3 Bedded limestone Tanka 1a SCa
Si19 3e41 0.1e1.5 Calcareous clay-limestone Darat 2a & 5 G SS
Si20 38e45 0.1e0.4 Bedded limestone Tanka 1a G
Si21 49 0.2e2.7 Bedded limestone Tanka 1a SCa G
Si22 40e59 1.5e13 Bedded limestone Thebes 4 Composite
Si23 60 1e10 Calcareous clay-limestone Darat 1b Composite
Si24 53e68 0.2e10 Calcareous clay-limestone Darat 2c & 5 SCa SS
Si25 75 2e4.7 Calcareous clay-limestone Tanka 5 SCa SS
Si26 75 0.8e9.5 Bedded limestone Tanka 1a SCa
Si27 80 1.5e9 Calcareous clay-limestone Darat 1a Composite
Si28 80 2e7 Massive limestone Thebes 3a SCa G
Si29 80 0.5 Massive limestone Thebes 1a G
Si30 80 1.2e6.5 Bedded limestone Tanka 1a&5 SCa
Si31 35e95 2.2e4.5 Calcareous clay-limestone Tanka 1a&2a SCa SS
Si32 95 0.4e7 Calcareous clay-limestone Thebes 1b Composite
Moderate offset
faults (1 me10 m)
Si33 100 1.3e12 Bedded limestone Tanka 1b CB SCa
Si34 105 2.9e13 Bedded limestone Tanka 1a&5 SCa SS
Si35 68e109 2.5e12 Calcareous clay-limestone Thebes 1a&5 SS Composite
Si36 3e114 0.2e13.1 Calcareous clay-limestone Tanka 4 SCa SS
Si37 115 2e13 Massive limestone Thebes 5 CB SCa
Si38 70e135 0.1e14 Calcareous clay-limestone Darat 2a & 5 CB SCa SS
Si39 120e140 1e15 Shale-limestone Darat 1a Composite
Si40 140 3e5 Bedded limestone Tanka 1b CB SS Composite
Si41 100e150 2e6 Shale-limestone Darat 2a SS
Si42 150 0.8e10 Massive limestone Thebes 1c, 5 G
Si43 150 0.1e6 Massive limestone Thebes 4 Sca G
Si44 150 1.8e5.1 Bedded limestone Tanka 2a&2b Sca G
Si45 190 2.3e5.8 Calcareous clay-limestone Darat 1a SCa SS
Si46 200 1.1e1.4 Massive limestone Thebes 2c G
Si47 220 1.3e13 Massive limestone Thebes 1a Sea
Si48 250 0.2e6 Massive limestone Thebes 1a CB G
Si49 250 2.6e11.5 Bedded limestone Tanka 1a SCa SS
Si50 100e260 0.6e55 Calcareous clay-limestone Darat 2c & 5 SCa SS G
Si51 150e280 3e50 Massive limestone Thebes 3b CFR
Si52 350 2e7 Limestone-clay-chert Thebes 2b & 5 CB SCa G
Si53 400 2e9.9 Shale-limestone Darat 1a SS Composite
Si54 450 5e50 Calcareous clay-limestone Darat 5 CB Composite
Si55 450e500 4.5e7.8 Bedded limestone Darat 1a Composite
Si56 560 1.3e27 Massive limestone Thebes 2b & 5 CB SCa G
Si57 600 10e25 Limestone-clay-chert Thebes 5 Composite
Si58 600 10.4 Shale-limestone Darat 1a SS
Large offset
faults (<10 m)
Si59 1000 6e28.6 Bedded limestone Tanka 1a&5 CB SCa
Si60 1600 1.4e15 Massive limestone Thebes 5,2a CB G
Si61 2800 15e20 Limestone-clay-chert Thebes 1a Composite
Si62 3000 1e38 Limestone-clay-chert Thebes 1a, 5 CB SCa G
Si63 3500 4e30 Massive limestone Thebes 5 CB SCa
Si64 5000 21e23 Limestone-clay-chert Thebes 5 CB
Si65 23000 70e195 Limestone-chalk-shale Sudr-Thebes 7 CB Composite
Si66 25000 15e110 Bedded limestone Tanka 7 CB
Si67 35000 1100 Shale (Esna) Sudr -Thebes 7 SS
Si68 40000 100e490 Limestone shale Tanka-Thebes 7 CB SS
E. Bastesen, A. Braathen / Journal of Structural Geology 32 (2010) 1609e1628 1615
Calcite vein membranes are generally fairly continuous along
the fault core. Their thickness varies from 0.5 to 13 cm, with the
largest thickness related to fault bends where lens shaped vein
bodies are formed. In most of these lenses the vein lamination
curves similar to the lens shape, and the central part of the calcite
lled lens exhibits open voids with sparry calcite.
In thin section, sheared calcite veins are seemed to consist of
a matrix of crystalline calcite cross-cut by fractures that, in some
Fig. 5. Examples of small and moderate offset faults observed in Sinai, Egypt. a) Fault with 20 cm displacement hosted in marly limestone, chert and chalk layers (Thebes
Formation). The fault has jogs, which are characterized by thick breccia pods next to chert layers. Thin slip zones characterize more planar parts of the fault in the limestone layers.
b) Fault with 1.5 m displacement showing a fault core of calcite veins (SCa) and shale supported carbonate breccia (SCB). c) Mosaic of photomicrographs that displays the
microstructures of a sheared calcite vein collected from a fault with w2 m displacement. The upper part is a cemented gouge (SCa G), neighbouring a zone of calcite crystals (SCa)
cut by numerous curved fractures (slip surfaces). The fractures are associated with thin slivers of gouge and protolith rock, probably derived from shearing along the vein-protolith
contact. The lowermost part is a cemented carbonate breccia (CCB). d) Fault with 4 m displacement cutting through marl and shale prone parts of the Darat Formation. The fault has
a prominent slip surface with well developed shale smears (SS), gouge (G) and breccia (CB) membranes. Note lens formation near the fault bend and the drag of layering in the
footwall and hanging wall.
E. Bastesen, A. Braathen / Journal of Structural Geology 32 (2010) 1609e1628 1616
cases, are coated with a thin (w1 mm) membrane of gouge (Fig. 5c).
Inside the crystalline calcite vein, <1 mm thin lenses/slivers of host
rock limestone can be observed.
Carbonate breccias are commonly found in lenses and typically
consist of coarse, clast-supported breccia that shows extensive
calcite cementation. These lenses are especially common near fault
bends related to chert bands of the Thebes Formation (Fig. 5a).
In shale prone carbonates (Darat and parts of Thebes formations),
small faults appear as bifurcated, splayed and bended (Fig. 5a). In
these faults, lenses occur as partly shattered limestones that
commonly are found along complex fault overlap zones. Shale
smears, together with fault rocks lenses and calcite veins, locally
form composite fault cores. The shale smears show variable
smearing potential, depending on the thickness of the shale proto-
liths. Shale smear factors (Lindsay et al., 1993) are typically around 4.
5.1.2. Moderate offset faults
In faults with moderate offset, calcite veins occur as broken up
lenses and/or juxtaposed with other fault core facies, such as
breccias and shale smears (Fig. 5b,d). Carbonate breccias are more
common compared to small faults, constituting 35% of the core
facies. The breccia matrix mostly consists of calcite gouge that is
cemented, but in the shale prone units breccias are also shale
supported. The carbonate breccias are arranged in semi-continuous
membranes or as highly elongated lenses, bound by slip zones,
often in association with shale smears and calcite veins.
5.1.3. Large offset faults
Faults featuring 10e50 m displacements were only observed in
the massive parts of the Thebes Formation. These are characterized
by relatively thin fault cores (2e50 cm) consisting of membranes of
ne grained (sub-mm clast size) carbonate gouge/breccias, thin
clay gouge, calcite veins and fault rock lenses (Fig. 6a). The lenses
are 2e5 m long and may be 20e50 cm thick, consisting of ne
grained to coarse grained breccias.
In large faults with more than 100 m displacement, shale smears
of the approximately 50 mthick Esna and Thal formations (between
Darat and Tanka formations) are common (Fig. 6b,c). These faults
may be totally dominated by shale smear, and the thickness of
observed shale layers in cores are up to 11 m in large faults juxta-
posing the upper Sudr to the Darat formations. However, common
thicknesses are around 0.5e3 m. In some cases these shale smears
are associated with thick layers of coarse grained fault breccias,
partly shattered limestone and calcite veins along slip surfaces. They
are classied as composite fault cores (Fig. 6b). In larger faults or in
positions far away from the source shale layer of the Esna shale, the
displacement exceeds the smearing potential of the shale. Fault
cores in these cases are dominated by breccia membranes and thin
and patchy lenses/pockets of shale (Fig. 6c). All large faults are
accompanied by a well-dened damage zone consisting of fractures,
small and partly moderate faults, and, locally stylolites.
5.2. Faults in central Oman
Faults studied at the Jebel Madar locality are mostly NeS and
NEeSW oriented, steep and dening horst-and-graben structures
parallel to the salt diapir fold crests (Fig. 7a,b). In this locality the
displacement of the studied faults varies between 10 cmand 300 m,
but may be much larger in other parts of the Jebel Madar (Fig. 7a and
Table 2a). In the Jebel Quasaybah area, data were collected from an
array of extensional faults, dipping 60
e80
e40
to 85
with an average
value of about 19
for
subparallel interacting sheared joints (Fig. 5a,b and Table 1a).
Terminal areas of a fault zone generally reect the incipient
stages of fault development. In this regard, Fig. 6 shows a portion of
a fault zone that has w65 cm maximum observable right-lateral
slip elucidating the transition from an echelon sheared-joints array
to a through-going fault formation. Similar to the cases shown in
Fig. 4, shearing of the initial joint systemresulted in splay fracturing
and continued shearing facilitated the formation of multiple sets of
sequential splays localizing into discontinuous pockets of high
density fractures, and, in places, fragmentation zones. The incipient
short slip surfaces eventually go through these pockets of weak-
ened rock at fault steps.
The photographs and maps in Fig. 7a,b showa well-exposed fault
of about 14 m left-lateral slip, which displays several characteristic
architectural elements common to all strike-slip faults in the study
area: Fault rock, slip surfaces, and damage zone. Fig. 7a shows
domains of different deformation zones and of fracture densities,
which allow one to see a simpler picture of elongated, noncolinear
bodies of the fault rock and the adjacent areas of high fracture
density. Fig. 7a,b also shows slip surfaces invarious orientations and
sizes, one of whichis continuous fromone endof the mappedarea to
the other going through the elongated bodies of fault rock. There are
also short slip surfaces terminating at an acute angle to the
rectilinear strings of fault rock (Fig. 7b). We interpret these diagonal
short slip surfaces as relics of the initial sheared joints and the high
Fig. 3. Idealized diagrams summarizing general trends of (a) mean segment length
and number of steps per kilometer, and (b) mean fault rock and mean damage zones
widths as the maximum fault slip increases.
Fig. 4. (a) Incipient right-lateral shearing (w2 cm) of a series of echelon joints with right steps. Splay fractures at high-angle to the sheared joints are localized near the tips of the
segments at the steps. From Myers and Aydin (2004). (b) Two sets of splay fractures at and around a right step along a strike-slip fault with about 80 cm right-lateral slip. The two
sets have a range of intersection angles from 30
to 60
),
and for closely spaced interacting subparallel faults and their splays (50
). From de
Joussineau et al. (2007). (b) The spacing range dened by the best t line to the
smallest end of the spacing distribution of the fault-related fractures in the study
area. From de Joussineau and Aydin (2007). The spacing values obtained by this
method are between w1 and 5 cm with the largest concentration between 1 and
2 cm (see inset 1). The angular differences between the scanline and fault-related
fractures for 14 m fault (inset 2 in which the bins represent intervals of 09, 1019,
etc.). More than 75% of the fractures make angles larger than 50
Interacting
faults
50
b
Fault Spacing
range
(cm)
Inset 2
80 cm 1.518
8 m 2.039
14 m 2.051
80 m 1.4110
Mixed
Scanline #1 2.052
Scanline #2 0.914
Scanline #3 5.038
A. Aydin, J.G. Berryman / Journal of Structural Geology 32 (2010) 16291642 1633
with 14 m slip referred to earlier in this manuscript (see Figs. 7 and
8). The spacing distribution for the fault-related fractures for this
fault has been determined along 8 scanlines perpendicular to the
fault trace in intervals about 26 m recording the distance,
orientation and length of the fractures with a resolution of 0.5 cm.
The spacing has been calculated as the distance between consec-
utive fractures. Of interest here is the spacing distribution on the
smaller end of the spectrum identied by the linear trend in the
spacing distribution plot which denes a range of spacing values
from 2 to 51 cm (see Fig. 10a). Fig. 10b shows the fracture spacing
distribution obtained froma single scanline across many faults with
aggregate slip on the order of a few hundreds of meters. Here the
range of the smallest linear spacing trend is 538 cm. In both cases,
we focus on the smallest ends of the ranges, 2 cm in Fig. 10a and
5 cm in Fig. 10b) which represent the smallest fracture spacing for
a statistically signicant number of data points measured near the
fault cores. Similarly, in Table 1b, the spacing ranges dened by the
linear ts to the smallest slopes in the distribution of data, and the
corresponding minimum spacing values, for three other faults and
two additional scanlines across a number of faults are given. As
shown in the histogram summary (inset 1) a majority of the
minimumspacing values falls between 1 and 2 cm. Considering the
minimum measurable spacing was 0.5 cm, these numbers are well
above the minimum resolvable spacing.
5. Analysis using effective medium models
The premise of this study is that a certain degree of high
intensity fracturing at fault steps and fault damage zones weakens
the rock masses thereby facilitating fault lengthening through
linkage and coalescence of neighboring segments and fault zone
widening by incorporation of the fractured and fragmented
material into the fault rock via a cataclastic process.
Next, we will use effective medium models to investigate the
degradation of the strength and reduction of resistance to
cataclastic deformation, which presumably pave the way for the
setting of through-going faults. Given the complexity of the
Fig. 6. Detailed map of the end of a small shear zone of about 65 cm observable
maximum right-lateral slip showing a set of slightly sheared and highly overlapped
echelon joints and sheared joints with many splay joints at high-angle to the sheared
joints. Gray shading marks narrow pockets of fragmentation and thick lines show
incipient through-going slip surfaces orientated at a small-angle to the sheared initial
echelon joints. Slightly revised from Davatzes and Aydin (2004).
Fig. 7. (a and b). Detailed maps of a strike-slip fault with about 14 m left-lateral slip. (b) shows the orientations, lengths, and intersections of damage zone fractures (splay joints and
sheared splay joints) around the fault core (Myers and Aydin, 2004) whereas (a) shows a new reinterpreted version of the same fault architecture in which noncolinear pockets of
fault rocks and highly fractured domains of occasionally fragmented damage zone are delineated in the eld. One through-going slip surface (thick solid line) and several short slip
surfaces (dotted lines) diagonal to the through-going slip surface are highlighted. The geometry and distribution of many of short diagonal slip surfaces resemble the initial echelon
sheared joints observed along faults with smaller slip in the area. The original map by R. Myers (1999); the present version was revised from Davatzes and Aydin (2004).
A. Aydin, J.G. Berryman / Journal of Structural Geology 32 (2010) 16291642 1634
fractures around faults, the problem is obviously rather difcult
and, at this stage, further simplication is desirable for applications
to an effective medium theory. We rst idealize the common
fracture patterns in terms of their orientation-intersection angle,
length, and density (Fig. 11a). We then study parametrically the
effective elastic moduli of such a conguration as a function of
fracture density for each idealized fracture pattern dened by the
angle between fracture sets in order to assess the degree of
degradations in the moduli values as the fracture density increases.
The details of the effective medium theory that we shall employ
have been recently described by Berryman and Aydin (in press) and
is based on the earlier work by Backus (1962), Schoenberg and Muir
(1989), and Berryman and Grechka (2006).
Fig. 11a shows an idealized fracture network which is consistent
in principle with the sequential formation of two fracture sets and
their ultimate pattern, the examples of which can be seen in Figs. 4,
6, 7, and 8. Here the lengths (l) of the fractures, the density (r) or
spacing (s) of the fractures, and angle (F
F
) between the two sets of
fractures characterize the pattern in a layer with a thickness, h.
One of the most commonly used fracture density concepts goes
back to Bristow (1960) and Budiansky and OConnell (1976). For
a set of rectangular at (or a ribbon-shaped) fractures, which is the
most pertinent to physical properties of fractured media such as
resistivity, uid ow, and elasticity, is
r nh
2
l (1)
where n N/V, with N and V being the number of fractures and the
rock volume, respectively. The volume, V is equal to lhs where l, h,
and s are average fracture length, height, and spacing, respectively.
Taking t as the average fracture thickness or fracture aperture, the
porosity of a system of rectangular at fractures with an average
Fig. 8. Damage zone characteristics around the strike-slip fault with w14 m left-lateral slip. (a) Two fractured domains were distinguished: The inner damage zone of high fracture
density right next to the fault core; and the outer damage zone of signicantly lower fracture density. (b) A detailed map of the inner damage zone shows that several generations of
splay fractures (marked by different color codes) ll in between the echelon sheared joints of the initial stage. (a,b) From de Joussineau and Aydin (2007). (c) Detail map showing
ne- and coarse-grained fault rock and major through-going slip surfaces along a left-lateral fault with about 25 m left-lateral slip. Some of the earlier fractures within the coarse-
grained fault rock can still be identied (dotted lines). Also important to point out triangular pockets of ne-grained fault rock protruding into the damage zone in some locations on
the left hand side of the fault core, where the fracture intensity is the highest.
A. Aydin, J.G. Berryman / Journal of Structural Geology 32 (2010) 16291642 1635
spacing value, s, is dened as a fraction of the spacing distance
occupied by the pores:
4 lht=V t=s (2)
Then, the fracture density is given as
r h=twh=s (3)
Based on Eqs. (2) and (3), we nd that the fracture density as
dened here is proportional to height or bed thickness over spacing
and is dimensionless. The fracture density of, for example, 1.0,
corresponds to a commonly observed average spacing for one set of
opening mode fractures in a bed, for which the average spacing, s,
scales with the bed thickness, h, for well-developed fracture
systems (Wu and Pollard, 1995; Bai and Pollard, 2000). Thus, for
a bed thickness of 5 cm, the average spacing is 5 cm. For two sets of
overlapping fracture systems of equal density in a bed, the value for
the density approaches 2.0, and the average spacing is h/2. For a bed
of 5 cm thick, the equivalent spacing for the two overlapping sets is
2.5 cm.
5.1. Compliance matrix and the corresponding Youngs and shear
moduli components
The quasi-static equation of elasticity using Voigt notation is
(Nye, 1985; Pollard and Fletcher, 2005):
Fig. 9. Conceptual model showing linkage and coalescence of fault segments or strands which result in longer segment lengths, reduced number of fault steps, larger step sizes, and
wider damage zones and fault rock zones with increasing fault slip. From de Joussineau and Aydin (2007).
Fig. 10. Fracture spacing distributions: (a) for the 14 m fault and (b) across an area which included many faults with an aggregate slip of a few hundred meters. The line ts to the
data at the smallest ends of the spacing range are also shown. The smallest spacings dened by these linear trends are taken as the critical values below which the systems are
thought to be unstable. Slightly changed from de Joussineau and Aydin (2007).
A. Aydin, J.G. Berryman / Journal of Structural Geology 32 (2010) 16291642 1636
0
B
B
B
B
B
B
@
3
11
3
22
3
33
3
23
3
31
3
12
1
C
C
C
C
C
C
A
0
B
B
B
B
B
B
@
S
11
S
12
S
13
S
14
S
15
S
16
S
21
S
22
S
23
S
24
S
25
S
26
S
31
S
32
S
33
S
34
S
35
S
36
S
41
S
42
S
43
S
54
S
45
S
46
S
51
S
52
S
53
S
54
S
55
S
56
S
61
S
62
S
63
S
64
S
65
S
66
1
C
C
C
C
C
C
A
0
B
B
B
B
B
B
@
s
11
s
22
s
33
s
23
s
31
s
12
1
C
C
C
C
C
C
A
(4)
where 3 and s are the six independent components of strain and
stress, respectively, and S is the symmetric 6-by-6 compliance
matrix. The numbers 1, 2, 3 always indicate Cartesian axes (say, x, y,
z respectively). Elastic extension in the x- or 1-direction is denoted
by 3
11
, etc., while a shearing (torsion or twisting) strain around the
x- or 1-axis is represented by 3
23
, etc. Similarly, the normal stress or
tension in the x-direction is s
11
, and the shear stress around the
x-axis is symbolized by s
23
, etc.
For any system, the full compliance matrix (Eq. (4)), or its
inverse, the stiffness matrix, has six eigenvectors, each of which is
a 1 6 matrix and is associated with a scalar eigenvalue. If S is the
matrix, v is the eigenvector, and c is the eigenvalue, then by de-
nition Sv cv. This means that when matrix S is multiplied by
vector v, the result is a vector proportional to v, and the constant of
proportionality is the eigenvalue c. There are always 6 distinct
eigenvectors. However, eigenvalues may or may not all be distinct.
For an isotropic system, ve of these eigenvalues are for shearing
modes and one is for pure compression/tension mode. Of the ve
shearing modes, three are the independent torsional or twisting
motions and/or the corresponding stresses; for example, in an
isotropic system, 3
23
, couples simply to s
23
, while all the off--
diagonal compliances and/or stiffnesses involving subscripts 4, 5, 6
vanish identically. Two other types of shear modes are eigenmodes
for an isotropic system; for example, when s
22
s
11
, we have
a comparable push-pull or pure shear mode resulting in the
eigen-response 3
22
3
11
for the strain. For the isotropic case,
there are three distinct versions of these pure shear behaviors that
give analogous results, but for nonisotropic systems usually only
one of these will actually be an eigenmode the most common
example of this behavior being for transversely isotropic systems.
In the presence of a set of perfectly parallel fractures in an
otherwise isotropic elastic medium, the elastic matrix becomes
transversely isotropic. The plane of the parallel fractures is the
plane of symmetry, and the direction perpendicular to this plane is
the axis of symmetry. Elastic behavior strictly within the plane of
symmetry (i.e., two-dimensional behavior in this plane) remains
isotropic, which is the origin of the term transverse isotropy
this type of isotropic behavior thus occurring transversely to the
axis of symmetry.
When analyzing such systems in three-dimensional space, it is
common to choose the axis of symmetry to coincide with one of the
spatial axes, x, y, and z, or 1, 2, and 3, respectively. This choice makes
no difference to the nal results but makes some difference to the
level of difculty in obtaining those results. In particular, making
a good choice of axes can simplify the matrix of elastic coefcients
somewhat, so that, for a set of parallel fractures, we have
a compliance matrix in the Voigt (Nye, 1985) 6 6 matrix form of
the elastic tensor notation:
S
0
B
B
B
B
B
B
B
B
B
@
1
E11
n
12
E11
n
13
E33
n
12
E11
1
E22
n
23
E33
n
13
E33
n
23
E33
1
E33
1
G44
1
G55
1
G66
1
C
C
C
C
C
C
C
C
C
A
(5)
Thus, for orthorhombic symmetry, the diagonal components of
the matrix; the Youngs moduli E
11
, E
22
, and E
33
and the shear
moduli G
44
, G
55
, and G
66
are inversely related to these diagonal
components. Note that the zero matrix elements were left blank
in Eq. (5) for simplicity as is standard practice. A particular
modulus that we call qGp, for quasi-pure shear mode is also
calculated because it is likely to play a role in the failure of the
systems we examine here, and it is well known in the geological
sciences. The mode qGp is actually an eigenvector of the system
considered, but its physical interpretation is not simple, because it
is not exactly any one of the six standard modes of a simple elastic
system pointed out earlier. However, its behavior is very close to
that of a pure shear mode and that is why the term quasi is used
here.
In calculating these components of the effective elastic moduli
for a medium with Poissons ratio of 0.4375 appropriate for sand-
stone, which has two fracture sets (Fig. 11a), we followan approach
based primarily on layer averaging methods of Backus (1962) and
Schoenberg and Muir (1989). The details of the mathematical
analysis of the effective properties of such a composite system are
given by Berryman and Aydin (in press). Basically, two different
layers each containing one set of fractures with the same density (r)
but possibly differing distributions, are considered for the effective
moduli calculations (Fig. 11b). After constructing one layer with one
of the fracture sets, this layer is rotated in such a way that the
combined fracture system will have the desired angle between the
Fig. 11. (a) A simplied and idealized fracture pattern at fault steps and within inner
damage zones. The pattern is dened by the angle (F
F
) between the two fracture sets
and lengths (l) and density (r) of the fracture systems. Fracture density is proportional
to the ratio of layer thickness or fracture height (h) to average spacing (s). (b,c) Block
diagrams showing the procedure to represent layers with each fracture set and the
congurations of the layers for averaging the effective properties in z- and x-directions
or 3- and 1-directions, respectively. The block diagrams represent stacking up layers
vertically in z-direction ((b) sandwich conguration) and arranging the layers side by
side in x-direction ((c) contiguous conguration).
A. Aydin, J.G. Berryman / Journal of Structural Geology 32 (2010) 16291642 1637
two fracture sets. This is done by rotating each layer plus/minus one
half of the angle between the two fracture sets. In this paper, we
investigate cases where the angles between the two fracture sets
(F
F
) are 15
, 30
, 45
and 60
, 30
, 45
, and 60
and
for a range of fracture densities from 0 to 0.2. This range of density
is constrained by the availability (from previous work of Berryman
and Grechka, 2006) of the fracture inuence coefcients which are
required for the effective property calculations. The results show
that the E
11
components of the Youngs moduli for all four fracture
congurations do not change at all (all four lines overlap along the
top blue line in Fig. 12a) with increasing fracture density up to 0.2.
This is because E
11
corresponds to uniaxial loading in the x- or
1-direction and the changes of the angles and densities of fractures,
as seen from this direction edge-on, make no difference on the
effective moduli. The E
22
and E
33
components show systematic
decrease for all congurations as the fracture densities increase. We
note that the E
22
for the conguration F
F
60
and 45
experiences
greater decreases for the range of densities than fracture sets with
other intersection angles, whereas E
33
shows greater decreases for
congurations with smaller intersection angles, F
F
15
and 30
.
For example, the effective Youngs modulus, E
33
, corresponding to
F
F
15
and 15
. However, this
decrease amounts to about 20% of the modulus for a medium
without any fractures. The curves for G
44
components have
a crossover at a fracture density, r, between 0.1 and 0.15. This
crossover is curious and remains to be investigated further.
Fig. 12c shows the variation in the effective quasi-shear modulus
for pure shear (qGp) as being one of the special cases. This
parameter shows a smaller variation of about 5% (with respect to
the modulus for the no-fracture state) for F
F
60
at the highest
fracture density (0.2) used in the calculations.
5.2. Extrapolations
Because the concept of elasticity is based on energy storage in
the elastic material/medium, there is an elastic conservative energy
associated with the elastic system. Each eigenvalue is a measure of
the elastic energy that can be stored in the system associated with
its elastic matrix. Since these stored energies must be positive
quantities, it follows that the eigenvalues themselves must all be
positive. If any elastic eigenvalue for a system vanishes, then this
means that it is impossible to store energy in this particular mode
and that the strain of the systemincreases without additional stress
if it is attempted to excite this mode. This condition denes
a mechanical instability in the system. So it is reasonable to use this
condition as one denition of mechanical system failure, and this is
why we look for the appearance of such failed modes in our analysis
of elastic system response. However, the vanishing values that we
want lie outside the range of values in our plots. This is because, as
pointed out earlier, the required fracture inuence coefcients are
not presently available for values outside of the range of densities
considered here.
The ways in which the shear and quasi-shear moduli vary with
increasing fracture density for each conguration in our models are
nearly linear. This may warrant extrapolations using the last
segment of the curves (for r between 0.1 and 0.2) to estimate the
critical fracture densities corresponding to the vanishing values of
shear moduli components. Hence, G
55
and G
44
plots for the inter-
section angle of 15
, 30
, 45
and 60
) and
a range of fracture densities (r from 0.0 to 0.2). (a) The three components of the Youngs moduli (E
11
, E
22
, and E
33
) with increasing fracture density. The nearly horizontal solid blue
line near the top represents all four E
11
plots for the density range and for all four angles between the fracture sets (the lines just overlap). (b) The three components of the shear
moduli (G
44
, G
55
, and G
66
) for the four angles with increasing fracture density. (c) The quasi-shear moduli for pure shear for the same four angular congurations and density range.
This moduli show the least decrease in magnitude with respect to that for no-fracture state.
A. Aydin, J.G. Berryman / Journal of Structural Geology 32 (2010) 16291642 1639
Schulson, 2001) of thin and slender rock slabs between a set of
fractures as a mechanism for through-going shear fracture forma-
tion. It is difcult to identify these mechanisms in the eld. As the
eld data show, most fracture-bounded blocks have diamond
shapes due to dihedral intersection angle between the fracture sets
and may not be favorable either for buckling or bending. Rather, the
triangular areas at the intersection of the fractures appear to be
most prone to further fracturing and fragmentation. The rotation
and translation of some fracture-bounded rock blocks with respect
to the neighboring blocks can be identied in advanced stages of
deformation, particularly within slivers preserved within fault
cores. However, the relative timing of these rotations and relative
motions with respect to the shear zone evolution cannot be
determined.
The presence of multiple sets of fractures formed by splaying
within fault steps and inner damage zones with intersection angles
different than 90
) used in the
effective medium models in our study. However, the effective
properties based on the high end of the intersection angles
(4560
NW) and fractures of various sizes and types, many of which are
lled with secondary minerals (Fig. 4). Only a small part of the damage zone (which is
many hundred metres thick) is seen in the photograph.
Fig. 4. Mineral-vein network in a part of the damage zone of the Husavik-Flatey Fault
(Fig. 3). Some 80% of the veins are pure extension fractures, driven open by uid
overpressure (Gudmundsson et al., 2002).
Fig. 5. View southeast, an example of the variation in fracture frequency with distance
from the core of a fault in Vaksdal, West Norway. The highest number of fractures is at
the contact between the core and the innermost part of the damage zone. From there,
the fracture number decreases, in an irregular fashion, towards the host rock (gneiss),
at about 10 m from the core. Many fracture frequency proles of this type are provided
by Simmenes (2002) and Larsen (2002).
A. Gudmundsson et al. / Journal of Structural Geology 32 (2010) 16431655 1645
a three-dimensional body with elastic properties that differ from
those of the host material. More specically, an elastic inclusion is
a body with material properties that contrast with those of the
surrounding material, commonly referred to as the matrix, to
which the inclusion is welded.
The concept of an elastic inclusion as described here is well
established in the classical elasticity and rock mechanics literature
(Eshelby, 1957; Savin, 1961; Jaeger et al., 2007). The more recent
literature on micromechanics, however, uses inhomogeneities
rather than elastic inclusions for the concept dened above
(Nemat-Nasser and Hori, 1999; Qu and Cherkaoui, 2006). Here an
elastic inclusion denotes a material body hosted by a larger body
with different elastic properties (Gudmundsson, 2006), so that any
rock body, such as a fault zone, hosted by a larger body with
different properties is regarded as an inclusion.
The presence of an elastic inclusion modies the regional stress
eld so as to generate a local stress eld that operates both within
the inclusion and in its vicinity (Fig. 7). This follows because the
elastic properties of the inclusion, particularly its Youngs modulus
or stiffness, differ from those of the host rock. Thus, during any
loading (stress, displacement, or pressure), the responses of
the rocks constituting the inclusion differ from those of the
surrounding rocks. For example, if the rocks that constitute the
inclusion (the fault zone) are stiffer (higher Youngs modulus) than
the host rock, then the inclusion takes on most of the loading and
becomes subject to either relative tensile stresses (if the loading is
in extension) or compressive stresses (if the loading is in
compression). By contrast, if the inclusion rocks are more
compliant or softer than the host rock, then most of the loading is
taken up by the host rock which, thereby, develops locally high
relative tensile or compressive stresses depending whether the
loading is in extension or compression.
As is indicated above, many, and perhaps most, fault zones are
composed of a core and a damage zone that are widely different in
mechanical properties (Figs. 25). In addition, the damage zone
itself is commonly composed of subzones with different mechan-
ical properties, partly attributable to variations in fracture
frequencies (Figs. 5 and 8). Thus, the local stresses are likely to vary
not only between the host rock and the fault zone, or between the
core and the damage zone, but also within the damage zone itself.
To take the difference in stiffness between the host rock and the
fault zone into account, and how these change the local stresses of
fault zones, consider rst the model in Fig. 9. This model is based on
a normal fault zone in Vaksdal, close to Bergen in West Norway
(Fig. 8). The model divides the fault zone into four main subzones.
In the centre there is the fault plane itself, modelled as an internal,
compliant elastic spring. Based on estimates from open fractures
Fig. 7. Fault zone modelled as a simple elastic inclusion (Fig. 6). The nite-element
(www.Ansys.com; Zienkiewicz, 1977) model can be viewed either as a sinistral strike-
slip fault (lateral section) or as a normal fault (vertical section). Youngs modulus of the
fault zone is 1 GPa, a typical generalised value (Figs. 10 and 11), and that of the host
rock 40 GPa. The horizontal tension is 5 MPa, a value close to the maximum tensile
strength of solid rocks (Haimson and Rummel, 1982; Schultz, 1995; Amadei and
Stephansson, 1997), and may thus be regarded as a typical loading before fault slip in
active rift zones. The trends of the stress trajectories of s
3
(the minimum principal
compressive, maximum tensile, stress) change at the contact between the fault zone
and the host rock, indicating that the fault zone has a local stress eld different from
the regional eld of the host rock.
Fig. 8. Fracture frequency as a function of distance from the core of the fault modelled
in Figs. 9 and 10. The measurements are from a major normal fault zone in Vaksdal,
close to Bergen in West Norway (Simmenes, 2002). The inner part of the damage zone
has 24 fractures per unit area but at a distance of 20 m from the core, the outer part of
the damage zone has 14 fractures per unit area. Then at 80 m from the core, the
fracture frequency has fallen to 4 per unit area and is the same at a distance of 110 m;
these two latter areas are thus regarded as part of the host-rock fracture frequency.
Fig. 6. Schematic illustration of a fault zone as an elastic inclusion (inhomogeneity).
Normally, the elastic properties of the fault rock (damage zone and core) differ from
those of the host rock, so that there will be stress concentration around the fault zone,
as well as a local stress eld inside it. It is this local stress eld that controls fracture
development and fault slip in the fault zone and, therefore, largely its permeability.
A. Gudmundsson et al. / Journal of Structural Geology 32 (2010) 16431655 1646
(Gudmundsson and Brenner, 2003), the stiffness of the spring is
taken as 6 MPa m
1
. The stiffness of an elastic spring is determined
from a stressdisplacement curve, whereas Youngs modulus is
determined from a stressstrain curve. Thus, while Youngs
modulus has the units of (M)Pa, the spring has the units of
(M)Pa m
1
.
The fault plane is surrounded by the fault core, whose Youngs
modulus is taken as 0.1 GPa. This value is based on typical Youngs
moduli of unconsolidated rocks as well as in situ measurements
from various fault cores worldwide with common values between
0.1 and 1 GPa (Hoek, 2000; Schon, 2004). The Youngs modulus of
the inner damage zone is 1 GPa, that of the outer damage zone
10 GPa, and that of the host rock 50 GPa. These values reect the
decreasing number of fractures (Fig. 8) with increasing distance
from the inner damage zone to the host rock. The rock itself is
gneiss, but in accordance with well-known effects of fractures and
other cavities on Youngs modulus (Farmer, 1983; Priest, 1993;
Nemat-Nasser and Hori, 1999; Sadd, 2005), Youngs modulus is low
for the highly-fractured inner damage zone (Fig. 9), somewhat
higher for the less-fractured outer damage zone, and highest for the
normally fractured host rock.
The stressconcentration results (Fig. 10) indicate that, because
of the lower Youngs modulus inside the fault zone than outside it,
for the given loading conditions there will be lower von Mises shear
stresses in the fault zone than in the host rock. This may seem
surprising given that, when generalised, the fault slip is mostly
conned to the fault zone rather than the host rock. However, the
von Mises shear stresses reach the typical stress drops/driving
stresses for seismogenic fault slip, which are mostly 112 MPa
(Scholz, 1990), and the slip would occur in the fault zone simply
because it already has a weak fault plane and, most likely, a much
higher pore-uid pressure than the host rock. It is well-known that
tectonic earthquakes are usually related to zones of high-uid
pressure, so that, using the modied Coulomb criterion, the driving
shear stress for seismogenic fault slip, s, becomes:
s = 2T
0
f (s
n
P) (1)
where T
0
is the tensile strength of the rock, f is the coefcient of
internal friction, s
n
is the normal stress on the fault plane, and P is
the total uid pressure on the fault plane at the time of slip. When
the uid pressure approaches or equals the normal stress, the term
f(s
n
P) approaches or equals zero (for a higher uid pressure the
term may, in fact, become negative), so that the driving shear stress
for slip becomes 2T
0
. Since the in situ tensile strength of rocks is
commonly in the range of 0.56 MPa (Haimson and Rummel, 1982;
Schultz, 1995; Amadei and Stephansson, 1997), it follows that, for
high-uid-pressure fault zones, the driving shear stress for slip
should be 112 MPa, which is in agreement with common stress
drops (Kasahara, 1981; Scholz, 1990). Thus, even if the low-Youngs
modulus in the damage zone and core results in comparatively low
shear stresses in many active fault zones, they tend to slip because
of the existing weak fault plane (or planes), the high-uid-pressure
(and thus low friction), and the low effective normal stress on the
fault plane.
The local stresses in a fault zone do not depend only on the
stress magnitudes, but also on the directions of the stress vectors,
as represented by the trajectories of the principal stresses. The
model in Fig. 7 considers the fault as a single zone, an inclusion, but
as we have seen (Figs. 25, 8) there is commonly a signicant
difference in mechanical properties between the core and the
damage zone, as well as between the various subzones of the
damage zone itself. This is taken into account in the model in
Fig. 10, and also in the model below (Fig. 11).
In the model in Fig. 11 the fault zone is divided into ve
subzones. One, in the centre, represents the core of the fault zone
and has a Youngs modulus of 1 GPa, similar to many compliant or
soft breccias and unconsolidated rocks (Hoek, 2000; Schon, 2004).
Then comes the inner part of the damage zone, on either side of the
core, with a Youngs modulus of 5 GPa. This is, again, similar to the
Youngs modulus of many fractured rocks, as is indicated above. The
other part of the damage zone has a stiffness of 10 GPa, which is
similar to many fractured rocks where the fractures are not very
dense (Gudmundsson and Brenner, 2003). Finally, the host rock has
a Youngs modulus of 40 GPa, which is typical for many solid rocks
(Bell, 2000; Nilsen and Palmstro m, 2000).
The loading is extension oblique to the fault (Fig. 11). Viewed in
a vertical section, the loading would be appropriate for a reverse
fault, whereas viewed in a lateral section, the loading would be
appropriate for a dextral strike-slip fault. In either case, the oblique
loading combined with the variation in stiffness (Youngs modulus)
towards the centre of the fault (through the damage zone and to the
Fig. 9. Set-up of the model in Fig. 10 is largely based on the internal structure of the
fault zone in Fig. 8. The fault trends N40
1 n
2
K
2
I
E
1 n
2
K
2
II
E
(1 n)K
2
III
E
(2)
where G
IIII
are the material toughnesses for the ideal crack-
displacement modes IIII (Broberg, 1999; Anderson, 2005), E is
Youngs modulus (compliance or stiffness), n is Poissons ratio,
and K
IIII
are the associated stress-intensity factors. The critical
value of the stress-intensity K
c
denotes the fracture toughness.
Eq. (2) assumes plane-strain conditions; in the case of plane-
stress, the term (1 n
2
) =1. By their nature and loading, fractures
that become deected into discontinuities or interfaces
(contacts) are generally of a mixed-mode (Hutchinson, 1996;
Xu et al., 2003).
As regards pure crack-displacements, the opening (extension)
mode is denoted by I, the in-plane shear mode by II, and the out-of-
plane (anti-plane) shear mode by III (Broberg, 1999; Anderson,
2005). In geology, a mode I crack model is suitable for extension
fractures whereas mode II is suitable for many dip-slip faults
(normal and reverse) and mode III for strike-slip faults. All of these
fracture types and modes, IIII, are common in the damage zones of
fault zones (Figs. 35).
If the subzones or layers on either side of a discontinuity have
the same mechanical properties, such as is sometimes approxi-
mately the case in parts of a faults zone, the condition for an
extension fracture to penetrate the discontinuity (Fig. 14B) is that
the strain energy release rate G
p
, (with subscript p for penetration)
reaches the critical value for fracture extension, namely the mate-
rial toughness of the layer, G
L
(with subscript L for rock layer). Thus,
from Eq. (2) the conditions become:
G
p
=
1 n
2
K
2
I
E
= G
L
(3)
By contrast, the fracture will kink at or deect into the discon-
tinuity if the strain energy release rate reaches the material
toughness of the discontinuity itself, G
D
(with superscript D for
discontinuity). Since the fracture propagates in a mixed-mode
(mode I and II) along the discontinuity (Hutchinson, 1996; Xu et al.,
2003; Wang and Xu, 2006), it follows from Eq. (2) that deection
into the discontinuity occurs if:
G
d
=
1 n
2
K
2
I
K
2
II
= G
D
(4)
where the stress-intensity factors K
I
K
II
now refer to the discon-
tinuity. From Eqs. (3) and (4), the extension fracture penetrates the
discontinuity if:
G
d
G
p
<
G
D
G
L
(5)
but becomes deected into the discontinuity if:
G
d
G
p
_
G
D
G
L
(6)
Equations (3)(6) are likely to control, partly at least, whether
a fracture penetrates or becomes deected along a discontinuity,
such as a fracture or a contact, in some fault zones.
When there is an abrupt change in the mechanical properties at
interfaces such as contacts or discontinuities (Figs. 25, 8), an
elastic mismatch, the assumption of the rock layers on either side of
the interface being with the same properties is not warranted. The
magnitude of the mechanical change across a discontinuity or an
interface is commonly indicated by the Dundurs (1969) elastic
mismatch parameters. The two Dundurs parameters, a and b, may
be given as (cf. He and Hutchinson, 1989; Hutchinson, 1996; Freund
and Suresh, 2003):
a =
E
*
1
E
*
2
E
*
1
E
*
2
(7)
b =
1
2
m
1
(1 2n
2
) m
2
(1 2n
1
)
m
1
(1 n
2
) m
2
(1 n
1
)
(8)
where m is shear modulus, n is Poissons ratio, and the plain strain
Youngs modulus is E
*
=E/(1 n
2
). The subscript 2 is used for the
modulus of the rock hosting the fracture and subscript 1 for the
material on the other side (the far side with respect to the fracture
tip) of the discontinuity. Generally, a is a measure of mismatch in
the extensional or uniaxial stiffness and b in the volumetric or areal
stiffness (Freund and Suresh, 2003).
The strain energy release rate associated with fracture pene-
tration into the layer above the discontinuity, G
p
, is given by (He
and Hutchinson, 1989; He et al., 1994):
Fig. 13. View north, a section of limestone and shale layers in the Bristol Channel in
Wales. The joints (extension fractures, many with mineral llings) differ in trends
between the limestone layers, as indicated. For example, where the person is sitting, the
joints in the layer at her feet differ by about 20
K
2
I
K
2
II
.
4 cos h
2
p3
(10)
with
K
2
I
K
2
II
= k
2
1
a
12l
h
[d[
2
[e[
2
2R
e
(de)
i
(11)
where d and e are non-dimensional complex-valued functions that
depend on the Dundurs parameters. The ratio G
d
/G
p
is independent
of k
1
as well as the fracture-segment length a (Fig. 14) and is given
by (He and Hutchinson, 1989):
G
d
G
p
=
1 b
2
1 a
[d[
2
[e[
2
2R
e
(de)
c
2
(12)
By analogy with Eqs. (5) and (6), the fracture is likely to pene-
trate the discontinuity or interface between the dissimilar layers if:
G
d
G
p
<
G
D
(j)
G
1
L
(13)
but more likely to become deected into the discontinuity (and
often arrested) if:
G
d
G
p
_
G
D
(j)
G
1
L
(14)
the subscript for the material toughness being for layer 1 (Figs. 14
and 17) and j is a measure of the relative proportion of mode II to
mode I, namely, j =tan
1
(K
II
/K
I
) so that j =0
, a s90
, or a s180
(Withjack
and Jamison, 1986). Furthermore, when a s90
, the maximum
extension direction and displacement direction differ: the
maximum extension direction lies midway between the displace-
ment direction and the normal to the trend of the deformation zone
(see Withjack and Jamison (1986) for details). Previous models of
oblique extension (e.g., Withjack and Jamison, 1986; Tron and Brun,
1991; Clifton et al., 2000) have shown that only normal faults form
when 45
a 135
and a >135
, normal,
oblique-slip and/or strike-slip faults develop.
All models inthis study have two phases of deformation(Fig. 2b).
During the rst phase, the mobile sheet moves outward at a rate of
4 cmh
1
(1.1 10
3
cms
1
) in a prescribed direction (a
1
45
) for
a prescribeddisplacement (3.5 cm). Inresponse, a pervasive(but not
continuous) fabric consisting of normal faults develops throughout
the deformation zone in the clay layer. During the second phase, the
mobile plate moves outward at a rate of 4 cmh
1
in a different
prescribed direction (a
2
135
, 120
, 105
, or 90
to 45
in our
models (Figs. 1, 2), whereas the angle between the rst- and second-
phase extension directions varies from 45
to 22.5
(Fig. 2b).
2.3. Model analysis
Photographs of the top surface of the clay layer, taken at regular
time intervals, record the surface deformation through time during
both phases of extension. To exclude edge effects, we analyze only
the central part of the top surface of the deformation zone. Offsets
of supercial linear markers on the top surface of the clay layer
indicate the sense of slip on faults during both phases of extension
(Fig. 3a). We conrm the sense of slip by observing corrugations on
the fault surfaces at the end of each experiment (Fig. 3b). Corru-
gations are grooves on fault surfaces that parallel the slip direction
(Granger, 2006; Granger et al., 2006, Granger et al., 2008). To
determine the numbers and orientations of fault segments, we t
straight lines to individual fault segments (Fig. 3c, d). Fitting lines to
fault segments, rather than drawing lines from fault tip to fault tip,
captures all segment orientations in areas where second-phase
faults have linked with rst-phase faults. To determine whether
deformation patterns changed signicantly with depth, we also
examined the bottom surface of a dried model with boundary
conditions identical to Model 1 (a
1
45
and a
2
135
).
Wet clay
Silicone polymer
3.5 cm 4 cm
8 cm Fixed rigid
sheet
Mobile rigid
sheet
Rubber sheet
a
60 cm
6
8
c
m
8 cm rubber sheet
Fixed
rigid
sheet
Mobile
rigid
sheet
b
Displacement
direction
Trend of
deformation zone
Fig. 1. Experimental set-up in (a) plan view and (b) cross section. Plan view set-up
shows denition of a (based on Withjack and Jamison, 1986).
A.A. Henza et al. / Journal of Structural Geology 32 (2010) 16561667 1657
3. Experimental observations
3.1. Large angle between extension directions
Model 1 has the maximum possible angular difference
between displacement directions (90
), many of the
rst-phase normal faults are reactivated with oblique slip (right-
lateral and normal components) (Fig. 4a
2
). The strike-slip compo-
nent is signicantly larger than the dip-slip component (Table 1).
New normal faults also develop during the second phase of
Fig. 4. Photographs of the top surface of the clay layer showing central region of the models after rst and second phases of extension. Arrows indicate extension direction, and
dashed box shows location of line drawings in Fig. 5. Fault scarps dipping toward the top of the page appear bright; fault scarps dipping toward the bottom of the page appear dark.
A.A. Henza et al. / Journal of Structural Geology 32 (2010) 16561667 1659
extension (Figs. 4a
2
, 5a
2
). Most new normal faults initiate at rst-
phase faults and propagate away from them (Fig. 6). The strike of
the new normal faults ranges from orthogonal to oblique to the
second-phase extension direction, and many new faults change
orientation during fault propagation. These faults are initially
orthogonal to the pre-existing fault, and become approximately
orthogonal to the second-phase extension direction as they prop-
agate away fromthe pre-existing fault (Fig. 6). Where a newnormal
fault encounters a pre-existing fault, the new fault either cuts
across the pre-existing fault or terminates against it (Fig. 7). In
Model 1, more newnormal faults cut across pre-existing faults (61%
Table 1
Fault-population statistics for second-phase deformation.
Model
number
Angle between extension
directions during 1st and
2nd phases
% of Fault population
that form during
2nd phase
% of Dip-slip (normal)
motion on reactivated
faults during 2nd phase
1 45
42 43
2 37.5
33 50
3 30
27 65
4 22.5
12 70
Fig. 5. Summary of modeling results. Line drawings of fault heaves are from the central region of the models (locations shown in Fig. 4). Rose diagrams show orientations of fault
segments obtained using the method shown in Fig. 3. Arrows on rose diagrams show extension directions for each phase. Bin size for rose diagrams is 5
and 30
in the rst-phase s
1
/
s
3
reference frame, have apparent dips between 51
(Model 1) and
58
1
2
3
4
51 54 56 58
102 108 112 116
51 54
56
58
45 37.5 30 22.5
100 200 300 400 500 600
100
200
n
New faults during 1
st
phase
Reactivated faults during 2
nd
phase
(number indicates model designation)
c
(Pa)
(Pa)
a
b
Pre-existing fault surface
Cross sections in (a)
= C +
n
=
s n
1
2
3
4
(see b)
* is the angle between and the normal to the fault plane
1
3
51
3
54
1
3
56
1
3
58
1
Model 1 Model 2 Model 3 Model 4
2
Cross section parallel to
3
during 2
nd
phase
Angle between
extension directions
Apparent dip of reactivated
frame during 2
nd
phase
faults in
1
/
3
reference
2 = 102
(faulting)
(frictional sliding)
Fig. 10. (a) Stress regime for reactivated faults at the beginning of the second phase of extension for all models. (b) Block diagram showing orientation of cross sections shown in (a).
(c) Mohr-circle diagram for stress state in the experimental models at the base of the 3.5-cm thick clay layer. Numbers 14 refer to the four different models. Black line shows the
MohrCoulomb failure envelope for wet clay, and the grey line shows the assumed frictional-sliding failure envelope for wet clay. The two failure envelopes differ only by the value
of cohesion.
A.A. Henza et al. / Journal of Structural Geology 32 (2010) 16561667 1663
and higher to be reactivated in our models (Fig. 10c). This agrees
with the observations that all rst-phase faults are reactivated as
oblique-slip faults during the second phase of extension in our
models.
4.3. Comparison to previous modeling studies
Bonini et al. (1997) and Keep and McClay (1997) simulated two
non-coaxial phases of extension using a layer of dry sand overlying
a basal layer of silicone polymer. Their models differ from our
models in terms of the modeling material (dry sand vs. wet clay)
and the width of the basal layer of silicone polymer (the entire base
of the model vs. a narrow zone). Despite these differences, the
modeling results are qualitatively similar. Their models show that
faults that form during the rst phase of extension inuence the
fault patterns that develop during the second phase of extension.
Specically, the reactivation of the rst-phase faults during the
second phase, the development of new second-phase faults, and
the slip and attitude of the second-phase faults depend on the
angle between the rst- and second-phase extension directions.
Reactivation is more likely when the angle is small, whereas new
fault development is more likely when the angle is large. Our
models support these conclusions. In addition, our models show
that the angle between the rst- and second-phase extension
directions controls the relative components of dip slip and strike
slip on the reactivated faults and the number of new faults that
form during the second phase of extension. Furthermore, our
models allow for the observation of small-scale features (such as
the interaction between pre-existing faults and newfaults) because
faults are more numerous, have smaller displacements, and are
more closely spaced in clay models than in sand models (e.g.,
Withjack et al., 2007). Our work, in combination with that of Bonini
et al. (1997) and Keep and McClay (1997), illustrates that a pre-
existing fault fabric substantially inuences the fault patterns that
develop during subsequent episodes of extension.
4.4. Natural examples of multi-phase extension
The fault interactions in our models are similar to those
observed in nature. In the Terra Nova region of the Jeanne dArc rift
basin of offshore Newfoundland, two main fault orientations (NS
and EW) are present (e.g., Enachescu, 1987; Sinclair, 1995a, b)
(Fig. 11a). Movement on N-striking normal faults occurred during
the Late Jurassic through the early Early Cretaceous, whereas
movement on E-striking normal faults occurred during the late
Early Cretaceous (Sinclair, 1995a). Well and 3D seismic-reection
data show that some of the younger, E-striking normal faults cut
and offset the older, N-striking normal faults, whereas others
terminate against the older, N-striking normal faults (McIntyre
et al., 2004) (Fig. 11b, c). The presence of both types of fault inter-
actions in the Jeanne dArc rift basin matches our experimental
observations that both types of fault interactions are likely to occur
with multiple phases of non-coaxial extension. In addition, in many
parts of the Jeanne dArc basin, displacement on the younger,
E-striking normal faults is greatest adjacent to the older, N-striking
normal faults (Fig. 11b). This displacement variation is common in
our models where new faults initiate at pre-existing faults and
propagate outward (Fig. 6). The modeling results suggest that the
5 km
a
b
Pre-existing faults
New faults
Atlantic
Ocean
c
c
b
Newfoundland
N
Fig. 11. (a) Map of the Terra Nova region of the Jeanne dArc rift basin, offshore Newfoundland, showing faults offsetting the B marker (Early Cretaceous). Map modied from
McIntyre et al. (2004). Insert shows location of the Jeanne dArc basin (box) relative to Newfoundland, Canada. (b) Younger, E-striking normal faults terminating against older, N-
striking normal faults. (c) Younger, ESE-striking normal faults cutting and offsetting older, N-striking normal faults. Locations are shown in (a).
A.A. Henza et al. / Journal of Structural Geology 32 (2010) 16561667 1664
younger, E-striking normal faults initiated at and propagated
outward from the older, N-striking normal faults.
Our modeling results also agree with natural observations of
normal-fault reactivation. Studies of abyssal-hill normal faults in
subduction zones (e.g., Masson, 1991; Moritera-Guitie rrez et al.,
2003; Billen et al., 2007) show that the formation of new outer-
slope normal faults (which form to accommodate bending-induced
extension of subducting plates) depends on the orientation of the
pre-existing abyssal-hill normal faults relative to the orientation of
the trench axis. Abyssal-hill normal faults format spreading centers
to accommodate extension of newly forming crust, becoming
permanent features of the oceanic crust (e.g., Rea, 1975; Kriner
et al., 2006; and references therein). Moritera-Guitie rrez et al.
(2003) and Billen et al. (2007) show that, if the angle between the
trench axis and the strike of the pre-existing abyssal-hill faults is
less than 25
to 22.5
, the strain
accommodation changes from fault reactivation and new normal-
fault formation (Model 3) to mainly fault reactivation (Model 4).
This transition, however, is gradual, occurring between 30
and
22.5
. (b) Abyssal-hill
normal faults reactivate and no new outer-slope normal faults form if the angle
between the trench axis and the strike of the abyssal-hill normal faults is less than 25
.
Modied from Billen et al. (2007).
2 cm
Inferred extension directions
Actual extension directions
45
75
Fig. 13. Line drawing of the top surface of the clay layer of Model 1 after both phases of
extension (location shown in Fig. 4). Arrows indicate inferred extension directions
determined by assuming that extension is orthogonal to the strike of the normal faults.
The angle between the inferred extension directions is 75
.
A.A. Henza et al. / Journal of Structural Geology 32 (2010) 16561667 1665
The angle between the strike of the pre-existing normal faults
and the second-phase extension direction controls the sense of
slip on the reactivated faults, with the component of dip slip
relative to strike slip decreasing as the angle between the rst-
and second-phase extension directions increases.
New normal faults form during the second phase of extension
in all models. These faults are generally shorter than faults that
formin models with no pre-existing fault fabric. The number of
new fault segments varies between models, with more new
fault segments forming as the angle between the rst- and
second-phase extension directions increases.
The orientations of new normal-fault segments are both
orthogonal and oblique to the second-phase extension direc-
tion. New normal faults with strikes oblique to the second-
phase extension direction likely reect local perturbations in
the stress/strain state near pre-existing faults.
New normal faults commonly initiate at pre-existing normal
faults and propagate outward. Displacement on these new
normal faults is greatest adjacent to the pre-existing faults.
Where new normal faults encounter pre-existing faults, the
newfaults either cut across the pre-existing faults or terminate
against them.
Fault interactions and fault reactivations in the models are
similar to those observed in nature, including the Jeanne dArc
rift basin of offshore Newfoundland and at outer highs near
subduction zones.
Interpretations of extensional histories based solely on fault
orientations and fault interactions represent only one of many
possible deformation histories. Additional information, such as
fault slip and the timing of fault activity (i.e., the presence or
absence of growth beds), is necessary to constrain the paleo-
strain state.
Acknowledgments
We thank our colleagues Michael Durcanin, Iain Sinclair, and
Judith McIntyre for valuable discussions and insights about
modeling and fault interactions and Hemal Vora for his assistance
in the laboratory. We also thank Marco Bonini, Chris Jackson, and
Bruno Vendeville for their detailed and helpful reviews of the
manuscript. We gratefully acknowledge the support of the Struc-
tural Modeling Laboratory at Rutgers University by the National
Science Foundation (EAR-0838462 and EAR-0408878), Husky
Energy Inc., and Petrobras S.A.
References
Anderson, E.M., 1951. The Dynamics of Faulting and Dyke Formation with Appli-
cations to Britain. Oliver & Boyd, Edinburgh.
Badley, M.E., Price, J.D., Dahl, C.R., Agdestein, T., 1988. The structural evolution of the
northern Viking Graben and its bearing upon extensional models of basin
formation. Journal of the Geological Society, London 145, 455472.
Bartholomew, I.D., Peters, J.M., Powell, C.M., 1993. Regional structural evolution of
the North Sea: oblique slip and the reactivation of basement lineaments. In:
Parker, J.R. (Ed.), Petroleum Geology of Northwest Europe: Proceedings of the
Fourth Conference. The Geological Society, London, pp. 11091122.
Bellahsen, N., Daniel, J.M., Bollinger, L., Burov, E., 2003. Inuence of viscous layers on
the growth of normal faults: insights from experimental and numerical models.
Journal of Structural Geology 25, 14711485.
Billen, M., Cowgill, E., Buer, E., 2007. Determination of fault friction from reac-
tivation of abyssal-hill faults in subduction zones. Geology 35, 819822.
Bonini, M., Souriot, T., Boccaletti, M., Brun, J.P., 1997. Successive orthogonal and
oblique extension episodes in a rift zone: laboratory experiments with appli-
cation to the Ethiopian Rift. Tectonics 16, 347362.
Brace, W.F., Kohlstedt, D.L., 1980. Limits on lithospheric stress imposed by labora-
tory experiments. Journal of Geophysical Research 85, 62486252.
Byerlee, J., 1978. Friction of rocks. Pure and Applied Geophysics 116, 615626.
Clifton, A.E., Schlische, R.W., Withjack, M.O., Ackermann, R.V., 2000. Inuence of rift
obliquity on fault-population systematics: results of clay modeling experi-
ments. Journal of Structural Geology 22, 14911509.
Destro, N., 1995. Release fault: a variety of cross fault in linked extensional fault
systems, in the Sergipe-Alagoas Basin, NE Brazil. Journal of Structural Geology
17, 615629.
Eisenstadt, G., Sims, D., 2005. Evaluating sand and clay models: do rheological
differences matter? Journal of Structural Geology 27, 13991412.
Enachescu, M.E., 1987. Tectonic and structural framework of the northeast
Newfoundland continental margin. In: Beaumont, C., Tankard, A.J. (Eds.),
Sedimentary Basins and Basin-Forming Mechanisms. Canadian Society of
Petroleum Geologists Memoir 12, 117146.
Frseth, R.B., 1996. Interaction of Permo-Triassic and Jurassic extensional fault-
blocks during the development of the northern North Sea. Journal of the
Geological Society, London 153, 931944.
Granger, A.B., 2006. Inuence of basal boundary conditions on normal-fault systems
in scaled physical models. Masters thesis, Rutgers University.
Granger, A.B., Withjack, M.O., Schlische, R.W., 2006. Undulations on normal-fault
surfaces: insights into fault growth using scaled physical models of extension.
Geological Society of America Abstracts with Program 38, p. 480.
Granger, A.B., Withjack, M.O., Schlische, R.W., September 2008. Fault surface
corrugations: insights from scaled experimental models of extension. In: Fault
Zones: Structure, Geomechanics, and Fluid Flow Conference, Abstracts
Volume. Geological Society of London. 38.
Handin, 1966. Strength and ductility. In: Clark, S.P. (Ed.), Handbook of Physical
Constants, 97. Geological Society of America Memoir, pp. 233289.
Homberg, C., Hu, J.C., Angelier, J., Bergerat, F., Lacombe, O., 1997. Characterization of
stress perturbations near major fault zones: insights from 2D distinct-element
numerical modelling and eld studies (Jura mountains). Journal of Structural
Geology 19, 703718.
Homberg, C., Angelier, J., Bergerat, F., Lacombe, O., 2004. Using stress deections to
identify slip events in fault systems. Earth and Planetary Science Letters 217,
409424.
Hubbert, M.K., 1937. Theory of scale models as applied to the study of geological
structures. Geological Society of America Bulletin 48, 14591519.
Kattenhorn, S.A., Aydin, A., Pollard, D.D., 2000. Joints at high angles to normal fault
strike: an explanation using 3-D numerical models of fault-perturbed stress
elds. Journal of Structural Geology 22, 123.
Keep, M., McClay, K.R., 1997. Analogue modeling of multiphase rift systems. Tec-
tonophysics 273, 239270.
Kriner, K.K., Pockalny, R.A., Larson, R.L., 2006. Bathymetric gradients of lineated
abyssal hills: inferring seaoor spreading vectors and a new model for hills
formed at ultra-fast rates. Earth and Planetary Science Letters 242, 98110.
Masson, D.G., 1991. Fault patterns at outer trench walls. Marine Geophysical
Researches 13, 209225.
McIntyre, J., DeSilva, N., Thompson, T., 2004. Mapping of key geological markers in
the Jeanne dArc basin based on 3-D seismic. Canadian Society of Petroleum
Geologists Annual Meeting.
McClay, K.R., White, M.J., 1995. Analogue modeling of orthogonal and oblique rift-
ing. Marine and Petroleum Geology 12, 137151.
Moritera-Guitie rrez, C.A., Scholl, D.W., Carlso, R.L., 2003. Fault trends on the
seaward slope of the Aleutian Trench: implications for a laterally changing
stress eld tied to a westward increase in oblique convergence. Journal of
Geophysical Research 108. doi:10.1029/2001JB001433.
Morley, C.K., Gabdi, S., Seusutthiya, K., 2007. Fault superimposition and linkage
resulting from stress changes during rifting: examples from 3D seismic data,
Phitsanulok Basin, Thailand. Journal of Structural Geology 29, 646663.
Morley, C.K., Harayana, C., Phoosongsee, W., Pongwapee, S., Kornsawan, A.,
Wonganan, N., 2004. Activation of rift oblique and rift parallel pre-existing
fabrics during extension and their effect on deformation style: examples from
the rifts of Thailand. Journal of Structural Geology 26, 18031829.
Nalpas, T., Brun, J.P., 1993. Salt ow and diapirism related to extension at crustal
scale. Tectonophysics 228, 349362.
Rea, D.K., 1975. Model for the formation of topographic features of the East Pacic
Rise Crest. Geology 3, 7780.
Renard, V., Nakamure, K., Angelier, J., Azema, J., Bourgois, J., Deplus, C., Fujioka, K.,
Hamano, Y., Huchon, P., Kinoshita, H., Labaume, P., Ogawa, Y., Seno, T.,
Takeuchi, A., Tanahashi, M., Uchiyama, A., Vigneresse, J.L., 1987. Trench triple
junction off Central Japan preliminary results of the FrenchJapanese 1984
Kaiko cruise, Leg 2. Earth and Planetary Science Letters 83, 243256.
Schellart, W.P., 2000. Shear test results for cohesion and friction coefcients for
different granular materials: scaling implications for their usage in analogue
modeling. Tectonophysics 324, 116.
Sinclair, I.K., 1995a. Transpressional inversion due to episodic rotation of exten-
sional stresses in Jeanne dArc Basin, offshore Newfoundland. In: Buchanan, J.G.,
Buchanan, P.G. (Eds.), Basin Inversion. Geological Society Special Publication 88,
pp. 249271.
Sinclair, I.K., 1995b. Sequence stratigraphic response to AptianAlbian rifting in
conjugate margin basins: a comparison of the Jeanne dArc Basin, offshore
Newfoundland, and the Porcupine Basin, offshore Ireland. In: Scrutton, R.A.,
Stoker, M.S., Shimmield, G.B., Tudhope, A.W. (Eds.), The Tectonics, Sedimenta-
tion, and Palaeoceanography of the North Atlantic Region. Geological Society
Special Publication 90, pp. 2949.
Sinclair, I.K., Withjack, M.O., 2008. Mid to Late Cretaceous structural and sedi-
mentary architecture at the Terra Nova oileld, offshore Newfoundland
implications for tectonic history of the North Atlantic. In: Brown, D.E. (Ed.),
Central Atlantic Conjugate Margins. Dalhousie University, Halifax, Nova Scotia,
pp. 125141.
A.A. Henza et al. / Journal of Structural Geology 32 (2010) 16561667 1666
Strecker, M.R., Blisniuk, P.M., Eisbacher, G.H., 1990. Rotation of extension direction
in the central Kenyan Rift. Geology 18, 299302.
Tankard, A.J., Welsink, H.J., 1987. Extensional tectonics and stratigraphy of Hibernia
oil eld, Grand Banks, Newfoundland. AAPG Bulletin 71, 12101232.
Tron, V., Brun, J.P., 1991. Experiments on oblique rifting in brittle-ductile systems.
Tectonophysics 188, 7184.
Weijermars, R., 1986. Flow behaviour and physical chemistry of bouncing putties
and related polymers in view of tectonic laboratory applications. Tectonophy-
sics 124, 325358.
Weijermars, R., Jackson, M.P.A., Vendeville, B.C., 1993. Rheological and tectonic
modeling of salt provinces. Tectonophysics 217, 143174.
Withjack, M.O., Callaway, J.S., 2000. Active normal faulting beneath a salt layer: An
experimental study of deformation in the cover sequence. AAPG Bulletin 84,
627652.
Withjack, M.O., Jamison, W.R., 1986. Deformation produced by oblique rifting.
Tectonophysics 126, 99124.
Withjack, M.O., Schlische, R.W., 2006. Geometric and experimental models of
extensional fault-bend folds. In: Buiter, S.J.H., Schreurs, G. (Eds.), Analogue and
Numerical Modelling of Crustal-Scale Processes. Geological Society (London)
Special Publication 253, pp. 285305.
Withjack, M.O., Schlische, R.W., Henza, A.A., 2007. Scaled experimental models of
extension: dry sand vs. wet clay. Houston Geological Survey Bulletin 49 (8), 3149.
A.A. Henza et al. / Journal of Structural Geology 32 (2010) 16561667 1667
Using empirical geological rules to reduce structural uncertainty in seismic
interpretation of faults
Brett Freeman
a,
*
, Peter J. Boult
b
, Graham Yielding
a
, Sandy Menpes
c
a
Badley Geoscience Ltd., North Beck House, North Beck Lane, Hundleby, Lincolnshire PE23 5NB, UK
b
GINKGO ENPGNG, 57, Seventh Avenue, St. Morris SA 5068, Australia
c
Palaeosearch, 41 Walker Avenue, Heatheld SA 5153, Australia
a r t i c l e i n f o
Article history:
Received 3 February 2009
Received in revised form
21 October 2009
Accepted 2 November 2009
Available online 23 November 2009
Keywords:
Normal fault
Displacement
Displacement gradient
Wall-rock strain
Seismic interpretation
a b s t r a c t
Good seismic interpretation of faults should include a workow that checks the interpretation against
known structural properties of fault systems. Estimates of wall-rock strains provide one objective means
for discriminating between correct and incorrect structural interpretations of 2D and 3D seismic data
implied wall-rock strain should be below a geologically plausible maximum. We call this the strain
minimisation approach. Drawing on the large body of published data for strike dimension and maximum
displacement for faults we suggest a realistic upper limit of wall-rock shear strain of 0.05, and 0.1 for
maximum longitudinal strain when measured in the displacement direction. Small-scale variation of
fault wall-rock strain also adheres to this rule, except in specic areas of strain localisation such as relay
zones. As a case study we review an existing structural interpretation of 2D seismic surveys. Mapping of
shear and longitudinal strain on the fault planes show values commonly greater than 0.05 and 0.1
respectively. Thus the model is deemed inadmissible. We then reinterpreted the area in an iterative
manner using the strain minimisation approach. By using regions of implied high wall-rock strain as an
indicator of high uncertainty in the interpretation, we were able to break out two self-consistent fault
sets, each of which had geologically plausible wall-rock strains, where previously there had only been
one fault set with highly implausible wall-rock strains.
2009 Elsevier Ltd. All rights reserved.
1. Introduction
It has been established for more than twenty years that the
displacement on geological fault surfaces varies in a smooth,
continuous and consistent manner. Rippon (1985) and Barnett et al.
(1987) rst demonstrated this for isolated normal faults from the
English coalelds. They used precise survey data from coal mine
plans to measure the throw (vertical component of dip separation)
for a number of coal seams at varying spatial locations. When the
throwvalues were plotted on a strike projection of the fault surface,
the contours of throw were concentric and sub-parallel, with
a maximumthrowclose to the centre of the fault surface. Moreover
the boundary, or tip, of the fault surface was approximately ellip-
tical. These important observations have stimulated an enormous
amount of research into the form and scaling relationships of
displacement distributions, the quantitative systematics of fault
geometry and speculation on fault growth mechanisms. The
simplicity of the observations implies both a consistency and a limit
to the strain in the wall rocks. Barnett et al. (1987), Bouvier et al.
(1989) (normal faults) and then Chapman and Meneilly (1991)
(reactivated normal fault with net reverse displacement) demon-
strated similar patterns from seismic interpretation. Although the
precision of the structural information from seismic data is
considerably poorer than the surveyed data of Rippon (1985), these
early examples of displacement distributions are also characteris-
tically continuous and smooth.
A priori knowledge of the shape/form of the displacement
distribution and its gradients can be useful as an aid to seismic
interpretation. Barnett et al. (1987) suggested that it could serve
both as a quality control metric and a means to predict quantita-
tively the location of lithological layers (horizons) and faults where
data is limited. In other words it can be used to manage interpre-
tation uncertainty. In faulted reservoirs, structural uncertainty
arises from two major sources of error: the systematic error of the
seismic method, and the human error of the interpreter. For good
quality 3D seismic data the order of error in lateral positioning of
structures is approximately the same as the error in the vertical
dimension and both are dominantly systematic. However, when
the spacing between fault traces from line samples (e.g. seismic
lines) is closer than the spacing between the line samples
* Corresponding author. Tel.: 44 (0) 1754 890390; fax: 44 (0) 1790 753527.
E-mail address: brett@badleys.co.uk (B. Freeman).
Contents lists available at ScienceDirect
Journal of Structural Geology
j ournal homepage: www. el sevi er. com/ l ocat e/ j sg
0191-8141/$ see front matter 2009 Elsevier Ltd. All rights reserved.
doi:10.1016/j.jsg.2009.11.001
Journal of Structural Geology 32 (2010) 16681676
themselves, the lateral correlation of faults is equivocal (e.g.
Freeman et al., 1990). So for 2D seismic data, reconnaissance
mapping, poor quality 3D seismic data and for small faults in 3D
seismic data, the pattern of faulting becomes a serious interpretive
issue. The balance of the error, or uncertainty, is then strongly one-
sided and the effects of systematic errors become secondary to
those inherent in the interpreters model. Freeman et al. (1990)
introduced a methodology that used displacement patterns to
distinguish likely fault plane correlations from possible and
impossible correlations. In a similar vein, Needham et al. (1996)
showed howthis kind of analysis was valuable for validating three-
dimensional structural models. Traditionally the analytical part of
the process has taken the form of visual inspection of the throw
contours. If the resulting pattern is smooth and continuous, the
fault may be judged as a valid interpretation, otherwise the
correlation is deemed to be suspect. Although ostensibly objective,
the effectiveness of the above approach is dependent on the skill or
experience of the interpreter in being able to identify bad contour
patterns. We know of no published work that actually quanties
what constitutes either a good or a bad contour pattern. In this
paper we suggest that the above basic validation procedure can be
improved by (1) quantifying the strains that are implied by the
contour patterns and (2) setting out reasonable limits for the
magnitudes of these strains.
We show that there is a simple relationship between strain and
the displacement gradient. Drawing from a large database of pub-
lished information on the shapes of displacement proles and the
scaling relationship between displacement and fault dimension, we
can suggest reasonable limits on the amount of strain that is
admissible in the walls of a fault. Implied strains that exceed these
empirical limits indicate aws in the structural model. The result-
ing strategy for interpretation leads to a structural model that
minimizes the strain attributable to faulting.
As an example we show how an analysis of a 2D seismic inter-
pretation from South Australia consistently implies erratic and
unrealistically large strains. An iterative structural reinterpretation
using our minimum strain approach provides a solution that is
geologically more feasible.
2. Displacement and wall-rock strain
There is a simple relationship between the displacement gradient
and the strain of the wall rocks in the plane parallel to the fault.
Fig. 1a shows the deformation associated with the faulting of a pre-
existing uniform horizontal layer (i.e. the fault is not a growth fault).
The element, E (Fig. 1a and b), is dened by the position of the layer
in the undeformed state with the top of the layer at p and the base of
the layer at q (Fig. 1b). For the sake of argument we assume that
displacement is in the dip direction of the fault, and that strain is
partitioned equally in the two walls of the fault. The layer is then
faulted such that, inthe dip direction (parallel to y), p moves to p
0
and
q moves to q
0
. The stretch in the dip direction is then
1 e
u
q
q
u
p
p
q p
(1)
where e is the unit extension, u
p
and u
q
are the absolute displace-
ments for one side of the fault (half the total, relative displace-
ment). This can be re-written as
1 e
1
1
2
Du
Dy
(2)
where the factor of 1/2 means that u refers to the total relative
displacement across the fault. At the limit as Dy approaches zero
1 e
1
1
2
vu
vy
(3)
In other words the unit extension is equal to half the displacement
gradient. Using an alternative formulation it is easy to show
that the stretch in the upthrown layer is the reciprocal of the stretch
in the downthrown layer and that the undeformed layer thickness
is the average of the upthrown and downthrown thicknesses
(cf. gure 1 from Barnett et al., 1987).
Referring back to Fig. 1b we can also see that, for each wall of the
fault, the strain g for shear in the dip direction is given by
g
u
s
u
p
=s p
1
2
Du
Dx
(4)
then as Dx approaches zero
g
1
2
vu
vx
(5)
Eqs. (3) and (5) are useful results because (1) they are independent
of the form of the displacement distribution, and (2) they give us
a direct way to measure and represent strain from information that
is almost universally available from seismic interpretations.
If we can place realistic limits on the strain values, we then have
a method for distinguishing between good and bad fault interpre-
tation that is entirely quantitative and objective.
Fig. 1. (a) Schematic of an idealized fault plane (strike dimension L, dip dimension L/2)
showing the absolute displacements from a horizontal, unfaulted layer, to the faulted,
upthrown and downthrown positions. The element E in the unfaulted state is trans-
lated and strained to E
0
. (b) Analysis of the change in shape of the rectangular element
E. x is the strike direction of the fault and y is the dip direction, u
p
, u
q
and u
s
are
absolute, dipslip displacements.
B. Freeman et al. / Journal of Structural Geology 32 (2010) 16681676 1669
2.1. Ideal displacement patterns for unrestricted faults
Although direct measurements of wall-rock strains is probably
beyond normal eld techniques, the simple fact that unrestricted
faults have tip lines means that the wall rocks must be differentially
strained. Eshelby (1957) and Pollard and Segall (1987) suggest the
slip on a dislocation in a linear elastic solid is characterised by
a semi-elliptical slip prole. In other words, a straight marker line
in a wall of a fault and initially perpendicular to the slip direction
will have a deformed shape of a semi-ellipse and reects directly,
the differential wall-rock strains. This type of slip prole equates to
a single earthquake event. However, Nicol et al. (1996) and many
others show that natural examples of unrestricted faults have
approximately linear normalized proles i.e. triangular. Further-
more Manzocchi et al. (2006) argue that this feature of geological
faults seems to hold irrespective of the growth mechanism or the
form of the slip prole for an individual event. This is a convenient
conclusion because it means the gradient of displacement on an
unrestricted fault surface is approximately constant.
2.2. Limits on displacement
Various compilations of data for D
max
/L (D
max
is the maximum
displacement, L is the strike dimension of a fault) have been pub-
lished (e.g. Bailey et al., 2005; Kim and Sanderson, 2005; Schultz
et al., 2008). Although there remains debate about the exact nature
of the power-law distribution of D
max
vs L, it seems that 0.1 repre-
sents the naturally observed upper bound for all types of faults over
all measured scales (Fig. 2). However, if we focus our attention on
the scale range imaged on seismic data, we can rene the limit to
0.05 (Fig. 2). Then if we assume for an unrestricted fault that the
displacement prole is triangular, (1/2) (D
max
/2)/(L/2) 0.05 places
a natural limit of 0.05 on the shear strain in each wall.
Unfortunately there is no similar database for the relationship
between displacement and the dip dimension of a fault. In this
respect we make a further assumption that the aspect ratio of our
unrestricted fault is 2 (e.g. Nicol et al., 1996), then (1/2) (D
max
/2)/
(L/4) 0.1 represents the limit of the longitudinal strain in each
wall. These suggested limits for shear and longitudinal strain are
consistent with detailed measurements of coaleld fault displace-
ment gradients by Walsh and Watterson (1989). For tip restricted or
half restricted faults the displacement prole is steepened towards
the tips (Nicol et al., 1996). Potentially this could increase the shear
strain by a factor of two or more.
3. Method
All the analysis has been performed using the TrapTester
software (www.badleys.co.uk/products/traptester.htm). Fault
planes have been generated as triangulated meshes from the
vertices of fault sticks picked on seismic sections. The horizon
cutoffs at the faults are calculated from seismic interpretation of
the horizons. It is rare for seismic horizon picks to tie exactly
with the fault picks and therefore some extrapolation is required
to make the cutoff lines. In fact we make three-dimensional
surface models of both the upthrown and downthrown sides of
the fault. Each of these models is extrapolated so that it extends
beyond the fault in both directions and the cutoff is calculated as
the intersection between the fault surface and the horizon
surface model. If faults are joined at branch lines, the horizon
surface model is based on the structurally coherent horizon data
that is conned to the appropriate parts of the interpretation as
bounded by all relevant fault planes. The quality of t of the
resulting polygons has been assessed visually and is, in all cases,
a fair representation of where an expert geoscientist might draw
the cutoff by hand.
Immediately prior to analysis all information for a fault plane
is referred to a coordinate frame that is specic to each particular
fault. We call this a natural coordinate system (NCS). The xy
plane (x along strike, y down-dip) of the NCS represents a best t
plane through all the data points that dene the fault plane
topography. The upthrown and downthrown cutoffs form
a polygon and this is the basis for the structural measurements.
We measure the dip separation on a set of sweep lines (constant x)
at an interval of 50 m in the x direction of the NCS. The
raw measurements are interpolated on to a 50 m50 m grid
using the multi-level B-Spline method of Lee et al. (1997). This
interpolation scheme honours the raw data and produces
a smoothly interpolated surface everywhere else. We calculate the
strike and dip gradients of dip separation using a central difference
formula.
It should be noted that the displacement gradient is measured
and recorded midway between the corresponding upthrown and
downthrown layers i.e. at the location in the undeformed state.
However, the implied strains refer to the layers themselves i.e. the
deformed state.
4. Interpretation example from South Australia
4.1. Gambier Embayment, Otway basin
The Otway Basin is a passive margin forming a large part of the
Eastern Australian Rift System that resulted from the separation of
Australia fromAntarctica during the Late Jurassic to Late Cretaceous
(Lovibond et al., 1995). Along with the rest of the Australian margin,
the Otway Basin has a complex Mesozoic to present day structural
history, with multiple rifting events causing principal stress/strain
directions and magnitudes to change repeatedly during its devel-
opment. The Gambier Embayment is a Tertiary sub-basin of the
Otway Basin. It is bounded to the north and east by the Tartwaup
Hinge Zone, to the west by the continental shelf and to the east it
Fig. 2. Compilation of maximum displacement (D
max
) and maximum fault length (L)
adapted from Schultz et al. (2008). The data are contoured in lines of constant D
max
/L.
The 0.05 contour is our putative upper bound for D
max
/L ratio at the range of scales of
normal faults imaged on seismic data.
B. Freeman et al. / Journal of Structural Geology 32 (2010) 16681676 1670
merges (in Victoria) into the Tertiary NWSE trending Portland
Trough.
Within the Gambier Embayment the Late Cretaceous section
comprises deep- to marginal-marine and deltaic sediments up to
2 km thick onshore and over 3 km offshore. The Tertiary section is
up to 1000 mthick in the Gambier Embayment and is up to 2700 m
in the Portland Trough. It comprises uvial to deltaic sediments
overlain by marl and limestone up to 400 m thick in the Gambier
Embayment and 900 m thick in the Portland Trough (Boult, 1999;
Boult and Hibburt, 2002).
4.2. Seismic data
The raw data for this study is based entirely on 2D time-
migrated seismic data with approximate line spacings of 1.5 km for
dip lines and 5 km for strike lines (see Fig. 3a). For the purpose of
piecewise depth conversion we have used an average velocity of
3000 ms
1
over the entire two way time (TWT) interval. There are
a number of different vintages of seismic and the quality of the
reection data is variable. For the most part fault offsets are clearly
imaged in the upper part of the sections but they become less easy
to interpret as the data reaches about 2000 ms. Similarly, seismic
reectors are relatively well imaged in the upper sections becoming
less well imaged at depth. An example of the variability of quality is
illustrated in Fig. 4.
4.3. Initial structural model
Interpretation (Essential Petroleum Resources Ltd., 2006) had
been conducted using a typical industry strategy on 2D seismic
panels and horizon maps and as far as we are aware, made no
deliberate attempt to adhere to the basic rules of displacement
continuity as outlined in the introduction to this paper. The
resulting structural model comprises a set of fault planes and ve
horizons over a TWT range of about 2500 ms. A summary of the
initial model at top reservoir level (Fig. 3b) shows an eastward
shallowing structure dissected by a set of large NWSE trending
normal faults (Essential Petroleum Resources Ltd., 2006). These
faults are spaced at about 2 km, they have maximum throws of
about 200 ms (approximately 300 m) and maximum dip separa-
tions of the order 400 m. Throughout the discussion of these data
we use dip separation as the measure of displacement in the fault
plane. The top reservoir has a TWT range of the order 300 ms about
an absolute TWT of 1900 ms. At this level and deeper, it becomes
more challenging to tie horizons from line to line and across faults.
It is therefore very important to make efcient use of the more
Fig. 3. (a) Base map of the 2D seismic shot lines from the Gambier Embayment, Otway Basin, South Australia. The line with the asterisks is shown in Fig. 4. (b) Prospect map at the
level of top reservoir. Colour code dark to light indicates shallow to deep.
B. Freeman et al. / Journal of Structural Geology 32 (2010) 16681676 1671
reliable structural information above the top reservoir in order to
constrain the fault model at depth. Fig. 5be illustrates that the
fault planes themselves are picked over the full depth range.
However, displacement information (from the horizon cutoffs) is
limited to a relatively small area of the fault planes (approximately
50%) since not all of the ve horizons are picked persistently on all
lines and there were no picks at the top of the fault plane. Again,
Fig. 5be shows the regions where the displacement information is
reliable i.e. in the close vicinity of horizon picks.
Bulk measurements for the faults fall in a geologically sensible
cluster with 0.1 <D
max
/L< 0.001. Because the maximum displace-
ments are small relative to the interpreted fault spacing we might
expect there to be minimal interference between the faults hence
we might also expect simple displacement contour patterns.
Notwithstanding these two observations, displacement mapping
clearly shows that all of the fault planes have erratic contour
patterns (Fig. 5a). They show multiple bulls-eyes (highs and lows)
and exhibit both sub-vertical and sub-horizontal valleys in the
displacement magnitudes. In terms of displacement gradients,
hence strain, all the faults exhibit multiple lateral swings in the sign
of shear strain and longitudinal strain (Fig. 5be). We expect shear
strain to be highest at the tips, positive at the left of the strike
projection and negative at the right. In fact we see the polarities
inverted, locally and globally, and the high shear strains concen-
trated towards the centre of the faults. Equally importantly, nearly
all the faults have areas where the magnitudes of the strains lie
outside our bounding, acceptable threshold values (red and
magenta in Fig. 5be). In general we take the (implied) high shear
strain anomalies to indicate a locationwhere the current fault plane
should either be split in to two separate faults or where two faults
join at a branch line. The anomalous (implied) longitudinal strains
and associated ips in polarity usually indicate that a horizon has
been picked persistently in the wrong part of the waveform on one
or other side of the fault.
4.4. Revised structural model
One of us (PJB) undertook a reinterpretation of the seismic
data with a view to producing (1) a model that was geologically
more acceptable and (2) a model that minimized the implied
strains. We picked the same ve horizons with the exception of
the top reservoir where, instead, we picked a reection 200 ms
above the original. (With sufcient picks the shape of the
displacement distribution should be independent of the actual
horizons that are chosen.) The majority of faults intersect at least
three of the ve horizons. In most cases the upper ends of the
traces (sticks) are observed to be at zero displacement so, in
addition to horizon-based displacement information, the upper
tips have been explicitly assigned displacements of zero.
Although all the original interpreted fault traces were reused
(they received minor lateral shifts on some of the lines) our
reinterpretation identied many more fault traces on each of the
sections. In all we more than doubled the amount of fault
information. A possible reason for the inadequacy of the original
model is that the interpretation strategy was driven by picking
the faults with the largest offsets and then forcing them into
common-trend, lateral, correlations. In fact it is equally impor-
tant to identify the fault traces with small offsets, since a fault
plane can be interpreted as nearing its tip, only if it is correlated
to traces with minimal displacement. In this work we have used
an interpretation strategy that is both iterative and incremental
inasmuch as the structure is validated as it evolves. The proce-
dure is summarized in Boult et al. (2008).
Fig. 6 summarizes the results of the reinterpretation and we
outline the major contrasts below:
(1) The new interpretation has a denser fault pattern. Unlike the
original, we identify two major fault sets: one set in a NWSE
orientation (c.f. Fig. 5a), the other trending, broadly, WNW
ESE.
(2) There are no faults that transect the entire survey area. In
general the new faults have smaller map dimensions than the
originals. We identify isolated, en echelon and linked structures.
(3) The displacement contours are, largely, smooth and continuous
and they are void of bulls-eyes.
(4) In general, the displacement gradients and strain patterns are
also smooth and lack the erratic polarity ips that we see in
Fig. 5. In particular, the implied strains are relatively low. The
majority of surface area on each fault has strain values well
below our upper bounds.
Fig. 4. Example of the 2D seismic data from line bu85-38_r9 (see Fig. 3 for location). The sub-vertical lines are fault segments, the sub-horizontal lines are horizon picks. All
segments and picks are from the revised interpretation. Dashed vertical lines mark the intersection with other seismic lines.
B. Freeman et al. / Journal of Structural Geology 32 (2010) 16681676 1672
(5) Values of longitudinal strains mostly lie in the range
0.1 <e <0.1 and they are tightly clustered around smaller
values. The signs are uniformly positive in the upper parts of
the fault and negative at depth. However, on several of the
faults we see persistent bands of high implied strain that lie
outside our upper limit (red colours in Fig. 6b and c). Initially
this appears to put the interpretation in question but in fact,
the top horizon is picked above an unconformity. This gives the
effect of anomalously high displacement gradients and we take
these high values to indicate that the faults were at least
partially active prior to the unconformity.
(6) The two fault trends (above) can also be distinguished on the
basis of the displacement and longitudinal strain patterns. The
NWSE set displays a symmetry that we associate with
displacement dying out towards both an upper and lower tip
while the WNWENE set has displacement increasing
downwards.
(7) The shear strain patterns show an obvious symmetry with low
values in the central part of the fault, increasing to maximum
values at each lateral tip.
(8) To a large degree the shear strain maps conform to our notion
of maximum strain. The polarities are mostly consistent and
the extents of regions where the strains are beyond our limits
of g 0.05 are conned to the fault tip regions. This increase in
gradient and strain towards the tips is likely to be a conse-
quence of interference between two or more faults.
5. Discussion and conclusions
We have shown that measurement of displacement and
mapping of displacement gradients leads to the notion of limits
to the wall-rock strain either side of a fault. Applying these limits
allows us to make objective judgements about the validity of
fault interpretation from seismic data and thus reduces the
uncertainty inherent in the interpretation process. In practice it is
the bounding values that are important and it is of little conse-
quence whether the actual metric that is used is the displace-
ment gradient or the strain. The displacement gradient is, of
course, the measured quantity but its meaning is slightly less
tangible than strain. Strain is a quantity that is more commonly
used in the literature and more likely to be linked, at least
intuitively, to other phenomenon. For example, one might predict
there to be a correlation between wall-rock strain and the degree
of fracture damage in the close vicinity of a fault.
The brief and simple analysis (Eqs. (3) and (5)) is based on dip
slip relative motion. Therefore its application is most suitable for
unrestricted, normal or reverse faults. However, dip separation is
always a minimumestimate of the true slip magnitude; for dipslip
faults it is exact but for oblique slip faults it is less than the true slip.
Similarly our estimate of both the longitudinal strain and the shear
strain will underestimate the values in the true slip direction. We
would suggest that for minor obliquity the dip slip thresholds may
still be usefully applied but it would be more accurate if either (1)
the dip slip values were to be corrected for the rake of the slip
vector or (2) the offsets were to be measured along the slip direc-
tion. But precise slip direction is almost always impossible to
determine from seismic reection data.
Given the generality of approximately triangular displacement
proles, we believe the upper limit we place on shear strain of
g 0.05 for isolated faults at the seismic scale is reasonably robust
since it is based on a large collection of D
max
/L data. As we have
already discussed, the limit of D
max
/L 0.05 seems to hold over the
scale range observed on seismic reection data. However, we
should note that other workers show that small faults (up to
L 5 m) in incompletely lithied sand have D
max
/L bounded at 0.1.
The contrast in mechanical properties with completely lithied
sandstone in elasto-plastic models of fault propagation can explain
such increases (Wibberley et al., 1999). Similarly, at the upper end
of the scale thrust faults also seem to be bounded by D
max
/L 0.1
(see Fig. 2), but this may be due to other processes such as tip
propagation blocking by basalt layers (Puentes Hills), or ductile
deformation mechanisms accommodating along-strike decolle-
ment strains (Rocky Mountains thrusts). Although there are issues
involving sampling of the principal axis of a fault, the strike
dimension and maximum offset are relatively straightforward to
measure. The main source of uncertainty is whether or not the
faults in the compilation (Fig. 2) are truly isolated. If the highest
D
max
/L values are due to faults that are horizontally restricted, then
our estimate of maximum shear strain will be too high. We should
also note that in all our natural examples described here the shear
strain gradient is lower at the centre of the faults than at the tips
and that the gradient is highest between the fault centre and the
tips i.e. the pattern is actually more of a bell shape than linear.
Moreover, in the vicinity of overlapping tips, we record higher
shear strains than the upper limit we expect for isolated faults. This
is consistent with other natural examples (Nicol et al., 1996) and
with the effect of mechanical interaction between two faults.
Relative to an unrestricted fault, the shear strain, where faults
overstep, increases with size of overlap, dip dimension and
decreasing distance between the faults (Willemse, 1997). Again,
these high strain zones may give an indication of sub-seismic scale
fracturing and thus have some impact on the local uid ow
behaviour.
Maximum allowable longitudinal strain in the dip slip direc-
tion is less robust. Schultz and Fossen (2002) argue that the
displacement scales not only with length but with aspect ratio
and that the highest D
max
/L corresponds to the smallest aspect
ratios. Soliva et al. (2005) report a similar phenomenon that for
a given constant fault height, D
max
/L decreases with length. Both
of these studies refer to outcrop scale observations where the
structures are conned to single layers. Beyond outcrop scale the
dip dimension of faults is, by comparison with the strike
dimension, more difcult to constrain. For example on seismic
reection data it is common for the upper tips to be truncated at
unconformities and/or the lower tips to be not imaged clearly.
Consequently there is little in the way of published data that
incorporates D
max
, L and aspect ratio at the scale of seismically
imaged faults. Nicol et al. (1996) show that the aspect ratios of
unrestricted faults, over the scale range of 10s of metres to 10s of
kilometers, lie between 1 and 3 but unfortunately they provide
no information on the maximum displacement for the same
faults. In setting our upper limit we have tried to embrace both
of these sets of observations, so a maximum longitudinal strain
of e 0.1 is based on D
max
/L 0.05 and the ellipticity of a fault
being a maximum of 2:1.
Our upper limit of dip-direction longitudinal strain leads to
a maximum layer thickness ratio of approximately 1.2 for corre-
sponding layers either side of the fault. Note that we are describing
fault-related strain of pre-faulting layers, not thickness changes in
syn-faulting layers (i.e. growth sequences). Increasing the aspect
ratio of the ellipse and maintaining the same displacement and
strike dimension (in contrast to Soliva et al. (2005), above)
dramatically increases the implied longitudinal strain. For example
an ellipticity of 3:1 implies a maximum longitudinal strain of
e 0.15 and differential thicknesses of 1.35. In our experience
apparent strains of this magnitude are always associated with
sedimentary growth or they are found near the tops of faults that
have been truncated by unconformities. In either case they are not
real strains. On the other hand choosing an aspect ratio of unity
reduces the upper limit by a factor of two to e 0.05. We believe
B. Freeman et al. / Journal of Structural Geology 32 (2010) 16681676 1673
that the routine examination of longitudinal strain should provide
a valuable quality control metric but we also suggest that the
published database for D
max
, L and aspect ratio needs to be
enhanced.
All our analysis and discussion is based on upper limits to strain.
It is also possible, from the point of view of 2D seismic interpre-
tation, for faults to be over-correlated laterally with large fault
strike lengths but only very small displacements. It seems that this
type of error is more difcult to quantify in terms of strain limits
since the range of known D
max
/L implies that strains could easily be
at least 100 times less than our upper bounds.
There remains the problem of anomalous implied strains in the
walls of faults which, in all other respects, have been interpreted
using this minimum strain approach. Depending on scale and
quality of the seismic data it is always possible that relatively small
structures are missed from the interpretation. In which case, these
anomalies may be the best indication of additional, real, structure.
Although we strongly caution against the invention of structure it
may be that such indirectly observed features could be incorpo-
rated into end member structural models. These would have
particular signicance if the three-dimensional structural model is
to be used for uid ow simulation.
Fig. 5. (a) Perspective viewof the faults from the original interpretation, colour coded and contoured in dip separation. The fault surfaces are displayed only where the displacement
information is present. Details of faults I, II, III, and IV are shown in strike projection in b through e. In each panel (be) the top image shows the entire fault surface (grey) and the
region where displacement information is present is colour coded in dip separation. Beneath is the map of shear strain, beneath that is the map of longitudinal strain. In the legend
the strains can be converted to displacement gradient by multiplying by two.
B. Freeman et al. / Journal of Structural Geology 32 (2010) 16681676 1674
5.1. Conclusions
(1) The problem of validating fault interpretation from seismic
data can be addressed using displacement and strain analysis.
(2) There is a simple relationship between instantaneous displace-
ment gradients and wall-rock shear and longitudinal strains.
(3) For unrestricted faults, reasonable natural upper limits to the
magnitudes of shear strain and longitudinal strain in the dip
slip direction are 0.05 and 0.1 respectively.
(4) Faults with strains that lie above these bounds are unlikely to
be correlated correctly.
(5) Fault interpretations that minimize the wall-rock strains
provide the most feasible geological solution.
Acknowledgements
We would like to thank the many colleagues at Badley Geo-
science Ltd and the Fault Analysis Group whose ideas have helped
to shape this work. The manuscript has beneted from editing by
Chris Wibberley and the careful and constructive reviews of Matt
Pachell and one anonymous reviewer.
References
Bailey, W.R., Walsh, J.J., Manzocchi, T., 2005. Fault populations, strain distribution
and basement fault reactivation in the East Pennines Coaleld, UK. Journal of
Structural Geology 27, 913928.
Barnett, J.A., Mortimer, J., Rippon, J.H., Walsh, J.J., Watterson, J., 1987. Displacement
geometry in the volume containing a single normal fault. American Association
of Petroleum Geologists Bulletin 71, 925937.
Boult, P.J., 1999. Maturity Modelling of the Casterton Formation and Killara Coals in
PEP 111 and PEP 101. Boral Energy Resources Ltd., Otway Basin, Victoria.
Unpublished report.
Boult, P.J., Hibburt, J.E. (Eds.), 2002. The Petroleum Geology of South Australia. Vol.
1: Otway BasinPetroleum Geology of South Australia Series, second ed., vol. 1.
Department of Primary Industries and Resources, South Australia.
Boult, P, Freeman, B., Yielding, G., Menpes, S., Diekman, L.J., 2008. A minimum-strain
approach to reducing the structural uncertainty in poor 2D seismic data,
Gambier Embayment, Otway Basin, Australia. In: Proceedings of the Third
Eastern Australasian Basins Symposium, Sydney, September 14th17th.
Fig. 6. (a) Perspective view of faults from the reinterpretation, colour coded and contoured in dip separation. (b) Shear strain (upper) and longitudinal strain (lower) maps of the
WNWESE fault set. (c) Shear strain (upper) and longitudinal strain (lower) maps of the NWSE fault set.
B. Freeman et al. / Journal of Structural Geology 32 (2010) 16681676 1675
Bouvier, J.D., Kaars-Sijpesteijn, C.H., Kluesner, D.F., Onyejekwe, C.C., Van der Pal, R.C.,
1989. Three-dimensional seismic interpretation and fault sealing investigations,
Nun River Field, Nigeria. American Association of Petroleum Geologists Bulletin
73, 13971414.
Chapman, T.J., Meneilly, A.W., 1991. The displacement patterns associated with
a reverse-reactivated, normal growth fault. In: Roberts, A.M., Yielding, G.,
Freeman, B. (Eds.), The Geometry of Normal Faults. Geological Society of Lon-
don, Special Publication, vol. 56, pp. 183191.
Eshelby, J.D., 1957. The determination of the elastic eld of an ellipsoidal inclusion and
relatedproblems. Proceedings of theRoyal Societyof London, SeriesA241, 376396.
Essential Petroleum Resources Ltd., 2006. Internal Presentation to the PEL 72 Joint
Venture Technical Committee.
Freeman, B., Badley, M.E., Yielding, G., 1990. Fault correlation during seismic
interpretation. First Break 8, 8795.
Kim, Y.-S., Sanderson, D.J., 2005. The relationship between displacement and length
of faults: a review. Earth-Science Reviews 68, 317334.
Lee, S., Wolberg, G., Shin, S.Y., 1997. Scattered data interpolation with multilevel B-
splines. IEEE Transactions on Visualization and Computer Graphics 3, 228244.
Lovibond, R., Suttill, R.J., Skinner, J.E., Aburas, A.N., 1995. The hydrocarbon potential
of the Penola Trough, Otway Basin. APEA Journal 35, 358371.
Manzocchi, T., Walsh, J.J., Nicol, A., 2006. Displacement accumulation from
earthquakes on isolated normal faults. Journal of Structural Geology 28,
16851693.
Needham, D.T., Yielding, G., Freeman, B., 1996. Analysis of fault geometry and
displacement patterns. In: Buchanan, P.G., Nieuwland, D.A. (Eds.), Modern
Developments in Structural Interpretation, Validation and Modelling. Geolog-
ical Society of London, Special Publication, vol. 99, pp. 189199.
Nicol, A., Watterson, J., Walsh, J.J., Childs, C., 1996. The shapes, major axis orienta-
tions and displacement patterns of fault surfaces. Journal of Structural Geology
18, 235248.
Pollard, D.D., Segall, P., 1987. Theoretical displacements and stresses near fractures
in rock: with applications to faults, joints, veins, dikes, and solution surfaces. In:
Atkinson, B.K. (Ed.), Fracture Mechanics of Rock. Academic Press, London, pp.
277349.
Rippon, J.H., 1985. Contoured patterns of the throw and hade of normal faults in the
Coal Measures (Westphalian) of north-east Derbyshire. Proceedings of the
Yorkshire Geological Society 45, 147161.
Schultz, R.A., Soliva, R., Fossen, H., Okubo, C.H., Reeves, D.M., 2008. Dependence of
displacementlength scaling relations for fractures and deformation bands on
the volumetric changes across them. Journal of Structural Geology 30, 1405
1411.
Schultz, R.A., Fossen, H., 2002. Displacementlength scaling in three dimensions:
the importance of aspect ratio and application to deformation bands. Journal of
Structural Geology 24, 13891411.
Soliva, R., Schultz, R.A., Benedicto, A., 2005. Three-dimensional displacement
length scaling and maximum dimension of normal faults in layered rocks.
Geophysical Research Letters 32, L16302. doi:10.1029/2005GL023007.
Walsh, J.J., Watterson, J., 1989. Displacement gradients on fault surfaces. Journal of
Structural Geology 11, 307316.
Wibberley, C.A.J., Petit, J.-P., Rives, T., 1999. Mechanics of high displacement gradient
faulting prior to lithication. Journal of Structural Geology 21, 251257.
Willemse, E.M.J., 1997. Segmented normal faults: correspondence between three-
dimensional mechanical models and eld data. Journal of Geophysical Research
102 (B1), 675692.
B. Freeman et al. / Journal of Structural Geology 32 (2010) 16681676 1676
Excavation induced fractures in a plastic clay formation: Observations at the
HADES URF
Philippe Van Marcke
*
, Wim Bastiaens
1
SCKCEN, Boeretang 200, 2400 Mol, Belgium
a r t i c l e i n f o
Article history:
Received 3 March 2009
Received in revised form
1 December 2009
Accepted 20 January 2010
Available online 13 February 2010
Keywords:
Geological disposal
Boom Clay
Tunnel excavation
Excavation induced fractures
Hydro-mechanical behaviour
Self-sealing
a b s t r a c t
The Belgian research programme for geological disposal of radioactive waste focuses on the Boom Clay as
the potential host rock formation. To examine the feasibility of constructing an underground repository
in this clay layer, an underground research facility HADES has been constructed in several stages since
1980. The two galleries most recently excavated, the Connecting Gallery in 2002 and the Praclay Gallery
in 2007, were constructed by means of an industrial method using a tunnelling machine. During these
excavations the hydro-mechanical response of the clay was characterised.
A fracture pattern was observed consistently during the excavation of both galleries. The extent of this
fractured zone was determined for the Connecting Gallery, but requires some further study. A strong
hydro-mechanical coupling and a clear time dependency were noticed, even at an unexpectedly large
distance from the excavation. Furthermore the Boom Clay responds in an anisotropic manner to the
excavation due to anisotropy in the in situ stress state and the Boom Clay characteristics. Self-sealing
processes were observed and appear to occur relatively fast.
2010 Elsevier Ltd. All rights reserved.
1. Introduction
The geological disposal of radioactive waste has been studied in
Belgium since the early seventies by the Belgian Nuclear Research
Centre (SCKCEN). The research is focused on the Boom Clay:
a plastic clay layer that is found from a depth of 190 m under the
site of SCKCEN in Mol (in the northeast of Belgium) where it has
a thickness of about 100 m. It has a low hydraulic conductivity (in
the order of 10
12
m/s) and displays a plastic behaviour which
results in self-sealing properties and a relatively high convergence
when excavating galleries in it.
In 1980 SCKCEN started the construction of the underground
facility HADES at a depth of 225 m (Fig. 1). Its main purpose was to
examine the feasibility of constructing such a repository and to
provide SCKCEN with an underground infrastructure for experi-
mental research on the geological disposal of radioactive waste. Not
much knowledge and experience on excavating in a deep plastic
clay formation were available at that time. The work during this
phase is therefore considered to be pioneering.
In 2002 the second shaft was connected with the existing
underground infrastructure by the Connecting Gallery (80 m
long and 4.8 m in external diameter). This was done in an
industrial manner by the use of a tunnelling machine. Several
measurement and research programmes were carried out before,
during and after the construction works to characterise the
hydro-mechanical response of the clay around the repository
(Bastiaens et al., 2003; Bernier et al., 2003). In particular the
fracture pattern resulting from the excavation was characterised.
In 2007 the Praclay Gallery (45 m long and 2.5 m in external
diameter) was constructed perpendicular to the Connecting
Gallery. Again, the hydro-mechanical response was measured
and characterised.
This paper discusses these measurements and observations.
First the Boom Clay characteristics are given, then the used exca-
vation technique is described after which the hydro-mechanical
observations are presented.
2. Boom Clay characteristics
The Boom Clay is a silty clay characterised by a structure of
bands that are several tens of centimetres thick, reecting mainly
cyclical variations in grain size (silt and clay content) due to uc-
tuations in the wave action on the sedimentation medium and to
variations in the carbonate and organic matter contents. Typical
concretions, known as septaria, are found in the marly bands
occurring throughout the thickness of the formation.
* Corresponding author. Tel.: 32 14 33 29 88; fax: 32 14 32 37 09.
E-mail addresses: pvmarcke@sckcen.be (P. Van Marcke), wbastiae@sckcen.be
(W. Bastiaens).
1
Tel.: 32 14 33 27 90; fax: 32 14 32 37 09.
Contents lists available at ScienceDirect
Journal of Structural Geology
j ournal homepage: www. el sevi er. com/ l ocat e/ j sg
0191-8141/$ e see front matter 2010 Elsevier Ltd. All rights reserved.
doi:10.1016/j.jsg.2010.01.010
Journal of Structural Geology 32 (2010) 1677e1684
The formation belongs to the Rupelian which is the geological
part of the Tertiary Period with an age between 36 and 30 million
years. It is found at a depth of about 190 m under the SCKCEN site
of Mol where it has a thickness of about 100 m (see Fig. 2). The
Boom Clay layer is almost horizontal (it dips 1e2% towards the NE)
and water bearing sand layers are situated above and below it.
Due to its vertical lithological heterogeneity the mineralogy of
the Boom Clay is characterised by a wide variation in the content of
clay minerals (from 30 to 70% volume, dry matter). In descending
Fig. 1. Construction history of the underground research facility HADES.
Fig. 2. Geological section under the Mol site.
Table 1
Undrained geomechanical characteristics of the Boom Clay at the depth of HADES.
Property Unit Value
Young's modulus tangential at the origin E 200e400 MPa
Poisson's ratio n 0.40e0.45
Unconned compressive strength UCS 2 MPa
Angle of friction 4 4
C)
Sliding
velocity
V (mm/s)
Final
shear
strain g
Apparent
compaction
(mm)
Final steady
state friction
coefcient m
MUS17 500 0.1 79.2 0.20 0.61
MUS18 500 3.7 63.7 0.18 0.68
MUS21 500 1 103.3 0.19 0.63
MUS04 500 1 ND ND ND
MUS05 300 1 36.9 0.12 0.70
MUS07 100 1 142.0 0.22 0.630.70
MUS11 500 1 133.4 0.19 0.89
MUS36 700 1 7.3 0.16 0.37
Normal stress-stepping experiments, V 1.0 mm/s, s
eff
20, 40, 60, 80, 100 MPa
Experiment Temperature
T (
C)
Final
shear
strain g
Final apparent
compaction
(mm)
Steady state
friction
coefcient m
Notes
MUS01 20 29.1 0.08 0.43
MUS02 150 25.0 0.08 0.43
MUS03 500 28.6 0.15 0.51 Stick-slip
MUS06 400 28.7 0.08 0.60
MUS08 300 ND ND 0.54
MUS09 600 27.4 0.06 0.74
MUS12 500 27.8 0.06 0.76 Stick-slip
MUS16 100 30.3 0.22 0.38
MUS22 700 31.7 0.14 0.55 Initial dilatation
MUS24 600 28.6 0.14 0.55 Initial dilatation
MUS28 225 22.6 0.16 0.37
MUS29 500 28.5 0.01 0.49 Initial dilatation
MUS31 300 29.0 0.20 0.38
MUS32 400 29.1 0.09 0.47 Initial dilatation
MUS33 500 33.5 0.14 0.53 Initial dilatation
Velocity-stepping experiments, s
eff
100 MPa, V 0.03, 0.1, 0.5, 1, 3.7 mm/s
Experiment Temperature
T (
C)
Final
shear
strain g
Final apparent
compaction
(mm)
Final steady
state friction
coefcient m
Notes
MUS13 500 32.6 0.12 0.90 Stick-slip, initial
dilatation
MUS14 300 37.0 0.14 ND
MUS19 700 31.0 0.26 0.43
MUS20 20 37.7 0.12 0.38
MUS25 600 42.5 0.16 0.56 Initial dilatation
MUS26 150 42.2 0.11 0.46
MUS30 500 44.8 0.14 0.62 Stick-slip, initial
dilatation
MUS34 225 39.2 0.10 0.47 Initial dilatation
MUS35 400 40.2 0.16 0.54
Compaction experiment, s
eff
100 MPa, P
f
100 MPa
MUS23 700 0 0.11
0
20
40
60
80
100
120
0 2 4 6 8 10 12 14 16 18
Shear displacement (mm)
S
h
e
a
r
-
o
r
n
o
r
m
a
l
s
t
r
e
s
s
(
M
P
a
)
0
0.1
0.2
0.3
0.4
0.5
0.6
0.7
0 5 10 15 20 25
Shear strain
C
o
m
p
a
c
t
i
o
n
d
u
r
i
n
g
s
h
e
a
r
(
m
m
)
Shear stress
Normal stress
Compaction
= 0.38 + 15.713
R
2
= 0.9993
= 0.38 +16.636
R
2
= 0.9985
0
10
20
30
40
50
60
0 20 40 60 80 100 120
Normal stress (MPa)
S
t
e
a
d
y
s
t
a
t
e
s
h
e
a
r
s
t
r
e
s
s
(
M
P
a
)
Step up Step down
Average = 0.38
a
b
Fig. 2. Typical example of a normal stress-stepping experiment (MUS31), T 300
C,
P
f
100 MPa, s
eff
20100 MPa, V 1.0 mm/s and g 29. (a) Shear stress, normal
stress and apparent compaction as function of shear displacement and shear strain.
Note that the apparent compaction is not corrected for apparatus distortion. (b)
Correlated shear stress vs. normal stress diagram. The slope of the linear best t
through the data points represents a measure for the mean steady state friction
coefcient of the sample. The intercept of the linear best t at zero effective normal
stress is the effect of the friction of the O-ring seals and the conning rings (Niemeijer
et al. (2008), see Section 2.4). These intercept values were used to correct the
measured shear stresses obtained from velocity-stepping experiments.
E.W.E. Van Diggelen et al. / Journal of Structural Geology 32 (2010) 16851700 1689
contribution s
f
exerted by the O-ring seals on the pressure-
compensated piston and by the conning rings surrounding the
sample (Fig. 1c and d). Friction calibrations performed using an
internal normal force gauge have shown that at measured normal
stresses of 100 MPa the contribution of the O-ring friction to the
normal stress is constant and negligible compared to the applied
normal stress s
m
, so that the normal stress on the sample s
s
s
m
.
Therefore, the measured friction coefcient (m
m
in Eq. (1) repre-
sents true sample behaviour m
s
s
s
=s
s
provided the O-ring
contribution to the measured torque is independent of normal
stress. We performed the following calibration experiments to
allow meaningful processing of the data obtained.
In this procedure, two annular PEEK (PolyEtherEtherKetone)
rings of w0.3 mm in thickness were inserted into the apparatus in
the place of the usual fault gouge samples. Friction experiments
were performed using these PEEK samples (i) with a stainless
steel outer/inner conning ring, (ii) with a PEEK outer/inner
conning ring, and (iii) without an outer conning ring. The
frictional behaviour of the contact between the two solid PEEK
samples was measured in all of these cases for effective normal
stresses in the range 20100 MPa. The data obtained show well-
dened linear relations between shear stress and normal stress,
indicating PEEK-on-PEEK friction coefcients m
s
in the range 0.03
0.06. No systematic change in m
s
was seen in relation to the type of
inner/outer conning ring used (steel, PEEK or no ring), implying
that the frictional forces generated by the conning rings are
minor. The m
s
values 0.030.06 are low compared to the value of
0.1 (peek on peek) quoted by the manufacturer (see Niemeijer
et al., 2008), but t well with the low end of the range of 0.050.5
given in the literature (Burris and Sawyer, 2006a, 2006b; Theiler
and Gradt, 2008). This agreement implies that the frictional forces
caused by the seals and conning rings are unlikely to be signif-
icantly inuenced by the normal stress. We, therefore, conclude
that linear ts to the shear stress vs. normal stress data measured
in our experiments on muscovite (Fig. 2) yield true friction coef-
cients m
s
for the fault gouge. This is consistent with previous
conclusions based on experiments in the same apparatus by
(Niemeijer et al., 2008).
From the data of Niemeijer et al. (2008), the cohesion C
s
of our
muscovite samples can be taken as near zero (Niemeijer et al., 2008,
who measured 0.16 MPa for muscovite gouge), meaning that C
m
in
Eq. (1) is made up almost entirely of O-ring friction C
f
. Values of C
m
have been determined for every normal stress-stepping test per-
formed. Although they show some variability (1323 MPa), no
systematic relationship with temperature was apparent. We,
therefore, used the average value obtained for C
m
(17 MPa) to
correct the measured shear stress s
m
for O-ring and (minor)
conning ring friction. Dividing the result by the applied effective
normal stress s
m
(cf. Eq. (1)) then yields the internal friction coef-
cient of the sample:
m
s
s
m
C
m
s
m
(2)
Curves of corrected shear stress (s
m
C
m
) vs. displacement can
thus easily be recast into m
s
vs. shear strain g curves for our samples.
Compaction and dilatation measurements were determined
from the position of the Instron loading ram. Since the heated
length of the loading pistons is very limited, we can assume that
temperature does not inuence elastic apparatus distortion. Axial
machine distortion, therefore, depends only on normal stress. We
did not correct for this. However, we compare measured compac-
tion/dilatation data for different experiments only fromthe onset of
shear deformation and at xed effective normal stress. Any differ-
ences between experiments, therefore, reect sample behaviour
and are unrelated to apparatus behaviour.
2.5. Sample handling and microstructural methods
After the experiments, most samples broke into arc-shaped
fragments during removal of the sheared gouge from the set-up.
The fragments varied in size between 3 and 14 mmin arc length, by
3 mm wide and w0.5 mm thick. All were dried for w45 min at
60
C after extraction. The dried samples were then vacuum
impregnated with epoxy resin for about 15 min and left to dry for
a period of 24 h. The impregnated fragments were nally sectioned
normal to the shear plane and parallel to the shear direction, and
polished. Mineral content and microstructure were investigated
using a Scanning Electron Microscope (SEM) equipped with Energy
Dispersive X-ray analysis system (EDX). Grain sizes and mineral
fractions in the deformed samples were obtained from point
counting analyses.
3. Mechanical and microstructural results
We performed 8 sliding experiments at xed conditions, 14
normal stress-stepping experiments, 9 velocity-stepping experi-
ments and 1 compaction experiment. The experiments and corre-
sponding conditions are represented in Table 1.
3.1. Effect of shear strain
Fig. 3a and b shows the evolution of friction coefcient with
shear strain for a number of samples deformed at 500
C and at
0
0.1
0.2
0.3
0.4
0.5
0.6
0.7
0.8
0 5 10 15 20 25 30
Shear displacement (mm)
F
r
i
c
t
i
o
n
c
o
e
f
f
i
c
i
e
n
t
0.0
0.5
1.0
1.5
2.0
2.5
3.0
3.5
4.0
0 5 10 15 20 25 30 35 40 45
Shear strain
(
y
t
i
c
o
l
e
v
g
n
i
d
i
l
S
)
s
/
m
Shear stress
Sliding velocity
V = 1.0 m/s 1.0 m/s 1.0 m/s
0.5 m/s
3.7 m/s
0.1 m/s0.03 m/s
0
0.1
0.2
0.3
0.4
0.5
0.6
0.7
0.8
0 10 20 30 40 50 60 70
Shear displacement (mm)
F
r
i
c
t
i
o
n
c
o
e
f
f
i
c
i
e
n
t
0 20 40 60 80 100
Shear strain
MUS17 V = 0.1 um/s
MUS18 V = 3.7 um/s
MUS21 V = 1.0 um/s
MUS30 V-stepping experiment
V = 0.1 m/s
V = 3.7 m/s V = 1.0 m/s
Unstable
stick-slip behaviour
Stable sliding
behaviour
a
b
Fig. 3. (a) Friction coefcient vs. shear displacement (i.e. shear strain) obtained from
three sliding experiments under constant conditions (non-stepping) at T 500
C,
P
f
100 MPa, s
eff
100 MPa and one velocity-stepping experiment under the same
pressure and temperature conditions for comparison. The samples deformed at 0.1 and
1.0 mm/s showunstable stick-slip behaviour until sliding stabilizes at g 16 and g 24,
respectively. The velocity-stepping experiment also shows unstable sliding in most
steps, while the sample deformed at 3.7 mm/s exhibited stable sliding throughout the
experiment. (b) Friction coefcient data from velocity-stepping experiment MUS30
compared with the steps in sliding velocity.
E.W.E. Van Diggelen et al. / Journal of Structural Geology 32 (2010) 16851700 1690
100 MPa effective normal stress. In the early stages of the experi-
ments (g up to 3), the friction coefcient strongly increased with
shear strain, then gradually approached a steady state value of
w0.6 at g of 3540. The samples compacted more or less contin-
uously during shear (Fig. 4a) with little effect of velocity (Fig. 4b).
Stick-slip behaviour was observed in some experiments at 400 and
500
C, but no systematic relationship with strain or velocity was
found. Stick-slip was not observed in tests at temperatures outside
this temperature range. Fig. 5 shows values of friction coefcient vs.
shear strain at g >7, as obtained from a representative set of our
velocity-stepping experiments. Note that data obtained at constant
temperature (i.e. in a given stepping test) show a minor increase in
friction coefcient with strain within the range g 1035, and little
effect of velocity.
Fig. 6ad showthe microstructures developed at shear strains of
w29, 40, 64 and 103, in samples deformed at 400 and 500
C,
100 MPa uid pressure and a sliding velocity of 1.0 mm/s. Of these
samples, three were deformed under 100 MPa normal stress and
one at 20100 MPa in a normal stress-stepping experiment
(MUS32). The microstructures developed at g 29 and g 40 at
400
C (Fig. 6a, b and e) are rather similar. They show a muscovite
foliation oblique to the shear direction and the presence of thin
(w2 mm) anastomosing shear bands, dominantly in Y-shear orien-
tation. Most muscovite grains have sizes of 520 mm, similar to the
starting material and some grains show occasional folding or kink-
ing. The thin bands contain grains of 12 mm, and are sometimes
bounded by thicker (310 mm) zones of 14 mm size grains. Signi-
cant porosity is visible at the tips of the muscovite grains, which are
often jagged and broken in appearance. About 5% quartz porphyr-
oclasts (520 mm) are present, often showing extensional fractures.
The effect of shear strain can best be assessed using Fig. 6c and
d for g 64 and g 103, both at 500
C and similar normal stress
and uid pressure. Shear bands are wider now than seen at lower
shear strain at 400
C, and their width increases with strain, up to
20 mm in the sample deformed to g 103. The bands anastomose
around dense lenses of coarser muscovite (Fig. 6f), showgrain sizes
<1 mm (Fig. 6g), and occupy roughly 2035% of the gouge volume.
The intervening lenses do not show a clear foliation, but do show
local folding and kinking in the muscovite grains and have a grain
size of 220 mm. The lenses also contain occasional quartz por-
phyroclasts with a grain size of 210 mm.
3.2. Effect of normal stress
Data on the measured shear stress, the effective normal stress
and apparent compaction (uncorrected vertical displacement) vs.
rotary displacement, obtained from a typical normal stress-step-
ping experiment (MUS31), are illustrated in Fig. 2a (see also Table
1). The corresponding shear stress vs. normal stress diagram, from
which the mean steady state coefcient of friction is calculated, is
shown in Fig. 2b. In individual stress steps, the shear stress initially
increased rapidly, but reached steady state within w1 mm of
displacement (Fig. 2a). Also, the apparent compaction increased
instantly when normal stress was stepped up, while an instant
expansion occurred when normal stress was stepped down. Rough
stiffness calibrations have shown that these vertical displacements
were largely due to elastic distortion of the apparatus, though net
compaction of the gouge was recorded at the end of the experi-
ments (Table 1). We do not take the apparent compaction obtained
from the normal stress-stepping experiments into account in the
interpretation of our results, since we cannot correct accurately for
apparatus stiffness. Nonetheless, friction coefcients obtained in
the normal stress-stepping tests ranged from 0.37 to 0.60
depending on temperature (see Section 3.4).
Fig. 6a and b show microstructures of samples that reached
different shear strains, at T 400
C, but that also differed in
effective normal stress. The micrograph shown in Fig. 6a is obtained
from an experiment in which the effective normal stress was
stepped from 20 MPa to 100 MPa and back (in 20 MPa-steps).
Fig. 6b represents a sample deformed at constant normal stress of
100 MPa. The microstructures of the two samples are rather similar
(see also Section 3.1), hence, there seems to be no signicant effect
a
b
-0.05
0
0.05
0.1
0.15
0.2
0.25
0 10 20 30 40 50 60 70
Shear displacement (mm)
)
m
m
(
r
a
e
h
s
g
n
i
r
u
d
n
o
i
t
c
a
p
m
o
C
MUS17 V = 0.1 um/s
MUS18 V = 3.7 um/s
MUS21 V = 1.0 um/s
MUS30 V-stepping experiment
V = 0.1 m/s
V = 3.7 m/s
V = 1.0 m/s
V- stepping
-0.05
0
0.05
0.1
0.15
0.2
0.25
0 5 10 15 20 25 30
Shear displacement (mm)
)
m
m
(
r
a
e
h
s
g
n
i
r
u
d
n
o
i
t
c
a
p
m
o
C
0.0
0.5
1.0
1.5
2.0
2.5
3.0
3.5
4.0
(
y
t
i
c
o
l
e
v
g
n
i
d
i
l
S
)
s
/
m
Compaction
Sliding velocity
V = 1.0 m/s 1.0 m/s 1.0 m/s
0.5 m/s
3.7 m/s
0.1 m/s 0.03 m/s
Fig. 4. (a) Compaction vs. shear displacement diagram obtained from the three
constant velocity experiments and a single velocity-stepping experiment as in Fig. 3.
Only the compaction recorded from the onset of shearing is indicated here, at constant
effective normal stress of 100 MPa. The irregularities in the data signal are due to a low
signal to noise ratio and are not related to stick-slip behaviour of the sample. (b)
Compaction data from velocity-stepping experiment MUS30 compared with the steps
in sliding velocity.
0
1 . 0
2 . 0
3 . 0
4 . 0
5 . 0
6 . 0
7 . 0
8 . 0
0 3 5 2 0 2 5 1 0 1 5 0
Shear displacement (mm)
F
r
i
c
t
i
o
n
c
o
e
f
f
i
c
i
e
n
t
5 4 0 4 5 3 0 3 5 2 0 2 5 1 0 1 5 0
Shear strain
MUS35, T = 400 C MUS30, T = 500 C
MUS19, T = 700 C MUS20, T = 20 C
20 C
700 C
400 C
500 C
Fig. 5. Diagram illustrating the effect of shear strain on the coefcient of friction
obtained from velocity-stepping experiments at different temperatures. At 700
C,
steady state shear stress levels were not reached within a single velocity step. The
reported values for this temperature are only represented as an indication and should
be considered with care.
E.W.E. Van Diggelen et al. / Journal of Structural Geology 32 (2010) 16851700 1691
of normal stress on the development of microstructures, at least not
at 400
C.
3.3. Effect of sliding velocity
From Figs. 35, it has already been noted that sliding velocity
does not have an important effect on the mechanical behaviour of
the gouges. This is largely conrmed in Fig. 7, where the steady
state friction coefcients obtained in our velocity-stepping tests are
explicitly plotted against sliding velocity. Clearly, there is no
signicant effect of velocity on friction coefcient at xed
temperatures up to 500600
C, though m does increase with
temperature (see Section 3.4). At 700
C, the generally lower fric-
tion coefcient seems to peak at 1.0 mm/s, even though steady state
shear stress levels were not reached under these conditions. We did
not observe a systematic effect of sliding velocity on the compac-
tion behaviour of the gouges (Fig. 4a and b).
The microstructures typifying samples deformed at constant
sliding velocities ranging from0.1 to3.7 mm/s are illustratedinFig. 8,
for T 500
C and for shear strains of about 6379 (samples MUS17,
18 and 21, Table 1). The microstructures developed under these
conditions are again characterized by ne grained, elongate lenses
and oval clasts measuring typically 2050 by 10 mm (Fig. 8a and c)
consisting of often folded and kinked muscovites some 220 mm in
length(Fig. 8bandd). Inthe SEM, the lenses showa lowporosityand
are separated by a through-going anastomosing network of ultra-
ne grained (1 mm) bands, oriented parallel to and, though less
well developed, at roughly 2040
to (001) or fracture at 90
km
1
and
a surface temperature of 25
C, an average crustal density of
2.7 gcm
3
, and a Byerlee friction coefcient of 0.75. A frictional
strength prole corresponding to m 0.2 is added to represent the
strength expected for the San Andreas fault zone on the basis of
heat ow measurements (Lachenbruch and Sass, 1980). The crystal
plastic ow strength of muscovite is added (Fig. 14) following
(Kronenberg et al., 1990) for shear strain rates of 10
14
, 10
12
, and
10
10
s
1
and assuming an activation enthalpy of 270 kJmol
1
(Mariani et al., 2006). The owstrength of quartz is added since this
is lower than the ow strength of muscovite under low-crustal
conditions (Fig. 14) following an empirical equation for dislocation
creep (Gleason and Tullis, 1995).
The shear strength of muscovite gouge inferred from our
experimental data, at depths of 510 km, is 3070 MPa, while
a strength of 1020 MPa is implied for the San Andreas fault zone
(Lachenbruch and Sass, 1980). Based on these results, it seems
Transcurrent fault
0
5
0 1
5 1
0 2
5 2
0 0 2 0 5 1 0 0 1 0 5 0
D
e
p
t
h
z
(
k
m
)
Byerlee's law
Muscovite
0 1
4 1 -
s u m
0 1
2 1 -
s u m
0 1
0 1 -
s u m
0 1
4 1 -
z t q
0 1
2 1 -
z t q
0 1
0 1 -
z t q
Muscovite
breakdown
Shear strength (MPa)
= 0.75
= 0.2
Fig. 14. Crustal strength prole comparing Byerlees Rule with the frictional strength
prole for muscovite, drawn using data presented in Fig. 9a. We assumed a geothermal
gradient of 30
Ckm
1
and a surface temperature of 25
C, an average crustal density of
2.7 gcm
3
, and a Byerlee friction coefcient of 0.75. The frictional strength prole
corresponding to m 0.2 is added to represent the strength expected for the San
Andreas fault zone on the basis of heat ow measurements (Lachenbruch and Sass,
1980). The plastic ow strength of muscovite is added following Kronenberg et al.
(1990) for various strain rates assuming an activation energy of 270 kJmol
1
after
Mariani et al. (2006). The plastic ow strength of wet quartz is added following
Gleason and Tullis (1995), since it is lower than the ow strength of muscovite under
low-crustal conditions.
E.W.E. Van Diggelen et al. / Journal of Structural Geology 32 (2010) 16851700 1698
unlikely that the presence of muscovite alone, with the rate-inde-
pendent frictional strength as taken in our study, can account for
the inferred weakness of large scale crustal fault zones. Muscovite,
however, shows stick-slip behaviour, which is the laboratory
equivalent of earthquakes, at temperatures between 400 and
500
C, corresponding to a depth of 1317 km in a transcurrent
fault zone. This unstable sliding behaviour might be of some
signicance in controlling fault behaviour towards the base of the
seismogenic zone.
5. Conclusions
The present study aimed to determine the frictional behaviour
of simulated muscovite fault gouge at high shear displacements
and under hydrothermal conditions in the temperature range 20
700
C, at shear strain rates w10
5
to 10
3
s
1
. The following
conclusions were reached:
(1) all our samples showed strong strain hardening at low shear
strains, followed by a more gradual increase in strength, until
steady state was reached. The steady state coefcient of friction
increased from 0.37 at room temperature to 0.56 at 300
C,
remaining around this value up to 600
C. At 700
C, the coef-
cient of friction decreased again, to a value of 0.38;
(2) all samples showed substantial grain size reduction at T up to
600
C due to pervasive and localized cataclasis, which resulted
in continuous compaction and hardening of the gouge. This
hardening is due to the progressive development of an anas-
tomosing network of ne grained, cataclastic shear bands,
gradually widening and hardening. Coarse grained relict lenses
between the cataclastic bands show folded and kinked
muscovite grains indicative of active ductile mechanisms;
(3) due to the presence of quartz impurities (<10%) in the gouge it
is possible that partial melting occurred at 700
C;
(4) on the basis of our results, it seems unlikely that the presence
of muscovite can signicantly contribute to the long-term
weakness of large scale crustal fault zones, unless its strength
dramatically decreases with decreasing sliding velocity or
shear strain rate.
Acknowledgements
This research was funded by the Netherlands Research Centre
for Integrated Solid Earth Science, project AM 2.1. We thank Thony
van der Gon Netscher and Gert Kastelein for constructing and
adjusting the hydrothermal rotary shear apparatus. We would also
like to thank Eimert de Graaff and Peter van Krieken for their
technical support and also Andre Niemeijer and Gill Pennock for
the useful discussions. Finally, we gratefully acknowledge Eli-
sabetta Mariani and Diane Moore for the constructive reviews.
References
Arancibia, G., Morata, D., 2005. Compositional variations of syntectonic white-mica
in low-grade ignimbritic mylonite. Journal of Structural Geology 27, 745767.
Balfour, N.J., Savage, M.K., Townend, J., 2005. Stress and crustal anisotropy in
Marlborough, New Zealand: evidence for low fault strength and structure-
controlled anisotropy. Geophysical Journal International 163, 10731086.
Blanpied, M.L., Lockner, D.A., Byerlee, J.D., 1995. Frictional slip of granite at hydro-
thermal conditions. Journal of Geophysical Research 100, 1304513064.
Bos, B., Peach, C.J., Spiers, C.J., 2000. Slip behavior of simulated gouge-bearing faults
under conditions favoring pressure solution. Journal of Geophysical Research
105, 1669916717.
Bos, B., Spiers, C.J., 2000. Effect of phyllosilicates on uid-assisted healing of gouge-
bearing faults. Earth and Planetary Science Letters 184, 199210.
Bos, B., Spiers, C.J., 2001. Experimental investigation into the microstructural and
mechanical evolution of phyllosilicate-bearing fault rock under conditions
favouring pressure solution. Journal of Structural Geology 23, 11871202.
Bos, B., Spiers, C.J., 2002. Frictional-viscous ow of phyllosilicate-bearing fault rock:
microphysical model and implications for crustal strength proles. Journal of
Geophysical Research 107, 2028. doi:10.1029/2001JB000301.
Burris, D.L., Sawyer, W.G., 2006a. Improved wear resistance in alumina-PTFE
nanocomposites with irregular shaped nanoparticles. Wear 260, 915918.
Burris, D.L., Sawyer, W.G., 2006b. A low friction and ultra low wear rate PEEK/PTFE
composite. Wear 261, 410418.
Byerlee, J., 1990. Friction, overpressure and fault normal compression. Geophysical
Research Letters 17, 21092112.
Byerlee, J.D., 1978. Friction of rocks. Pure and Applied Geophysics 116, 615626.
Carpenter, B.M., Marone, C., Saffer, D.M., 2009. Frictional behavior of materials in
the 3D SAFOD volume. Geophyscial Research Letters 36. doi:10.1029/
2008GL036660.
Chatterjee, N.D., Johannes, W., 1974. Thermal stability adn standard thermodynamic
properties of synthetic 2M1-muscovite, KAl2[AlSi3O10(OH)2]. Contributions to
Mineralogy and Petrology 48, 89114.
Chester, F.M., 1995. A rheologic model for wet crust applied to strike slip faults.
Journal of Geophysical Research 100, 1303313044.
Collettini, C., Barchi, M.R., 2002. A low-angle normal fault in the Umbria region
(Central Italy): a mechanical model for the related microseismicity. Tectono-
physics 359, 97115.
Collettini, C., Holdsworth, R.E., 2004. Fault zone weakening and character of slip
along low-angle normal faults: insights from the Zuccale fault, Isle of Elba, Italy.
Journal of the Geological Society of London 161, 10391051.
Deer, W.A., Howie, R.A., Zussman, J., 1962. An Introduction to the Rock Forming
Minerals: Micas. Longman Group Limited, London.
Di Toro, G., Hirose, T., Nielsen, S., Pennacchioni, G., Shimamoto, T., 2006. Natural and
experimental evidence of melt lubrication of faults during earthquakes. Science
311, 647649.
Dieterich, J.H., 1978. Time-dependent friction and the mechanics of stick-slip. Pure
and Applied Geophysics (Historical Archive) 116, 790806.
Evans, J.P., Chester, F.M., 1995. Fluid-rock interaction in faults of the San Andreas
system: inferences from San Gabriel fault rock geochemistry and microstruc-
tures. Journal of Geophysical Research 100, 1300713020.
Faulkner, D.R., Rutter, E.H., 2001. Can the maintenance of overpressured uids in
large strike-slip fault zones explain their apparent weakness? Geology 29,
503506.
Gleason, G.C., Tullis, J., 1995. A ow law for dislocation creep of quartz aggregates
determined with the molten salt cell. Tectonophysics 247, 123.
Goetze, C., Evans, B., 1979. Stress and temperature in the bending lithosphere as
constrained by experimental rock mechanics. Geophysical Journal International
59, 463478.
He, C., Wang, Z., Yao, W., 2007. Frictional sliding of gabbro gouge under hydro-
thermal conditions. Tectonophysics 445, 353362.
He, C., Yao, W., Wang, Z., Zhou, Y., 2006. Strength and stability of frictional sliding of
gabbro gouge at elevated temperatures. Tectonophysics 427, 217229.
Hickman, S.H., 1991. Stress in the lithosphere and the strength of active faults.
Review of Geophysics 29, 759775.
Hickman, S.H., Sibson, R.H., Bruhn, R., 1995. Introduction to special section:
mechanical involvement of uids in faulting. Journal of Geophysical Research
100, 1283112840.
Holdsworth, R.E., 2004. Weak faults rotten cores. Science 303, 181182.
Holdsworth, R.E., Stewart, M., Imber, J., Strachan, R.A., 2001. The structure and
rheological evolution of reactivated continental fault zones: a review and case
study. In: Miller, J.A., Holdsworth, R.E., Buick, I.S., Handy, M.R. (Eds.), Continental
Reactivation and Reworking, vol. 184. Geological Society of London, pp. 115137.
Huang, W.L., Wyllie, P.J., 1974. Melting relations of muscovite with quartz and
sanidine in the K2O-Al2O3-SiO2-H2O system to 30 kilobars and an outline of
paragonite melting relations. American Journal of Science 274, 378395.
Ikari, M.J., Saffer, D.M., Marone, C., 2007. Effect of hydration state on the frictional
properties of montmorillonite-based fault gouge. Journal of Geophysical
Research 112. doi:10.1029/2006JB004748.
Imber, J., Holdsworth, R.E., Butler, C.A., Lloyd, G.E., 1997. Fault-zone weakening
processes along the reactivated Outer Hebrides Fault Zone, Scotland. Journal of
the Geological Society of London 154, 105109.
Jefferies, S.P., Holdsworth, R.E., Shimamoto, T., Takagi, H., Lloyd, G.E., Spiers, C.J.,
2006a. Origin and mechanical signicance of foliated cataclastic rocks in the
cores of crustal-scale faults: examples from the Median Tectonic Line, Japan.
Journal of Geophysical Research 111. doi:10.1029/2005JB004205.
Jefferies, S.P., Holdsworth, R.E., Wibberley, C.A.J., Shimamoto, T., Spiers, C.J.,
Niemeijer, A.R., Lloyd, G.E., 2006b. The nature and importance of phyllonite
development in crustal-scale fault cores: an example from the Median Tectonic
Line, Japan. Journal of Structural Geology 28, 220235.
Kanagawa, K., 2002. Frictional behavior of synthetic gouge-bearing faults under the
operation of pressure solution. Earth Planets and Space 54, 11471152.
Kanagawa, K., Cox, S.F., Zhang, S., 2000. Effects of dissolutionprecipitation
processes on the strength and mechanical behavior of quartz gouge at high-
temperature hydrothermal conditions. Journal of Geophysical Research 105,
11,11511,126.
Kronenberg, A.K., Kirby, S.H., Pinkston, J., 1990. Basal slip and mechanical anisotropy
of biotite. Journal of Geophysical Research 95, 1925719278.
Lachenbruch, A.H., Sass, J.H., 1980. Heat ow and energetic of the San Andreas fault
zone. Journal of Geophysical Research 85, 61856223.
Lehner, F.K., Bataille, J., 1984. Nonequilibrium thermodynamics of pressure solution.
Pure and Applied Geophysics 122, 5385.
E.W.E. Van Diggelen et al. / Journal of Structural Geology 32 (2010) 16851700 1699
Logan, J.M., Rauenzahn, K.A., 1987. Frictional dependence of gouge mixtures of
quartz and montmorillonite on velocity, composition and fabric. Tectonophysics
144, 87108.
Mares, V.M., Kronenberg, A.K., 1993. Experimental deformation of muscovite.
Journal of Structural Geology 15, 10611075.
Mariani, E., Brodie, K.H., Rutter, E.H., 2006. Experimental deformation of muscovite
shear zones at high temperatures under hydrothermal conditions and the
strength of phyllosilicate-bearing faults in nature. Journal of Structural Geology
28, 15691587.
Meike, A., 1989. In situ-deformation of micas: a high-voltage electron-microscope
study. American Mineralogist 74, 780796.
Miller, S.A., Nur, A., Olgaard, D.L., 1996. Earthquakes as a coupled shear stress-high
pore pressure dynamical system. Geophysical Research Letters 23, 197200.
Miller, S.A., Olgaard, D.L., 1997. Modeling seismicity clustering and fault weakness
due to high pore pressures. Physics and Chemistry of The Earth 22, 4348.
Moore, D.E., Lockner, D.A., 2004. Crystallographic controls on the frictional behavior
of dry and water-saturated sheet structure minerals. Journal of Geophysical
Research 109. doi:10.1029/2003JB002582.
Moore, D.E., Lockner, D.A., 2007. Comparative deformation behavior of minerals in
serpentinized ultramac rock: application to the slabmantle interface in
subduction zones. International Geology Review 49, 401415.
Moore, D.E., Lockner, D.A., Shengli, M., Summers, R., Byerlee, J.D., 1997. Strengths of
serpentinite gouges at elevated temperatures. Journal of Geophysical Research
102, 1478714801.
Moore, D.E., Rymer, M.J., 2007. Talc-bearing serpentinite and the creeping section of
the San Andreas fault. Nature 448, 795797.
Morrow, C.A., Moore, D.E., Lockner, D.A., 2000. The effect of mineral bond strength
and adsorbed water on fault gouge frictional strength. Geophysical Research
Letters 27, 815818.
Morrow, C.A., Radney, B., Byerlee, J.D., 1992. Frictional strength and the effective
pressure law of montmorillonite and illite clays. In: Evans, B., Wong, T.-F. (Eds.),
Fault Mechanics and Transport Properties of Rocks. Academic, San Diego, CA,
pp. 6988.
Nakatani, M., Scholz, C.H., 2004. Frictional healing of quartz gouge under hydro-
thermal conditions: 1. Experimental evidence for solution transfer healing
mechanism. Journal of Geophysical Research 109. doi:10.1029/2001JB001522.
Niemeijer, A.R., Spiers, C.J., 2005. Inuence of phyllosilicates on fault strength in the
brittle-ductile transition: insights from rock analogue experiments. In:
Bruhn, D., Burlini, L. (Eds.), High Strain Zones: Structure and Physical Properties,
vol. 245. Geological Society of London, pp. 303327.
Niemeijer, A.R., Spiers, C.J., 2006. Velocity dependence of strength and healing behav-
iour in simulated phyllosilicate-bearing fault gouge. Tectonophysics 427, 231253.
Niemeijer, A.R., Spiers, C.J., 2007. A microphysical model for strong velocity weak-
ening in phyllosilicate-bearing fault gouges. Journal of Geophysical Research
112. doi:10.1029/2007JB005008.
Niemeijer, A.R., Spiers, C.J., Bos, B., 2002. Compaction creep of quartz sand at 400
600 C: experimental evidence for dissolution-controlled pressure solution.
Earth and Planetary Science Letters 195, 261275.
Niemeijer, A.R., Spiers, C.J., Peach, C.J., 2008. Frictional behaviour of simulated
quartz fault gouges under hydrothermal conditions: results from ultra-high
strain rotary shear experiments. Tectonophysics 460, 288303.
OHara, K., 2007. Reaction weakening and emplacement of crystalline thrusts:
diffusion control on reaction rate and strain rate. Journal of Structural Geology
29, 13011314.
Ranalli, G., 1995. Rheology of the Earth. Chapman & Hall. 413p.
Rawling, G.C., Goodwin, L.B., 2003. Cataclasis and particulate ow in faulted, poorly
lithied sediments. Journal of Structural Geology 25, 317331.
Rutter, E.H., Holdsworth, R.E., Knipe, R.J., 2001. The nature and tectonic signicance
of fault zone weakening: an introduction. In: Holdsworth, R.E., Strachan, R.A.,
Magloughlin, J.F., Knipe, R.J. (Eds.), The Nature and Tectonic Signicance of Fault
Zone Weakening, vol. 186. Geological Society of London, pp. 111.
Rutter, E.H., Maddock, R.H., 1992. On the mechanical properties of synthetic kao-
lintie/quartz fault gouge. Terra Nova 4, 489500.
Rutter, E.H., Maddock, R.H., Hall, S.H., White, S.H., 1986. Comparative microstruc-
tures of natural and experimentally produced clay-bearing fault gouges. Pure
and Applied Geophysics (Historical Archive) 124, 330.
SAFOD core atlas, v.., 2007. Available from: www.earthscope.org.
Saileswaran, N., Panchanathan, V., 1973. Compaction of grains. General parameter
evaluation. Powder Technology 8, 1926.
Sammis, C.G., King, G., Biegel, R.L., 1987. The kinematics of gouge deformation. Pure
and Applied Geophysics 125, 777812.
Schleicher, A.M., Tourscher, S.N., van der Pluijm, B.A., Warr, L.N., 2009a. Constraints
on mineralization, uid-rock interaction, and mass transfer during faulting at
23 km depth from the SAFOD drill hole. Journal of Geophysical Research 114.
Schleicher, A.M., van der Pluijm, B.A., Warr, L.N., 2008. What Controls Creep on the
San Andreas Fault at the SAFOD Drillhole? Eos Transactions AGU 89 (53) Fall
Meet. Suppl. Abstract T13A1910.
Schleicher, A.M., Warr, L.N., van der Pluijm, B.A., 2009b. On the origin of mixed-
layered clay minerals from the San Andreas Fault at 2.53 km vertical depth
(SAFOD drillhole at Parkeld, California). Contributions to Mineralogy and
Petrology 157, 173187.
Scholz, C.H., 2002. The Mechanics of Earthquakes and Faulting, second ed. Cam-
bridge University Press Cambridge, UK. 471pp.
Scruggs, V.J., Tullis, T.E., 1998. Correlation between velocity dependence of friction
and strain localization in large displacement experiments on feldspar, musco-
vite and biotite gouge. Tectonophysics 295, 1540.
Shea, W.T., Kronenberg, A.K., 1992. Rheology and deformation mechanisms of an
isotropic mica schist. Journal of Geophysical Research 97, 1520115237.
Shimamoto, T., Logan, J.M., 1981. Effects of simulated clay gouges on the sliding
behavior of Tennessee sandstone. Tectonophysics 75, 243255.
Sibson, R.H., 1983. Continental fault structure and the shallow earthquake source.
Journal of the Geological Society of London 140, 741767.
Sibson, R.H., 2004. Controls on maximum uid overpressure dening conditions for
mesozonal mineralisation. Journal of Structural Geology 26, 11271136.
Sleep, N.H., 1995. Ductile creep, compaction, and rate and state dependent friction
within major fault zones. Journal of Geophysical Research 100, 1306513080.
Takahashi, M., Mizoguchi, K., Kitamura, K., Masuda, K., 2007. Effects of clay content
on the frictional strength and uid transport property of faults. Journal of
Geophysical Research 112. doi:10.1029/2006JB004678.
Tembe, S., Lockner, D.A., Solum, J.G., Morrow, C.A., Wong, T.-F., Moore, D.E., 2006a.
Frictional strength of cuttings and core from SAFOD drillhole phases 1 and 2.
Geophyscial Research Letters 33. doi:10.1029/2006GL027626.
Tembe, S., Lockner, D.A., Wong, T.-F., 2006b. Strength of SAFOD fault gouge under
hydrothermal conditions. Eos Transactions AGU 87 Abstract S23C-0181.
Theiler, G., Gradt, T., 2008. Inuence of the temperature on the tribological
behaviour of PEEK composites in vacuum environment. Journal of Physics:
Conference Series 100. doi:10.1088/1742-6596/100/7/072040.
Townend, J., Zoback, M.D., 2004. Regional tectonic stress near the San Andreas fault
in central and southern California. Geophysical Research Letters 31. doi:10.1029/
2003GL018918.
White, S.H., Bretan, P.G., Rutter, E.H., 1986. Fault-zone reactivation: kinematics and
mechanisms. Philosophical Transactions of the Royal Society of London A317,
8197.
Wibberley, C.A.J., 1999. Are feldspar-to-mica reactions necessarily reaction-soft-
ening processes in fault zones? Journal of Structural Geology 21, 12191227.
Wintsch, R.P., Christoffersen, R., Kronenberg, A.K., 1995. Fluid-rock reaction weak-
ening of fault zones. Journal of Geophysical Research 100, 1302113032.
Wu, F.T., Blatter, L., Roberson, H., 1975. Clay gouges in the San Andreas Fault System
and their possible implications. Pure and Applied Geophysics (Historical
Archive) 113, 8795.
Zhang, S., Tullis, T.E., Scruggs, V.J., 2001. Implications of permeability and its
anisotropy in a mica gouge for pore pressures in fault zones. Tectonophysics
335, 3750.
Zoback, M.D., Hickman, S.H., Ellsworth, W., Kirschner, D., Pennell, N.B., Chery, J.,
Sobolev, S., 2007. Preliminary Results from SAFOD Phase 3: implications for the
state of stress and shear localization in and near the San Andreas Fault at depth
in central California. Eos Transactions AGU 88 (52) Fall Meet. Suppl. Abstract
T13 G-03.
Zoback, M.D., Zoback, M.L., Mount, V.S., Suppe, J., Eaton, J.P., Healy, J.H.,
Oppenheimer, D., Reasenberg, P., Jones, L., Rayleigh, C.B., Wong, I.G., Scotti, O.,
Wentworth, C., 1987. New evidence on the state of stress of the San Andreas
fault system. Science 238, 11051111.
E.W.E. Van Diggelen et al. / Journal of Structural Geology 32 (2010) 16851700 1700
The effect of the intermediate principal stress on fault formation
and fault angle in siltstone
Bezalel Haimson
a,
*
, John W. Rudnicki
b
a
University of Wisconsin, Madison, WI 53706, USA
b
Northwestern University, Evanston, IL 60208, USA
a r t i c l e i n f o
Article history:
Received 6 February 2009
Received in revised form
5 July 2009
Accepted 25 August 2009
Available online 11 September 2009
Keywords:
Bifurcation
Fault angle
Faulting
Siltstone
MohrCoulomb
Shear localization
Strength criterion
True triaxial test
a b s t r a c t
We conducted true triaxial compression tests on specimens prepared from two siltstone core sections,
one above and one below the Chelungpu Fault, Taiwan. For different constant s
2
and s
3
magnitudes, the
maximum principal stress (s
1
) was raised until a post failure stage was reached, and a through-going
fault had developed. Despite differences between the properties of the two cores, in all tests peak s
1
increased as s
2
was set at higher levels than s
3
, in contrast to MohrCoulomb condition predictions. The
faultnormal vector was aligned with the s
3
direction and made an angle (q) with s
1
direction. The angle
q, which corresponds to fault dip in case of normal faulting, increased monotonically with s
2
for xed s
3
,
a variation that is also inconsistent with MohrCoulomb theory.
The results of shear band localization theory are used with fault angles observed for axisymmetric
compression and deviatoric pure shear to infer properties of the inelastic constitutive behavior. These
properties are signicantly different for the two cores. Using them to predict q for other deviatoric stress
states yields good agreement with the observations for core II and acceptable agreement for core I. The
results are used to predict the angle variation for constant mean normal stress (q decreases as the
deviatoric stress state varies from axisymmetric extension to axisymmetric compression) and at xed
deviatoric stress state (q decreases monotonically with increasing mean normal stress).
2009 Elsevier Ltd. All rights reserved.
1. Introduction
Laboratory experiments simulating compressive failure and
faulting in rocks are typically conducted on cylindrical specimens
subjected to constant lateral conning pressure and a rising axial
load until brittle fracture occurs. These conventional triaxial tests
replicate only a special case of crustal condition, that inwhich two of
the principal stresses are equal. Conventional triaxial tests on rocks
were conducted as early as the turn of the last century (Von Ka rma n,
1911). They gained acceptance because of the relatively simple
equipment, specimen preparation, and testing procedure. The
ubiquity of conventional triaxial testing can also be traced to the
assumption that the intermediate principal stress (s
2
) has a negli-
gible effect on rock failure characteristics as expressed, for example,
in the Mohr or MohrCoulomb failure criteria (Jaeger et al., 2007).
However, indications from the three major types of faulting
encountered in the eld, and results of numerous in situ stress
measurements at depths reaching several kilometers (McGarr and
Gay, 1978; Brace and Kohlstedt, 1980), point to a state of stress in
the earths crust that is fully three-dimensional (s
1
ss
2
ss
3
).
Murrell (1963), Handin et al. (1967), and Mogi (1967) compared
results of conventional triaxial compression tests with those of
conventional triaxial extension and deduced that the differences in
rock resistance to faulting between the two modes of loading were
due to the different magnitudes of s
2
applied. Inspired by this
evidence, Mogi (1971) introduced a true triaxial testing machine in
which rectangular prismatic specimens were subjected to three
different principal stresses. He found that indeed s
2
affects the
stress level at which faulting occurs, and hence the rock strength
criterion, as well as the angle at which the fault develops. Little
follow-up on Mogis seminal work took place until Haimson and co-
workers carried out similar true triaxial tests in igneous rocks
(Westerly granite, Haimson and Chang, 2000; Pohang rhyolite,
Chang and Haimson, 2007); a metamorphic rock (KTB amphibolite,
Chang and Haimson, 2000; Haimson and Chang, 2002) and a sedi-
mentary rock (siltstone, Oku et al., 2007). They found that strength,
deformability, and fault angle, were affected by s
2
in all the tested
rocks. An exception was found in tests of Long Valley (California)
* Corresponding author. Tel.: 1 608 262 2563.
E-mail address: bhaimson@wisc.edu (B. Haimson).
Contents lists available at ScienceDirect
Journal of Structural Geology
j ournal homepage: www. el sevi er. com/ l ocat e/ j sg
0191-8141/$ see front matter 2009 Elsevier Ltd. All rights reserved.
doi:10.1016/j.jsg.2009.08.017
Journal of Structural Geology 32 (2010) 17011711
ultra ne-grained hornfels and metapelite (Chang and Haimson,
2005), unusual rocks that appear to have no dilatancy, and develop
no visible microcracks.
Independently, building upon the antecedents of Hadamard
(1903), Mandel (1966), Thomas (1961) and Hill (1962), Rudnicki
and Rice (1975) (also, Rice, 1976; Besue lle and Rudnicki, 2004)
suggested a description of failure as a bifurcation from homoge-
neous (spatially uniform) deformation that predicted a strong
dependence on s
2
(via the deviatoric stress state). More specically,
Rudnicki and Rice (1975) established conditions for which a solu-
tion corresponding to concentrated deformation in a planar band
was an alternative to continued homogeneous deformation. The
appearance of this localized mode of deformation is often essen-
tially coincident with failure by development of a through-going
fault or fracture, but in other cases it may be the precursor to a more
extended evolution of localized deformation that ultimately
requires signicant additional strain for failure. Analysis based on
this approach yields a relation among constitutive parameters
required for the onset of bifurcation and, hence, depends strongly
on how the homogeneous deformation prior to bifurcation, espe-
cially the inelastic portion, is modeled.
The predictions of the failure stress by the bifurcation approach
depend strongly on certain details of the constitutive behavior that
are difcult to determine experimentally and, for various other
reasons, are difcult to compare with experimental observations of
failure (Besue lle and Rudnicki, 2004). However, the prediction for
the fault angle is much less sensitive to these details and is more
easily compared with observations in terms of the constitutive
parameters for homogeneous deformation just prior to bifurcation.
Rudnicki (2008a,b) has used and extended results from the bifur-
cation theory to interpret observations of failure plane inclinations
in true triaxial tests of Westerly Granite (Haimson and Chang,
2000) and to infer aspects of the constitutive behavior.
In this paper we describe two series of true triaxial tests con-
ducted on samples of siltstone taken fromthe hanging wall and the
footwall of the Chelungpu fault, Taiwan. The tests reveal a clear
dependency of strength and fault angle on the magnitude of s
2
. The
results are compared with predictions based on shear localization
theory incorporating a yield surface and plastic potential that
depend on three stress invariants (rather than two, as in Rudnicki
and Rice (1975)). Dependences of the yield surface and plastic
potential on mean stress are inferred fromthe fault angles observed
in axisymmetric compression and deviatoric pure shear. These
dependences are used to compare the predicted fault angles with
observations for other deviatoric stress states and to predict the
variation that would be observed with mean stress for xed
deviatoric stress state and with deviatoric stress state for xed
mean stress.
2. Rocks tested
Rock specimens used in the true triaxial tests described here
came from core recovered from the scientic hole A, near the
northern end of the Chelungpu fault. The hole was drilled as part of
the Taiwan Chelungpu Fault Drilling Project (TCDP). The project
was undertaken to study the faulting mechanism behind the
destructive Chi-Chi earthquake (1999; M
w
7.6), characterized as
a thrust motion across the North-South striking Chelungpu fault
(Shin and Teng, 2001; Lin et al., 2003). Core made available to us
came from short sections centered at the depth of 891 m (core I)
and 1252 m (core II), straddling the active fault, which was inter-
cepted at 1111 m. Core I is a siltstone belonging to the early Pleis-
tocene Cholan Formation, which persists to a depth of 1013 m; core
II is also a siltstone, belonging to the Pliocene Chinshui Formation,
which prevails at depths of 10131313 m. Thus, core II is
representative of the rock traversed by the active Chelungpu fault.
Core I comes from a somewhat younger formation. As shown
below, there are distinct differences in mineral content, as well as in
mechanical behavior between the two siltstones. We cannot tell
whether the differences are related to their juxtaposition with
respect to the fault and its activity or, what appears more likely, the
consequence of their different deposition ages.
The siltstone in core I contains 68% quartz, 19.5% clay, 9.5%
feldspar, and 3% biotite; the siltstone in core II consists of 65%
quartz, 25.5% clay, 7.5% feldspar, and 2% biotite. The only major
difference is the amount of clay, and this may have a role in the
disparity between the two core sections in some of their physical
and mechanical properties (Table 1).
The 891 m siltstone is both stronger and stiffer in compression
than the 1252 m core. Also notable is that core I rock has signi-
cantly larger grain size than core II. Although the siltstone is
a sedimentary rock, no bedding planes were visible, and neither
core showed signs of inhomogeneity. We examined the degree of
anisotropy by running unconned (uniaxial) compression tests in
which cylindrical specimens (25.4 mm in diameter) were drilled
out of the vertical core I at inclinations of 0
, 30
, 60
, and 90
. The
compressive strengths in specimens from the four inclinations
were all within 3% of the mean 79.4 MPa. Hence, we considered the
siltstone to be basically isotropic with respect to strength. With
respect to elastic modulus, the largest difference from the mean
13.7 GPa was less than 7%, a sign of a rather mild degree of
anisotropy. These results enabled us to consider the siltstone
practically isotropic.
3. Experiment setup and procedure
The apparatus used in our tests was a recently fabricated true
triaxial testing machine (Haimson and Chang, 2000). As stated
above, it had been successfully employed to characterize mechan-
ical properties under the most general 3Dstate of stress of Westerly
granite, KTB amphibolite, Long Valley hornfels and metapelite, and
Pohang rhyolite. The true triaxial cell facilitates the application of
three principal stresses to a rectangular prismatic specimen (size
19 19 38 mm
3
) by use of three independent servo-controlled
units. The application of the maximum principal stress (s
1
) in the
axial direction of the specimen and of the intermediate principal
stress (s
2
) in one of the two lateral directions is carried out by use of
two pairs of hydraulically driven pistons. The minimum principal
stress (s
3
) is directly applied to the other pair of lateral faces by
conning uid pressure inside the cell. Details of the testing
system, and its calibration, can be found in Haimson and Chang
(2000).
In selected specimens strain gages for measuring strains in the
direction of s
1
and s
2
were afxed to the faces subjected to s
3
loading. Strain in the s
3
direction was measured using a beryllium
copper strain-gaged beam mounted on one of the specimen s
3
faces.
Table 1
Some physical and mechanical properties of core I and core II.
Property Core I (891 m) Core II (1252 m)
Mean grain size, mm 56 (200) 44 (200)
Dry density, kg/m
3
2594 (28) 2587 (28)
Effective porosity, % 6.9 (15) 6.1 (15)
Unconned compressive strength
(UCS), MPa
79.5 (2) 63.4 (2)
Youngs modulus, GPa 13.7 (2) 9.2 (2)
Poissons ratio 0.13 (2) 0.2 (2)
Brazilian tensile strength, MPa 5.4 (11) 4.2 (3)
Note: Numbers in parenthesis refer to the amount of measurements conducted.
B. Haimson, J.W. Rudnicki / Journal of Structural Geology 32 (2010) 17011711 1702
The tests reported here were conducted under dry conditions.
For that purpose the s
3
faces, as well as the spaces between the
edges of the anvils applying the other two loads were coated with
a thin layer of polyurethane so that the conning uid could not
penetrate into the specimen. The scarcity of available core pre-
vented us from testing saturated siltstone under pore pressure.
However, that condition would have only affected the strength
magnitudes, not the general mechanical behavior.
Testing procedure consisted of raising s
1
at a constant strain rate
of 8 10
6
s
1
while holding s
2
and s
3
at their preset magnitudes,
until a post failure stage was reached, at which s
1
had declined
about 10% from its peak level. Upon unloading, tested specimens
were sectioned along the s
1
s
3
plane in order to record fault atti-
tude. In selected samples, sections were also prepared for SEM
inspection.
4. Triaxial strength
4.1. Conventional triaxial strength
Conventional triaxial tests (s
2
s
3
) were rst carried out in
order to establish the Mohr strength criterion in terms of principal
stresses, for later comparison with the true triaxial compressive
strength (s
2
>s
3
). The Mohr criterion was obtained by tting
a monotonically increasing power function to the experimental
data (Fig. 1). In core I the best-tting criterion is expressed by
a power function:
s
1
76:89 9:34s
0:80
3
R 0:997 (1a)
In core II the Mohr criterion takes the form:
s
1
73:23 6:36s
0:84
3
R 0:998 (1b)
The tting is excellent in both cases, with core II showing lower
conventional triaxial strength. The experimental results in both
cores can also be tted by linear regression (MohrCoulomb
criterion):
s
1
94:9 3:6s
3
R 0:992 for core I (2a)
s
1
84:2 3:0s
3
R 0:992 for core II (2b)
Again, core II proves to be weaker throughout the range of s
3
(Fig. 2).
4.2. True triaxial strength
True triaxial tests were conducted for several constant magni-
tudes of s
3
. For a given s
3
, a series of tests were run, each at
different s
2
between s
2
s
3
and s
2
approaching peak s
1
. The
results in terms of peak s
1
versus s
2
for the different preset s
3
are
plotted in Fig. 3a and b. The solid line in the plots represents the
Mohr criterion determined from tests in which s
2
s
3
. Dashed
curves showthe trend of the true triaxial strength for each constant
s
3
. The plots reveal a common characteristic of gradually increasing
strength with the rise in s
2
until a top level is reached, followed by
a gradual decline, as predicted theoretically by Wiebols and Cook
(1968) and conrmed by tests in other rocks by Mogi (1971),
Haimson and Chang (2000), and Chang and Haimson (2000, 2007).
The higher compressive strength when s
2
>s
3
as compared with
that at s
2
s
3
reveals the inadequacy of the Mohr (or Mohr
Coulomb) criterion to predict rock failure under the most general
state of stress.
We attempted to determine the true triaxial strength criteria for
siltstones by rst employing Nadais (1950) proposed relationship.
He suggested a true triaxial strength criterion for brittle materials,
such as rocks, in terms of the two stress invariants, octahedral shear
stress (s
oct
) and octahedral normal stress, or mean stress (s
oct
). He
related these invariants by a function (f) dependent on the rock
material properties in the form of s
oct
f s
oct
. Plotting all test
results shown in Fig. 3 in the Nadai domain (Fig. 4) yields the
following best-t power function curves:
s
oct
5:27s
0:60
oct
R 0:927 for core I (3a)
0
100
200
300
400
500
0 20 40 60 80 100 120
1
(
M
P
a
)
2
=
3
(MPa)
1
= 76.89 + 9.338
3
0.80
R = 0.997
Core I
1
= 73.23 + 6.36
3
0.84
R = 0.998
Core II
Fig. 1. Maximum compressive principal stress (s
1
) at failure as a function of the least
principal stress (s
3
) under conventional triaxial stress condition (s
2
s
3
), and the best-
tting power function strength criterion (Mohr) for each of the two siltstone cores.
0
100
200
300
400
500
0 0 2 40 60 80 100 120
1
(
M
P
a
)
2
=
3
(MPa)
1
= 84.2 + 3.0
3
R = 0.992
Core II
1
= 94.9+ 3.6
3
R = 0.992
Core I
Fig. 2. Maximum compressive principal stress (s
1
) at failure as a function of the least
principal stress (s
3
) under conventional triaxial stress condition (s
2
s
3
), and the best-
tting linear strength criterion (MohrCoulomb) for each of the two siltstone cores.
B. Haimson, J.W. Rudnicki / Journal of Structural Geology 32 (2010) 17011711 1703
s
oct
5:10s
0:58
oct
R 0:915 for core II (3b)
where
s
oct
_
_
s
1
s
2
2
s
2
s
3
2
s
3
s
1
2
_
1=2
__
3 (4)
and
s
oct
s
1
s
2
s
3
=3 (5)
The Nadai criterion appears to t the data reasonably well, albeit
with considerable scatter.
Mogi (1971) adjusted the Nadai criterion to correspond with the
mode of compressive failure in brittle rock, which does not occur
over the entire volume (as expressed by Eqs. (3)), but is restricted to
faulting along a plane aligned with the s
2
direction. For this reason
he replaced s
oct
with the mean stress acting on the plane of failure
(s
m,2
). All data points in Fig. 3 are well-tted by a power function in
Mogis domain (Fig. 5):
0
100
200
300
400
500
600
(
M
P
a
)
2
(MPa)
Mohr criterion (
2
=
3
)
2
=
1
Uniaxial compressive strength C
0
25 MPa
40 MPa
60 MPa
100 MPa
3
= 10 MPa
Core I
0
100
200
300
400
500
600
0 100 200 300 400 500 0 100 200 300 400 500
(
M
P
a
)
2
(MPa)
Mohr criterion (
2
=
3
) 2
=
1
3
= 10 MPa
40 MPa
60 MPa
100 MPa
Uniaxial compressive strength C
0
Core II
a b
Fig. 3. Variation of peak compressive stress s
1
as a function of s
2
for different constant values of s
3
(a) in the core I and (b) in core II.
0
50
100
150
200
0 50 100 150 200 250 300 350
o
c
t
(
M
P
a
)
oct
(MPa)
oct
(MPa) =5.27
oct
0.60
R = 0.927
Core I
oct
(MPa) = 5.10
oct
0.58
R = 0.915
Core II
Fig. 4. A true triaxial strength criterion for the siltstone (cores I and II), based on all the
experimental results shown in Fig. 3, in the Nadai (1950) domain of s
oct
versus s
oct
.
0
50
100
150
200
0 50 100 150 200 250 300 350
o
c
t
(
M
P
a
)
m,2
(MPa)
oct
(MPa) = 2.32
m,2
0.75
R = 0.995
Core I
oct
(MPa) = 2.86
m,2
0.69
R = 0.991
Core II
Fig. 5. A true triaxial strength criterion for the siltstone (cores I and II), based on all the
experimental results shown in Fig. 3, in the Mogi (1971) domain of s
oct
versus s
m,2
.
B. Haimson, J.W. Rudnicki / Journal of Structural Geology 32 (2010) 17011711 1704
s
oct
2:32s
0:75
m;2
R 0:995 for core I (6a)
s
oct
2:86s
0:69
m;2
R 0:991 for core II (6b)
where
s
m;2
s
1
s
3
=2 (7)
Comparison between Figs. 4 and 5 makes it clear that Mogis
relationship (Eqs. (6)) better represents the true triaxial strength
criteria for both cores. A physical interpretation of Eqs. (6) is that
brittle failure, or faulting, occurs when the distortional strain
energy reaches a critical value that increases monotonically with
the mean normal stress on the failure plane.
5. Fault angle
Failure in true triaxial tests took the form of single or conju-
gate through-going faulting. The faultnormal direction and the
angle (q) between it and s
1
were carefully measured at the
conclusion of each test. Under conventional triaxial stresses fault
normal direction was random since s
2
s
3
. Conventional triaxial
strength data appear to be tted well by the MohrCoulomb
criterion (Fig. 2). The linear relationship represented by this
criterion enables the computation of the unique angle of internal
friction f:
f 34:6
34:6=2 62:3
30:3=2 60:2
to 59
to 52
in
core II, as s
2
s
3
increases from 10 MPa to 100 MPa (Fig. 6). A
similar trend of fault angle decrease with conning pressure
increase has been documented in other rocks, such as Fontaine-
bleau sandstone (Haied et al., 2000), Westerlygranite (Haimson and
Chang, 2000), and KTB amphibolite (Chang and Haimson, 2000).
This variation of the fault angle with lateral stress contradicts the
prediction based on the MohrCoulomb criterion (Eqs. (9)).
In true triaxial tests the trend in both cores is for the fault angle
to increase with s
2
for constant s
3
. The total increase is limited to
about 10
27
p
J
3
=2s
3
where J
3
dets
ij
is the thirdinvariant of the deviatoric stress. (The
difference in sign from Rudnicki (2008a,b) occurs because stresses
are taken as positive in compression here). In the space of principal
stresses, q
L
is the angle in planes that are normal to the hydrostat
s
1
s
2
s
3
and for which s
oct
is constant. The angle q
L
is zero for
deviatoric pure shear s
2
s
1
s
3
=2, and varies between
p=630
1=2
where de
p
ij
is the
deviatoric portion of the inelastic strain increment. The direction of
inelastic strain increments is taken to be normal to the projection of
the yield surface in deviatoric planes (planes normal to the
hydrostatic axis) because there are neither observational evidence
nor physical grounds indicating otherwise. As a result, the devia-
toric parts of the normals to the yield surface and plastic potential
are identical,P
0
ij
Q
0
ij
, where the prime denotes the deviatoric part,
and G
s
F
s
and G
q
L
F
q
L
.
With these denitions and assumptions the fault angle is given
by
q p=4 1=2arcsin a (10)
where
a
1=31 nP
kk
Q
kk
Q
0
2
1 2n
4
_
Q
0
ij
Q
0
ij
_
3
_
Q
0
2
_
2
_ (11)
v is Poissons ratio and Q
0
2
is the intermediate principal deviatoric
value of Q
ij
Eq. (11) can be expressed in terms of derivatives of the
yield function and plastic potential as
a
1 ncos jF
s
G
s
=F
s
1 2nsinj q
L
=
3
p
cosj q
L
(12)
by using the following relations:
50
55
60
65
70
75
80
F
a
u
l
t
a
n
g
l
e
(
d
e
g
)
2
(MPa)
3
= 10 MPa
25 MPa
40 MPa
60 MPa
100 MPa
core I
50
55
60
65
70
75
80
0 50 100 150 200 250 300 350 400 0 50 100 150 200 250 300 350
F
a
u
l
t
a
n
g
l
e
(
d
e
g
)
2
(MPa)
3
= 10 MPa
40 MPa
60 MPa
100 MPa
Core II
a b
Fig. 7. Measured fault angle (q) variation with s
2
for different magnitudes of constant s
3
(a) for core I, and (b) for core II.
50 100 150 200 250 300 350
52
56
60
64
68
72
76
core I
10 MPa
25 MPa
60 MPa
40 MPa
100 MPa
F
a
u
l
t
A
n
g
l
e
(
d
e
g
)
Mean Compressive Stress,
0ct
(MPa)
core II
10 MPa
40 MPa
60 MPa
100 MPa
Fig. 8. Measured fault angle (q) variation with s
oct
for different magnitudes of constant
s
3
(different symbols) for core I (open) and core II (solid).
B. Haimson, J.W. Rudnicki / Journal of Structural Geology 32 (2010) 17011711 1706
P
kk
3G
s
; Q
kk
3F
s
; tan j
F
q
L
=s
F
s
; 2Q
0
ij
Q
0
ij
_
F
q
L
=s
_
2
F
s
2
(13)
and
Q
0
2
Q
0
ij
_
Q
0
ij
=2
2
3
p sinj q
L
(14)
Eq. (12) reduces to the Rudnicki and Olsson (1998) form of the
Rudnicki and Rice (1975) expression by setting j 0, noting that
F
s
=F
s
3m, G
s
=F
s
3b and sin q
L
3
p
=2N, where
N s
2
=s is the deviatoric stress state parameter used by Rudnicki
and Rice (1975) (again, the minus sign occurs here because stresses
are positive in compression).
To use Eq. (12) to make predictions for comparisonwith data it is
necessary to adopt a specic form for the yield condition. Rudnicki
(2008a) showed that although limited agreement with data on
Westerly granite could be achieved with a two invariant constitu-
tive model (no dependence on q
L
), better agreement was possible
with a yield condition of the following form:
4
27
_
A sin3q
L
_
s
s
0
_
3
_
s
s
0
_
2
1 0 (15)
where 0 A 1. If A 0, then the shape of the yield surface in the
deviatoric plane reduces to a circle with the radius determined by
s
0
s
oct
. When A 1, the shape reduces to a triangular Rankine
type model, in which yield is determined by a critical value of the
least compressive stress. Rudnicki (2008b) has discussed how
assuming that s
0
is proportional to s
oct
and particular choices for
the constant of proportionality and A duplicate the forms suggested
by Lade and Duncan (1975) and Matsuoka and Nakai (1974) (also
described in Borja et al., 2003).
In general, A, in addition to s
0
, could depend on the octahedral
normal stress, s
oct
, but we use a constant value of 0.7 here. Limited
experimentation indicated no strong dependence on A (as long as it
was not near its limits) but we made no attempt to optimize the
choice to agree with the data.
Because the rst termin Eq. (15) vanishes for q
L
0, s
0
(s
oct
) gives
the mean stress dependence of the yield stress in deviatoric pure
shear s
2
s
1
s
3
=2. Evaluating the third of Eqs. (13) for Eq.
(15) yields an expression for tan j and the ratio F
s
=sF
s
is given by
52
56
60
64
68
72
76
80
F
a
u
l
t
a
n
g
l
e
(
d
e
g
)
2 sin(
L
) 2 sin(
L
)
10 MPa
25 MPa
40 MPa
60 MPa
100 MPa
core I
-0.8 -0.4 0.0 0.4 0.8 1.2 -1.2 -0.8 -0.4 0.0 0.4 0.8 1.2
52
56
60
64
68
72
76
10 MPa
40 MPa
60 MPa
100 MPa
F
a
u
l
t
a
n
g
l
e
(
d
e
g
)
Core II
a
b
Fig. 9. Measured fault angle (q) variation with deviatoric stress state, from axisymmetric extension (1, left) to axisymmetric compression (1, right) for different magnitudes of
constant s
3
(different symbols) for core I (a) and core II (b).
50 100 150 200 250 300
60
64
68
72
76
Axisym Comp
(
L
= /6))
76.7 - 0.089 * x
F
a
u
l
t
a
n
g
l
e
(
d
e
g
)
F
a
u
l
t
a
n
g
l
e
(
d
e
g
)
Mean Compressive Stress,
oct
(MPa)
Core I
Dev Pure Shear
(
L
= 0)
79.7 - 0.051 * x
50 100 150 200 250 300 350
52
56
60
64
68
72
76
80
Dev Pure Shear
L
= 0
84.2 - 0.087*x
Axisym Com
L
= /6
84.6 - 0.160*x
Core II
Mean Compressive Stress,
oct
(MPa)
a
b
Fig. 10. Fault angle (q) variation with s
oct
for axisymmetric compression (q
L
p/6) and deviatoric pure shear (q
L
0) for core I (a) and core II (b).
B. Haimson, J.W. Rudnicki / Journal of Structural Geology 32 (2010) 17011711 1707
F
s
sF
s
s
s
0
ds
0
s
oct
ds
oct
(16)
because s s
0
for pure shear q
L
0, Eq. (16) gives a friction coef-
cient (which will be different for other deviatoric stress states, e.g.,
axisymmetric compression or extension). A similar ratio could be
inferred from the data shown in Fig. 4, but would pertain to the
slope of the failure surface rather than to the yield function. The
failure surface can be quite different from yield surface (even
evaluated at failure) (see, e.g., Holcomb and Rudnicki (2001) or
Besue lle and Rudnicki (2004)).
A consequence of the requirement that normality be satised in
the deviatoric plane (see end of paragraph preceding Eq. (10)) is
that the yield function and plastic potential can differ only by
a function of s s
oct
that we denote by H(s). A dilatancy factor b,
the ratio of the inelastic increment of volume strain to inelastic
increment of shear strain (dened earlier) is given by
b
3
2
1
1 A
2
=3
_
_
2
ds
0
ds
oct
s
0
H
0
s
oct
_
(17)
As for Eq. (16), this expression applies for pure shear and would be
slightly different for other deviatoric stress states.
7. Application to TCDP data
Because the yield surface and predictions for the band angle are
expressed in terms of invariants of the stress rather than the
principal stresses themselves, we plot the observed fault angle
versus the mean compressive stress s
oct
(in Fig. 8) and the Lode
angle q
L
(in Fig. 9). If both the mean stress and Lode angle are
known then the value of s is determined by the yield condition. In
Fig. 8, both core I and core II showan approximately linear increase
in fault angle with s
oct
for xed values of least compressive stress,
but an overall decrease in band angle with increasing mean stress.
Figs. 9a (core I) and Fig. 9b (core II) showthe fault dip angle plotted
against 2 sin q
L
which varies from 1 on the left for axisymmetric
extension, through 0 for deviatoric pure shear, to 1 for axisym-
metric compression on the right. Data from both cores appear to
show a slightly decreasing dip angle for increasing Lode angle, but
this is more evident for core II.
As already noted, the tests were conducted for xed values of
the least compressive stress. Consequently, neither s
oct
nor q
L
is
constant in the tests and it is difcult to infer the dependence on
either from Figs. 8 and 9. Nevertheless, within both data sets are
several tests for axisymmetric compression s
2
s
3
; q
L
p=6
and near pure shear s
2
s
1
s
3
=2; q
L
0. Here near
means j2 sin q
L
j 0:05 for core I and 0.08 for core II. Fig. 10a (core I)
and Fig. 10b (core II) show the fault angles for axisymmetric
compression and pure shear against s
oct
and linear ts through the
data for each deviatoric stress state. All the slopes are negative, but
the magnitudes are greater for core I than for core II. For each core
the magnitude of the slope is larger for axisymmetric compression
than for deviatoric pure shear, by nearly a factor of 2 for core II.
Substituting the linear relations shown in Fig. 10 into Eq. (12),
using Eq. (15), and evaluating for pure shear (q
L
0) and axisym-
metric compression (q
L
p/6) yields two linear equations for the
unknown functions of s: ds
0
=ds (Eq. (16)) and H
0
s (Eq. (17)).
Fig. 11 shows the solutions for the two data sets. For both the
variation is greater for core II than for core I.
Values of ds
0
=ds, which has the interpretation of a friction
coefcient, are reasonable, though, perhaps, on the low side, for
both core I and core II. The slight increase with mean stress is,
however, surprising as friction coefcient nearly always decreases
with increasing mean stress. For comparison, a friction coefcient
-50 50 1001 50 200 250 300 350 400
-1.2
-0.8
-0.4
0.0
0.4
0.8
1.2
Mean Compressive Stress,
oct
(MPa)
d
0
/d
Dilatancy Factor
0
Fig. 11. Solutions for ds
0
=ds, Eq. (16), and the dilatancy factor b, Eq. (17), as functions
of s
oct
based on linear ts in Fig. 10. Solid lines show results for core I and dashed for
core II.
60
64
68
72
76
10 MPa
25 MPa
40 MPa
60 MPa
100 MPa
F
a
u
l
t
A
n
g
l
e
(
d
e
g
)
Mean Compressive Stress (MPa)
Core I
50 100 150 200 250 300 350 50 100 150 200 250 300
50
55
60
65
70
75
80
10 Mpa
40 MPa
60 MPa
100 MPa
F
a
u
l
t
A
n
g
l
e
(
d
e
g
)
Mean Compressive Stress (MPa)
Core II
a
b
Fig. 12. Comparison of predicted (open symbols) with observed (lled symbols) fault angles plotted against mean compressive stress (s
oct
) for core I (a) and core II (b).
B. Haimson, J.W. Rudnicki / Journal of Structural Geology 32 (2010) 17011711 1708
calculated as the slope of the curve for core I in Fig. 4 decreases from
0.66 at 50 MPa to 0.30 at 350 MPa. That for core II decreases from
0.57 to 0.25 over the same range of compressive stress. The increase
shown in Fig. 11 is, however, not large and given the idealizations of
the model and variation in the data, it has questionable
signicance.
The values for the dilatancy factor decrease substantially with
mean stress. This is as expected but the magnitude of the values at
lowand high conning stresses are surprisingly large. The values at
low mean stress are around 1 and indicate very strong dilatancy.
Values for core I decrease to a small negative value, about 0.3 at
the highest mean stress. Although a decrease of dilatancy with
mean stress is reasonable, it is less likely that compaction would
occur, even at the highest mean stress, given the relatively low
porosity and the history of signicant shear. The very strong
compaction predicted for core II at high mean stress is not realistic.
Some aspects of the functions plotted in Fig. 11 are in accord
with expected behavior, but it seems difcult to assign them more
than qualitative signicance. It is, however, interesting that the
behavior inferred for the two cores is signicantly different. Despite
the questionable aspects of the plots in Fig. 11, they are based on the
observed decrease in band angle with mean stress for pure shear
and axisymmetric compression shown in Fig. 10 and will, of course,
reproduce this behavior. This is a signicant improvement over
constitutive models based on only two of the stress invariants. In
addition, the predicted band angle (see Eq. (12)) depends primarily
on the sum of the two functions. For this reason, it seems that the
slight increase with mean stress predicted for the friction coef-
cient is compensated by an excessive decrease in the dilatancy
factor. At this point, it is unclear what improvement might lead to
a more reasonable division.
The functions of mean stress shown in Fig. 11 are then re-
substituted into Eq. (12) using Eq. (15) and used to evaluate the
predicted band angles for all the Lode angles in the measured data.
Fig. 12a (core I) and Fig. 12b (core II) show the fault angle data and
predictions plotted against the mean compressive stress (with the
open symbols showing the predictions). Clearly, the agreement is
much better for core II (correlation 95%) than for core I (correlation
80%). Fig. 13 plots the same data against the Lode angle parameter
2 sin q
L
. Not surprisingly, for both the agreement is better for
positive values since the mean stress variation used the data for
pure shear q
L
0 and axisymmetric compression (q
L
p/6).
Figs. 14 and 15 show the predicted fault angles if tests were
conducted at xed values of mean stress or of Lode angle q
L
. Fig. 14
plots the fault angle against mean compressive stress for ve values
of q
L
corresponding to xed deviatoric stress states. As expected,
-0.8 -0.4 0.0 0.4 0.8 1.2
52
56
60
64
68
72
76
80
F
a
u
l
t
A
n
g
l
e
(
d
e
g
)
2 sin(
L
) 2 sin(
L
)
10 MPa
25 MPa
40 MPa
60 MPa
100 MPa
core I
-1.2 -0.8 -0.4 0.0 0.4 0.8 1.2
50
55
60
65
70
75
80
10 Mpa
40 MPa
60 MPa
100 MPa
F
a
u
l
t
A
n
g
l
e
(
d
e
g
)
Core II
a
b
Fig. 13. Comparison of predicted (open symbols) with observed (lled symbols) fault angles plotted against deviatoric stress state 2 sin q
L
for core I (a) and core II (b).
-50 0 50 100 150 200 250 300 350 400
40
50
60
70
80
90
Axisym Comp (
L
= )
L
=
Pure Shear (
L
= )
L
=
Axisym Ext (
L
= )
F
a
u
l
t
A
n
g
l
e
(
d
e
g
)
Mean Compressive Stress (MPa)
Core I
-50 0 50 100 150 200 250 300 350 400
20
30
40
50
60
70
80
90
AxiSym Comp (
L
/6)
L
= /12
Pure Shear (
L
= 0)
L
= /12
AxiSym Ext (
L
/6)
F
a
u
l
t
A
n
g
l
e
(
d
e
g
)
Mean Compressive Stress(MPa)
Core II
a b
Fig. 14. Predicted variation of the fault angle against mean compressive stress (s
oct
) for constant deviatoric stress states 2 sin q
L
for core I (a) and core II (b).
B. Haimson, J.W. Rudnicki / Journal of Structural Geology 32 (2010) 17011711 1709
the fault angles decrease with increasing mean stress. The straight
lines for axisymmetric compression and pure shear reect the use
of these data in the tting. For both cores dilation bands (perpen-
dicular to least compressive stress; extension fractures) are pre-
dicted for axisymmetric extension at the lowest mean normal
stresses, though stress states corresponding to q <q
L
are not well
populated by the data. For axisymmetric compression at the
highest mean stresses the angles extend down to slightly less than
50
, suggesting the
possibility of compaction band formation at mean stresses some-
what higher than achieved in the tests. This prediction is also
consistent with the strong compaction inferred for core II at higher
mean stresses (Fig. 11b).
Fig. 15 plots the fault angle against the deviatoric stress state
parameter for ve constant values of the mean compressive stress.
The band angle decreases as the deviatoric stress state varies from
axisymmetric extension (left side) to axisymmetric compression
(right side). This trend is consistent with the predictions of the
simpler constitutive relation used by Rudnicki and Rice (1975)
although the decrease here is not so large. Because the fault angle
also decreases with mean compressive stress, the variations with
mean stress and deviatoric stress states can offset or augment each
other. The results for core II predict a dilation band for the lowest
mean stress. Although neither result shows a compaction band, the
fault angle does decrease rapidly approaching axisymmetric
compression for the highest mean stress values shown. Rudnicki
(2004) has suggested this rapid decrease as a possible reason for
the infrequent observation of dip angles in this range.
8. Conclusions
True triaxial tests on two siltstone cores from the TCDP hole A
reveal that the intermediate principal stress is a signicant
contributor to their compressive strength, and bring into question
the suitability of the Mohr and MohrCoulomb strength criteria,
which neglect the effect of s
2
. Rather, strength criteria in terms of
the invariants octahedral shear stress and the 2-D mean stress in
the s
1
s
3
plane t well the experimental data.
The angle of the fault created upon brittle fracture, is also
strongly affected by s
2
. It is found that the angle rises with s
2
for
constant s
3
, further questioning the adequacy of the Mohr-type
criteria, which predict a fault angle independent of s
2
.
The variation of the fault angle with mean compressive stress
and deviatoric stress state is modeled using localization theory
with a three invariant form for the yield function and plastic
potential. Calibration of the results using subsets of the data for
deviatoric pure shear and axisymmetric compression yields two
inelastic properties that are functions of the mean stress. Incor-
porating these inferred functions, the predictions agree well with
the entire data set for core II and acceptably with that for core I. The
results are then used to predict the variation for the band angle for
true triaxial tests conducted at constant mean stress and xed
deviatoric stress state. The fault angle at constant mean normal
stress is predicted to decrease as the deviatoric stress state varies
from axisymmetric extension to axisymmetric compression. The
fault angle at xed deviatoric stress state is predicted to decrease
monotonically with increasing mean normal stress.
Although the inferred inelastic constitutive functions are
reasonable, there is no independent verication of their form. Nor
are there true triaxial tests available at constant mean stresses or
deviatoric stress state. Nevertheless the variation of observed fault
angle is clearly not well-described by the simple MohrCoulomb
type theory. Comparison with the more elaborate theory, despite
the uncertainties involved, yields predictions that could be evalu-
ated by further testing and offers insight into the inelastic consti-
tutive behavior of the rock and its relation to failure.
Acknowledgements
Partial nancial support for JWR was provided by the US Dept. of
Energy, Ofce of Science, Basic Energy Sciences, Geosciences
Program through grant DE-FG02-93ER14344/A016 to North-
western University. Partial nancial support for BHwas provided by
NSF grant EAR-0346141. Thanks are extended to H. Oku for carrying
out the laboratory experiments and to Florent Gimbert for many
helpful discussions about three invariant constitutive relations.
Core fromthe TCDP was obtained courtesy of Professor Sheng-Rong
Song, National Taiwan University. The authors also thank Philip
Benson and an anonymous reviewer for comments that improved
the paper.
References
Besue lle, P., Rudnicki, J.W., 2004. Localization: shear bands and compaction bands.
In: Gue guen, Y., Boute ca, M. (Eds.), Mechanics of Fluid Saturated Rocks. Inter-
national Geophysics Series, 89. Academic Press, London, pp. 219321.
Borja, R.I., Sama, K.M., Sanz, P.F., 2003. On the numerical integration of three-
invariant elastoplastic constitutive models. Computer Methods in Applied
Mechanics and Engineering 192, 12271258.
-1.2 -0.8 -0.4 0.0 0.4 0.8 1.2
30
40
50
60
70
80
90
F
a
u
l
t
A
n
g
l
e
(
d
e
g
)
2 sin(
L
)
Core I
50 MPa
150 MPa
250 MPa
350 MPa
450 MPa
-1.2 -0.8 -0.4 0.0 0.4 0.8 1.2
0
10
20
30
40
50
60
70
80
90
F
a
u
l
t
A
n
g
l
e
(
d
e
g
)
2 sin(
L
)
Core II
50 MPa
150 MPa
250 MPa
350 MPa
450 MPa
a b
Fig. 15. Predicted variation of the fault angle against deviatoric stress state 2 sin q
L
at ve constant values of s
oct
for core I (a) and core II (b).
B. Haimson, J.W. Rudnicki / Journal of Structural Geology 32 (2010) 17011711 1710
Brace, W.F., Kohlstedt, D.L., 1980. Limits on lithospheric stress imposed by labora-
tory experiments. Journal of Geophysical Research 85 (B11), 62486252.
Chang, C., Haimson, B.C., 2000. True triaxial strength and deformability of the KTB
deep hole amphibolite. Journal of Geophysical Research 105, 1899919014.
Chang, C., Haimson, B.C., 2005. Nondilatant deformation and failure mechanism in
two Long Valley Caldera rocks under true triaxial compression. International
Journal Rock Mechanics and Mining Sciences 42, 402414.
Chang, C., Haimson, B., 2007. Effect of uid pressure on rock compressive failure in
a nearly impermeable crystalline rock: implication on mechanism of borehole
breakouts. Engineering Geology 89, 230242.
Hadamard, J., 1903. Leons sur la Propagation de Ondes et Les Eqs. de LHy-
drodynamique, Paris.
Haied, A., Kondo, D., Henry, J.P., 2000. Strain localization in Fontainebleau sand-
stone. Mechanics of CohesiveFrictional Materials 5, 239253.
Haimson, B., Chang, C., 2000. A new true triaxial cell for testing mechanical prop-
erties of rock, and its use to determine rock strength and deformability of
Westerly granite. International Journal of Rock Mechanics and Mining Sciences
17, 285296.
Haimson, B., Chang, C., 2002. True triaxial strength of the KTB amphibolite under
borehole wall conditions and its use to estimate the maximum horizontal in
situ stress. Journal of Geophysical Research 107 (B10) ETG 15-1 to 14.
Handin, J., Heard, H.C., Magouirk, J.N., 1967. Effect of the intermediate principal
stress on the failure of limestone, dolomite, and glass at different temperature
and strain rate. Journal of Geophysical Research 72, 611640.
Hill, R., 1962. Acceleration waves in solids. Journal of the Mechanics and Physics of
Solids 19, 116.
Holcomb, D.J., Rudnicki, J.W., 2001. Inelastic constitutive properties and shear
localization in Tennessee marble. International Journal for Numerical and
Analytical Methods in Geomechanics 25, 109129.
Jaeger, J.C., Cook, N.G.W., Zimmerman, R., 2007. Fundamentals of Rock Mechanics,
fourth ed. Blackwell Publishers, 475 pp.
Lade, P.V., Duncan, J.M., 1975. Elasto-plastic stressstrain theory for cohesionless
soil. Journal of the Geotechnical Engineering Division American Society of Civil
Engineers 101 (GT10), 10371053.
Lin, C.-W., Lee, Y.-L., Huang, M.-L., Lai, W.-C., Yuan, B.-D., Huang, C.-Y., 2003. Char-
acteristics of surface ruptures associated with the Chi-Chi earthquake of
September 21, 1999. Engineering Geology 71, 1330.
Mandel, J., 1966. Conditions de stabilte et postulat de Drucker. In:
Kravtchenko, J., Sirieys, P.M. (Eds.), Rheology and Soil Mechanics. Springer
Verlag, pp. 5868.
Matsuoka, H., Nakai, T., 1974. Stressdeformation and strength characteristics of soil
under three different principal stresses. Proceedings of Japan Society of Civil
Engineers 232, 5970.
McGarr, A., Gay, N.C., 1978. State of stress in the earths crust. Annual Review of
Earth and Planetary Sciences 6, 405436.
Mogi, K., 1967. Effect of the intermediate principal stress on rock failure. Journal of
Geophysical Research 72, 51175131.
Mogi, K., 1971. Fracture and owof rocks under high triaxial compression. Journal of
Geophysical Research 76, 12551269.
Murrell, S.A.F., 1963. A criterion for brittle fracture of rocks and concrete under
triaxial stress, and the effect of pore pressure on the criterion. In: Fairhurst, C.
(Ed.), Proceedings of the Fifth Symposium on Rock Mechanics. Pergamon Press,
pp. 563577.
Nadai, A., 1950. Theory of Flowand Fracture of Solids, vol. 1. McGraw-Hill, New York.
Oku, H., Haimson, B., Song, S.R., 2007. True triaxial strength and deformability of the
siltstone overlying the Chelungpu fault (Chi-Chi earthquake), Taiwan.
Geophysical Research Letters 34, L09306. doi:10.1029/2007GL029601.
Ottosen, N.S., Runesson, K., 1991. Properties of discontinuous bifurcation solutions
in elasto-plasticity. International Journal of Solids and Structures 27, 401421.
Perrin, G., Leblond, J.B., 1993. Rudnicki and Rices analysis of strain localization
revisited. Journal of Applied Mechanics 60, 842846.
Rice, J.R., 1976. The localization of plastic deformation. In: Koiter, W.T. (Ed.), Theo-
retical and Applied Mechanics, Proceedings of the 14th International Congress
on Theoretical and Applied Mechanics. North-Holland Publishing Company,
Delft, The Netherlands, pp. 207220.
Rudnicki, J.W., 2004. Shear and compaction band formation on an elliptic yield cap.
Journal of Geophysical Research 109. doi:10.1029/2003JB002633.
Rudnicki, J.W., Olsson W.A., 1998. Reexamination of fault angles predicted by shear
localization theory. In: Proceedings of Third North American Rock Mechanics
Symposium (NARMS98), Rock Mechanics in Mining, Petroleum and Civil
Works, 35 June, 1998, Cancun, Mexico. Extended Abstract in International
Journal of Rock Mechanics and Mining Sciences 35(415), 512513.
Rudnicki, J.W., Rice, J.R., 1975. Conditions for the localization of deformation in
pressure-sensitive dilatant materials. Journal of the Mechanics and Physics of
Solids 23, 371394.
Rudnicki, J.W., 2008a. Localized failure in brittle rock. In: Shao, J.F., Burlion, N. (Eds.),
Thermo-Hydromechanical and Chemical Coupling in Geomaterials and Appli-
cations, Proceedings of Third International Symposium GeoProc2008. Wiley,
pp. 2540.
Rudnicki, J.W., 2008b. Failure of Brittle Rock in the Laboratory and in the Earth, To
Appear in Proceedings of XXII International Congress on Theoretical and
Applied Mechanics, Adelaide, Australia, 2430 August.
Shin, T.-C., Teng, T.-L., 2001. An overview of the 1999 Chi-Chi, Taiwan, earthquake.
Bulletin of the Seismological Society of America 91, 895913.
Thomas, T.Y., 1961. Plastic Flow and Fracture in Solids. Academic Press.
Von Ka rma n, T., 1911. Festigkeitsversuche Unter all Seitigem Druck. Z. Verin Deut.,
Ingr. 55, 17491759.
Wiebols, G.A., Cook, N.G.W., 1968. An energy criterion for the strength of rock in
polyaxial compression. International Journal of Rock Mechanics and Mining
Sciences 5, 529549.
B. Haimson, J.W. Rudnicki / Journal of Structural Geology 32 (2010) 17011711 1711
Porosity and particle shape changes leading to shear localization
in small-displacement faults
Jafar Hadizadeh
a,
*
, Reza Sehhati
b, c
, Terry Tullis
d
a
Department of Geography and Geosciences, University of Louisville, Louisville, KY, USA
b
Department of Civil and Environmental Engineering, Washington State University, Pullman, WA, USA
c
Berger/ABAM Engineering Inc., Federal Way, WA 98003, USA
d
Department of Geological Sciences, Brown University, Providence, RI, USA
a r t i c l e i n f o
Article history:
Received 26 January 2009
Received in revised form
17 August 2010
Accepted 23 September 2010
Available online 1 October 2010
Keywords:
Shear localization
Gouge porosity
Particle shape and size
Particle size distribution
a b s t r a c t
A microstructural study of shear localization in fault gouge was carried out in small-displacement faults
so there would be minimum masking effects from a complex deformation history. We studied particle
size, shape, and porosity changes in gouge adjacent to zones of shear localization in natural and synthetic
gouges subjected to shear displacements d, of up to 1.2 m. Scanning electron microscope images were
used for estimating image porosity F
I
, and measuring particle size of the deformed and undeformed
gouges. The particle size data were used for calculating simulated porosity F
S
from computer-generated
simple fractal gouge model of each sample. Modeled microstructures contained round grains and
a fractal distribution matched to that of the measured natural samples. Changes in F
I
, F
S
, and F
I
/F
S
with
increasing d were used for tracking changes in particle shape and porosity of the gouges precursory to
shear localization. The F
I
and F
S
values for the natural and synthetic gouges converge at d w 0.1 m,
suggesting that gouge particles adjacent to shear localization sites tend to become rounded. Porosity for
such densied regions of the gouge adjacent to Y-shear zones was determined to be <1% at large
displacements. In the same regions, the porosity reductions were also associated with decreased sorting
coefcient and fractal dimensions D > 2.6. The study suggests that brittle shear localization may involve
favorably-oriented micro porous pockets of gouge that result from competing changes in particle shape
and particle size, which tend to affect gouge porosity in different ways.
2010 Elsevier Ltd. All rights reserved.
1. Introduction
Brittle shear localization microstructures such as slip surfaces,
shear bands, and cataclastic foliation are commonly found within
the core of many natural fault zones. The localization process
results in mechanical weakening of the fault and is often associated
with the most comminuted and densied regions in fault gouge
(e.g. Evans and Chester, 1995; Chester and Chester, 1998; Boullier
et al., 2004; Hayman, 2006; Rawling and Goodwin, 2006;
Rockwell and Ben-Zion, 2007; Tanaka et al., 2007; Brogi, 2008).
Shear localization has been observed in relatively unaltered small-
displacement faults with displacements typically <1 m as in most
experimental faults as well as in mature natural fault zones
involving a variety of alteration products. Experimental fault gouge
studies have shown that 0.1 m of shear displacement results in
shear localization, although microstructural changes leading to
shear localization are not well understood. A number of studies of
natural, experimental and computer simulated gouge deformation
conclude that shear localization is primarily a particle size and
particle-size distribution driven process (Dieterich, 1981; Marone
and Scholz, 1989; Biegel et al., 1989; Logan et al., 1992; Gu and
Wong, 1994; Billi, 2007; Keulen et al., 2007; Sammis and Ben-
Zion, 2008). Experimental studies by Mandl et al. (1977),
Vardoulakis (1980), Marone and Scholz (1989), Mair and Marone
(1999), and Mair et al. (2002) indicate that gouge attains a critical
strain or particle size distribution prior to shear localization. Mandl
et al. (1977) based on experimental data, suggested that shear
localizes in favorably-oriented bands of gouge that have achieved
a critical PSD through conned comminution. This is possible
because an increase in the proportion of ne particles in gouge,
while not reducing cohesive forces, might reduce the friction.
Subsequently, at a lowthreshold value of internal friction there will
be a drastic reduction in boundary shear by development of a slip
plane and growth of a shear zone. In the powder industry this
* Corresponding author.
E-mail address: hadizadeh@louisville.edu (J. Hadizadeh).
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Journal of Structural Geology 32 (2010) 1712e1720
phenomenon is attributed to rounding and size equalizing of
particles which tend to reduce interlocking resistance of a granular
material (Lowrison, 1974). The microstructural aspects of the model
described above draws support from a number of observations.
Gouge deformation simulations by Morgan and Boettcher (1999)
showed that a sharp drop in sliding contacts accompanied local-
ized failure as fewer particles were involved in the zone of defor-
mation. Mair and Marone (1999) studied the controlling effect of
particle size on shear localization and noted that PSD evolution in
ne and coarse gouge differed only by a shear strain g of 3.4. They
suggested that higher fracture toughness might have inhibited
further comminution of ne gouge as its PSD became more
uniform. Scarpelli and Wood (1982), Moore et al. (1989), Logan
et al. (1992) reported the same sequence of microstructural
development preceding shear localization in calcite and halite
gouges. In a model presented by Shipton and Cowie (2001) a critical
amount of comminution or strain was necessary for slip surface
nucleation, but the slip surfaces accommodated further strain
without appreciable amount of comminution. The experimental
study of Marone and Scholz (1989) concluded that transition from
pervasive to localized shear occurs at a critical strain or PSD in
addition to the inuence of gouge density.
This study attempts to provide a microstructural model for shear
localization by investigating the combined effect of particle shape
and particle size on porosity changes that precede shear localiza-
tion in small-displacement natural and experimental faults. The
effect of porosity changes on shear localization in granular material
with varied particle size and PSD has been studied previously. The
work of Mead (1925) and Frank (1965) shows that since in conned
comminution dilatancy is suppressed, shear localization must
occur in regions of gouge with least dilatancy rate. Marone and
Scholz (1989) reported dilatant behavior at the onset of shear
localization in their experimental quartz gouge. The actual shear
localization occurred on R
1
Riedel shear bands, and the dilatancy
was believed to be the result of unpacking of over consolidated
gouge. Marone and Scholz (1989) noted that particles within their
Fig. 1. Natural and synthetic gouges used in the study. (a) Typical small-displacement splay fault in the Aztec sandstone. Outcrop picture is labeled with the apparent displacement
vector (d
a
), and trend of the slickenside lineation on the fault plane (dotted line). Compaction bands serve as displacement markers. Typical undeformed texture of the sandstone is
shown on the right. (b) Rotary shear sample ring consisting of a 2 mm gouge layer (not to scale). The close-up view of the gouge layer on the left is an actual section across
undeformed simulated Westerly granite gouge compacted to 25 MPa pressure. The gouge is held between granite forcing blocks g, forming the shear zone boundaries; Q quartz;
K potassium feldspar; P plagioclase feldspar; lightest shade particles are phyllosilicates. Textures are back-scattered SEM images.
J. Hadizadeh et al. / Journal of Structural Geology 32 (2010) 1712e1720 1713
experimental shear bands did not follow a fractal size distribution,
but the relationship between the shear band microstructures and
the unpacking process was not discussed.
2. The studied fault gouges and methods of data collection
2.1. The natural gouge
Gouge samples were collected from splay faults of the Lonewolf
fault zone in the Valley of Fire State Park (VOF), Nevada, U.S.A. The
splay faults are believed to be sheared joints, the coalescence of
which lead to the development of the Lonewolf fault zone (Flodin
and Aydin, 2004; Myers and Aydin, 2004; Davatzes et al., 2003).
The splay faults ranging in measurable displacement fromw0.01 m
to 1.5 m were found in the upper domain of the Aztec sandstone
Formation described in detail by Flodin et al. (2003). Due to soft
iron oxide/clay mineral cementation of the sandstone and weath-
ering effects, in-situ impregnation by clear resin was necessary
before the samples could be lifted from the desired locations along
the splay faults. The sampling method thus preserved the gouge
microstructures. The Aztec sandstone in the sampling area is an
aeolian feldspathic quartz arenite made of rounded to well-
rounded quartz grains forming 95% of the rock with average
undeformed grain size of 230 mm. The measured porosity (Flodin
et al., 2003 via Helium porosimetry) ranged from 21.2% to 23.8%
with the presence of a small amount of pore-lling quartz or iron
oxide cement. Our sampling area corresponds to porosity samples
54e56 in Table 2, Flodin et al. (2003).
The strike-slip and normal faulting in the VOF area had taken
place during the Miocene Basin and Range tectonic activity. It has
been suggested that in the initial phase of the activity the Aztec
Formation had been buried by at least 1.6 km of sediments
(Bohannon, 1983), andpossibly byanadditional 1e4kmof overlying
Sevier-related thrust sheet (Brock and Engelder, 1977). Based on the
overburden thickness estimates and studies by Flodin and Aydin
(2004) the possible range of overburden pressure at the time of
faulting could be 10e40 MPa assuming an overburden density of
2700 kg/m
3
. The effective pressure for the VOF gouges might have
been closer to the middle of the range because for approximately
same shear strains the overall degree of cataclasis in the VOF shear
bands was similar to those observed in our experimental shear
bands deformed at 25 MPa. We found no optical evidence of crystal
plastic deformation in quartz grains of the sampled gouge, which
suggested temperatures <200
C for the observed deformation.
Compaction bands that formed prior to faulting in the Aztec
sandstone (Flodin et al., 2003) were offset by the studied splay faults
and provided excellent displacement markers (Fig. 1a). The measured
apparent displacement d
a
(the band offset), together with the pitch R
of slickenside lineation on the fault plane was used to determine true
displacement d, as d
a
/cos R. This relationship holds for the studied
faults with a strike-slip displacement component since the displace-
ment markers (compaction bands) were oriented vertical, or near
vertical (see Fig. 1a). The apparent displacement varied along the
strike, but the faults were sampled exactly where the structural
measurements were made. The true displacements thus calculated
were17mm, 116mm, 348mmand1219mmfor thefour samples used
inthis study. The gouge zone thicknesses measured inthinsections at
right angles to the shear zone borders ranged from1.2 0.08 mm to
15.5 5 mm, positively correlating with the displacement values. A
summary of the sample data is presented in Table 1.
2.2. The synthetic gouge
Synthetic gouge samples (WGK) with particle size 88 mm
consisting of 28% quartz, 35% microcline, 32% plagioclase, 5% mica,
and <1% opaque were prepared by grinding Westerly granite. The
study uses deformed synthetic gouge froma previous experimental
study. The gouge had been deformed at room temperature and
25 MPa normal stress in a rotary shear apparatus. The sliding
velocity was stepped between 1 mm/s (for 1 mm distance), and
10 mm/s (for 10 mmdistance) in all experiments. Technical specics
of the apparatus are described elsewhere (Tullis and Weeks, 1986;
Beeler et al., 1996). For each experiment approximately 1 g of the
material was packed in a ring-shaped sample holder forming
a w2 mm layer of gouge with w35% initial porosity (Fig. 1b). A
compaction test showed that raising normal stress to 25 MPa at the
beginning of each experiment reduces the gouge layer thickness by
about 5.5% (w110 mm). The compaction run also provided refer-
ences for the initial texture, PSD, and porosity. The gouge layers
were deformed to 44 mm, 79 mm, and 387 mm of shear
displacement. While the simulated gouge was almost entirely
velocity weakening up to the largest displacements, the effect of
sliding rate on shear localization was not tested.
2.3. Particle size measurements
The thin section areas selected for imaging, and the subsequent
particle size and porosity measurements, were thoroughly inspec-
ted for surface damage and plucked grains. The particle size
measurements other than for characterizing the undeformed
materials were made on areas of gouge adjacent to zones of shear
localization. It was assumed that the intensely deformed gouge
within 500 mm distance either side of Y-shear zones reect the
particle shape, particle size distribution, and porosity that existed at
the onset of shear localization regardless of shear strain in the
samples. It has been shown that shear localization as a result of
comminutionoccurs during the early increments (g
0
s 1e10) of shear
strain (Marone and Scholz, 1989; Gu and Wong, 1994; Mair and
Marone, 1999; Wolf et al., 2003). The minimum shear strain in
VOF and WGK gouges were 14 and 22 respectively. Furthermore, we
assume that as the gouge approaches the shear localization stage
the regions around potential shear zone deform more intensely
thaninthe bulk gouge. This maybe so because shear localizes where
gouge is weakening. Thus unlike the bulk gouge, microstructures
near the localized shear zone are expected to record the process of
weakening by developing a somewhat different set of textural
attributes. We avoided selecting areas for microstructural
measurements where relatively undeformed gouge (e.g. coarse
gouge near the fault zone margins in Fig. 2a and b) was in contact
with shear bands. Such contacts might indicate the possibility of
shear localization due to some preexisting microstructural
inhomogeneity. The PSD data sets were used in computer simula-
tions, and for determining fractal dimension of the gouges. The
measurements were conducted on polished petrographic thin
sections cut perpendicular to the shear planes and viewed in back-
scattered mode in a Zeiss Supra-35VP scanning electron micro-
scope (SEM). Optical microscopy included transmitted and reected
Table 1
A summary of structural data for the studied splay faults of the Lonewolf fault zone
in the Aztec sandstone.
Sample Fault plane attitude and
sense of shear
d
a
, mm R
d, mm T, mm
VOF4A 284, 39NE; RL 16.5 17 17 1.2
VOF2A 340, 74SW; Rl 70 53 116 1.4
VOF01 342, 72NE; LL 90 75 348 3.6
VoF5A 015, 76NW; LL 1275 17 1219 15.5
d
a
apparent displacement; R
.
Measurements were carried out using SigmaScan Pro and MAT-
LAB Image Processing Toolbox applications. To gain a wider
particle size range we pooled size data for each deformed gouge
sample from 2 (VOF samples) or 3 (WGK samples) sets of images
taken telescopically at increasing magnication. Asingle SEMimage
each of the undeformed VOF sandstone and WGK simulated gouge
was used for reference particle size measurements. In telescopic
imaging, the microscope magnicationwas varied by a factor of two
from 250 to 32000 depending on the particle size. Duplicate
particle sizes (same values to 3 decimal places) in two consecutive
images in each set were discarded. As mentioned earlier, particle
size measurement of the deformed gouge was based on selected
areas adjacent to Y-shear zones (Fig. 2). Particle size data for this
study consisted of 3469 measurements (particle outline traces)
acquired from 9 SEM images for the VOF samples and 5868
measurements acquired from10 SEMimages for the WGK samples.
The difference in number of measurements (including the unde-
formed materials) reects the smaller average particle size of the
WGK gouge that resulted in a larger number of particles per unit
area compared to that for the VOF gouge.
The size-number data was used for determining the fractal
dimension D, as the slope of the logelog size (S) e number (N)
distribution given by NS cS
D
, where c is a constant (Turcotte,
1986). We present D values here as the 3D, or the volume fractal
dimension by adding 1.0 to the calculated D values (Falconer, 1985).
The changes in PSD with increased shear displacement were
analyzed using the sorting coefcient Q [S
q1
/S
q3
]
, where S
q1
and
S
q3
are the rst quartile (25% of distribution >S
q1
) and the third
quartile (75% of the distribution >S
q3
) in a given cumulative size
Fig. 2. Typical areas of gouge used for particle size and porosity measurements. The dotted line labeled Y is trace of the Y-shear determined from a larger area of the sample than
shown here. Whole shear zones are shown on top with insets showing location of enlarged areas at the bottom. (a) Shear zone in the Aztec sandstone sample with 348 mm of
displacement (VOF01 in Table 1). (b) Synthetic Westerly granite gouge layer after 44 mm of displacement (sample WGK258).
J. Hadizadeh et al. / Journal of Structural Geology 32 (2010) 1712e1720 1715
distribution respectively (Krumbein and Sloss, 1951). Decreasing Q
values indicate reduced particle size range, improved sorting, and
increased porosity while increasing Q values indicate widening of
the size range, poor sorting, and decreased porosity (Rogers and
Head, 1961; Marone and Scholz, 1989). The gouge porosity here
referred to as 2D image porosity F
I
, is the total porosity estimated
via digital analysis of the SEM images. Pore spaces as conrmed by
secondary electron images appeared as dark featureless areas
between particles on backscattered SEM images (see Fig. 1). Thus
F
I
[SA
(pore space)
100]/A
(image)
, where A is the measured areas on
the image. The image processing method and its validation with
respect to volume porosity in actual rock material has been dis-
cussed elsewhere (e.g. Antonellini et al., 1994; Anselmatti et al.,
1998; Solymar and Fabricius, 1999; Talukdar et al., 2002;
Johansen et al., 2005). The sources of error in area measurements
included image quality and magnication, and the thresholding
process. Thresholding refers to a digital manipulation of the image
that involves differentiating a certain range of pixel intensities from
the full image pixel intensity distribution. For our purpose, the pore
space pixel range was isolated by thresholding binarized images
using SigmaScan Pro application. The thresholding average
values were 0.1% and 2% of the total measured pore areas at the
highest (64000) and the lowest (250) image magnications
respectively. After thresholding, but prior to the measurements,
pore-like pixel areas within particles were removed from the
threshold copy of the image. Since the particle-pore borders con-
sisted of only a zone of 2e5 pixels, an average of the highest and
lowest pore space area was used in individual images. The image
porosity was then estimated by averaging F
I
values fromindividual
images in a telescopic series that represented each sample. A 3-
image telescopic set was used for the highest displacement samples
(WGK262 and VOF5A) while a 2-image telescopic set, taken at
250and 1K, was used in all other samples; 9 (VOF) and 7 (WGK)
SEM images were used.
2.4. Porosity from computer-generated models
A computer program for generating simple fractal gouge (SFG)
models was written in VC. The basic objective of the simulations
was to obtain reasonable estimates for the limiting values of
porosity as gouge particles become more rounded with increased
comminution. The program was not written to simulate realistic
gouge microstructures; it was intended for providing estimates of
porosity in a hypothetical gouge with spherical particles and fractal
particle size distribution. The F
S
values, generated based on our real
gouge PSD data, were an approximation which we considered more
accurate than a hypothetical PSD. The nature of the PSD approxi-
mation was that we expected a deviation to occur from the so-
called steady-state fractal size distribution (D 2.6) within the
shear bands. The approximation is reasonable since the PSD data
used for the simulations were collected from selected areas adja-
cent to the shear bands.
The algorithmtakes as input a PSD data set acquired fromgouge
images as described earlier. The size and total number of particles
remained unchanged from the original distribution, and the
particles were simulated as perfect circles. As a fractal distribution,
the particles for each simulated texture were assembled such that
least number of same size particles touched. This was achieved by
packing particles according to tangent circle solutions for neigh-
boring particles. The program produced maximum packing density
for the given data set under the described conditions. The packing
density of the SFG models could exceed the maximum packing
density achievable with uniform size circles in 2D space given by
p/O12 (Hecht, 2004). The algorithm was capable of processing
an unlimited number of particles with unlimited size range in
a descending order. The 2D porosity of the SFG was calculated by
summing up the void space areas within the perimeter of the entire
simulated mass. Although in this study only 2D results are reported
we note that the 3D porosity tends to be higher than the 2D
Fig. 3. Examples of the simple fractal gouge (SFG) model based on particle size data from gouge areas adjacent to Y-shear zones. Models are for (a) Aztec sandstone gouge area
shown in Fig. 2a, N 817 particles, D 3.306, and (b) Synthetic gouge area shown in Fig. 2b, N 1386 particles, D 3.284. Scale bars on images are approximately true for the
models.
J. Hadizadeh et al. / Journal of Structural Geology 32 (2010) 1712e1720 1716
porosity for the same packing arrangement. The main source of
error in the simulation originated in the manual particle outline
tracing operations that provided the particle size data set. This was
in turn dependent upon magnication and image quality. The
programnumeric output included SFGporosity, F
S
, and a simulated
texture generated via a visualizer program (examples shown in
Fig. 3). We note that both F
S
and F
I
should be considered reason-
able proxies for the gouge porosity rather than actual porosity.
3. Results
The results are shown in plots of the Figs. 4e6. The measured
porosity of 21e23% in the undeformed Aztec sandstone is compa-
rable to its SFG model porosity of 28.55%. The difference is in part
due to non-fractal PSD of the undeformed sandstone and the
presence of pore-lling cement. In the deformed sandstone gouge,
F
I
and F
S
values rst diverge at d > 20 mm and then converge at
d > 300 mm (Fig. 4a). This transient deviation from continuous
porosity reduction is due to comminution of the round undeformed
particles into angular particles during initial increments of shear
displacement on the fault. The porosity of the sandstone gouge
drops with further comminution and rounding of the cataclastic
particles. For the synthetic Westerly granite gouge with initially
angular particles, F
S
and F
I
values have their largest difference in
the undeformed gouge, but the difference is reduced monotonously
with increasing displacement (Fig. 4b). In terms of particle shape,
Fig. 4c shows that conned comminution in the synthetic gouge
results in rounding of the particles as well as size reduction,
although an aggregate with initially round particles will do so by
rst transforming to an angular aggregate. Although in both gouge
types the porosity is signicantly reduced by comminution, and F
I
and F
S
values tend to converge at d > 0.1 m, the porosity reduction
in the synthetic gouge occurs at a higher rate with respect to
displacement presumably due to higher normal stresses. For
example, at dw360 mm F
I
is w0.45% and w1.7% for the synthetic
gouge and the sandstone gouge respectively. We note that as
expected, dF
S
/dd is nearly identical for both gouge types.
The evolution of porosity with shear displacement in the two
gouge types is compared by presenting the changes as F
I
/F
S
ratio in
C
b
a
Fig. 4. Changes in porosity with shear displacement of areas adjacent to Y-shear zones
in (a) The Aztec sandstone gouge and (b) Westerly granite synthetic gouge. The
porosity is represented by the 2D image porosity F
I
, and its corresponding SFG model
porosity F
S
. Undeformed gouge in both cases is assigned a nominal displacement value
of 1 mm (c). SEM backscattered image of simulated Westerly granite gouge texture
showing abundance of particles with round to sub rounded shapes bordering the main
Y-shear zone in the experiment with 387 mm shear displacement.
Fig. 5. Plot comparing change in F
I
/F
S
ratio with shear displacement in Aztec sand-
stone and Westerly granite synthetic gouges. The undeformed gouge is assigned
a nominal displacement value of 1 mm.
J. Hadizadeh et al. / Journal of Structural Geology 32 (2010) 1712e1720 1717
Fig. 5. The F
I
/F
S
for the two deformed gouge types converge with
increasing shear displacement while remaining >1, which indicates
a higher porosity for the gouges compared to their SFG models. The
plots in Fig. 5 conrm that lower porosities adjacent to zones of
shear localization result from rounding as well as ning of the
particles. The particle shape changes are associated with increasing
D values and decreasing sorting coefcient Q (decreasing size
range). The interrelationship between sorting, porosity, and D,
based on our data, is shown in Fig. 6. The plots indicate that
densication of gouge at potential shear localization sites is
concurrent with particle rounding and improved sorting.
4. Discussion
The studied gouges had not undergone appreciable hydro-
thermal alterations or pressure solution and lack signicant phyl-
losilicates fractions, all of which are shown to be important factors
in deformation of gouge in mature fault zones. As suggested by
studies of clayequartz gouge mixtures (e.g. Crawford et al., 2008),
the microstructural and mechanical effects of phyllosilicates in
quartz-feldspatic gouge are minimal if phyllosilicates make 5% of
the gouge contents. Experimental studies also show that the
process of shear localization is affected by the magnitude of the
mean stress mostly in the dilatational phase (Marone and Scholz,
1989; Gu and Wong, 1994; Besuelle, 2001). This is mainly because
the strain per fracture rule applicable to a fractal PSD breaks down
within the shear bands (Sammis et al., 1987) and there is less
number of contact points in a non-fractal aggregate. The shear
localization model discussed here is relevant to microstructures
within the shear bands in the post dilatational stage of the shear
localization.
The results provide both visual and numeric conrmation of the
correlation between shape of particles in the studied gouges and
the porosity values calculated from their corresponding SFG
models. The undeformed Aztec sandstone with round particles has
F
I
/F
S
of w1, while the undeformed synthetic Westerly granite
gouge with highly angular particles has F
I
/F
S
of w9. The changes in
F
I
and F
S
values and F
I
/F
S
ratio with increasing shear displacement
showthat using the model results to infer changes in particle shape
of the deformed gouges is also reasonable. In the regions adjacent
to Y-shear zones, porosities calculated from SFG models change
from 2% to 0.2% with increasing displacement as the corresponding
gouge porosity is reduced from 5% to 0.3%. Furthermore, we
showed that despite mineralogical differences of the two gouge
types the microstructural attributes of the gouge adjacent to shear
bands converge with increased comminution.
The particle size distribution adjacent to Y-shear zones in both
gouge types consistently yields D values greater than what is
considered to be an ideal gouge PSD of 2.6 produced through
conned comminution (Sammis et al., 1987; Sammis and Biegel,
1989; Blenkinsop, 1991). The gouge regions with D > 2.6, having
already achieved a high packing density, might be viewed as having
a critical PSDin connectionwith the shear localization processes. To
further dene the condition we consider porosity (packing density)
and sorting characteristics of the gouge from the studied regions.
The sorting ratio of the gouge was shown to drop continuously with
increasing displacement as the average ratio of the largest to
smallest size particles within the zones was reduced from 3 to 1.3
prior to shear localization. However, to interpret the decreasing Q
as an increasing porosity appears to contradict the decreasing F
I
and F
s
values we report in the same gouge over the same range of
shear displacements. A possible explanation is that porosity
reductions through particle rounding are offset by the porosity
gains through decreasing Q. The SFG models clearly showthat prior
to shear localization the rounding effect continues with increasing
shear displacement. The local dilation rate, therefore, may depend
on net porosity gain or loss due to the competing effects of the
changes in size and shape of particles within a densifying gouge.
A shear localization model based on the presented data and
arguments above appears to be in general agreement with the
localizationmodel discussedbyMarone andScholz (1989). For shear
localization in gouge with the critical PSD, a sufcient density of
regions with lower dilation rate dF/dg, is required. The shear
localization model illustrated in Fig. 7 thus involves unpacking (loss
of cohesion at micro-scales) of densied gouge along a zone of
favorably oriented pockets of relatively porous (microporous)
gouge. Based on our analyses of the critical PSD above, the net
porosity gain within a shear band region would occur only if the
comminution process (including particle boundary attrition and
transgranular fracture) becomes more effective in eliminating
particle size differences than rounding the particles. The micro-
mechanics of the competing effect could not be ascertained from
this study, but the data indicates that such microstructural state is
reached at fractal dimensions D > 2.6 as depicted in the schematic
Fig. 6. Change in image porosity F
I
and sorting ratio Q of gouge with increasing fractal
dimension. Vertical error bars represent range of porosity measurements about the
average. (a) Aztec sandstone gouge (b) Synthetic gouge. Higher Q values for sandstone
reect its higher average grain size compared to synthetic gouge. Data points for
undeformed material are shown for reference on the left side in each case.
J. Hadizadeh et al. / Journal of Structural Geology 32 (2010) 1712e1720 1718
plot of Fig. 7a. It is likely, however, that shear localization in the
densied gouge involves a range of D values at D > 2.6 rather than
requiring a unique D value for the entire length of a potential shear
band. This is simply because of the non-uniform nature of commi-
nutionintensity withina deforming gouge (Hadizadehand Johnson,
2003). The model gouge microstructure at the onset of shear local-
ization includes pockets or streaks of porous gouge, some of which
are favorablyorientedparallel to Y-shear orientationas illustratedin
Fig. 7b. Based on a sorting criterion Marone and Scholz (1989)
showed that in the bulk gouge with fractal size distribution sorting
remained poor and porosity remained low since small particles
tended to ll in between the larger particles. Within the bulk gouge,
regions with non-fractal PSD (e.g. within shear bands) tend to have
reduced particle size range, improved sorting and a higher porosity.
It is assumed that regions with lower rates of porosity reduction
(represented by dashed segment of porosity curve in Fig. 7a) are
initially highly localized, and that these regions need not to spread
throughout a signicant thickness of the gougebeforetheunpacking
occurs. This assumption is consistent with the common observation
that thickness of slipsurfaces andshear bands oftenconstituteavery
small fractionof the total gouge zone thickness. At particle-scale, the
model is supported by previous work that shows comminution in
shear localization sites tends to eliminate larger particles (Marone
and Scholz, 1989; Blenkinsop, 1991; Mair and Marone, 1999) prob-
ably because the growing number of smaller particles of all shapes
require larger stresses to fracture (Kendall, 1978). A uniformly
distributed porosity in the densied gouge would assist the shear
localization process by providing weak links between the porous
pockets.
The shear strength of an incipient unpacking surface is depen-
dent upon surface roughness (Biegel et al., 1992), which in this case
is mainly determined by the maximum size and spacing of the
pores in the dense gouge. Shearing of the sub-micron asperities
along unpacking surfaces may explain presence of thin bands of
extremely ne particles observed within shear bands and along
slickenside surfaces (e.g. Yund et al., 1990; Power and Tullis, 1989).
5. Conclusions
1. Conned comminution generates similar particle shape and
size distributions with increasing shear displacement in the
two studied gouge types with different mineral composition
and initial textures.
2. Adjacent to sites of shear localization the gouge particles are
rounder and particle size range is signicantly narrowed with
fractal dimensions D >2.6. This observation suggests that shear
localization must involve unpacking of a densied gouge.
3. The microstructural data and simple fractal gouge models
indicate that unpacking of the gouges might be the result of
highly localized porosity variations within the densied gouge,
caused by the competing effects of changes in particle shape
and particle size.
Acknowledgements
We wish to thank the editors of the JSG for their valuable
comments. Lori Kennedy and the anonymous reviewers of this
manuscript are thanked for their constructive comments. Discus-
sions with Judy Chester and Joseph C. White resulted in signicant
improvements in the manuscript. We wish to thank David Goldsby
and Anoaur Koncachbaev for their assistance with the rotary shear
experiments at Brown Geosciences Department, and Joseph Wil-
liams for his help with electron microscopy at the University of
Louisville. This research was partially supported by the US National
Science Foundation grant NSF-EAR-0229654 to Jafar Hadizadeh.
References
Antonellini, M.A., Aydin, A., et al., 1994. Microstructure of deformation bands in
porous sandstones at Arches National Park, Utah. J. Struct. Geol. 16, 941e959.
Anselmatti, F.S., Luthi, S., Eberli, G.P., 1998. Quantitative characterization of
carbonate pore systems by digital image analysis. AAPG Bull. 82, 1815e1836.
Beeler, N.M., Tullis, T.E., Weeks, J.D., 1996. Frictional behavior of large displacement
experimental faults. J. Geophys. Res. 101, 8697e8715.
Besuelle, P., 2001. Evolution of strain localization with stress in sandstone: brittle
and semi-brittle regimes. Phys. Chem. Earth (A) 26, 101e106.
Fig. 7. A model for shear localization based on competing changes in particle shape
and particle size of the gouge. (a) Behavior of microstructural variables porosity F, and
sorting coefcient Q with increased fractal dimension of gouge shown in an ideal
extrapolation of data. Shear localizes in regions of gouge where porosity increase by
narrowing particle size range (decreasing Q) is greater than porosity reduction by
particle rounding (decreasing F). Porosity is expected to drop sharply within shear
localization microstructures such as shear bands. It is assumed that changes in Q and F
take place at D > 2.6. (b) Highly schematic representation of gouge microstructure at
the onset of shear localization in gouge (dashed line segments in a) in an approxi-
mately 20 15 mm area. Pockets of well-sorted porous gouge (angular particles) are
bordered by dense gouge (round particles) with well distributed micro porosity. The
PSD in this region of gouge has D > 2.6. The particle shapes are simplied and exag-
gerated, and sense of shear is arbitrary.
J. Hadizadeh et al. / Journal of Structural Geology 32 (2010) 1712e1720 1719
Biegel, R.L., Wang, W., Scholz, C.H., Boitnott, G.N., Yoshioka, N., 1992. Micro-
mechanics of rock friction, 1. Effects of surface roughness on initial friction and
slip hardening in Westerly granite. J. Geophys. Res. 97, 8951e8964.
Biegel, R.L., Sammis, C.G., et al., 1989. The frictional properties of a simulated gouge
having fractal particle distribution. J. Struct. Geol. 11, 827e846.
Billi, A., 2007. On the extent of size range and power law scaling for particles of
natural carbonate fault cores. J. Struct. Geol. 29, 1512e1521.
Blenkinsop, T.G., 1991. Cataclasis and the processes of particle size reduction.
PAGEOPH 136, 59e86.
Bohannon, R.G., 1983. Mesozoic and Cenozoic tectonic development of the muddy,
north muddy, and northern black mountains, Clark County, Nevada. Geol. Soc.
Am. Mem. 157, 125e148.
Boullier, A.-M., Fujimotob, K., Ohtanib, T., Roman-Rossa, G., Lewinc, E., Itob, H.,
Pezardd, P., Ildefonsed, B., 2004. Textural evidence for recent co-seismic
circulation of uids in the Nojima fault zone, Awaji Island, Japan. Tectonophys.
378, 165e181.
Brock, W.G., Engelder, T., 1977. Deformation associated with the movement of the
muddy Mountain overthrust in the Bufngton Windows, S. Nevada. Geol. Soc.
Am. Bull. 88, 1667e1677.
Brogi, A., 2008. Fault zone architecture and permeability features in siliceous
sedimentary rocks: insights from the Rapolano geothermal area (Northern
Apennines, Italy). J. Struct. Geol. 30, 237e256.
Chester, F.M., Chester, J.S., 1998. Ultracataclasite structure and friction processes of
the Punchbowl fault, San Andreas system, California. Tectonophys. 295,
199e221.
Crawford, B.R., Faulkner, D.R., Rutter, E.H., 2008. Strength, porosity, and perme-
ability development during hydrostatic and shear loading of synthetic
quartzeclay fault gouge. J. Geophys. Res. 113 (B03207).
Davatzes, N.C., Aydin, A., Eichhubl, P., 2003. Overprinting faulting mechanisms
during the development of multiple fault sets in sandstone, Chimmney rock
fault array, Utah, USA. Tectonophys. 363 (1e2), 1e18.
Dieterich, J.H., 1981. Constitutive properties of faults with simulated gouge.
Mechanical behavior of crustal rocks, the Handin volume. AGU. Monog. 24,
103e120.
Evans, J.P., Chester, F.M., 1995. Fluid-rock interaction in faults of the San Andreas
system: inferences from San Gabriel fault rock geochemistry and microstruc-
tures. J. Geophys. Res. 100, 13007e13020.
Falconer, K.J., 1985. Fractal Geometry: Mathematical Formulations and Applications.
Cambridge University Press.
Flodin, E., Aydin, A., 2004. Evolution of a strike-slip fault network, Valley of Fire
state park, southern Nevada. Geol. Soc. Am. Bull. 116, 42e59.
Flodin, E., Prasad, M., Aydin, A., 2003. Petrophysical constraints on deformation styles
in Aztec sandstone, southern Nevada, USA. Pure Appl Geophys 160, 1589e1610.
Frank, F.C., 1965. On dilatancy in relation to seismic sources. Rev. Geophys. 3,
485e553.
Gu, Y., Wong, T.-F., 1994. Development of shear localization in simulated quartz
gouge: effect of cumulative slip and gouge particle size. Pure Appl Geophys 143,
387e423.
Hadizadeh, J., Johnson, W.K., 2003. Estimating local strain due to comminution in
experimental cataclastic textures. J. Struct. Geol. 25, 1973e1979.
Hayman, N.W., 2006. Shallow crustal fault rocks from the Black Mountain detach-
ments, Death Valley, CA. J. Struct. Geol. 28, 1767e1784.
Hecht, C.A., 2004. Geomechanical models for clastic grain packing. Pure Appl
Geophys 163, 331e349.
Johansen, T.E.S., Fossen, H., Kluge, R., 2005. The impact of syn-faulting porosity
reduction on damage zone architecture in porous sandstone: an outcrop
example from the Moab Fault, Utah. J. Struct. Geol. 27, 1469e1485.
Kendall, K., 1978. The impossibility of comminuting small particles by compression.
Nature 272, 710e711.
Keulen, N., Heilbronner, R., Stnitz, H., Boullier, A.-M., Ito, H., 2007. Grain size
distributions of fault rocks: a comparison between experimentally and natu-
rally deformed granitoids. J. Struct. Geol. 29, 1282e1300.
Krumbein, W.C., Sloss, L.L., 1951. Stratigraphy and Sedimentation. Freeman.
Logan, J.M., Dengo, C.A., Higgs, N.G., Wang, Z.Z., 1992. Fabrics of experimental fault
zones: their development and relationship to mechanical behavior. In: Fault
Mechanics and Transport Properties of Rocks. Academic Press Ltd., pp. 33e66.
Lowrison, G.C., 1974. Crushing and Grinding: The Size Reduction of Solid Materials.
Butterworth Publishers, London, 286 pp.
Mair, K., Marone, C., 1999. Friction of simulated fault gouge for a wide range of
velocities and normal stresses. J. Geophys. Res. 104, 28899e28914.
Mair, K., Frye, K.M., et al., 2002. Inuence of grain characteristics on the friction of
granular shear zones. J. Geophys. Res. 107 (ECV4), 1e9.
Mandl, G., De Jong, L.N.J., Maltha, A., 1977. Shear zones in granular materials. Rock
Mech. 9, 95e144.
Marone, C., Scholz, C.H., 1989. Particle-size distribution and microstructures within
simulated fault gouge. J. Struct. Geol. 11, 799e814.
Mead, W.J., 1925. The geologic role of dilatancy. J. Geol. 33, 685e698.
Moore, D.E., Summers, R., et al., 1989. Sliding behavior and deformation textures of
heated illite gouge. J. Struct. Geol. 11, 329e342.
Morgan, J.K., Boettcher, M.S., 1999. Numerical simulations of granular shear zones
using the distinct element method 1. Shear zone kinematics and micro-
mechanics of localization. J. Geophys. Res. 104, 2703e2719.
Myers, R., Aydin, A., 2004. The evolution of faults formed by shearing across joint
zones in sandstone. J. Struct. Geol. 26, 947e966.
Power, W.L., Tullis, T.E., 1989. The relationship between slickenside surfaces in ne-
grained quartz and the seismic cycle. J. Struct. Geol. 11, 879e894.
Rawling, G.C., Goodwin, L.B., 2006. Structural record of the mechanical evolution of
mixed zones in faulted poorly lithied sediments, Rio Grande rift, New Mexico,
USA. J. Struct. Geol. 28, 1623e1639.
Rockwell, T.K., Ben-Zion, Y., 2007. High localization of primary slip zones in large
earthquakes from paleoseismic trenches: observations and implications for
earthquake physics. J. Geophys. Res. 112 (B10304), 1e12.
Rogers, J.J.W., Head, W.B., 1961. Relationship between porosity, median size, and
sorting coefcients of synthetic sands. J. Sediment. Petrol. 31, 467e470.
Sammis, C.G., Ben-Zion, Y., 2008. Mechanics of grain-size reduction in fault zones.
J. Geophys. Res. 113.
Sammis, C.G., Biegel, R.L., 1989. Fractals, fault-gouge, and friction. Pure Appl Geo-
phys 131, 255e271.
Sammis, C.G., King, G., Biegel, R., 1987. The kinematics of gouge deformation. Pure
Appl Geophys 125, 777e812.
Scarpelli, G., Wood, D.M., 1982. Experimental observations of shear band patterns in
direct shear tests. In: Vermeer, P.A., Luger, H.J. (Eds.), Deformation and Failure of
Granular Materials. IUTAMDelft, Balkema, Rotterdam, Netherlands, pp. 473e484.
Shipton, Z.K., Cowie, P.A., 2001. Damage zone and slip-surface evolution over
micron to km scales in high-porosity Navajo sandstone, Utah. J. Struct. Geol. 23,
1825e1844.
Solymar, M., Fabricius, I.L., 1999. Image analysis and estimation of porosity and
permeability of Arnager Greensand, upper Cretaceous, Denmark. Phys. Chem.
Earth (A) 24, 587e591.
Talukdar, M.S., Torsaeter, O., Ioannnidis, M.A., Howard, J.I., 2002. Stochastic recon-
struction of chalk from 2D images. Transport Porous Media 48, 101e123.
Tanaka, H., Omura, K., Matsuda, T., Ikeda, R., Kobayashi, K., Murakami, M.,
Shimada, K., 2007. Architectural evolution of the Nojima fault and identication
of the activated slip layer by Kobe earthquake. J. Geophys. Res. 112 (B07304),
1e20.
Tullis, T.E., Weeks, J.D., 1986. Constitutive behavior and stability of frictional sliding
of granite. Pure Appl Geophys 124, 384e414.
Turcotte, D.L., 1986. Fractals and fragmentation. J. Geophys. Res. 91, 1921e1926.
Vardoulakis, I., 1980. Shear band inclination and shear modulus of sand in biaxial
tests. Int. J. Numer. Anal. Meth. Geomech. 4, 103e119.
Wolf, H., Konig, D., et al., 2003. Experimental investigation of shear band patterns in
granular material. J. Struct. Geol. 25, 1229e1240.
Yund, R.A., Blanpied, M.L., Tullis, T.E., Weeks, J.P., 1990. Amorphous material in high
strain experimental fault gouges. J. Geophys. Res. 95, 15589e15602.
J. Hadizadeh et al. / Journal of Structural Geology 32 (2010) 1712e1720 1720
Field evidences for the role of static friction on fracture orientation in extensional
relays along strike-slip faults: Comparison with photoelasticity and 3-D
numerical modeling
Roger Soliva
a,
*
, Frantz Maerten
a, b
, Jean-Pierre Petit
a
, Vincent Auzias
c
a
Universite Montpellier II, Lab. Geosciences Montpellier, UMR 5243, Place E. Bataillon, 34095 Montpellier cedex, France
b
IGEOSS, Parc Euromedecine, 340 rue Louis Pasteur, 34790 Grabels, France
c
BERKINE SONATRACH ANADARKO, Rte de Cina, 16001 Hassi Messaoud, Algeria
a r t i c l e i n f o
Article history:
Received 25 May 2009
Received in revised form
14 January 2010
Accepted 18 January 2010
Available online 28 January 2010
To the memory of Maurice Mattauer,
professor at the University of Montpellier II,
who left us in April 2009
Keywords:
Fault
Friction
Relay
Wing cracks
Damage zone
a b s t r a c t
Fault friction is a parameter that is difcult to assess along fault zones since its determination depends on
the knowledge of any factor controlling the state of stress around faults. In brittle homogeneous rocks,
a limited number of these factors, such as the shape of the fault surface, the vicinity of fault tips or the
remote stress ratio, are crucial to constrain for this determination. In this paper, we propose to analyse
a eld example in which all these properties are met and where the nature of the slipped structure
suggest differences in static friction. We compare the orientations of branching fractures at strike-slip
relay zones between en echelon stylolites and en echelon joints both reactivated in shear. The eld data
are compared with both photoelastic and 3-D numerical models that consider the remote stress
conditions and the role of the geometry of the strike-slip segments. Based on eld observations, these
analyses quantitatively demonstrate the signicant role of fault friction on the local stress eld orien-
tation and subsequent fracture formation. This work points out that estimations of fault friction based on
analyses of fracture patterns or in situ stresses must be accompanied with a thorough investigation of the
3-D fault shape, its segmentation and the remote stress state.
2010 Elsevier Ltd. All rights reserved.
1. Introduction
Static friction along faults is an extremely important parameter
for the understanding of the seismic cycle, the distribution of
stresses, fracture patterns and damage zones around faults. In the
past decades many efforts have been made to estimate fault friction
along natural faults (e.g. Hanks, 1977; Zoback and Zoback, 1980;
Brace and Kohlstedt, 1980; Lachenbruch and Sass, 1980; Zoback
and Healy, 1984; Mount and Suppe, 1987; Brudy et al., 1997;
Zoback et al., 1987; Scholz, 2000). The measure of static friction
estimated using laboratory tests on fault gouges is scale-limited, i.e.
on gouge samples from a bore hole cutting crossing the fault, and
therefore may not represent the frictional state of the whole
surface. Other approaches, based on the analyses of the heat ow
(Brune et al., 1969; Lachenbruch and Sass, 1980; dAlessio et al.,
2003) or numerical modeling (e.g. Parsons, 2002; Lovely et al.,
2009), allow discussion on the state of friction along the fault but
are quite indirect. The analysis of in situ stresses from bore hole
measurements or fracture patterns are considered as the best
indicator of the frictional state along a fault, (Zoback and Healy,
1984; Zoback et al., 1987; Scholz, 2000).
Assuming that fault cohesion can be close to zero on an active
fault (Byerlee, 1978), the static friction has been approximated by
Amontons rst law, in which the frictional cfcient (m) is
expressed as a function of the shear (F) and normal (N) components
of the forces applied to a frictional surface.
F m*N (1)
This law states that the friction coefcient of an innitely long
fault surface is directly related to the orientation and the
magnitude of the stresses close to this surface (Fig. 1a). This
reveals that the analysis of the stress eld around a fault can be
used to determine the static friction along a fault, in cases where
the remote ratio of stresses applied to the sliding surface is
known. Therefore, any indicators of the stress eld around faults
* Corresponding author.
E-mail addresses: roger.soliva@gm.univ-montp2.fr (R. Soliva), fmaerten@igeoss.
com (F. Maerten), vincent_auzias@berkine.com (V. Auzias).
Contents lists available at ScienceDirect
Journal of Structural Geology
j ournal homepage: www. el sevi er. com/ l ocat e/ j sg
0191-8141/$ see front matter 2010 Elsevier Ltd. All rights reserved.
doi:10.1016/j.jsg.2010.01.008
Journal of Structural Geology 32 (2010) 17211731
(e.g. bore hole analysis, faults or fracture patterns) provide the
opportunity to quantify the static friction. However, this analytical
approach, based on Amontons rst law, assumes that the fault
plane is rectilinear and that the fault tips are innitely far from
the study area. Such a rst order approximation is quite unreal-
istic for natural faults having tips, being irregular, segmented or
more complex in shape (Fig. 1bd). The local orientation and
magnitude of the stress eld around a fault does not rely only on
fault friction, which makes its determination non-unique unless
we have knowledge of the other factors perturbing the local stress
eld.
In homogeneous rocks, the rst parameter that has been
considered as acting on the local stress eld, and more precisely on
the crack angle to a fault, is the static friction coefcient (e.g. Petit
and Barquins, 1988; Barquins et al., 1997; Martel, 1997; Ohlmacher
and Aydin, 1997; Willemse and Pollard, 1998; Zhou, 2006; Mutlu
and Pollard, 2008). However, the remote stress angle has been
considered as very important (Barquins et al., 1992; Ohlmacher and
Aydin, 1997) as well as the remote stress ratio (Auzias et al., 1997;
Katternhorn et al., 2000; Zhou, 2006). Others factors more related
to the geometry and behaviour of the fault surface also seem to be
very inuential, as the 3-D geometry of the faults (e.g. Segall and
Pollard, 1980; King et al., 1994; Willemse, 1997; Maerten et al.,
2002; Bourne and Willemse, 2001), its spatial/temporal evolution
(Willson et al., 2007; Lunn et al., 2008; Moir et al., 2009), and fault
opening (Kattenhorn and Marshall, 2006). Therefore, any analysis
of fault zones that aims to estimate the role of fault friction on the
stress eld, or in contrast to determine the state of friction from
stresses analysis, must know any of these factors that can perturb
the local stress eld.
In this paper, we analyse a eld example in which these factors
can be estimated. Drastic differences in fracture orientation
between reactivated frictional stylolites (i.e. structures of high
friction coefcient) and frictionless joints (i.e. structures of low
friction coefcient) suggest that friction is a prominent property
inuencing the stress perturbation at the close vicinity of a fault.
We chose to study fracture orientation at extensional relay zones
because the stress orientation has been described as quite stable in
space along a relay zone (compared to outside) due to the juxta-
position of the two extensive fault quadrants (see Fig. 1c) (e.g.
Auzias et al., 1997; Ohlmacher and Aydin, 1997). We compare the
eld data to photoelastic and 3-Dnumerical models to demonstrate
and quantify the signicant role of static friction on the stress and
fracture orientation at extensional relay zones.
= /
n
a b
Rectilinear fault
without tips
c
Rectilinear not frictional
segmented fault with tips
Rectilinear not frictional
fault with tips
1
1st Amonton's law
Stress trajectory = 0
= 0
= 0
= 0
n
= 0
=
0
= 0
n = 0
n = 0
n = 0
n = 0
n = 0
Stress trajectory = 0
d
Complex not frictional
fault shapes with tips
Stress trajectory = 0
1
Fig. 1. Comparison between the stress perturbation due to fault friction (a) and the stress perturbation due to different examples of fault geometry (b, c and d). (b), (c) and (d) are s
1
stress patterns inferred from photoelastic modeling (Joussineau et al., 2003). The remote stress applied is uniaxial. Dots represent fault tips. This gure shows that even for m 0, the
orientation of s
1
can be oblique and even parallel to the fault surface, rather than perpendicular as suggested by Amontons rst Law.
R. Soliva et al. / Journal of Structural Geology 32 (2010) 17211731 1722
2. Field data
2.1. Geological setting
The studied exposure, located close to Les Matelles (15 kmNorth
of Montpellier, France, Fig. 2), is a suitable site for the study of
brittle tectonics in limestones and stress perturbations around
meso-scale faults (Rispoli, 1981; Fletcher and Pollard, 1981; Petit
et al., 1999). The brittle tectonic structures observed (Fig. 3a)
were formed during multiphase compressive tectonics allowing the
formation of joints and stylolites. These structures of similar
dimension and orientation have been reactivated as slip surfaces
during a late tectonic event (Petit and Mattauer, 1995). Because of
their different roughness, joints and stylolites are expected to be of
different frictional properties during slip. Most of them show
secondary fracturing and linkage at relay zones (Fig. 3b), that can be
used as indicators of the palaeostress orientation (e.g. Rispoli et al.,
1981; Petit and Mattauer, 1995). It is therefore worthwhile to
address with particular care on the geological setting and history of
the brittle structures that will be used to constrain the role of fault
friction on fracture orientation.
The studied exposure has been fully described in a number of
previous studies (e.g. Rispoli et al., 1981; Taha, 1986; Petit and
Mattauer, 1995; Petit et al., 1999). This area is located in the
vicinity of a fault branch called the Lirou fault (Fig. 2b). The Matelles
fault zone, like many faults in the area, had both left-lateral strike
slip related to the Pyrenean shortening and normal slip related to
the Oligocene rift extension in the Languedoc. Middle cretaceous
normal slip along the MatellesCorconne fault zone is also expec-
ted during the Durancian tectonic events.
The brittle deformation sequence described by Petit and
Mattauer (1995) begins by a vertical jointing stage of the lime-
stone layers with two principal trends, N020 and N140. The second
stage is a rst generation of stylolite formation oriented N040. The
third stage, the most important for our study, is the reactivation of
the previous structures as sinistral and dextral strike slips due to
a last shortening creating wing cracks, en echelon veins and
a second generation of stylolites around the reactivated defects. As
shown by this last generation of joints and stylolites formed, the
last shortening stage occurs with the maximumprincipal stress (s
1
)
oriented NorthSouth. As suggested by rock experiments, photoe-
lastic models, numerical and analytical solutions (see Wawersik
and Brace, 1971; Petit and Barquins, 1988; Barquins and Petit,
1992; Chaker and Barquins, 1996; Lunn et al., 2008) the presence
of wing cracking around reactivated defects (see Fig. 3b) implies
remote stress conditions close to uniaxial loading (s
1
/s
3
10).
Conditions close to horizontal uniaxial stresses are possible at
shallowdepths, i.e. for little conning pressure. The expected depth
of faulting in the upper Jurassic limestone was probably less than
the thickness of the lower cretaceous series (w200 m), which was
potentially yet well eroded during the Pyrenean shortening. This
local stress state condition (high ratio of maximum to minimum
principal stresses, s
H
/s
h
) and reorientation of s
1
axis has been
related to a restraining bend along the Les Matelles fault during
Pyrenean strike-slip movements (Rawnsley et al., 1992; Petit et al.,
1999).
2.2. Extensional relay geometries
The last stage event provides the opportunity to analyse the
geometry of branching at relay zones between slipped overlapping
stylolites vs. slipped overlapping joints (Figs. 4 and 5a). The angle b,
dened as the angle between the orientation of remote s
1
relative
to the joints or stylolites reactivated in shear (Fig. 5c), is quite
variable (variation of w40
to 90
, 45
and 70
.
Here b is the angle between the axial loading and the planar defects
forming the relay zone.
The models are subjected to uniaxial conditions in order to be
consistent with the eld conditions expected (see Section 2). The
axial compressive load is imposed by an electromechanic testing
machine (Davenport 30 kN) and no lateral pressure is added. To
prevent bending of the PMMA plate under vertical loading, the
samples are maintained between vertical tighteners.
3.2. Experimental results of extensional relay stress pattern
The two types of models (frictional vs. not frictional) show
signicant differences in their local stress eld distribution. Fig. 7
presents stream lines of s
1
for not frictional (a) and frictional
conguration (b) subjected to a vertical loading with b 20
. The
stress eld is less perturbed in orientation for the frictional case.
Without friction, the orientation of s
1
is normal to the faults
(deviation of 70
. Additional tests,
not presented here and done for variable overlap and constant
spacing between the defects, showthe same maximumvalues of s
1
deviation and a better stability of s
1
orientation in the relay zone as
the overlap increases.
All the studied tests show results generally consistent with eld
observations. Fig. 8 exhibits the compilation of the a and b angles
data for all the tests done with constant relay geometry. Note that
a here corresponds to the angle between the slipped defect and s
1
at the center of the relay zone. Tests with no friction lie in the graph
area of high a and relatively lowb angles, which corresponds to the
zone of slipped joints (of low friction compared to slipped stylo-
lites). In contrast, frictional tests data lie in the area of lower a and
relatively high b angles, which corresponds to the zones of slipped
stylolites (high friction).
4. Numerical modeling
The numerical code used to investigate fault friction is a 3-D
Boundary Element Method (BEM) called Poly3D (Thomas, 1993). It
relies on the analytical solution of an angular dislocation in
a homogeneous elastic whole- or half-space (Comninou and
Dundurs, 1975). As opposed to the Okadas code (Okada, 1985),
Fig. 3. Field photographs of the study area. (a) Outcrop overview showing the layered upper Jurassic mudstones damaged by a dense pattern of calcite sealed fractures and
stylolites. (b) First generation stylolites reactivated as sinistral strike slips showing wing cracks and branched stylolites (see Rispoli, 1981). The length of the scale bar is 20 cm.
R. Soliva et al. / Journal of Structural Geology 32 (2010) 17211731 1724
which uses rectangular elements, Poly3D discretizes faults and
fractures using triangular elements, and therefore avoids the
creation of overlaps and gaps between adjacent elements which
perturb the solution (Maerten et al., 2005). Mixed traction
displacement boundary conditions can be used for each constitu-
tive element of the model (tractions are shear and normal stresses
resolved on the fault surface). When traction boundary conditions
are specied, we have to solve for the corresponding unknown
displacement discontinuity according to the initially prescribed
traction values. As soon as all displacement discontinuities are
known (i.e. the slip patches), strain, stress and displacement can be
computed at any observation point within the elastic eld. Note
that transient variations in friction coefcient or the dynamic stress
eld are not considered (e.g., Poliakov et al., 2002).
Inorder tohave a frictional behaviour, the codehas beenextended
to support inequality constraints on traction and displacement.
Specically, the static Coulomb friction has been implemented as
a traction inequality constraint and validated by comparison with
analytical and numerical solutions (Maerten et al., 2009). For a given
fault surface, thecoefcient of frictionandcohesioncanbeprescribed
globally onto a fault surface or locally, each constitutive element
having their own coefcients. Traction boundary conditions are
imposed along the three axis of each triangular element local coor-
dinate system (dip, strike and normal directions).
Fig. 4. (a) Examples of joints (left side) and stylolites (right side) reactivated as left-lateral strike slips. (b) Interpretation in terms of stress orientation using the remote orientation
of syn-kinematic joints and stylolites. The remote stress orientation is slightly different in the two cases because they were not measured at the same location and that the larger
Lirou fault probably modify the stress eld orientation at this NorthSouth last compressive stage.
R. Soliva et al. / Journal of Structural Geology 32 (2010) 17211731 1725
For a model subjected to a compressive far eld stress, inter-
penetration of the elements has to be avoided. This is achieved by
using the displacement inequality constraint u
z
0, where u
z
represents the computed normal displacement of a triangular
element. Again, traction boundary conditions are imposed along
the three axes of each triangular element local coordinate system
(Maerten et al., 2009).
4.1. Model set up
Fig. 9a and b depicts the model congurations used for the BEM
modeling for joints and stylolites, respectively, and are built upon
eld observations (Fig. 4). All the veins and stylolites traces in the
vicinity of the zone of interest have been carefully mapped and
then vertically extruded in depth all along the limestone layer
thickness, giving rise to the 3-D triangulated surfaces. As the
Coulomb friction relates the shear to the normal components of the
forces applied to a frictional surface, traction boundary condition
for the three local axes (dip, strike and normal) of each constitutive
triangular element is used.
4.2. Modeling of joints reactivated in shear
The joints model, depicted in Fig. 9a, is based on the eld
observations shown in Fig. 4, left side. These rst generation joints
are subjected to a far eld remote stress with uniaxial compressive
condition and s
1
(in this area) oriented N170 as suggested by the
presence of surrounding joints and stylolites (third brittle defor-
mation stage, see Section 2). In order to display the stress orien-
tation within the extensional relay resulting from the computed
displacement discontinuities, an observation grid is placed close to
the top of the model (Fig. 9). Then, two simulations are performed:
a rst one, with a constant coefcient of friction m 0.6 for all
discontinuities and no cohesion, and a second without any friction
but with non-interpenetration as a unique constraint. The elastic
material properties used for the surrounding limestone are n 0.25
and E 1 GPa (see Hatheway and Kiersch, 1989).
Fig. 10a and b displays the frictional and not frictional models,
respectively. The orientation of s
1
axis t better with the strike of
the branching fracture in the case where the frictional coefcient
equals zero. Since s
1
should be parallel to the strike of the
branching fracture, these models suggest that, at the initiation of
the linkage, the slipping joints were preferably not frictional. This is
consistent with the absence of macroscopic irregularities along
these rectilinear structures.
4.3. Modeling of stylolite reactivated in shear
For the stylolites model depicted in Fig. 9b, the uniaxial
compressive far eld stress is oriented N015, as proved by the
Fig. 5. Variation in the geometry of extensional jogs between stylolites and joints reactivated in shear. (a) and (b) are eld examples of reactivated stylolites and joints, respectively.
Coloured dashed lines represent the orientation of fractures in the relay zones. (c) Graph of a and b angles for all the reactivated stylolites and joints measured in the eld. a and
b angles are represented in a small scheme at the left. Field observations show higher b angles for reactivated stylolites.
R. Soliva et al. / Journal of Structural Geology 32 (2010) 17211731 1726
presence of surrounding joints and stylolites (see Fig. 4, right hand
side, and Fig. 9b in Petit and Mattauer, 1995). Since this model is
composed of two relays, two observation grids are placed in the
vicinity of them close to the top of the model. A rst simulation is
done using only the non-interpenetration constraint (i.e. with
m 0), whereas a second one employs a constant coefcient of
friction m 0.6 without cohesion.
Fig. 11a and b displays the results on the two observation grids
for the not frictional and frictional models, respectively. As opposed
to the previous joint modeling, the linking structures are more
consistent with high friction stress orientations.
4.4. Parametric analysis
A series of models have been done for variable friction and
constant fault geometry consistent with the overlapping segments
of the experimental PMMA model. The 3-D shape of the model is
shown in Fig. 12a. The results are analyzed on the observation grid
which allows to compare a conguration close to the eld and the
photoelastic modeling (Fig. 12b). The models were performed with
variation of static friction coefcient and b angles as shown in
Fig. 12c. The elastic material properties used are the same than
above since a large part of the eld data used for comparison were
measured in limestone (Fig. 5).
The results are in good agreement both with eld and experi-
mental data. Fig. 13 exhibits the compilation of the a and b angles
data for all the tests done. As for the experimental analysis,
a corresponds to the angle between the slipped surface and s
1
at
the center of the relay zone. The numerical models with no or little
friction lay in the graph area of high a and relatively low b angles,
which corresponds to the zone of slipped joints (frictionless
structures). In contrast, frictional models t in the area of lowa and
relatively high b angles, which corresponds to the zone of slipped
stylolites (frictional structures).
Fig. 6. (a) Experimental device of the photoelastic modeling. (b) Example of isoclinic and isochromic fringes obtained in a vertical uniaxial loading experiment of slipping over-
lapping open defects.
Fig. 7. Drawing of s
1
obtained from the analysis of isoclinic fringes for (a) uniaxial
vertical loading of open defects, i.e. non-frictional and (b) closed defects, i.e. frictional.
Open defects (no friction)
Closed defects (frictional)
Limestone (Liu, 1983)
Limestone (Taha, 1986)
Granite (Raynaud, 1978)
80
70
60
50
40
30
20
20 40 60 80
10
Slipped stylolites
Slipped joints
()
()
Experimental data Field data
Frictional
Frictionless
Present
study
Present study
Fig. 8. Comparison between a and b angles obtained by photoelastic modeling with
the dataset measured in the eld.
R. Soliva et al. / Journal of Structural Geology 32 (2010) 17211731 1727
5. Discussion
5.1. Stress perturbation and friction of the slipping defects
The models are in good agreement with the eld observations,
however they do not cover the entire range of data, especially for
the low a angles (Fig. 13). This point can be discussed with respect
to a limited number of unconstrained factors that may inuence the
stress eld around the slipping defects.
Inelastic deformations can modify the magnitude of residual
stresses in the host rock around faults, but this is probably not the
explanation of the scatter observed in Fig. 13 for two reasons. First,
with respect to the brittle subsurface conditions of deformation,
the studied limestone probably has negligible inelastic behaviour
preceding its shear yielding strength (Rispoli, 1981; Petit and
Mattauer, 1995). Second, inelastic deformation around fault, if
any, has probably a larger inuence on the stress magnitude than
on the orientation (Bu rgmann and Pollard, 1994), which one close
to the fault must be directly related to fault friction. This suggests
that elastic models are relevant to simulate residual stresses related
to fault slip in this geological context, and that the spreading of eld
data compared to the model is mainly due to others factors.
The effect of 3-D fault geometry, especially the fault aspect ratio,
on the stress distribution around fault has probably little inuence.
Soliva et al. (2006) show that the dimension of the area of stress
perturbation around a fault scales linearly with the fault length
since the fault growth is radial, and tends to be limited to a certain
distance when the fault height reaches a constant value. This means
that for vertically restricted fault by strata bounds, the 3-D shape
can inuence the orientation of the stress eld at about a distance
around the fault equivalent to the layer thickness. In the present
study, for all the measures of angles made in the eld (in the relay
zones), the distance of the fractures around the faults is always
lower than the layer thickness potentially restricting the faults
(tens of cm). We therefore work in a window around the faults
where the stress perturbation should not be inuenced by the fault
aspect ratio (3-D shape), and that all the eld measurements could
be compared to 2-D photoelastic models or 3-D full space models
proposed.
The reason of the difference of scatter between the eld data
and the models is potentially purely geometric. For all the model
results presented in Fig. 13 (both parametric and photoelastic), the
fault congurations are idealized as two planar surfaces with
constant overlap and spacing, whereas the eld data are fromfaults
more complex in shapes, with variable overlap and spacing, curved,
with multiple segments and potentially more complex in 3-D.
We have shown that friction is the main factor controlling the
stress perturbation and the orientation of linking fractures.
However the physical cause for this variation in friction needs to be
discussed. Obviously, this cause can be reasonably ascribed to the
difference of surface roughness between the stylolites and the
joints. However, the analysis of the roughness of the slipped defects
is not very relevant on faulted stylolites since after faulting they
show a smoothed irregularity that is certainly different than the
initial one. The measure of roughness has been done on non-
reactivated stylolites (see Delair and Leroux, 1978; Raynaud and
Carrio-Schaffhauser, 1992, for the quantitative analysis of stylo-
lites roughness in the same study area). However, these stylolites
were not reactivated potentially because of a threshold friction,
then different than the initial state on the faulted stylolites. On the
other hand, efforts in measuring the surface roughness of the
slipped defects cannot be very conclusive since it represents
the nite strain.
5.2. Estimation of fault friction and upscaling
From three different approaches: (1) eld study, (2) experi-
mental modeling and (3) numerical modeling, we have shown that
the angle of fracture branching in strike-slip relay zones is highly
dependent on the frictional state of the overlapping faults.
It is worthwhile to note that the static friction estimated at the
relay zone corresponds to the friction of the faults in the vicinity of
the relay zone and at the time of the fault interaction through the
relay zone. This frictional property may therefore have evolved
through time and space with the progression of fault coalescence.
The quasi-static friction estimated must be therefore considered
as the time integrated friction of the period of fault interaction
through their stress eld. We also must keep in mind that this
Fig. 9. 3-D view of the geometry of the defects reactivated in shear used in Poly3D for the numerical simulation. (a) 3-D geometry of the reactivated joints shown in Fig. 4 with the
observation grid (dark square) on which the stresses are represented in Fig. 10. The position of the observation grid corresponds to the position of the top of the limestone layer
observed in the eld. (b) 3-D geometry of the reactivated stylolites shown in Fig. 4. The observation grids (not shown here), on which the stress eld is presented in Fig. 11, are
placed at the same level than in (a).
R. Soliva et al. / Journal of Structural Geology 32 (2010) 17211731 1728
approach is only suitable along faults if the remote stress conditions
are well known, since the ratio of s
H
/s
h
is very important for the
stress orientation in the relay zone (Auzias, 1995). Any indicator of
the remote and relay zone stress eld are worth considering.
With regard to estimation of fault friction on large scale active
faults, particular care must be taken with the stress eld deter-
mined from data close to the Earths surface. For depth shallower
than 300 m, the orientation of the principal stress may be different
than the tectonic stresses predominant at depth (e.g. Engelder,
1993). On large faults (kilometric scale length), it seems therefore
more appropriate to provide an estimation of fault friction based on
the tectonic stress orientation or the fracture patterns measured in
deep bore holes, which are more representative of the brittle crust
stress state.
This approach, based on eld observation and numerical
modeling at the relay zone, seems therefore relevant for the esti-
mation of fault friction along active fault segments interacting
through their stress eld. Its main advantage compared to rock test
measures, is the in situ estimation of the friction in its own
geological context. We integrate a large part of the fault surface
around the relay zone (and it can be done outside as well), as
opposed to tests done on fault rocks, which correspond to a specic
location of the fault surface crossed by the bore hole. Moreover, this
approach provides an overall value of friction of the entire active
fault zone, which may be composed of compartment with various
fault rocks, as for example coarse cataclasites or gouges which can
be difcult to analyse in laboratory tests. On the other hand,
a limitation of this method is that permanent deformation (e.g.
measured by GPS) can not be used to estimate friction with a quasi-
static elastic model. Visco-elastic simulation of the lithosphere
could be more appropriated if it allows to simulate the precise
geometry of the fault segments.
6. Conclusion
In situ static friction can be estimated along a fault plane if its
shape, the far eld stress conditions and the stresses at its vicinity
are well known. Joints and stylolites reactivated in shear show
roughly different angles of linking fractures at their extensional
relay zones. The irregular shape of the stylolites and the rectilinear
trace of the joints suggest that different frictional behaviour may
explain these differences in branching angles. Photoelastic and
numerical modeling conrm this phenomenon. For the same
remote stress conditions, variation of the static friction along
simulated faults explains a wide part of the range of branching
angle measured at relay zones. In particular, our paper reveals four
main points:
Fig. 10. Model results for the joints reactivated in shear shown in Fig. 9. (a) Modeling
result for m 0. (b) Modeling result for m 0.6. The small arrows on the observation
grid show the local orientation of s
1
.
Fig. 11. Model results for the stylolites reactivated in shear shown in Fig. 9. (a) Model
result for a m 0. (b) Model result for a m 0.6. The small arrows on the observation
grid show the local orientation of s
1
.
R. Soliva et al. / Journal of Structural Geology 32 (2010) 17211731 1729
(1) A simple Amontons rst law cannot be used systematically to
infer the static friction along natural faults,
(2) To discuss the amount of friction along a fault, the analysis of
the local stress eld must be compared to elasto-static
approaches that integrate the effect of mechanical interac-
tions along ended faults, irregular in shape, segmented or more
complex,
(3) Both eld data, photoelasticity and numerical modeling show
that wide variations of friction can explain a large part of the
variation in the angle of secondary fracturing in the relay zones,
(4) Shear-reactivated joints have lower estimated static friction
than shear-reactivated stylolites.
Acknowledgments
We which to particularly thank and dedicate this paper to
Maurice Mattauer who left us in April 2009. He discovered the
studied outcrop and recently participated to discussions about this
work at the laboratory and also in the eld. The eld work from
Roger Soliva was supported by an Action Structurante 2006 grant
from the laboratory Geosciences Montpellier UMR5243. W. Ashley
Grifth and Roy Schlische are thanked for their helpful comments.
References
dAlessio, M.A., Blythe, A.E., Bu rgmann, R., 2003. No frictional heat along the San
Gabriel fault, California: evidence from ssion-track thermochronology.
Geology 31, 541544.
Auzias, V., 1995. Photoelastic modeling of stress perturbations near faults and of the
associated fracturing: petroleum industry application, II: Mechanism of 3D joint
development in a natural reservoir analogue: the at-lying Devonian Old Red
Sandstone of Caithness (Scotland). Ph.D. thesis, Universite Montpellier II, p. 311.
Auzias, V., Rives, T., Rawnsley, K.D., Petit, J.-P., 1997. Fracture orientation modeling in
the vicinity of a horizontal well. Bulletin Elf Aquitaine Production F64018,
381397.
Barquins, M., Chaker, C., Petit, J.-P., 1997. Inuence du frottement sur le branche-
ment de ssures a` partir de de fauts obliques soumis a` une compression uni-
axiale. Compte Rendu de LAcademie des Sciences T324, 2936.
Barquins, M., Petit, J.-P., 1992. Kinetic instabilities during the propagation of
a branch crack: effects of loading conditions and internal pressure. Journal of
Structural Geology 14, 893903.
Barquins, M., Petit, J.-P., Maugis, D., Ghalayini, K., 1992. Path and kinetics of
branching from defects under uniaxial and biaxial compressive loading. Inter-
national Journal of Fracture 54, 139163.
Bourne, S.J., Willemse, E.J.M., 2001. Elastic stress control on the pattern of tensile
fracturing around a small fault network at Nash Point, UK. Journal of Structural
Geology 23, 17531770.
Brace, W.F., Kohlstedt, D.L., 1980. Limits on lithospheric stress imposed by labora-
tory experiments. Journal of Geophysical Research 85, 62486252.
Brudy, M., Zoback, M.D., Fuchs, K., Rummel, F., Baumgartner, J., 1997. Estimate of the
complete stress tensor to 8 km depth in the KTB scientic drill holes: impli-
cations for crustal strength. Journal of Geophysical Research 102, 18,45318,475.
Brune, J.N., Henyey, T.L., Roy, R.F., 1969. Heat ow, stress, and rate of slip along the
San Andreas fault, California. Journal of Geophysical Research 74, 38213827.
Bu rgmann, R., Pollard, D.D., 1994. Strain accommodation about strike-slip fault
discontinuities in granitic rock under brittle-to ductile conditions. Journal of
Structural Geology 16, 16551674.
Byerlee, J.D., 1978. Friction of rocks. Pure and Applied Geophysics 116, 615626.
Chaker, C., Barquins, M., 1996. Sliding effect on branch crack. Physics and Chemistry
of the Earth 21, 319323.
Comninou, M., Dundurs, J., 1975. The angular dislocation in a half space. Journal of
Elasticity 5, 203216.
Delair, J., Leroux, C., 1978. Me thodes de quantication de la disparition de matie` re
au niveau de stylolites tectoniques et me canismes de la deformation cassante
des calcaires. Bulletin de la Socie te Ge ologique de France 7, 137144.
Engelder, T., 1993. Stress Regimes in the Lithosphere. Princeton University Press,
Princeton, New Jersey, U.S.A., 475 pp.
Fletcher, R.C., Pollard, D.D., 1981. Anticrack model for pressure solution surfaces.
Geology 9, 419424.
Hanks, T.C., 1977. Earthquake stress drops, ambient tectonic stress, and the stresses
that drive plate motion. Pure and Applied Geophysics 115, 441458.
Hatheway, A.W., Kiersch, G.A., 1989. Engineering properties of rock. In:
Carmichael, R.S. (Ed.), Practical Handbook of Physical Properties of Rocks and
Minerals. CRC Press, Boca Raton, FL, pp. 672715.
Hete nyi, M., 1966. Handbook of Experimental Stress Analysis. Wiley, New York.
Joussineau, G., Petit, J.-P., Gauthier, B.D.M., 2003. Photoelastic and numerical
investigation of stress distributions around fault models under biaxial
compressive loading conditions. Tectonophysics 363, 1943.
Kattenhorn, S.A., Aydin, A., Pollard, D.D., 2000. Joints at high angles to normal fault
strike: an explanation using 3D numerical model of fault perturbated stress
eld. Journal of Structural Geology 22, 123.
Kattenhorn, S.A., Marshall, S.T., 2006. Fault-induced perturbed stress elds and
associated tensile and compressive deformation at fault tips in the ice shell of
Europa: implications for fault mechanics. Journal of Structural Geology 28 (12),
22042221.
King, G.C.P., Stein, R.S., Lin, J., 1994. Static stress changes and the triggering of
earthquakes. Bulletin of the Seismological Society of America 84 (3), 935953.
Lachenbruch, A., Sass, J., 1980. Heat ow and energetics of the San Andreas fault
zone. Journal of Geophysical Research 85, 61856222.
Open defects with contact points
(few friction)
Closed defects (frictional)
Limestone (Liu, 1983)
Limestone (Taha, 1986)
Granite (Raynaud, 1978)
80
70
60
50
40
30
20
20 40 60 80
10
Slipped stylolites
Slipped joints
= 1
()
()
= 0.8
= 0.6
= 0.4
= 0.2
= 0
Numerical data
Experimental data Field data
Frictional
Frictionless
Present
study
Present study
Present study
Fig. 13. Comparison of a and b angles between the data obtained by numerical
modeling, photoelastic modeling and the dataset measured in the eld. A wide part of
a and b spreading can be explained by variations in frictional coefcient of the slipped
defects.
Fig. 12. Input conditions for the parametric modeling. (a) Conguration of the 3-D
model geometry with an example of computed displacement contours in color. (b)
Horizontal view of the model conguration showing the angles a and b. (c) Variables
used in the parametric study.
R. Soliva et al. / Journal of Structural Geology 32 (2010) 17211731 1730
Lovely, P.J., Pollard, D.D., Mutlu, O., 2009. Regions of reduced static stress drop near
fault tips for large strike-slip earthquakes. Bulletin of the Seismological Society
of America 99, 16911704.
Lunn, R.J., Willson, J.P., Shipton, Z.K., Moir, H., 2008. Simulating brittle fault growth
from linkage of preexisting structures. Journal of Geophysical Research 113,
B07403. doi:10.1029/2007JB005388.
Maerten, L., Gillepsie, P., Pollard, D.D., 2002. Effect of local stress perturbation on
secondary fault development. Journal of Structural Geology 24, 145153.
Maerten, F., Resor, P.G., Pollard, D.D., Maerten, L., 2005. Inverting for slip on three-
dimensional fault surfaces using angular dislocations. Bulletin of the Seismo-
logical Society of America 95, 16541665.
Maerten, F., Maerten, L., Cooke, M., 2009. Solving 3D boundary element problems
using constrained iterative approach. Computational Geosciences. doi:10.1007/
s10596-009-9170-x.
Martel, S.J., 1997. Effects of cohesive zones onsmall faults andimplications for secondary
fracturing and fault trace geometry. Journal of Structural Geology 19, 835847.
Moir, H., Lunn, R.J., Shipton, Z.K., Kirkpatrick, J.D., 2009. Simulating brittle fault
evolution from networks of pre-existing joints within crystalline rock. Journal
of Structural Geology. doi:10.1016/j.jsg.2009.08.016.
Mount, V., Suppe, J., 1987. State of stress near the San Andreas fault: implications for
wrench tectonics. Geology 115, 11431146.
Mutlu, O., Pollard, D.D., 2008. On the patterns of wing cracks along an outcrop scale
aw: a numerical modeling approach using complementarity. Journal of
Geophysical Research 113, B06403. doi:10.1029/2007JB005284.
Ohlmacher, G.C., Aydin, A., 1997. Mechanics of veins, fault and solution surface
formation in the Appalachian valley, U.S.A.: implications for fault friction, state
of stress and uid pressure. Journal of Structural Geology 19, 927944.
Okada, Y., 1985. Surface deformation due to shear and tensile faults in a half-space.
Bulletin of the Seismological Society of America 75, 11351154.
Parsons, T., 2002. Nearly frictionless faulting from unclamping in long-term inter-
action models. Geology 30, 10631066.
Petit, J.-P., Barquins, M., 1988. Can natural faults propagate under mode II condi-
tions? Tectonics 7, 12431256.
Petit, J.-P., Mattauer, M., 1995. Palaeostress superimposition deduced from meso-
scale structures in limestone: the Matelles exposure, Languedoc, France. Journal
of Structural Geology 17, 245256.
Petit, J.P., Wibberley, C.A.J., Ruiz, G., 1999. Crack-seal, slip: a new fault valve
mechanism? Journal of Structural Geology 21, 11991207.
Poliakov, A.N.B., Dmowska, R., Rice, J.R., 2002. Dynamic shear rupture interactions
with fault bends and off-axis secondary faulting. Journal of Geophysical
Research 107 no. B11, 2295. doi:10.1029/2001JB000572, ESE 6-1 6-18.
Rawnsley, K.D., Rives, T., Petit, J.P., Hencher, S.R., Lumsden, A.C., 1992. Joint development
in perturbed stress elds near faults. Journal of Structural Geology 14, 939951.
Raynaud, S., Carrio-Schaffhauser, E., 1992. Rock matrix structures in a zone inu-
enced by a stylolite. Journal of Structural Geology 14, 973980.
Rispoli, R., 1981. Stress elds about strike-slip faults inferred from stylolites and
tension gashes. Tectonophysics 75, T29T36.
Scholz, 2000. Evidence for a strong San Andreas fault.
Segall, P., Pollard, D.D., 1980. Mechanics of discontinuous faulting. Journal of
Geophysical Research 85, 43374350.
Soliva, R., Benedicto, A., Maerten, L., 2006. Spacing and linkage of conned faults:
the importance of mechanical thickness. Journal of Geophysical Research 111,
B01402. doi:10.1029/2004JB003507.
Taha, M., 1986. Apport de la microtectonique cassante au proble` me des trajectoires
de contraintes et de leurs perturbations. Exemples du Nord de Montpellier,
the` se dE
sm
s
m
d
L
(2)
where m
d
is a prescribed dynamic friction value and L is the slip-
weakening distance. Once the element has slipped the length of L,
the frictional strength of the element will remain at the dynamic
friction value. Slip-weakening friction is a simplistic friction law in
the sense that it does not capture velocity dependence or memory
effects on friction (Dieterich, 1979; Ruina, 1983). However, slip-
weakening friction is adequate for dynamic rupture simulations
where earthquake cycles are not relevant and has been used
previously to study off-fault fracturing (Dalguer et al., 2003).
2.2. Model setup
We chose our boundary conditions to reect the conditions on
a fault at seismogenic depths subjected to shear stresses that are
close to the frictional strength of the center patch of the fault
(reecting that this part of the fault is critically stressed). Our
boundary conditions simulate simple shear conditions with
constant shear and normal displacements along the top of the body
and linearly decreasing displacements along the sides of the body,
H.M. Savage, M.L. Cooke / Journal of Structural Geology 32 (2010) 17321741 1733
with zero displacement along the bottom edge (Fig. 1). Normal
displacements are calculated to reect the lithostatic stress asso-
ciated with a fault buried approximately 5 km (125 MPa).
Displacement is applied in one step and changes in stress and
displacement in the ensuing iterations represent the system
reaching convergence. At each iteration, the friction coefcient
evolves along the fault (Eq. (2)), elements along the fault slip (Eq.
(1)) and new elements are added to simulate damage production.
The fault is 15.1 m long and horizontal. The fault and the bound-
aries are discretized into equal-length 5 cm long linear elements.
This element length provides satisfactory resolution of rupture
advancement while limiting computational load. Along the fault,
we prescribe a 3.1 m long center patch that is slightly frictionally
weaker, so that the center patch fails rst. In this way, we create
a nucleation patch along which the rupture begins and propagates
toward either end of the fault. The center patch is long enough to
induce unstable sliding along the entire fault. The shear displace-
ments along the boundaries of the model are chosen so that the
center weak patch (m 0.28) is at failure. The subsequent reduction
of friction coefcient from0.28 to 0.2 along the center patch during
slip provides a shear stress drop of 10 MPa. The sides of the fault
have higher prescribed friction coefcient, 0.32, and slip in
response to the stress drop on the center patch.
Because rocks are weakest in tension, we choose to investigate
areas likely to produce opening-mode cracks. New tensile fractures
grow where the tangential normal stresses along a fault element
exceed the tensile strength of the host rock, prescribed here as
15 MPa. Tensile stresses along the fault occur due to slip gradients
between elements. Because our model does not have a pre-existing
mesh, fractures are free to form at any orientation and nucleate
perpendicular to the local maximumtensile direction. Fractures can
form at the nodes between every other element and can grow by
one element length during each iteration of frictional slip. The
minimum spacing of off-fault fractures is 10 cm in these models.
Propagation continues until the stress intensity factor (K
I
) at the tip
of the fractures is less than the fracture toughness (K
Ic
), prescribed
here as 2.5 MPa m
1/2
. Newfractures are permitted to open and slide
with frictional resistance equal to the static friction value of the
fault element that spawned the new fracture. Fractures are not
allowed to interpenetrate.
2.3. Analysis of mechanical work
Total work of the fault system describes all of the energy
expended during tectonic deformation of the fault and the
surrounding host rock. Energy is consumed during deformation
from work against gravity (W
grav
), propagation of new surfaces
(W
prop
), work to overcome frictional resistance to sliding along the
fault (W
fric
), work that promotes ground motion in the form of
seismic radiation (W
seis
), and nally work that goes into off-fault
deformation which we refer to as an internal strain energy (W
int
).
The total work reects the summation of each of these
components:
W
tot
W
grav
W
prop
W
fric
W
seis
W
int
(3)
Each component of the total work done on the fault system
can be evaluated from our model. In our analyses, we do not
consider the effects of gravity because our fault is horizontal and
our surface has no topography. The deformational work budget
can be delineated in a variety of ways. Here we follow that used
by Mitra and Boyer (1986), Cooke and Murphy (2004), Del Cas-
tello and Cooke (2007) and Ismat (2008). The result is very
similar to the energy budget delineated by Kanamori and Heaton
(2000) and Abercrombie and Rice (2005). Both approaches
consider the same energy budget but divide the energy terms up
in slightly different ways based on a difference in observables. As
we describe each term we point out the differences in the
notations.
The external work represents the amount of work applied at the
external boundaries of our system. The complete external work
term is integrated along both the boundary and the applied
displacements (u
j
);
W
ext
ZZ
s
ij
u
j
; x
u
j
du
j
dx (4)
where s
ij
is the stress along the boundary due to u
j
and x is position
along the external boundary. In a closed system, the total work of
the system must equal the external work. In our models, the
boundaries are not permitted to move so the external work does
not change during rupture propagation; the only changes are the
partitioning of work amongst the different work components
within the system.
In order for a tectonic fault to slip, the shear stress along the
fault must overcome the frictional strength of the fault. The work
done against frictional resistance at a single fault segment is
calculated as:
W
fric
s
N
msA (5)
where s
N
is normal stress, m is the coefcient of friction, s is slip and
A is the ruptured areas of the fault. When stresses along the fault
are tensile so that normal stresses are zero or positive, the work
done against friction is zero. The complete frictional work in two
dimensions is integrated over both the loading path and the length
of the fault, l:
W
fric
ZZ
s
N
u
i
; lmsu
i
; ldu
i
dl: (6)
Frictional work depends on the coefcient of friction, which in
our slip-weakening model changes with increasing displacement.
Until the slip-weakening distance is reached, the m in the frictional
work termis a function of displacement. After a displacement equal
to L has been achieved, m is equal to the dynamic friction value. W
fric
is similar to the E
F
notation used by Kanamori and Heaton (2000) to
describe the frictional energy loss except that our frictional work
integrates over the decrease in shear stress as slip increases from
zero to L. E
F
only considers the frictional work done under the
dynamic shear stress. Consequently, W
fric
is equivalent to E
F
E
G
of
Kanamori and Heaton (2000) notation, where E
G
represents the
energy consumed along the fault as slip increases to the critical slip
distance. With our delineation of work terms, the frictional work
produced by the rupture is expected to depend on the slip-weak-
ening length of the fault, while the seismological frictional work
delineation does not.
The work done in the creation of new surfaces through the
nucleation and propagation of off-fault tensile cracks is a function
of the surface energy of a crack, G
c
, and the total area of new
fracture surface created, S.
E = 20 GPa;
=
0.2
u
n
= 2.81 cm; u
s
= 2.1 cm
2 m
Fig. 1. Schematic diagram of model setup. Black horizontal line is the 3.1 m long
critically stressed portion of the fault, which is otherwise shown as grey line.
H.M. Savage, M.L. Cooke / Journal of Structural Geology 32 (2010) 17321741 1734
W
prop
G
c
S (7)
The surface energy for rocks has been empirically estimated in
a variety of ways through laboratory (e.g. Wong, 1982; Cox and
Scholz, 1988) and eld analyses (e.g. Olgaard and Brace, 1983;
Chester et al., 2005). These estimations provide a wide range of
values. Analytically, the propagation energy can be directly calcu-
lated from the prescribed fracture toughness because the fractures
only grow when the stress intensity factor exceeds the fracture
toughness of the rock. The plane strain relationship between
energy release rate G
Ic
and K
Ic
provides a means to calculate W
prop
.
W
prop
SG
Ic
S
1 n
2
K
2
Ic
E
(8)
where E and n are the Youngs Modulus and Poissons ratio
respectively of the material. For the material property values
chosen for this modeling study (E 20 GPa; n 0.2;
K
Ic
2.5 MPa m
1/2
), the surface energy is 300 J/m
2
. This formulation
of W
prop
differs fromthat of the fracture energy parameter E
G
in the
Kanamori and Heaton (2000) formulation because we explicitly
solve for the surface energy involved in creating the off-fault frac-
tures in the damage zone.
The energy lost to ground shaking during an earthquake is
proportional to the shear stress drop during slip. Although our
quasi-static model cannot explicitly account for the energy that
would go into the seismic waves, we can approximate this term
based on the stress drop that occurs during slip-weakening. This
stress drop represents the release of some portion of the stored
elastic strain that accumulates as a fault is stressed, however stress
drop may represent only a small fraction of the total shear stress on
the fault. We approximate the seismic energy released during a slip
event as:
W
seis
ZZ
Ds
u
j
; l
u
j
; l
du
j
dl (9)
where Ds is shear stress drop during slip-weakening. W
seis
is
similar to the E
R
in the notation used by Kanamori and Heaton
(2000). We will investigate if faults with different roughnesses
release different amounts of seismic energy during rupture prop-
agation. Stress drop in a fully dynamic model maybe higher for the
given conditions than our model results, but the trends we see in
the seismic work for changing slip-weakening distance should be
applicable for a fully dynamic model.
The internal work of the fault system is measured strain energy
density. Timoshenko and Goodier (1951) derived the total strain
energy for a two-dimensional system to be the sum of stress
multiplied by strain over an innitely small increment of strain. The
integral of the strain energy over the entire two-dimensional body
yields:
W
int
ZZ
1
2
s
xx
3
xx
s
zz
3
zz
2s
xz
3
xz
dxdz (10)
Although the internal strain energy represents elastic (and
therefore recoverable) strain, the internal work term also repre-
sents the energy available for consumption by inelastic processes
such as the production of off-fault damage. Prior to any slip along
the fault, W
fric
W
prop
W
seis
0 so that W
int
equals external
work. We expect that W
int
will decrease with slip and damage
production along the modeled faults. In our study, the internal
work is sampled at observation points distributed throughout the
model. These observation points often fall within areas of
concentrated stresses near the tips of the off-fault damage. Near the
displacement discontinuity elements, the local stress singularity is
overestimated (i.e. r
1
instead of r
1/2
) so that sampling in these
regions produces articially high internal work. A more reliable
method of calculating W
int
is to subtract the other work terms from
W
ext
.
3. Results
3.1. Slip-weakening distance and off-fault fracture patterns
We compare the off-fault fracture patterns and slip proles
generated along faults with varying slip-weakening distances due
to the application of displacements at the boundaries of our model
(Fig. 2). The fracture patterns are shown for each fault at the time
when the rupture reaches the tip of the modeled fault. New frac-
tures develop perpendicular to the direction of greatest tensile
stress and at positions along the fault where local tensile stress
exceeds tensile strength. This occurs at the rupture front where an
element that slipped juxtaposes an element that has not; the high
slip gradient produces locally high tension on one side of the fault.
Fractures form mostly in the tensile quadrants of the rupture tip
and sub-perpendicular (approximately 7085
) to the fault. In
some cases, a few cracks develop in the overall compressive
quadrants of the fault when local tensions arise during rupture. As
new fracture tips continue to propagate, they grow in various
directions, highlighting the locally changing stress elds due to the
presence of other nearby fractures. The resulting sawtooth fracture
trace is element size dependent, however the average fracture
angle for a given crack is not. The lack of perfect symmetry in the
fracture pattern arises from slight asymmetry in boundary condi-
tions to prevent rigid body motion. Once a small degree of fracture
asymmetry is introduced, the asymmetry of the model is further
enhanced.
Faults with the longest L (Fig. 3A) show fracture patterns
resembling static friction fault models (Martel, 1997) where
fractures are located in the tensile quadrants at the fault tips.
Decreasing the slip-weakening length creates more fracturing
inboard of the fault tip, resembling fully dynamic models of
tensile fracture zones, with the fractured area forming a wedge
shape that tapers towards the center of the fault when new
fractures are allowed to continue to grow after the rupture has
reached the fault tip (Fig. 3B; Dalguer et al., 2003). This same
wedge-shaped pattern is predicted by analytical models that
predict zones of activated off-fault damage but do not explicitly
generate off-fault fractures (Andrews, 2005; Rice et al., 2005;
Templeton and Rice, 2008). However for our comparison, we
restrict our analysis of the fracture patterns to the iteration at
which the rupture reached the fault tip. Because the models
presented here, as well as other models of off-fault damage, do
not allow for the fault tip to propagate when the rupture reaches
the fault tip, the resulting off-fault damage pattern may not be
meaningful past this iteration.
The initial fault roughness (i.e. the slip-weakening distance)
has a large effect on resultant fracture density and clustering of
the damage zone (Fig. 2). Fracture density along the length of
the fault decreases as a function of increasing slip-weakening
distance (Fig. 4). The fractures form in clusters, with the number
of clusters decreasing with larger slip-weakening distances. The
clusters represent deviations in the slip prole from ellipticity.
For longer slip-weakening distances, the slip prole along the
fault maintains a mostly elliptical shape (Fig. 2). However, as L
decreases, small toes of slip extend from the rupture front that
represent the number of elements whose frictional strength is
falling from the static to the dynamic value in that iteration
(Fig. 4B; online supplementary material). The clusters form
between the element closest to the tip that has weakened to its
dynamic friction value and the elements along which friction is
H.M. Savage, M.L. Cooke / Journal of Structural Geology 32 (2010) 17321741 1735
falling. When the slip-weakening distance is small, fewer
elements are in transition between static and dynamic friction,
resulting in damage clusters that are closer together. The growth
length of fractures in a single iteration and spacing of the
clusters depend on element size, but not the pattern of
clustering.
Smaller slip-weakening distance means that elements reach
dynamic friction levels more quickly and the rupture reaches the
fault tip in fewer iterations for the smoother faults (Fig. 5). The
increased time spent slipping at higher coefcients of friction slows
the rupture speed on the rougher faults. The fault rupture propa-
gates on the order of centimeters per iteration, which can be
thought of as a unit of time. The growth of mode I fractures can
grow one element per iteration, which in these models is 5 cm.
Therefore the Mode II rupture speed is similar to the Mode I
propagation speed in these models.
An interesting point to note is that although aspects of the slip
patterns vary while the rupture is propagating, the nal slip
proles, as well as average and maximum slip values, are very
similar (Fig. 2). According to our models, faults with similar slip
0.4
0.0
0.4
P
o
s
i
t
i
o
n
(
m
)
0
A
B
2 4 6 8 10 12 14 16 18 20
Position (m)
L = 0.1 mm, Iteration 050
0.4
0.0
0.4
P
o
s
i
t
i
o
n
(
m
)
0 2 4 6 8 10 12 14 16 18 20
Position (m)
L = 1 mm, Iteration 070
Fig. 3. Faults allowed to continue fracturing after the rupture reached the fault tip show that faults with small slip-weakening distances resemble models of static friction (A) and
faults with longer slip-weakening distances resemble models of dynamic slip (B).
0.4
0.0
0.4
P
o
s
i
t
i
o
n
(
m
)
0 2 4 6 8 10 12 14 16 18 20
Position (m)
L = 1 mm, Iteration 029
0.0
0.5
S
l
i
p
(
c
m
)
0.4
0.0
0.4
P
o
s
i
t
i
o
n
(
m
)
0 2 4 6 8 10 12 14 16 18 20
Position (m)
L = 0.1 mm, Iteration 025
0.0
0.5
S
l
i
p
(
c
m
)
0.4
0.0
0.4
P
o
s
i
t
i
o
n
(
m
)
0 2 4 6 8 10 12 14 16 18 20
Position (m)
L = 0.01 mm, Iteration 024
0.0
0.5
S
l
i
p
(
c
m
)
0.4
0.0
0.4
P
o
s
i
t
i
o
n
(
m
)
0 2 4 6 8 10 12 14 16 18 20
Position (m)
L = 0.001 mm, Iteration 023
0.0
0.5
S
l
i
p
(
c
m
)
Fig. 2. Fracture patterns and slip proles generated along faults of varying slip-weakening distances: A) 0.001 mm, B) 0.01 mm, C) 0.1 mm, and D) 1 mm. Faults are compared at the
iteration at which the rupture front reaches the end of the fault. More off-fault damage occurs on smoother faults.
H.M. Savage, M.L. Cooke / Journal of Structural Geology 32 (2010) 17321741 1736
proles but different initial surface roughness would have very
different fracturing intensity within the damage zone, at least for
the initial stages of damage zone development. The maximum
width of the damage zone is similar between all models, and
damage zones are wider per unit of slip than predicted by fault
scaling models (Scholz, 2002).
3.2. Mechanical work
In an effort to assess the mechanical efciency of the fractured
fault zones, we analyze the mechanical work associated with
a variety of slip-weakening distances along the modeled faults. We
examine both the components of work when the rupture reaches
the tip of each fault, as well as how different components of the
total work change as the rupture propagates.
3.2.1. Change in work with rupture propagation
We investigate how the work against frictional resistance and
seismic energy release components of the work budget evolve
throughout the rupture process. Fig. 6 shows how frictional and
seismic work increase over the course of a slip event in two models
with differing slip-weakening lengths. At the onset, no slip has
A
B
2
3
4
5
6
5.0
10
15
20
25
0 0.5 1 1.5
Density
Clusters
L(mm)
N
u
m
b
e
r
o
f
C
l
u
s
t
e
r
s
D
e
n
s
i
t
y
(
f
r
a
c
t
u
r
e
s
/
m
)
0.4
0.0
0.4
P
o
s
i
t
i
o
n
(
m
)
0 2 4 6 8 10 12 14 16 18 20
Position (m)
Iteration 015
0
1
F
r
i
c
t
i
o
n
0.0
0.5
S
l
i
p
(
c
m
)
0.4
0.0
0.4
P
o
s
i
t
i
o
n
(
m
)
0 2 4 6 8 10 12 14 16 18 20
Position (m)
Iteration 014
0
1
F
r
i
c
t
i
o
n
0.0
0.5
S
l
i
p
(
c
m
)
Toe
Fig. 4. A) Along-strike fracture density decreases as a function of slip-weakening distance. This trend results in less continuous damage along the fault, so that there are fewer
clusters of fractures. B) Clustering is a function of slip-weakening distance because it represents the length of the transition zone from locked to slipping, and therefore the region
that is subjected to local tensile stress. This region can be seen as a toe on the slip prole in Iteration 15.
H.M. Savage, M.L. Cooke / Journal of Structural Geology 32 (2010) 17321741 1737
occurred and therefore no work has gone into friction or seismic
energy; all of the external work is expressed as internal work
within the material surrounding the fault. As slip progresses, W
fric
and W
seis
increase at the expense of internal work. The smoother
fault produces greater seismic work and greater frictional work at
each iteration of fault slip (Fig. 6). The seismic work and frictional
work are generated not just on the main fault but also along the off-
fault fractures as they slide. This accounts for irregularities along
the curves in Fig. 6. Because the rougher fault (L 1 mm) takes 4
more iterations than the smoother fault (L 0.01 mm) to reach the
fault tip, the frictional work on the rough fault slightly exceeds that
of the smooth fault once the rupture is at the tip of both modeled
faults.
3.2.2. Sensitivity of work components to slip-weakening distance
The total work of fracture propagation scales with the length/
area of new fracture surface produced (Eq. (9)). Faults with longer
slip-weakening distance produce less off-fault damage and
consume less work of fracture propagation (Fig. 7). The anoma-
lously large damage produced by the model with L 1 mmis due to
tensile fractures that develop within the compression quadrant of
this fault.
The work against frictional resistance to sliding and the seismic
radiated energy are plotted for faults when the rupture has just
reached the ends of the modeled faults. At this point, each of the
models has similar slip prole. The frictional work increases
modestly with increasing slip-weakening distance. Although faults
with L 1 mm have greater off-fault damage than faults with
L 2 mm, the L 1 mm faults require less frictional work. Fric-
tional work depends on both slip and friction coefcient. Because
the fault with longer slip-weakening distance slips while the fric-
tion coefcient is higher than the fault with shorter L, the fault with
longer L requires greater frictional work. Rough faults may require
slightly greater frictional work to slip than smoother faults;
however these differences may be small and impossible to discern
in the eld. In addition to slip and friction coefcient, frictional
work also depends on normal stress (Eq. (6)). While the normal
stress is the same for the faults modeled, it may differ signicantly
for faults in the eld.
The seismic radiated energy calculated from shear stress drops,
along both the primary fault and along the off-fault damage,
decreases sharply with increasing slip-weakening distance
(Fig. 7b). With smooth faults, the coefcient of friction and subse-
quently the shear stress can have greater drop between iterations
than along faults with long slip-weakening distance. The larger
shear stress drops between rupture propagation iterations
produces larger seismic radiated energy. Augmenting this trend is
the tendency for faults with smaller L to produce greater damage
(Fig. 7c). The development of off-fault fractures provides a means to
transfer stored internal work to seismic radiated energy. Together
these processes imply that rupture along smoother faults should
produce more shaking than ruptures along rougher faults.
3.3. Second generation of damage
Our model results suggest that fault surface evolution should
be accompanied with greater production of damage, slightly
lesser frictional work and signicantly greater seismic energy
release. These trends neglect the inuence of pre-existing off-
fault damage. To begin to address this issue we investigate the
propagation of rupture and development of damage along a fault
that already has some off-fault damage. We use the fracture
pattern from the L 1.5 mm fault for the initial damage pattern
and reapply the boundary conditions. Before we allow the
rupture to start along the central weak patch, we rst apply the
boundary displacements and let the faults slip and the cracks
grow to their full extent under the applied loading. Once this is
complete, the friction on all faults is brought to the static friction
levels and the friction along the central patch lowered to induce
rupture.
Table 1 presents the number of iterations to reach the modeled
fault tip and work values for the rst and second rupture episodes
along the L 1.5 mm fault. The number of slip iterations for the
rupture to reach the fault tip increases when the fault is anked by
existing fractures. The pre-existing damage deforms as the rupture
propagates along the fault, slowing down the rupture. Fracture
clusters fromthe rst episode are made longer but fewnewclusters
form (Fig. 8; online supplement). The growth of fractures per unit
slip also increases; whereas the total slip on the fault doubles when
adding the two episodes, the length of the longest fractures
quadruples. The total length of damage increases, which is reected
in the near doubling of the work of fracture propagation. The
presence of off-fault damage also increases the seismic radiated
Fig. 5. Iterations to the modeled fault tip for faults with differing slip-weakening
distance, L. Rupture propagates more slowly on rougher faults and takes more itera-
tions to reach the fault tip.
0.8
0.6
0.4
0.2
0.0
w
o
r
k
(
M
J
)
iterations of slip
10 20 0 30
W
seis
W
fric
W
seis
W
fric
L=1 mm L=0.1 mm
Fig. 6. Increase of work against frictional resistance to sliding (W
fric
) and work of
seismic energy release (W
seis
) during iterations of slip along two faults with different
slip-weakening distance, L. At each iteration of rupture propagation, W
fric
and W
seis
for
the smoother fault exceed the work of the rough fault. Once rupture reaches both
modeled fault tips, the ctional work is slightly greater for the rougher fault.
H.M. Savage, M.L. Cooke / Journal of Structural Geology 32 (2010) 17321741 1738
energy by a small amount compared to the rst episode of slip. The
internal work calculated by subtracting all the other work terms
from the external work decreases slightly with the second episode
of slip.
4. Discussion
4.1. Fracture patterns
The qualitative analysis of the off-fault damage pattern resulting
from fault slip has some interesting applications for eld
observations. The models show that the initial surface roughness
on the fault will have a considerable effect on the continuity of the
damage zone along strike. Faults with smaller slip-weakening
distances would have a more continuous damage zone, whereas
faults with large slip-weakening distances would have a less dense
network of fractures. However, we should note that a fault that has
greater geometric roughness (non-planarity) would concentrate
damage at asperities. The formation of off-fault damage diminishes
slip on the adjacent fault patch from an expected elliptical slip
distribution. The absolute magnitude of the critical slip distance is
difcult to assess for tectonic faults. In the laboratory, where
surfaces in general are smooth and gouge zone thicknesses in the
millimeter range, critical slip distance values are generally
measured to be 10 s of microns with the expectation of scaling up to
earthquake faults. Seismic slip inversions have estimated the crit-
ical slip distance on the order of 10100 cm. Although critical slip
distance should generally decrease as fault motion wears down
asperities (Sagy et al., 2007), the difculty of measuring the
evolution of this parameter in nature hampers our understanding
of these processes.
a b c
Fig. 7. Sensitivity of (a) frictional work, (b) and seismic energy and (c) propagation energy with slip-weakening distance along the modeled faults. We calculate the error of the
work values by examining the difference in work between these models and models with twice the element size. Open symbols denote models that produced tensional cracks
within the contractional quadrants of the fault. Frictional work increases slightly with L whereas seismic work decreases dramatically with increasing slip-weakening distance. The
amount of damage decreases with slip-weakening distance.
Table 1
Pre-existing fractures slow the speed of rupture and alter the work budget. More
work goes into seismic radiation and fracture propagation, whereas damage
decreases frictional work.
#Iterations
to fault tip
W
seis
(MJ) W
fric
(MJ) W
prop
(MJ) W
int
(MJ)
L 1.5 mm rst 34 0.10 0.86 0.0049 41.34
L 1.5 mm second 48 0.14 0.84 0.0075 41.31
1
0
1
P
o
s
i
t
i
o
n
(
m
)
0 2 4 6 8 10 12 14 16 18 20
Position (m)
L = 1.5 mm, Iteration 049
0.0
0.5
S
l
i
p
(
c
m
)
B
1
0
1
P
o
s
i
t
i
o
n
(
m
)
0 2 4 6 8 10 12 14 16 18 20
Position (m)
L = 1.5 mm, Iteration 034
0.0
0.5
S
l
i
p
(
c
m
)
A
Fig. 8. Fault with pre-existing damage subjected to second episode of slip shows enhanced damage in areas where damage had already localized, but little new fracturing.
H.M. Savage, M.L. Cooke / Journal of Structural Geology 32 (2010) 17321741 1739
The models demonstrating the second generation of slip imply
that faults in the eld with more than one slip episode will have
high density, clustered fracture patterns, due to the ease of prop-
agating a fracture relative to forming a new one. This results in
a larger ratio between damage zone width and slip than a fault that
has one slip episode where total slip equaled the sum of the two
smaller episodes. However, using damage zone width per unit of
slip to estimate number of slip events would need to be limited to
comparing faults within the same eld site. Additionally, signicant
interseismic healing of fractures through processes such as mineral
precipitation would mitigate the effect of pre-existing damage on
subsequent ruptures.
4.2. Mechanical work analysis
Within the models of this study, we hold the external work
constant; no additional work is added to the system during the
propagation of the rupture. Because the left hand side of Eq. (3) is
constant during slip, the changes in work that we observe in the
models reects the transfer of work from internal stored work to
the non-conservative frictional heating, seismic radiated energy
and propagation energy. Of these non-conservative terms, the
greatest work is consumed in frictional heating along the primary
fault and its damage structures. In contrast, the work of fracture
propagation is several orders of magnitude smaller than the other
work terms, however processes such as rock pulverization could
require much more propagation energy than the cracks formed in
these models (Wilson et al., 2005). The total work of the system is
about 42.3 MJ so by far the greatest component of work is the
internal work stored within the system. Within these models only
2% of the internal work is converted to non-conservative work
terms. Our models showthat internal work decreases slightly upon
the second episode of slip along the fault. This suggests that further
ruptures along the fault could continue to transfer stored energy
within the host rock into frictional heating, seismic radiated energy
and the creation of new fault surfaces. With successive rupture
events we expect the slip-weakening distance along the fault to
generally decrease. Smoother faults are more effective than rough
faults at transferring work from internal work to the non-conser-
vative work terms.
5. Conclusions
Two-dimensional linear elastic models of frictional faults
suggest that the frictional slip-weakening distance (L) has signi-
cant effects on the tensile, off-fault damage pattern in a fault zone.
Faults with smaller slip-weakening distance have more continuous
along-strike damage whereas faults with large slip-weakening
distances concentrate fracturing at fault tips. Fractures form along
the fault in small clusters, with the number of clusters increasing as
function of L, thereby making along-strike fracturing more
continuous. Slip-weakening distance also affects how work is
consumed within the fault zone, with longer slip-weakening
distance resulting in more work done against friction and less
radiated seismic energy. Pre-existing damage further localizes
fracturing and consumes more internal work. Because initial
damage may be related to fault roughness, this implies that incip-
ient fault roughness controls along-strike fracture density even
after many episodes of slip.
Acknowledgements
We would like to thank two anonymous reviewers whose
insights greatly improved the paper. This work was supported by
NSF Grant EAR-0349070 to Michele Cooke.
Appendix. Supplemental data
Supplemental data associated with this article can be found in
online version at doi:10.1016/j.jsg.2009.08.014.
References
Abercrombie, R.E., Rice, J.R., 2005. Can observations of earthquake scaling constrain
slip-weakening? Geophysical Journal International 162. doi:10.1111/j.1365-
246X.2005.02579.x.
Andrews, D.J., 2005. Rupture dynamics with energy loss outside the slip zone.
Journal of Geophysical Research 110. doi:10.1029/2004JB003191.
Broberg, K.B., 1999. Cracks and Fractures. Academic Press, San Diego.
Brock, W.G., Engelder, J.T., 1977. Deformation associated with the movement of the
Muddy mountain overthrust in the Bufngton window, southeastern Nevada.
Bulletin of the Geological Society of America 88, 16671677.
Burgmann, R., Pollard, D., Martel, S., 1994. Slip distributions on faults: effects of
stress gradients, inelastic deformation, heterogeneous host-rock stiffness, and
fault interaction. Journal of Structural Geology 12, 16751690.
Caine, J.S., Evans, J.P., Forster, C.B., 1996. Fault zone architecture and permeability
structure. Geology 24, 10251028.
Chester, F.M., Logan, J.M., 1986. Implications for mechanical properties of brittle
faults from observations of the Punchbowl Fault, California. Pure and Applied
Geophysics 124, 77106.
Chester, J.S., Chester, F.M., Kronenberg, A.K., 2005. Fracture surface energy of the
Punchbowl fault, San Andreas system. Nature 437. doi:10.1038/nature03942.
Cooke, M.L., 1997. Fracture localization along faults with spatially varying friction.
Journal of Geophysical Research 102, 22,425422,434.
Cooke, M.L., Murphy, S., 2004. Assessing the work budget and efciency of fault
systems using mechanical models. Journal of Geophysical Research 109.
doi:10.1029/2004JB002968.
Cowie, P.A., Shipton, Z.K., 1998. Fault tip displacement gradients and process zone
dimensions. Journal of Structural Geology 20, 983997.
Cox, S.J.D., Scholz, C.H., 1988. Rupture initiation in shear fracture of rocks: an
experimental study. Journal of Geophysical Research 93 (B4), 33073320.
Crouch, S.L., Stareld, A.M., 1990. Boundary Element Methods in Solid Mechanics.
Unwin Hyman, Boston, Mass.
Dalguer, L.A., Irikura, K., Riera, J.D., 2003. Simulation of tensile crack generation by
three-dimensional dynamic shear rupture propagation during an earthquake.
Journal of Geophysical Research 108. doi:10.1029/2001JB001738.
Del Castello, M., Cooke, M.L., 2007. The underthrusting-accretion cycle: work
budget as revealed by the boundary element method. Journal of Geophysical
Research 112, B12404. doi:10.1029/2007JB004997.
Dieterich, J.H., 1979. Modeling of rock friction: 1. Experimental results and consti-
tutive equations. Journal of Geophysical Research 84, 21612168.
Faulkner, D.R., Mitchell, T.M., Healy, D., Heap, M.J., 2006. Slip on weak faults by the
rotation of regional stress in the fracture damage zone. Nature 444. doi:10.1038/
nature05353.
Ismat, Z., 2008 Energy budget during fold tightening of a multilayer fold. Journal of
Structural Geology 31. doi:10.1016/j.jsg.2008.10.006.
Kanamori, H., Heaton, T.H., 2000. Microscopic and macroscopic physics of earth-
quakes. In: Rundle, J., Turcotte, D., Klein, W. (Eds.), Geocomplexity and the
Physics of Earthquakes. Geophysical Monograph, vol. 20. American Geophysical
Union, Washington D.C., pp. 127141.
Kim, Y.-S., Peacock, D.C.P., Sanderson, D.J., 2004. Fault damage zones. Journal of
Structural Geology 26, 503517.
Maerten, L., Willemse, E.J.M., Pollard, D.D., Rawnsley, K., 1999. Slip distributions on
intersection normal faults. Journal of Structural Geology 21, 259271.
Manighetti, I., King, G., Sammis, C.G., 2004. The role of off-fault damage in the
evolution of normal faults. Earth and Planetary Science Letters 217. doi:10.1016/
S0012-821X(03)00601-0.
Marone, C., Kilgore, B., 1993. Scaling of the critical slip distance for seismic faulting
within shear strain in fault zones. Nature 362. doi:10.1038/362618a0.
Marshall, S.T., Cooke, M.L., Owen, S.E., 2008. Effects of non-planar fault topology and
mechanical interaction on fault slip distributions in the Ventura basin, CA.
Bulletin of the Seismological Society of America 98, 11131127. doi:10.1785/
0120070159.
Martel, S.J., 1997. Effects of cohesivezones onsmall faultsandimplications for secondary
fracturing and fault trace geometry. Journal of Structural Geology 19, 835847.
Mitra, G., Boyer, S.E., 1986. Energy balance and deformation mechanisms of
duplexes. Journal of Structural Geology 8, 291304.
Okubo, C.H., Schultz, R.A., 2005. Evolution of damage zone geometry and intensity
in porous sandstone: insight gained from strain energy density. Journal of the
Geological Society of London 162, 939949.
Olgaard, D.L., Brace, W.F., 1983. The microstructure of gouge from a mining-induced
seismic shear zone. International Journal of Rock Mechanics and Mining
Sciences and Geomechanics Abstracts 20, 1119.
Ohnaka, M., 2000. A physical scaling relation between the size of an earthquake and
its nucleation zone size. Pure and Applied Geophysics 157, 22592282.
Rice, J.R., Sammis, C.G., Parsons, R., 2005. Off-fault secondary failure induced by
a dynamic slippulse. Bulletinof theSeismological Societyof America 95, 109134.
Rispoli, R., 1981. Stress elds around strike-slip faults inferred from stylolites and
tension gashes. Tectonophysics 75, T29T36.
H.M. Savage, M.L. Cooke / Journal of Structural Geology 32 (2010) 17321741 1740
Ruina, A., 1983. Slip instability and state variable friction laws. Journal of
Geophysical Research 88, 1035910370.
Sagy, A., Brodsky, E.E., Axen, G.J., 2007. Evolution of fault surface roughness with
slip. Geology 35, 283286.
Savage, H.M., Cooke, M.L., 2004. An investigation into the role of fault interaction on
fold pattern. Journal of Structural Geology 26, 905917.
Scholz, C.H., 2002. The Mechanics of Earthquakes and Faulting. Cambridge
University Press, New York.
Shipton, Z.K., Cowie, P.A., 2003. Aconceptual model for theoriginof fault damagezone
structures inhigh-porositysandstone. Journal of Structural Geology 25, 333344.
Sibson, R.H., 1977. Fault rocks and fault mechanisms. Journal of the Geological
Society 133, 191213.
Stein, R.S., 1999. The role of stress transfer in earthquake occurrence. Nature 402,
605609.
Templeton, E.L., Rice, J.R., 2008. Off-fault plasticity and earthquake rupture
dynamics: 1. Dry materials or neglect of uid pressure changes. Journal of
Geophysical Research 113. doi:10.1029/2007JB005529.
Timoshenko, S.P., Goodier, J.N., 1951. Theory of Elasticity. McGraw-Hill, New York.
Willemse, E.J.M., Pollard, D.D., Aydin, A., 1996. Three-dimensional analyses of slip
distributions on normal fault arrays with consequences for fault scaling. Journal
of Structural Geology 18 295209.
Wilson, B., Dewars, T., Reches, Z., Brune, J., 2005. Particle size and energetics of
gouge in from earthquake rupture zone. Nature 434, 749752.
Wong, T.-f., 1982. Shear fracture energy of westerly granite from postfailure
behavior. Journal of Geophysical Research 87, 9901000.
Yamashita, T., 2000. Generation of microcracks by dynamic shear rupture and its
effects on rupture growth and elastic wave radiation. Geophysical Journal
International 143, 395406.
H.M. Savage, M.L. Cooke / Journal of Structural Geology 32 (2010) 17321741 1741
Simulating brittle fault evolution from networks of pre-existing joints
within crystalline rock
Heather Moir
a,
*
, Rebecca J. Lunn
a
, Zoe K. Shipton
b
, James D. Kirkpatrick
b
a
Department of Civil Engineering, University of Strathclyde, Glasgow, Scotland, UK
b
Department of Geographical and Earth Sciences, University of Glasgow, Glasgow, Scotland, UK
a r t i c l e i n f o
Article history:
Received 9 February 2009
Received in revised form
17 August 2009
Accepted 20 August 2009
Available online 23 September 2009
Keywords:
Numerical modelling
Fault-zone evolution
a b s t r a c t
Many faults grow by linkage of smaller structures, and damage zones around faults may arise as a result
of this linkage process. In this paper we present the rst numerical simulations of the temporal and
spatial evolution of fault linkage structures from more than 20 pre-existing joints, the initial positions of
which are based on eld observation. We show how the constantly evolving geometry and local stress
eld within this network of joints contribute to the fracture pattern. Markedly different fault-zone trace
geometries are predicted when the joints are at different angles to the maximum compressive far-eld
stress ranging from evolving smooth linear structures to complex stepped fault-zone trace geometries.
We show that evolution of the complex fault-zone geometry is governed by: (1) the strong local vari-
ations in the stress eld due to complex interactions between neighbouring joints; and (2) the orien-
tation of the initial joint pattern with respect to the far-eld stress.
2009 Elsevier Ltd. All rights reserved.
1. Introduction
Several authors have proposed that faults evolve under imposed
stress by the linkage of pre-existing structures (Segall and Pollard,
1983; Martel, 1990; Bergbauer and Martel, 1999; Pachell et al.,
2003). The pre-existing structures from which faults nucleate are
commonly open or mineral-lled joints that are weaker than the
surrounding rock (Segall and Pollard, 1983; Bergbauer and Martel,
1999; Pachell and Evans, 2002). When pre-existing features expe-
rience compressive loading, stress concentrations (both tensile and
shear) develop around the tip of the feature. Shearing of these pre-
existing features often results in the formation of secondary frac-
tures at (or near) the tip of the feature. These secondary fractures
have different names including: tail cracks/fractures (Cruikshank
and Aydin, 1994; Willemse et al., 1997), splay fractures (Pachell and
Evans, 2002; Myers and Aydin, 2004), horsetail fractures (Granier,
1985; Kim et al., 2004) and wing cracks (Crider and Peacock, 2004).
In this paper all fractures (tension or shear) associated with faulting
at (or near) the tip of a pre-existing feature are termed wing cracks.
Conceptual models of fault evolution through the development of
wing cracks (Martel, 1990; Martel and Boger, 1998) are supported
by eld observations of wing crack evolution from single joints or
faults (Kattenhorn and Marshall, 2006; Joussineau et al., 2007) and
by observations of linking fractures that have developed between
pairs of isolated faults (Peacock and Sanderson, 1995; Kim et al.,
2004). Wing cracks developing with shear displacement are also
commonly observed in eld data, for instance the sheared dyke in
Fig. 1a and the large splay faults in Kirkpatrick et al. (2008, their
Fig. 8).
In this paper, we focus on fault-zone development in crystalline
rocks. Natural exposures of fault-zone traces within crystalline
rocks can have many geometries, from smooth, approximately
planar features (Fig. 1a) where faults appear to develop along strike,
to complex stepped structures (Fig. 1b) where adjacent faults are
linked at stepovers, or a combination of both (Fig. 1c). Key questions
are: what governs the geometry of the evolving fault-zones? How
are fractures within the fault-zone linked?
A series of numerical models simulating fault growth, support
these conceptual models for fault-zone evolution. These models
have simulated the evolution of wing cracks from the tips of pre-
existing structures (Shen and Stephansson, 1993; Bu rgmann et al.,
1994; Kattenhorn et al., 2000; Willson et al., 2007) or the linkage of
pairs of faults with extensional and contractional geometries (Du
and Aydin, 1995; Bremaecker and Ferris, 2004; Lunn et al., 2008).
These simple, two-dimensional (2D) models have enabled predic-
tion of the orientation of linkage fractures and their mode of failure,
for a single fracture or pair of fractures in an ideal homogeneous
medium. However, these simulations, derived from one or two
fractures, are not sufcient to understand the range of complex
geometries observed in the eld (Fig. 1). Within this paper we
* Corresponding author at: Department of Civil Engineering, University of
Strathclyde, 16 Richmond Street, Glasgow G1 1XQ, Scotland, UK.
E-mail address: heather.moir@strath.ac.uk (H. Moir).
Contents lists available at ScienceDirect
Journal of Structural Geology
j ournal homepage: www. el sevi er. com/ l ocat e/ j sg
0191-8141/$ see front matter 2009 Elsevier Ltd. All rights reserved.
doi:10.1016/j.jsg.2009.08.016
Journal of Structural Geology 32 (2010) 17421753
extend current knowledge by simulating fault-zone evolution in
granite from a network of more than 20 joints. We show that
evolution of the resulting fault-zone geometry is governed by:
(1) the strong local variations in the stress eld due to complex
interactions between neighbouring joints; and (2) the orientation
of the initial joint pattern with respect to the far-eld stress.
2. Methodology
We use the computer code Modelling Of Permeability Evolution
in the Damage Zone surrounding faults (MOPEDZ) (Willson et al.,
2007) to simulate spatial and temporal evolution of complex
patterns of linking fractures. MOPEDZ was developed using the
commercially available nite-element software COMSOL which is
called from within the MATLAB code. The COMSOL nite-element
routines assume plane strain during the simulations. MOPEDZ is
a 2D nite-element model which solves Naviers equation in
a series of quasi steady-states and uses a combined Mohr Coulomb
and tensile failure criteria. Elements within the nite-element
mesh are either intact host rock or fractured host rock. Elements
which contain fractures (including the initial joints) are repre-
sented by lower effective material values (10% of the host) for
Youngs modulus, Poissons ratio and material strength, in a similar
approach toTang (1997). Representing the accumulation of damage
within each element by altering that elements material properties
is consistent with other damage mechanics models (Jing, 2003).
The initial conguration for all MOPEDZ simulations is similar to
that illustrated in Fig. 2 with the host rock (granodiorite) having the
Fig. 1. Field examples of mapped sections from fault-zones. (a) A segment of the outcrop map from NE of Neves lake in the Italian Alps showing a section of fault-zone with smooth
planar features (Pennacchioni and Mancktelow, 2007). (b) A segment of the outcrop map from the Waterfall region in the Sierra Nevada, California (Martel, 1990). (c) Map of
fractures in an exposure of the Lake Edison granodiorite in the Bear Creek region in the Sierra Nevada, California, UTM coordinates are: 0333075 4136569.
H. Moir et al. / Journal of Structural Geology 32 (2010) 17421753 1743
properties listedinTable 1andanyelements containingpre-existing
joints having reduced material properties (10% of the host) (Willson
et al., 2007). The simulated maximum compressive far-eld stress
direction, s
1
, is parallel totheyaxis (i.e. top-to-bottominall MOPEDZ
gures) and the minimum, s
3
, is parallel to the x axis (i.e. left-to-
right inall MOPEDZgures) (Fig. 2). Note that inthe eldboths
1
and
s
3
are horizontal. Initially all boundaries are displaced inward
holding s
1
2s
3
, (in compression) however following the rst
failure (either Mohr Coulomb or tensile) the s
3
boundaries are held
constant andfromthis point ononly the s
1
boundaries are displaced
towards each other, i.e. s
1
progressively increasing with s
3
held
constant. All simulations presented here are in compression.
Throughout this paper s
1
and s
3
refer to the far-eld stress imposed
by the boundaries of the nite-element model and s
1
Local
and s
3
Local
refer to the local stress eld around damaged cells. All simulations
use square nite-elements; the number of mesh elements in the
simulations presentedinthis paper varies from6400to136,500; the
size of each cell is approximately 13 mm
2
.
As an element fails (in either shear or tension) its material
properties are altered. Although the rst failures are triggered by
displacement of the boundaries, the alteration of the material
properties of those failed cells causes a change in both the direction
and magnitude of s
1
Local
and s
3
Local
(Lunn et al., 2008). This alteration
of the local stress may be sufcient to trigger additional failures
without any further displacement of the model boundaries. These
subsequent failures can be adjacent to previous failures, i.e. rep-
resenting the lengthening of a macroscopic fracture, or they can
occur in locations that are disconnected from any previous failure,
or they may be further fracturing of the same element or any
combination of these. MOPEDZ iteratively reduces the values of the
material properties as elements are predicted to fail; this reects
increasing damage to the host rock (host rock elements containing
pre-existing joints start with the lowest values, 10% of host rock).
Each element can fail up to a maximum of six times (resulting in
a reduction of strength, Youngs modulus and Poissons ratio) in
a geometric sequence (Willson et al., 2007) until they reach the
lowest value permitted (equivalent to those elements containing
the initial joints). We emphasise that each element in the mesh
may represent, at a sub-element scale, any number of micro or
macroscopic failures in the eld (in these simulations a sub-
element scale is smaller than 13 mm
2
).
Fig. 2. Typical initial setup showing the orientation of s
1
and s
3
(simulated far-eld
stress). Gray area is host rock, black is host rock containing joints (n.b. the pixellated
nature of the pre-existing joints is a product of the model). The model boundaries (red)
are under displacement control, following the initial failure only the top and bottom
boundaries are displaced. To avoid consideration of structures generated at the
boundary in the large simulations, only the central window (within the blue box) was
presented in the results. For all small simulations no window was taken and all results
within the red model boundaries are presented. The number of mesh elements varies
from 6400 to 136,500 depending on the size of the simulation. (For interpretation of
the references to colour in this gure legend, the reader is referred to the web version
of this article.)
Table 1
MOPEDZ simulation parameters for brittle rock.
Rock property Value Reference
Host Rock Youngs
modulus
60 GPa Martin (1997)
Turcotte and Schubert (2002)
Host rock Poissons
ratio
0.2 Turcotte and Schubert (2002)
Youngs modulus of
fractured element
1.2 GPa Segall and Pollard (1983)
Poissons ratio of
fractured element
0.02
Co (shear strength) 130 MPa Martin (1997)
m (coefcient of friction) 0.6 Byerlee (1967)
To (tensile strength) 10 MPa Martin (1997)
Number of cells permitted
to fail in any one step of
the MOPEDZ code
6
Where relevant the right hand column contains the reference from which the value
of the mechanical property was derived.
Fig. 3. Cartoon showing the evolution of restraining and releasing bends for a pair of
overlapping and under-lapping pre-existing joints with either contractional or
extensional relationship, (i) is the initial orientation of the joints, and sequences (iiiv)
show evolution of the predicted structure. All slipped joints are left lateral.
H. Moir et al. / Journal of Structural Geology 32 (2010) 17421753 1744
Fig. 4. (a) Small section from map shown in Fig. 1c. (b) Mapped joints oriented at 60
to s
1
. Linkages at Location A are in a different
orientation to the rest of the simulation (see Fig. 7). At Location B overlapping joints in an extensional geometry link in a similar way as those in Fig. 3c. At Location C under-lapping
joints in a contractional geometry link in a similar way as those shown in Fig. 3b. At Location D two closely spaced joints in a contractional geometry link with a third more distant
joint which is in an extensional geometry. At Location E joints under-lapping and in an extensional geometry link in a similar way as those shown in Fig. 3a. All slipped joints are left
lateral.
H. Moir et al. / Journal of Structural Geology 32 (2010) 17421753 1745
Once a steady-state solution has been achieved for a given
boundary displacement, the top and bottom boundaries undergo
further displacement towards each other and the whole solution
process is repeated. Duringanyoneiterationof the code, onlya small
number of elements (<6inthesesimulations) are permittedtofail to
ensure stability of the model solution and provide an estimation of
the temporal propagation of the fractures.
Earlier research using MOPEDZ to examine failure from a single
joint (Willson et al., 2007) shows that fracture-trace geometries are
not sensitivetotheinitial mechanical properties of thehost rock(Table
1), e.g. if Youngsmodulus of thecrystallinehost rockis lower, thesame
tracegeometries areformedbut at lower values of the displacement of
the boundaries. Fracture-trace geometries are principally determined
bylocal variations intheYoungs modulus(i.e. damagedelements), the
orientation of the pre-existing joint to the far-eld maximum
compressive stress and the ratio of s
1
s
3
. In simulations where
s
1
[s
3
, failure was predominately in tension, for those where s
1
is
close to 2s
3
, simulated failure was predominately in shear. The mode
of failure results in different orientations of the evolving linkage
fractures (Lunn et al., 2008). MOPEDZ simulations of fault linkage
involving just two initial joints (Lunn et al., 2008) showed that frac-
ture-geometries develop in a predictable way summarised in Fig. 3.
Four initial stepover geometries were modelled: (a) under-lapping
extensional; (b) under-lapping contractional; (c) overlapping exten-
sional; and (d) overlapping contractional. The geometries of linkage
structures are governed by three key factors: (1) the ratio of s
1
s
3
; (2)
the initial relative positions of the joints, specically, contractional vs.
extensional geometries and overlapping vs. under-lapping joints; and
(3) the orientation of the most compressive principal stress direction
(s
1
) relative to the initial pair of joints.
In the following simulations we explore fault-zone evolution
through a large population of over 20 initial joints with gradually
increasing displacement of the s
1
boundaries of the model while s
3
boundaries are held constant. We start from an initial condition for
Fig. 6. Small simulation (80 80 nite-element mesh) with the joints in the same
orientation as Location E. (a) Surface plot of the principal stress prior to failure showing
interaction of the compressional quadrants of both joints. (b) Initial joint pattern
entered into MOPEDZ (overlap between j
1
and j
3
of 38 mm). (c) Damage plot of the
nal structure obtained.
Fig. 7. Small simulation with joints in the same orientation as Location A. (a) The
spatial and temporal evolution of the linking fractures predicted by MOPEDZ; black
represents elements of the of the nite-element mesh which contains fractures. (b)
Plots of the norm of the strain tensor which give a scalar representation of the
magnitude of the strain tensors; the darker the colour the higher the strain.
H. Moir et al. / Journal of Structural Geology 32 (2010) 17421753 1746
far-eld stresses of s
1
2s
3
. Within the initial joint population, all
four congurations of joint stepovers (as seen in Fig. 3) are locally
present. The simulations are conducted withthe joints at two angles
to s
1
(approximately 60
and 30
to s
1
. Stress around pre-existing joints which
intersect the boundaries may be inuenced by the proximity of the
boundary. For the simulations with>20 joints, to avoidconsideration
of anystructures whichmight result fromboundaryeffects, results are
displayed and discussed only for an internal area in the centre of the
nite-element mesh, the edges of which are dened in Fig. 4b. Small-
scale simulations are also presented, to investigate behaviour at
specic locations within the larger mesh. These smaller scale simu-
lationsdisplayresults over thewholemodel domain(i.e. nowindowis
taken). In each case, initial damage predictions for the small mesh
were compared with those within the larger mesh to conrm that
predicted structures were similar, and hence that boundary condi-
tions were not having a substantial effect on model results.
2.2. Presentation of simulation results
Simulation results are illustrated using three types of maps. (1)
Damage plots show the elements that have failed grey indicates
intact host rock and black indicates fractured host rock (since
individual elements may fail multiple times in shear and/or
tension, modes of failure are not shown). (2) Strain plots show the
Euclidean norm of the strain tensor, which is one of the methods of
representing the scalar magnitude of a strain tensor (Mathews and
Fink, 2004). Plots of the normof the strain tensors for each element
elucidate a more detailed structure than the damage plots, since
they also highlight elements which are under a high strain but that
have not yet failed. The norm of the strain tensor presented here
may not be appropriate for direct comparison with eld data since
we start all simulations from an initial condition of zero strain. (3)
Surface plots of the local principal stress show the spatial distri-
bution of s
1
Local
; these plots have the same colour scale to alloweasy
comparison between simulations. Note that surface plots of the
local principal stress were produced within COMSOL in which
compression is negative and tension positive (the opposite
convention is usually adopted within the geological literature).
3. Results
3.1. Development of linkage structures
The spatial and temporal evolution of the fracture develop-
ment and linkage predicted by MOPEDZ, for the joint pattern in
Fig. 4c, is shown in Fig. 5 as a damage plot. The initial joints are at
approximately 60
to s
1
(Fig. 5i). At rst wing cracks begin to
develop on some but not all joints (Fig. 5ii). The orientation of the
propagating wing cracks are similar to those predicted for iso-
lated joint pairs in Fig. 3. As the simulation continues (Fig. 5iiivi)
several types of linkage structures are observed which are similar
to those in Fig. 3; note that only six frames are shown from
a simulation consisting of 350 steps. At Location A the linking
structure is similar to that for an overlapping pair of joints in an
extensional orientation. At Location B the structure is similar to
that for under-lapping joints in a contractional geometry (the
joints under-lap by a single mesh element). At Location C the
Fig. 9. Surface plot of the principal stress prior to failure for individual pre-existing structures in the same orientation as those at Location D, Fig. 5.
H. Moir et al. / Journal of Structural Geology 32 (2010) 17421753 1748
structure is the same as that for under-lapping joints in an
extensional geometry. At Location D stepover geometries for both
overlapping extensional joints and under-lapping contractional
joints are represented, the predicted linkage structure is similar
to that for overlapping joints in an extensional orientation.
At Location E the linkage structure that develops is different to
that predicted for an isolated pair of under-lapping contractional
joints (Fig. 3b); at E initial failure occurs in the host rock between
the two joints as opposed to propagating from the joint tips.
Because processing time for the large simulation (<20 joints) was
56 days, a small (6400 element mesh) simulation investigated
local behaviour at Location E using three joints in the same relative
positions; both the physical size represented by each nite-element
and the boundary conditions (progressive displacement of the s
1
boundaries starting from an initial value of s
1
2s
3
) remain the
same as that for the large simulation in Fig. 5. The stress eld
(Fig. 6a) shows that the relative positions of the pre-existing joints
facilitates interaction of the compressional quadrants of the two
joints (j
1
and j
3
), which results in linkage due to shear failure. In the
simulation joints j
1
and j
3
shown in Fig. 6b overlap by 38 mm. If the
tip of joint j
3
is adjusted (by at least 38 mm either way) either to
clearly over- or under-lap j
1
, linkage geometries are similar to those
expected for over- or under-lapping contractional joints (Fig. 3b;
Fig. 12 in Segall and Pollard, 1980).
Three joints circled at Location A form two stepovers. The left
stepover is extensional and the right stepover is contractional. The
damage evolution (Fig. 7a) and strain evolution (Fig. 7b) in a small-
scale simulation (6400 elements) with the joints in the same
relative locations shows two types of linkage structure. Initially, the
pair on the left behaves as the extensional geometry in Fig. 3c.
However, as the fracture propagating from the middle of the upper
joint lengthens, it begins to interact with the joint on the right of
the gure, changing its orientation and eventually resulting in
linkage of the pair of joints on the right that are in a contractional
geometry (Fig. 3d).
Location B evolves a linkage structure similar to that predicted
in Fig. 3b for under-lapping contractional geometries (the joints
are displaced from being collinear by approximately 1 cm). At
Location D, despite the upper two joints being closer together
than those at Location B and in a more pronounced under-lapping
contractional geometry, a similar linkage structure does not
develop. Instead, a wing crack propagates from the much more
distant, extensionally-related joint below. This geometry illus-
trates the effect of neighbouring joints, which is investigated by
a small-scale simulation of the three joints at Location D. The
initial joint conguration is shown in Fig. 8a and the magnitude of
s
1
prior to failure is shown in Fig. 8b. Comparison of Fig. 8b with
the stress elds which are predicted for each of the three joints if
simulated separately (Fig. 9) shows that having all three joints
present reduces the magnitude and extent of the region of
compressional stress surrounding the interacting tips, and
increases the magnitude and extent of the tensional stress, most
notably between joints p
2
and p
3
. This explains the linkage
structure that develops between the more distant extensional pair
of joints, evident from the strain evolution (Fig. 8c). The resulting
fracture pattern is similar to that seen in a geometrically similar
conguration of starter joints in Segall and Pollard (1980, their
Fig. 12).
3.1.1. The inuence of joint length on linkage
The extent of the local stress perturbation associated with
a fault has previously been related to its trace length (e.g. Segall
and Pollard, 1980). To explore the effect of joint length on evolving
fault linkage, small-scale simulations were performed adjusting
the joint conguration at Location D. The lengths of the upper
Fig. 10. Spatial and temporal evolution of strain predicted by MOPEDZ with pre-
existing structures in the same relative positions as Location D but with (a) the initial
length of p
1
doubled (p
1.1
) and (b) initial length of p
1
and p
2
were both doubled (p
1.1
and p
1.2
respectively). Note that wing cracks only develop on p
3
when the upper wing
cracks reach the boundary; had it not done so the growth of wing cracks fromp
3
would
have been suppressed. (All faults are left lateral.)
H. Moir et al. / Journal of Structural Geology 32 (2010) 17421753 1749
joints (p
1
and p
2
) were increased, keeping the relative position of
the adjacent joint tips constant (Fig. 10). When the length of
either one (Fig. 10a) or both (Fig. 10b) of the upper joints are
doubled, joints p
1
and p
2
link in a manner similar to that for
under-lapping contractional joint pairs. Previously, in Fig. 8,
linkage was between the extensional pair of joints p
2
and p
3
. From
these simulations it is apparent that the length of the joints
affects the magnitude and extent of the stress eld at the joint
tips, thus enhancing or decreasing the likelihood of linkage
between a pair of joints.
3.1.2. Separation between joints
Small-scale simulations consisting of up to three joints were
used to explore the effect of joint separation; this was investigated
by increasing the distance between joints in the y direction. The
position of the lower joint (based on the conguration at Location
D) was adjusted until it was separated enough to allow the upper
two (p
1
and p
2
) to link (Fig. 11a). The stress perturbations due to
each joint were then explored by systematically removing each
fromthe simulation (leaving only two joints in the simulation). This
produced varying joint linkage geometries from joints linking
rapidly (Fig. 11b) through joints that fail to link, but show an
evolving structure within the linkage zone (Fig. 11c) to joints which
do not link (Fig. 11d). Note that in this nal simulation (Fig. 11d) the
wing crack which develops fromp
3
does so at a different angle (40
from s
1
) than for the original simulation (24
from s
1
). Simulations
illustrated in Fig. 11 show that the proximity of neighbouring joints
affects both the location and orientation of the linkage structures
that develop.
3.2. Exploring the effect of orientation of the regional stress eld
Simulations of linkage from isolated pairs of joints in Lunn
et al. (2008) showed that one of the key factors controlling the
fault-zone geometry was the orientation of the joints to s
1
. In the
case of joints at a low angle to s
1
wing cracks were found to
propagate back into the compressional quadrant, similar to
structures observed in the eld (Vermilye and Scholz, 1998). To
explore what effect the orientation of s
1
can have on fault-zone
evolution from a complex joint pattern, the joints in Fig. 4 were
oriented at an angle of approximately 30
to s
1
(Fig. 12i). The
predicted evolution of linkage structures through time is shown in
Figs. 12iivi; a visual comparison of Fig. 12vi and Fig. 5vi shows
the nal geometries to be very different. Critically, joint traces
that were approximately co-linear now progressively link up
along strike to form long smooth linear fault traces (e.g. Locations
Fig. 11. (a) Spatial and temporal evolution of strain predicted by MOPEDZ with joints in the same relative positions as Location D but with p
3
displaced away from the upper two
joints. Temporal evolution of strain predicted by MOPEDZ (b) if p
3
is removed, (c) if p
1
is removed and (d) if p
2
is removed. Note that the angle of the wing crack propagating from
the p
3
is 24
).
These factors that control the perturbations in the local stress
eld have an inuence on the different styles of linkage structures
which develop between initially co-linear joints.
Three key observations are apparent from a comparison of the
simulation results for pre-existing joints at 60
and 30
to s
1
:
For the simulation at 60
to s
1
(Fig. 5), the linking wing cracks
are at a much wider variety of angles (e.g. see Locations B and D
in Fig. 5) than those at 30
to s
1
(Fig. 12).
For the simulation at 60
to s
1
(Fig. 5) many joints develop
multiple wing cracks at individual joint tips and linkage
structures tend to exhibit more damage than those at 30
to s
1
(e.g. compare Location D in Figs. 5 and 12).
With structures at 30
to s
1
, approximately 60% of the joints
link along strike forming smooth linear features which span
the model domain (Fig. 12); these features do not form for the
simulations at 60
to s
1
,
4. Discussion
The simulations show three key ndings: (1) local spatial and
temporal variations in the stress eld have a signicant effect on
the location, orientation and timing of wing crack development,
resulting in signicantly different patterns than those predicted
from consideration of single fractures or pairs of fractures. (2) A
signicant difference in resulting fault-zone geometry is predicted
when s
1
is oriented at 30
to s1. Locations AE indicated on (i) correspond to those in Fig. 5i. Simulation was carried out
with the same initial conditions as that shown in Fig. 5. All faults are left lateral.
H. Moir et al. / Journal of Structural Geology 32 (2010) 17421753 1751
geometry and proximity of neighbouring features within the
network. Hence, predictions of the location and orientation of
fracture zones, such as those by Maerten et al. (2002) and Mick-
lethwaite and Cox (2004) may be improved by incorporating
simulation of the constantly evolving local stress.
4.2. Orientation of the maximum compressive far-eld stress
For a simulated s
1
at a high angle to the original joints (Fig. 5),
faults are principally formed by slip on pre-existing joints which
then grow in length as wing cracks evolve and link originally
discontinuous adjacent fault traces, resulting in a stepped fault-
zone geometry. A comparison of these results with Fig. 1 shows
them to be similar to the complex fault-zone geometry in Fig. 1b
(Martel, 1990); a large number of wing cracks and linkage struc-
tures are present, at a variety of angles, with few through-going
features. Our simulations suggest that the faults from the Waterfall
region of the Sierra Nevada (Fig. 1b) were formed under s
1
at
approximately 60
30
to s
1
link up approximately along strike, those at 60
and 60
304:8=Dt
p
3:23
(2)
The discrepancy may be explained by the dataset used to
establish Eq. (2), which is mostly based on younger, mechanically
weaker Tertiary age shales from the North Sea.
2.4. Material properties for reservoir layers
The reservoir layers are modelled as drained, poro-elastic
formations characterized by the shear modulus G, Poissons ratio n,
and bulk modulus of the grains K
s
. The drained bulk modulus of the
rock (K) is expressed as:
K 2G
1 n
31 2n
(3)
and Biots coefcient a as:
a 1 K=K
s
(4)
Table 1
Geometrical characteristic of two vertical sections through Brent fault used in two-
dimensional numerical analyses.
Section 1 Section 2
Average fault inclination (deg.) 54 47
Thickness of Brent Group
Footwall Hanging wall (m)
37183 204212
Thickness of Statfjord Formation
Footwall Hanging wall (m)
283319 268314
F. Cuisiat et al. / Journal of Structural Geology 32 (2010) 17541767 1757
The values of K, a, G and n are inferred from triaxial tests per-
formed in the laboratory on reservoir core samples. A nearly
continuous core of some 720 m through the Brent and Dunlin
Groups, and the Statfjord Formation was taken for rock mechanic
studies. Special care was taken to minimise the damage caused by
the coring process (Hettema et al., 2002). The interpreted values of
the elastic properties are summarized in Table 2.
2.5. Properties of faults
In this section, the main geological characteristics of the faults
are briey presented, together with the approach used for model-
ling faults in the numerical analyses.
2.5.1. Fault geometry
The fault is dened by its throw, the fault core thickness, and the
thickness of the damage zones on footwall and hanging wall sides
(Sperrevik et al., 2002). Fault throw, inferred from seismic data,
together with empirical correlations between damage zone thick-
nesses, fault core thickness and fault throw developed from outcrop
studies (Beach et al., 1997, 1999), were used to estimate geometrical
properties for the Brent Fault and the faults bounding the horst
structure. These data are tabulated inTable 3 for two vertical sections
throughthe Brent Fault and inTable 4 for one vertical sectionthrough
thehorst structure. Inaddition, the ShaleGouge Ratio(SGR) proposed
by Yieldinget al. (1997), whichis a measure of the percentage of shale
or clay in the slipped interval, is given in the tables.
Fig. 3. Finite Element model of Brent Fault Section 2, with detailed view of reservoir section and fault.
Fig. 4. Details of Finite Element model for horst structure Section 1. Note the faults Z0 (Snorre side) and Z5 (Statfjord side) dening the horst. Dp is the pressure reduction
applied in the model.
F. Cuisiat et al. / Journal of Structural Geology 32 (2010) 17541767 1758
Variation in the fault thickness is expected throughout the
sandstone shale sequence, with thicker damage zones observed
in competent shale units than in sandstone units (Sperrevik et al.,
2002). Signicant drag of the adjacent layers to the fault due to
shear deformation may also occur. However, in the absence of data
related to the drag, in particular of the relationship between drag
and fault throw, the drag component of deformation is not included
in the base case numerical model. As shown later in a simplied
parametric study, fault shear mobilisation is less when drag effects
are included, such that the base case is more conservative.
2.5.2. Fault mechanical properties
Shale Gouge Ratio (SGR) is used as an indication of the volume of
clay within the fault plane. This is a simplication of fault zone
complexity as SGR does not represent the detailed internal struc-
ture of a fault where clay smear, cataclasites, etc, might be present
together (Wibberley et al., 2008). Table 3 indicates that for the
Brent Group, the sandsand juxtaposition window is sealed by
a high clay concentration. Petrophysical and thin section analyses
on fault rock material from the Brent Group reservoir show most
fault rock may be classied as disaggregation zones/proto-cata-
clasites, indicating deformation at a relatively early stage in the
reservoirs burial history (Sverdrup and Bjrlykke, 1997; Fisher and
Knipe, 1998). In the Statfjord Formation, in view of the low SGR
values in the sandsand juxtaposition window (Table 3), it is
assumed that the Brent Fault zone, which is acting as a seal, consists
of a cataclastic fault rock, possibly with quartz cementation and
higher shear strength. In the horst structure, due to high SGR values
in the Statfjord Formation, low shear strength cannot be excluded
(Table 4).
The shear strength of the clay-rich fault core zone is dened
from the results of undrained triaxial tests carried out on core
samples from Viking Group and Burton Formation shale. Due to
limited data, a simple MohrCoulomb failure criterion is used. The
drained peak shear strength is found to vary between lower and
upper bounds as dened by the values of effective cohesion c
0
and
friction angle 4
0
given by (c
0
3 MPa, 4
0
24
) and (c
0
8 MPa,
4
0
23
m
2
=s
(6)
where: k Intrinsic fault core permeability (m
2
) K
u
Undrained bulk
modulus (MPa) K Drained bulk modulus (MPa) G Shear modulus
(MPa) m
f
Dynamic viscosity (MPa s) a Biot coefcient.
The Biot coefcient is given by Eq. (4). The undrained bulk
modulus can be inferred from the BiotGassmann relationship
(Mavko et al., 1998; Hettema and de Pater, 1998) written as:
K
u
K
a
2
K
s
K
f
nK
s
a nK
f
(7)
where K
s
is the bulk modulus of the solid grains constituting the
fault core, K
f
is the uid bulk modulus and n the fault core porosity.
Fault core specimens are rarely available for specic offshore
elds. In the absence of fault core material from the Statfjord Field,
the results from laboratory measurements of permeability and
diffusion coefcient on intact samples of Burton Formation shale
(Dunlin Group) were used to estimate the consolidation time in the
fault core. For a rst-order estimate, an average of the four tests
made parallel and normal to the lamination is used, disregarding
the effect of anisotropy. This gives a permeability (k) and diffusivity
coefcient (c) equal to 3.8 10
21
m
2
and 4.75 10
8
m
2
/s (1.5 m
2
/
yr), respectively.
Assuming a fault core zone thickness between 0.2 mand 0.45 m,
Eq. (5) gives t
100
between 10 and 50 days. Hence, in view of the
reservoir production timescale, segments of the fault core zone
which juxtapose sand-rich reservoirs can be considered fully
drained, with a pore pressure distribution through the fault (core)
zone varying approximately linearly between the uid pressures in
the adjoining formations.
Table 4
Characteristics of Section 1 of horst structure through Brent Group, Dunlin and
Statfjord Formation. Data for Fault Z0 is written in plain, data for Fault Z5 in bold
letters. Data from Statoil.
Formation/group Throw (m) Fault zone
thickness (m)
Damage zone
thickness (m)
SGR
Brent 1 6578
5665
0.050.15
0.050.10
28
24
<35
1025
Brent 2 6572
5662
0.050.15
0.050.10
28
24
>35
2550
Dunlin 6594
6070
0.050.15
0.100.15
32
30
>50
>50
Statfjord 1 9194
6078
0.050.15
0.100.15
32
30
>50
>40
Statfjord 2 2694
78102
0.050.15
0.100.20
1832
32
>50
>40
Fig. 6. Consolidation in a fault. The pressure in the reservoir on the left-hand side of the fault is depleted by Dp, while the pressure in the right hand side reservoir is unchanged. c is
the diffusivity coefcient of the fault, t is time since depletion started, p
o
the initial pressure and p
ss
the steady state pressure distribution in the fault (i.e. at end of consolidation).
F. Cuisiat et al. / Journal of Structural Geology 32 (2010) 17541767 1760
Note that t
100
is inversely proportional to the permeability (k) of
the fault core. The permeability of the fault core zone is highly
uncertain, but a value much smaller than that of the intact shale is
thought to be unrealistic. Even if the fault core permeability is ten
times smaller than the value assumed above, the time to reach full
consolidation (100500 days) is still smaller than production time
of several decades.
For sandclay juxtaposition, i.e. for segments of the fault where
the reservoir is juxtaposed to (undrained) shale-rich formations, it
is assumed that the fault is fully drained, with pressure in the fault
equal to the reservoir pressure.
2.6. Pore pressure history used for modelling
The pore pressure histories used in the study are obtained from
history matched reservoir simulations using the black-oil reservoir
simulation tool Eclipse.
For modelling of the Brent Fault, pore pressure histories from
2005 (present day situation) are used. The resulting pore pres-
sure reductions on both sides of the Brent Fault in the Brent Group
and Statfjord Formation are given in Table 5. For modelling of the
horst structure, pore pressure prognoses for the year 2020 (Late
life) are considered. The resulting pore pressure reductions at both
sides of the horst structure in the Brent Group and Statfjord
Formation are given in Table 6.
The shale formation is assumed undrained during the depletion,
i.e. no pore uid ow takes place. The pore pressure distribution
through the fault (core) zone is assumed to vary linearly between
the pressures in the reservoir layer at both sides of the fault. The
pore pressure in the damage zone is assumed to be equal to the
pore pressure in the reservoir layer on the same side of the fault
zone. The pore pressure within the horst structure is assumed to be
constant during reservoir depletion.
2.7. Boundary conditions
Boundary conditions for all models are such that displacements
along the bottom boundary are fully-xed, the top boundary is free
to move, while displacements along the side boundaries are xed in
the horizontal direction and free in the vertical direction.
3. Results from global analysis and parametric study
3.1. Global model
The global 2D model is used to assess the effects of boundary
conditions in local models on the stress changes in the fault zone
during depletion.
Acomparison between a local and a global model shows that the
displacement eld is inuenced by the close boundary conditions
in a local model (Fig. 7). However, the shear and normal stresses
along the fault are unaffected by the boundary conditions in the
local model (Fig. 8). Hence, our local models may be used to study
the stress changes in the fault during reservoir depletion, without
signicant effect of the close boundary conditions.
3.2. Parametric study of stress changes in fault core zone
The objective of the parametric study is to investigate the
sensitivity of the stress response in a fault core zone to variations in
fault geometry and material stiffness parameters. The two-
dimensional nite element model used for the parametric analysis
resembles the Brent Fault model (Section 2) at the Statfjord
Formation depth.
From a base case scenario, variation of several parameters has
been performed which can be grouped into:
- reservoir stiffness properties;
- overburden and intra-reservoir shale stiffness properties;
- fault geometry (inclination, thickness, drag and juxtaposition);
- pressure distribution and drainage of fault core zone.
The results presented in Table 7 are analysed in terms of
maximumshear stress s
max
s
1
s
3
=2 and effective octahedral
stress s
0
oct
s
0
1
s
0
2
s
0
3
=3 in the critical location of the fault
core zone. The results are averaged from the ten most critical
integration points representing approximately 10 m of the fault
length.
The closeness of the stress state to a MohrCoulomb failure line
may be dened as the degree of shear mobilisation or CF (e.g.
Templeton and Rice, 2008) given as:
Table 5
Pore pressure reduction applied during modelling of the Brent fault (pressure
histories for 2005).
Pressure reduction Dp (MPa) Brent Field Statfjord Field
Brent Group 30 8
Statfjord Formation 26 7
Table 6
Pore pressure reduction applied during modelling of the horst structure (pressure
histories for year 2020).
Pressure reduction Dp (MPa) Statfjord Field Snorre Field
Brent Group 20 0
Statfjord Formation. 20 10
Fig. 7. Calculated contours of horizontal displacements (in centimetres) around the Brent Fault at Year 2005 (Present day conditions). Global 2D model (Far BC) and local model
(Close BC). Note that color contours on the left and right gures have a slightly different scale.
F. Cuisiat et al. / Journal of Structural Geology 32 (2010) 17541767 1761
CF
s
max
sin f
0
$a s
0
m
(8)
where s
0
m
s
0
1
s
0
3
=2 is the in-plane mean effective stress, a the
attraction and f
0
the effective friction angle of the fault core
material. The attraction is related to the friction angle and the
cohesion c
0
by a c
0
=tan f
0
. By using the values of cohesion and
friction angle for clay/shale material presented in Section 2.5
(c
0
3 MPa, 4
0
24
s
0
oct
1
3
s
0
V
2s
0
h
(9)
After reservoir depletion, the maximum shear stress in the fault
core zone is concentrated in the area with sandsand juxtaposition
(Fig. 9); the most critical point being at the bottom of the depleted
reservoir layer, except for a reverse fault (Case 12 in Table 7) where
the most critical point is located at the top of the depleted reservoir.
The degree of mobilisation (CF) increases from circa 0.6 to 1.3 (Base
case) after depletion, indicating that shear failure may occur in the
fault core zone. Note that the value of CF and the occurrence of
shear failure depend on the actual shear strength. In fact, for the
higher strength estimate, CF increases from circa 0.5 initially to 0.9
after depletion; this indicates no shear failure at all.
28
29
30
31
32
33
34
35
12 14 16 18 20 22
Effective normal stress fault [ MPa]
]
m
[
h
t
p
e
D
CLOSE BC
FAR BC
28
29
30
31
32
33
34
35
4 6 8 10 12
Shear stress fault [ MPa ]
]
m
[
h
t
p
e
D
CLOSE BC
FAR BC
Brent gr. Brent gr.
Dunlin fm. Dunlin fm.
Statfjord fm. Statfjord fm.
0
0
1
x
0
0
1
x
Fig. 8. Calculated effective normal and tangential (shear) stress along the Brent Fault (at the Brent Field side of the fault) at Year 2005. Global 2D model (Far BC) and local model
(Close BC).
Table 7
Stress conditions at critical point in fault during reservoir depletion results from
parametric study. Figures in bold indicate an increase in shear stress mobilisation
after variation of one input parameter from base case. E
res
: Youngs modulus of
reservoir layers, n
res
: Poissons ratio of reservoir layers, E
shale
: Youngs modulus of
shale material, E
fault
: Youngs modulus of fault, FZT: fault zone thickness, L
juxt
: Sand
sand juxtaposition length, b: fault inclination.
Case ID Description s
max
(MPa) s
0
oct
(MPa) CF
Initial conditions before depletion 6.7 16.6 0.64
1 Base case
E
res
10 GPa, n
res
0.2,
E
shale
4 GPa,
E
fault
4 GPa, FZT 10 m,
L
juxt
50 m, b 45
, c
res
3 MPa) are
introduced in a predened area of the fault core zone and b) a Mohr
Coulomb model with tension cut-off T 0 MPa is utilised. The details
of the modelling are not presented here due to space limitation. The
main conclusions from these two cases are: 1) if shear failure is
initiated, the failure zone propagates along the fault dip direction
rather than in the cross-fault direction and 2) if tension failure occurs
it initiates on the non-depleted side of the fault core zone. Propa-
gation of tension failure through the fault is not feasible because the
fault is drained and the effective mean stress increases on the
depleted side of the fault core zone. Furthermore, it is found that
there is no signicant tendency for a tension zone to propagate along
the fault dip direction on the non-depleted side.
4. Stress conditions of the Brent Fault and horst structure
4.1. Present stress conditions of the Brent Fault
Fig. 10 shows the calculated effective principal stresses at the
year 2005 around the Brent Fault in Section 1 (left) and Section 2
(right). There is little rotation of the principal stresses along the
fault. The maximum shear stress after pressure reduction is equal
to circa 16 MPa at the most critical points located in Section 2 at
the bottom of the Statfjord Formation or Brent Group (Fig. 11
right, juxtaposition zones of Brent_2 to Brent_1 and Statfjord_2
to Statfjord_1). The corresponding maximum shear stress is
about 10 MPa (Brent Group) and 14 MPa (Statfjord Formation) in
Section 1.
The stress paths for two critical points in the fault core zone for
both Section 1 and Section 2 are shown in Fig. 12, together with the
upper and lower bounds for the peak shear strength of the shale.
During pressure reduction in the reservoir, the critical point within
the Brent Fault moves towards shear failure. Shear failure may be
expected in the Brent Group as the shear failure criterion is
exceeded, especially for the low shear strength estimate. Shear
failure is less probable in the Statfjord Formation, although the
modelling indicates stress levels in excess of the shear failure
criterion. This is because the actual strength for the Statfjord
Formation is probably signicantly higher than the estimate
plotted on the gure due to its low clay content and presumably
cataclastic fault rock material. Tensile failure may also occur as
indicated by the negative minimum principal stress (i.e. tensile
stress) at the critical points in Section 2.
4.2. Stress conditions within the horst structure during the late life
of the Statfjord Field
The results for the horst structure are shown for Section 1. This is
found to be the most critical section investigated. Only the stress
state at the end of production from the Statfjord Field (end of year
2020) is considered.
Critical stress paths in Brent Fault
0
5
10
15
20
25
30
35
-2 0 2 4 6 8 10 12
Effective min. principal stress (MPa)
p
i
c
n
i
r
p
.
x
a
m
e
v
i
t
c
e
f
f
E
a
)
a
P
M
(
s
s
e
r
t
s
Sec1,Statfjord fm.
Sec1,Brent fm.
Sec2, Statfjord fm.
Sec2, Brent fm.
Low strength
high strength
Critical stress paths in Brent Fault
0
2
4
6
8
10
12
14
16
18
0 2 4 6 8 10 12 14 16 18 20 22
Effective mean stress (MPa)
)
a
P
M
(
s
s
e
r
t
s
r
a
e
h
s
.
x
a
M
Sec1, Statfjord fm.
Sec1, Brent fm.
Sec2, Statfjord fm.
Sec2, Brent fm.
high strength
Low strength
Fig. 12. Stress paths for critical points in Sections 1 and 2 through the Brent Fault. Effective mean stress vs. maximum shear stress (left) and minimum principal effective stress vs.
maximum principal effective stress (right). Upper and lower bounds of shear strength are indicated for the Brent Group.
Fig. 13. Calculated effective principal stress crosses at the top of the horst structure (left) and at Fault Z5 in the juxtaposition of the Statfjord Formation (right) for Section 1. Note
that the purpose of the gure is to illustrate principal stress re-orientation rather than stress magnitudes for which contour plots should be used. Hence no scale is linked to the
crosses.
F. Cuisiat et al. / Journal of Structural Geology 32 (2010) 17541767 1764
Fig. 13 (left) shows the calculated principal effective stress
redistribution after pressure reduction at the top of the section
through the horst structure. Fig. 13 (right) shows the same for Fault
Z5 (Statfjord side of the horst), around the juxtaposition windowof
the Statfjord Formation (i.e. area where Statfjord_1 juxtaposes to
Statfjord_2). Vertical tensile cracks may develop at the top sand-
stone within the horst structure, as the minimum effective prin-
cipal stress is close to zero and its orientation is nearly horizontal.
The extent of the zone with tensile stresses increases with the pore
pressure depletion in the Statfjord Field (Fig. 14). The development
of tensile stresses along the fault only occurs inside the horst
structure. Tensile stresses are prevented fromdeveloping inside the
fault towards the Statfjord Field due to the drained behaviour of the
fault and effective stress increase during depletion. However,
a zone with tensile stresses also develops from a singular point
above the horst structure. In this area, a continuous drainage path
would not be expected since tensile stresses are prevented from
developing to the west of the horst structure (i.e. the depleted
Statfjord Field side).
Stress paths for two critical points in Fault Z5 show that shear
failure may occur in the Brent Group and the Statfjord Formation, as
the low strength failure criterion is exceeded (Fig. 15). Tensile
failure may also occur in the Brent Group as indicated by the
negative minimum principal stress (i.e. tensile stress) at the critical
points. The maximum shear stress is less than 10 MPa and
concentrated in the juxtaposition windows (Fig. 16).
5. Discussion
5.1. Present stress conditions associated with the Brent Fault
Numerical analyses of the present (Year 2005) stress condition
in the Brent Fault with signicantly larger pore pressure depletion
in the Brent Field compared to the Statfjord Field shows that shear
failure may have occurred in the fault at the juxtaposition window
for the Brent Group (i.e. Brent_2 against Brent_1 in Fig. 11). The
peak shear strength in this zone is assumed to be similar to that of
the Burton Formation shale, which is between 10 MPa and 18 MPa
for an effective octahedral stress between 14 MPa and 20 MPa. This
strength depends rst of all on the orientation of the lamination
compared to the orientation of the critical shear stress. The lowest
strength is obtained when the critical stress is oriented parallel to
the lamination. The actual orientation of the lamination within the
fault core zones is, however, unknown.
As shown by the modelling, if shear failure occurs, it is initiated
on the Statfjord Field side of the Brent Fault, i.e. the side with the
lowest pore pressure depletion and thus also the lowest effective
octahedral stress and corresponding strength. When the strength is
reduced to the residual strength (strain softening), the failure zone
may propagate along the fault but not across the fault. Tensile
failure may occur on one side of the fault in areas of low effective
horizontal stress (i.e. with the highest pore pressure). However, it
cannot propagate across the fault zone due to the increased
Fig. 14. Calculated zones with tensile stresses in Section 1 through the horst structure. Pore pressure depletion of 20 MPa (left) and 30 MPa (right) in the Statfjord Field side of the
horst structure.
Stresses in Z5 fault
Section 1
0
2
4
6
8
10
12
14
0 2 4 6 8 10 12 14
Effective mean stress (MPa)
)
a
P
M
(
s
s
e
r
t
s
r
a
e
h
s
.
x
a
M
Statfjord fm.
Brent fm.
Low strength
High strength
Stresses in Z5 fault
Section 1
0
5
10
15
20
25
-2 0 2 4 6 8 10 12
Effective min. principal stress (MPa)
l
a
p
i
c
n
i
r
p
.
x
a
m
e
v
i
t
c
e
f
f
E
)
a
P
M
(
s
s
e
r
t
s
Statfjord fm.
Brent fm.
Low strength
Fig. 15. Stress paths for critical points in Fault Z5, Section 1 through horst structure. Effective mean stress vs. maximum shear stress (left) and minimum principal effective stress vs.
maximum principal effective stress (right). Upper and lower bounds of shear strength for the Brent Group are indicated.
F. Cuisiat et al. / Journal of Structural Geology 32 (2010) 17541767 1765
effective horizontal stress on the other side (i.e. depleted side with
the lowest pore pressure). Hence, a continuous zone with shear or
tensile failure cannot develop across the fault core zone. As a result,
only minor changes in hydraulic properties of the fault zone are
expected as a result of present stress changes caused by pore
pressure depletion on the Brent Field side. This conclusion agrees
well with eld observations, indicating no signicant changes in
the seal integrity of the Brent Fault, which currently acts as
a hydraulic barrier between the Brent and Statfjord Fields.
5.2. Late life stress conditions on the horst structure
Analyses of the horst structure show similar stress changes in
the fault zone between the Statfjord Field and the horst structure
(Fault Z5) as identied for the Brent Fault. However, the corre-
sponding maximum shear stress in Fault Z5 is generally smaller
than in the Brent Fault, such that the predicted late life stress
situation is considered to be less critical than that already expe-
rienced by the Brent Fault. Hence, if mechanical and sealing
properties for the two faults are similar, it can be concluded that
the sealing integrity of the horst structure should not be altered
signicantly after full depletion of the Statfjord Field. Further-
more, the results of the modelling show that if failure takes place
in the horst structure, then the failure zones (either shear or
tensile failure) would propagate along and not across the fault (at
the side with lowest or no depletion). The development of
microcracks (and/or opening of pre-existing microcracks) is also
inhibited on the depleted side of the horst structure (towards
Statfjord) due to effective octahedral stress increase which occurs
during depletion.
It should be pointed out that only mechanical effects due to
effective stress changes were considered in this study. Other
mechanisms (e.g. capillary effects) might contribute to change the
hydraulic resistance of faults during pressure depletion.
6. Conclusions
We have presented the results of geomechanical analyses of
fault behaviour at the Statfjord Field as part of Statfjord Late Life
project. The objective was to assess the potential for developing
hydraulic communication between the Statfjord and Snorre elds
through a horst structure, during nal depressurisation of the
Statfjord Field. Two-dimensional plane strain geomechanical
analyses were carried out to calculate the deformation and
maximum shear stresses along the faults bounding the horst
structure, resulting fromcompaction and horizontal deformation of
the Statfjord Field due to pore pressure reduction.
The stress conditions at the Brent Fault separating the Brent and
Statfjord Fields were rst considered. According to pressure data,
the fault is acting as a pressure seal between the two elds. The
results of the modelling show that the calculated stress changes in
the horst structure are equal to or less critical than the calculated
present stress changes in the Brent Fault. It is therefore concluded
that the mechanical effects (i.e. stress changes) associated with the
planned depressurisation of the Statfjord Field during late life will
not affect signicantly the hydraulic resistance of the horst
structure.
Only two-dimensional plane strain models and simple fault
geometries were considered in this study, as the focus of the
analyses was to identify failure modes and mechanisms rather than
to predict absolute values of stress changes. The error by using
a two-dimensional approach instead of three-dimensional model-
ling is less critical as the results are used quantitatively to compare
two faults under similar conditions. This approach would not have
been valid if the geometry of one fault had been very different to
that of the other.
In light of the complexity and uncertainty associated with fault
parameters, a parametric study was conducted to investigate the
sensitivity of the modelled stress changes. This modelling can help
to test several geological scenarios (e.g. presence of drag, thickness
of fault) and assess the relative importance of particular mecha-
nisms (e.g. drainage conditions). It is found that the maximum
stress changes are not very sensitive to geometrical variations and
uncertainties in stiffness distributions. The largest uncertainty
relates to the peak shear strength of the fault (core) zone.
It should be pointed out that only the mechanical effects due to
effective stress changes were considered in this study. Other
mechanisms (e.g. capillary effects) might contribute to variations in
the hydraulic resistance of faults during pressure depletion.
Furthermore, the faults were modelled as single plane of weak-
nesses, which is clearly an over-simplication. Although the impact
of damage zone was not properly modelled in this study, it was
shown that its impact might not be so important for maximum
shear stress distribution in the fault.
The integrity of the horst structure, as presented in this paper,
was assessed relatively to the situation at the Brent Fault. Absolute
predictions in terms of changes in hydraulic resistance were
beyond the scope of this study. Further work would be required to
develop petrophysical models which could relate the hydraulic
properties of faults (e.g. permeability, capillary entry pressure) to
mechanical changes (e.g. fault dilation, grain and pore volume,
fracturing, tortuosity).
Instead, the approach chosen in this study assumed that the
sealing conditions of the Brent Fault could be used to calibrate the
methodology applied to the horst structure. This assumption is
supported by geological understanding of the structures. However,
the biggest uncertainty pertained to the strength of fault zones.
Without specic fault data, but from expected clay content and
deformation products in the fault zone, this study suggests that it is
acceptable to use the residual shear strength of shales as repre-
sentative of the strength of the fault (core) zone. Further work is
required to improve the determination of the strength of faults. For
instance, a systematic database could be developed by testing the
strength, stiffness and hydraulic properties of fault core samples in
the laboratory, for various fault deformation products sampled
from core from offshore elds. In that way, reliable data for future
analyses could be gathered.
Fig. 16. Contours of maximum shear stress in MPa (from 6 to 12 MPa with increments
of 1 MPa) for Section 1 through the horst structure at Year 2020 (Base Case). Areas with
maximum values are indicated with gures and arrows.
F. Cuisiat et al. / Journal of Structural Geology 32 (2010) 17541767 1766
Acknowledgements
The authors acknowledge the license partners of the Statfjord
Field (Statoil, ExxonMobil, Shell, ConocoPhilips, Enterprise oil and
Centrica) for their permission to publish this paper. We thank Rune
Holt and Dave Dewhurst for reviewing the manuscript and
contributing with valuable comments, and Chris Jackson for nal
editorial comments.
References
Beach, A., Brown, J.L., Welbon, A.I., McCallum, J.E., Brockbank, P.J., Knott, S.D., 1997.
Characteristics of fault zones in sandstones from NW England: application to
fault transmissibility. In: Meadows, N.S., Trueblood, S.P., Hardman, R., Cowan, G.
(Eds.), Petroleum Geology of the Irish Sea and Adjacent Areas. Geological
Society, London, Special Publications, vol. 124, pp. 315324.
Beach, A., Welbon, A.I., Brockbank, P.J., McCallum, J.E., 1999. Reservoir damage
around faults: outcrop examples form the Suez rift. Petroleum Geoscience 5,
109116.
Boge, R., Lien, S.K., Gjesdal, A., Hansen, A.G., 2005. Turning a North Sea Oil Giant into
a Gas Field Depressurization of the Statfjord Field SPE 96403, Offshore Europe,
69 September, Aberdeen, United Kingdom.
Cocco, M., Rice, J.R., 2002. Pore pressure and poroelasticity effects in Coulomb stress
analysis of earthquake interactions. Journal of Geophysical Research in Solid
Earth 107, 2030.
Fisher, Q.J., Knipe, R.J., 1998. Fault sealing processes in siliciclastic sediments. In:
Jones, G., Fisher, Q.J., Knipe, R.J. (Eds.), Faulting, Fault Sealing and Fluid Flow in
Hydrocarbon Reservoirs. Geological Society, London, Special Publications, vol.
147, pp. 117134.
Fisher, Q.J., Knipe, R.J., 2001. The permeability of faults within siliciclastic petroleum
reservoirs of the North Sea and Norwegian Continental Shelf. Marine and
Petroleum Geology 18, 10631081.
Grasso, J.R., 1992. Mechanics of seismic instabilities induced by the recovery of
hydrocarbons. Pure and Applied Geophysics 139, 507534.
Hesthammer, J., Jourdan, C.A., Nielsen, P.E., Ekern, T.E., Gibbons, K.A., 1999. A tec-
tonostratigraphic framework for the Statfjord Field, northern North Sea.
Petroleum Geoscience 5, 241256.
Hettema, M.H.H., de Pater, C.J., 1998. The poromechanical behaviour of Felser
Sandstone stress and temperature-dependent SPE/ISRM 47270, SPE/ISRM
Rock Mechanics in Petroleum Engineering, 810 (July, Trondheim, Norway).
Hettema, M.H.H., Hanssen, T.H., Jones, B.L., 2002. Minimizing Coring-Induced
Damage in Consolidated Rock. SPE/ISRM 78156, SPE/ISRM Rock Mechanics
Conference, 2023 October, Irving, Texas.
Horsrud, P., 2001. Estimating mechanical properties of shale from empirical
correlations. SPE 56017. SPE Drilling and Completion 16, 6873.
Maury, V.M.R., Grasso, J.R., Wittlinger, G., 1992. Monitoring of subsidence and
induced seismicity in the lacq gas-eld (France) the consequences on gas-
production and eld operation. Engineering Geology 32, 123135.
Mavko, G., Mukerji, T., Dvorkin, J., 1998. The Rock Physics Handbook. Tools for
Seismic Analysis in Porous Media. Cambridge University Press.
Odling, N.E., Harris, S.D., Knipe, R.J., 2004. Permeability scaling properties of
fault damage zones in siliclastic rocks. Journal of Structural Geology 26,
17271747.
Plaxis Users Manual. Plaxis 2D Prof. v.8.2, Build 811, 2004. www.plaxis.nl.
Raaen, A.M., Horsrud, P., Kjorholt, H., Okland, D., 2006. Improved routine esti-
mation of the minimum horizontal stress component from extended leak-off
tests. International Journal of Rock Mechanics and Mining Sciences 43,
3748.
Rice, J.R., Cleary, M.P., 1976. Some basic stress-diffusion solutions for uid-saturated
elastic porous media with compressible constituents. Reviews of Geophysics
and Space Physics 14, 227241.
Sperrevik, S., Gillespie, P.A., Fisher, J.F., Halvorsen, T., Knipe, R., 2002. Empirical
Estimation of Fault Rock Properties. In: Norwegian Petroleum Society (NPF)
Special Publication, vol. 11 109125.
Sverdrup, E., Bjrlykke, K., 1997. Fault properties and the development of cemented
fault zones in sedimentary basins: eld examples and predictive models. In:
Mller-Pedersen, P., Koestler, A.G. (Eds.), Hydrocarbon Seals: Importance for
Exploration and Production. Norwegian Petroleum Society (NPF) Special
Publications, vol. 7, pp. 91106.
Templeton, E.L., Rice, J.R., 2008. Off-fault plasticity and earthquake rupture
dynamics: 1. Dry materials or neglect of uid pressure changes. Journal of
Geophysical Research 113, B09306.
Wibberley, C.A.J., Yielding, G., Di Toro, G., 2008. Recent advances in the under-
standing of fault zone internal structure: a review. In: Wibberley, C.A.J.,
Kurz, W., Imber, J., Holdsworth, R.E., Collettini, C. (Eds.), The Internal Structure
of Fault Zones: Implications for Mechanical and Fluid-Flow Properties.
Geological Society, London, Special Publications, vol. 299, pp. 533.
Willson, S.M., Last, N.C., Zoback, M.D., Moos, D., 1999. Drilling in South America:
a wellbore stability approach for complex geologic conditions SPE 53940, Latin
American and Caribbean Petroleum Engineering Conference, 2123 April,
Caracas, Venezuela.
Wiprut, D., Zoback, M.D., 2000. Fault reactivation and uid ow along a previ-
ously dormant normal fault in the northern North Sea. Geology 28,
595598.
Yielding, G., Freeman, B., Needham, D.T., 1997. Quantitative fault seal prediction.
American Association of Petroleum Geologists Bulletin 81, 897917.
Zhang, X., Sanderson, D.J., 1998. Numerical study of critical behaviour of deforma-
tion and permeability of fractured rock masses. Marine and Petroleum Geology
15, 535548.
F. Cuisiat et al. / Journal of Structural Geology 32 (2010) 17541767 1767
Structural controls on leakage from a natural CO
2
geologic storage site:
Central Utah, U.S.A.
Ben Dockrill
a,
*
, Zoe K. Shipton
b
a
Department of Geology, Trinity College Dublin, Dublin 2, Ireland.
b
Department of Geographical and Earth Sciences, University of Glasgow, Glasgow G12 8QQ, Scotland
a r t i c l e i n f o
Article history:
Received 26 May 2009
Received in revised form
26 November 2009
Accepted 19 January 2010
Available online 28 January 2010
Keywords:
Fault zone
Damage zone
Fluid ow
Co
2
Hydrocarbons
a b s t r a c t
Faults and associated fracture networks can signicantly inuence regional ow of groundwater,
hydrocarbons and other uids. The distribution of CO
2
springs and seeps along the Little Grand Wash
fault and Salt Wash faults in central Utah is controlled by along-fault ow of CO
2
-charged groundwater
from shallow aquifers (<1 km deep). The same faults are the likely conduits that charge the shallow
aquifers with CO
2
from depth. We document fault zone trace geometry and architecture, and evidence
for palaeo-uid ow within the footwalls of both faults. Evidence for palaeo-uid ow consists of
extensive bleaching of sandstones and some siltstones, mineralisation of carbonates and celestine veins
and minor hydrocarbon staining. The eld evidence shows that the pathways for multiple phases of uid
ow were structurally controlled utilising the fracture network developed in the damage zone of the
faults. To investigate the likely effect of these faults on the regional uid-migration pathways at depth,
a 3D model of the faulted systemwas generated and a fault seal analysis applied to predict the cross-fault
sealing capabilities of the studied faults. Due to the scarcity of subsurface data, the results are not
conclusive but suggest probable multiple cross-fault leak points for uids to migrate across the fault, in
contrast to the eld observations that indicate fault-parallel ow. This comparison of eld observations
to the modelling approach demonstrates the inability of conventional seal analysis techniques to predict
fault-parallel uid leakage and highlight the effects fracture networks in the damage zone, especially at
structural complexities along the fault, have in producing pathways for vertical ow. Multiple uids have
utilised similar fault-parallel pathways over geological time demonstrating that such pathways have the
potential to cause long-term leakage from hydrocarbon reservoirs and CO
2
storage sites.
2010 Elsevier Ltd. All rights reserved.
1. Introduction
Faults and associated fracture networks can play a signicant
role in the subsurface migration of various uids. Focussing of ow
related to fault geometry has been demonstrated in geothermal
elds (Curewitz and Karson, 1997; Rowland and Sibson, 2004),
hydrothermal/epithermal systems (Breit and Meunier, 1990;
Micklethwaite, 2009) and petroleum systems (Chan et al., 2000;
Garden et al., 2001; Gartrell et al., 2004). In most studies, ow is
concentrated in the fracture network (commonly referred to as the
damage zone) that is developed around a main zone of slip.
Complexities along a fault related to terminations and/or linkages
between fault segments are commonly domains of high fracture
density and connectivity and are therefore likely to focus ow
(Curewitz and Karson, 1997; Anderson and Fairley, 2008; Eichhubl
et al., 2009). Models that predict fault properties based on the
throwand host rocks cut by the fault (Yielding et al., 1997; Yielding,
2002; Bretan et al., 2003) generally rely on oversimplied fault
geometries and complexities, leading to the possibility of under-
estimating likely leakage points due to fault throwpartitioning and
simplied fault linkages (Childs et al., 1996, 1997), and do not
account for along-fault ow.
This study investigates a natural leaking CO
2
-rich system at the
northern end of the Paradox Basin in central Utah, United States. In
this locality, the Little Grand Wash fault and northern fault of the
Salt Wash graben provide lateral barriers to present-day cross-fault
ow, but provide pathways through the cap rock via damage zone
fractures to allow CO
2
and additional uid regimes to leak to the
surface (Shipton et al., 2004, 2005). Multiple mineralisation and
diagenetic products associated with past and present migration of
uids along both faults demonstrate structural controls inuencing
multiple uid regimes to vertically migrate through a thick inter-
bedded sandstoneshale stratigraphy. By combining outcrop
* Corresponding author. Present address: Chevron Australia, 250 St Georges
Terrace, Perth, WA, Australia. Tel.: 61 8 9216 4141; fax: 61 8 9216 4103.
E-mail address: ben.dockrill@chevron.com (B. Dockrill).
Contents lists available at ScienceDirect
Journal of Structural Geology
j ournal homepage: www. el sevi er. com/ l ocat e/ j sg
0191-8141/$ see front matter 2010 Elsevier Ltd. All rights reserved.
doi:10.1016/j.jsg.2010.01.007
Journal of Structural Geology 32 (2010) 17681782
analysis of the fault zone and associated uid-migration products
with analysis of fault rock integrity through modelling, we have
examined constraints on leakage from shallow reservoirs. Faults
can form barriers to ow either by juxtaposing reservoir rock
against low-permeability clay-rich non-reservoir rock to create
a juxtaposition seal, or when processes of fault rock generation
form a low-porosity and low-permeability fault rock. By analysing
the fault rocks and the processes and factors that contribute to the
failure of this system, we can identify and highlight the roles faults
can play in trapping and transmitting uids. These results
demonstrate the impact fault zones can have on fault-parallel
leakage from a robust structural trap and highlight the potential
risks when assessing seal integrity for structural traps in the
hydrocarbon and emerging CO
2
geologic storage industries.
2. Geological setting
The eld area is located at the northern end of the Paradox Basin
in the Colorado Plateau region of the United States (Fig. 1), a late
Palaeozoic to Mesozoic intracratonic basin inlled with a thick
sequence of evaporite, carbonate and clastic sediments (Hintze,
1993). The basin is dened by the areal extent of the Pennsylva-
nian Paradox Formation, which contains nearly 2 km of evaporates
(Doelling et al., 1988). The basin has been investigated for hydro-
carbons (Peterson, 1973, 1989; Hansley, 1995; Huntoon et al., 1999)
and mineral resources (Breit and Meunier, 1990; Morrison and
Parry, 1986). Multiple reservoirs have accumulated CO
2
for
extended periods of times (Allis et al., 2001, 2005; Moore et al.,
2005; Shipton et al., 2004, 2005; White et al., 2005). Some of
these are now currently being exploited, predominantly for
enhanced oil recovery (i.e. Bravo and McElmo domes Allis et al.,
2001). Other reservoirs leak CO
2
due to the inuence of faults
(Springerville-St Johns Moore et al., 2005; Hurricane fault
Nelson et al., 2009) and/or boreholes (Woodside Doelling, 1994).
The stratigraphy in the Paradox basin ranges from the Penn-
sylvanian Paradox Formation to the Mid Cretaceous Mancos Shale,
though only the Upper Jurassic to Mid Cretaceous succession crops
out in the eld area (Fig. 1). The Pennsylvanian and Permian
formations consist of marine carbonates and shales that are
potential sources of the CO
2
(Heath et al., in press; Wilkinson et al.,
2008) and hydrocarbons (Peterson, 1973, 1989; Huntoon et al.,
1999; Chan et al., 2000; Garden et al., 2001). The aeolian reser-
voir sandstones of the Permian White RimSandstone are capped by
uvial and lacustrine shales of the Triassic Moenkopi and Chinle
formations. The aeolian Lower Jurassic Wingate and Navajo sand-
stones are important regional aquifers separated by the uvial
Kayenta Formation aquitard. Forming a seal above the Navajo
Sandstone is the Mid-Jurassic Carmel Formation, a complex sabkha
sequence of sandstone, siltstone, mudstone, limestone, anhydrite
and gypsum. The youngest reservoir units in the basin are the Mid-
Jurassic Entrada and Curtis aeolian to marginal marine sandstones,
overlain by marine siltstones and shales of the Middle Jurassic
Summerville Formation. The remaining stratigraphic sequence is
dominated by shales with small, disconnected reservoir units. The
Upper Jurassic Morrison Formation consists of stacked uvial
sandstone channels, interspersed and overlain by lacustrine shales.
The lower Cretaceous Cedar Mountain Formation lacustrine shales
are overlain by conglomeritic uvial channels from the lower
Cretaceous Dakota Formation. The youngest rocks exposed in the
eld area are marine marls of the Middle Cretaceous Mancos Shale.
The shallowly north- to northwest-plunging, open Green River
anticline is one of a series of northwest-trending folds that have
growth histories related to salt movement in the Paradox
Formation since the Permian (Doelling et al., 1988). The Green River
anticline is cut by the Little Grand Wash fault and the Salt
Wash graben. Timing of movement along both faults is poorly
constrained, with the youngest faulted stratigraphy being the Mid
Cretaceous Mancos Shale, though evidence presented by Pevear
M
o
a
b
f
a
u
l
t
S
a
l
t
V
a
l
l
e
y
C
o
u
r
t
h
o
u
s
e
s
y
n
c
l
i
n
e
Green River
Anticline
g
r
a
b
e
n
Green
River
0 10 Kilometres
K
uJ
uJ
uJ
uJ
uJ
uJ
mJ
mJ mJ
mJ
lJ
lJ
lJ
Fault (tick on downthrown side)
Cretaceous
Mancos and Cedar Mountain
K
Upper Jurassic
Morrison, Summerville, Curtis
uJ
Mid Jurrasic
Entrada and Carmel
mJ
Lower Jurassic
Navajo, Kayenta and Wingate
lJ
Tr
P
Q
Q
K
K
Fig. 2a
Fig. 2b
2
3
3
4
5
2
3
5
6 3
4
7
25
b
c
a
Ten Mile
graben
Salt Wash
graben
Little Grand
Wash fault
30 20 10 0
100
200
300
0
distance along strike (km)
t
h
r
o
w
(
m
)
260m
30 20 10 0
100
200
300
0
distance along strike (km)
t
h
r
o
w
(
m
)
400 366m
165m
R
R
210m
154m
B
UTAH
Fold (arrow in plunge direction)
Triassic
Chinle and Moenkopi
Tr
Permian
Cutler
P
Quaternary Q
N
110 15
o
110 00
o
39 00
o
38 45
o
Fig. 1. Geological map of the study area in central Utah, USA. Structural data from eld work and McKnight (1940) and Doelling (2001, 2002). Dashed boxes show the positions of
Fig. 2. The letters a, b, c indicate position of the fault zone strip maps displayed in Fig. 3. Top right inset, throw distribution along the Little Grand Wash fault zone. Centre inset,
throw distribution along the northern (black circles) and southern (grey triangles) Salt Wash graben. Note that relay ramps along the Salt Wash graben are coincident with throw
minima.
B. Dockrill, Z.K. Shipton / Journal of Structural Geology 32 (2010) 17681782 1769
et al. (1997) and Shipton et al. (2004) suggest possible early Tertiary
to Quaternary movement. Where the fold axis of the Green River
anticline is cut by the faults, CO
2
-charged groundwater effuses from
a series of geysers and springs (Doelling, 1994; Shipton et al., 2004,
2005). A series of actively-forming spring deposits and remains of
ancient travertine deposits are found along or immediately to the
North (i.e. in the footwall) of both fault systems (Fig. 2). All of the
springs originate in the footwall though some of the travertine
deposits drape over the fault into the hanging-wall. Additional
CO
2
-charged geysers occur where hydrocarbon or water boreholes
have penetrated the footwall reservoirs, though some boreholes
that penetrate the footwall reservoirs at depth do not leak (Shipton
et al., 2005).
The faulting of the Green River anticline has created a series of
stacked three-way anticlinal closures in the footwall reservoirs of
the Little Grand Wash fault and northern fault of the Salt Wash
graben. Southeast-directed regional groundwater ow (Hood and
Patterson, 1984) suggests that CO
2
-charged meteoric uids have
been focussed up the north-plunging Green River anticline towards
the central sections of the faults. Carrier beds for the CO
2
-charged
water include the Lower Jurassic Wingate and Navajo sandstones
and the Mid-Jurassic Entrada and Curtis sandstones (see Heath
et al., in press). The cap rocks for these structural closures are
provided by clay-rich and evaporitic beds in the sequence, while
lateral barriers to ow are provided by the faults (Shipton et al.,
2004, 2005). Stable isotope data indicate that travertine deposits
along both faults have resulted from a common CO
2
-rich uid
(Shipton et al., 2005; Dockrill et al. in prep). Furthermore, leakage is
conned to the footwall of both faults with no isotopically similar
carbonates located elsewhere in the area. The distribution of
springs and travertines therefore suggests that the studied faults
are providing a long-lived pathway for a signicant proportion of
CO
2
-rich uids to migrate vertically through multiple cap rocks in
the footwalls of both faults.
3. Fault geometry
The eastwest trending south-dipping Little Grand Wash fault
juxtaposes late Jurassic and Cretaceous siliciclastics at the surface
(Fig. 1). The fault has a 30 km long, arcuate surface trace that splits
into two dominant fault strands in the central part of the structure,
extending from 3.2 km east to 0.1 km west of the Green River
(Fig. 2). The two sub-parallel fault strands introduce structural
complexities to the fault with a series of prominent fault bends,
branch points and relay ramps. Between and surrounding the two
main strands, a complicated array of minor faults and relays has
Fig. 2. Relationship between Little Grand Wash fault and Salt Wash graben structures and locations of CO
2
leakage (modern and ancient travertines) and evidence for past uid
ow. (a) Central part of the Little Grand Wash fault where it splits into two main strands and cuts the hinge of the Green River anticline. Dashed boxes showthe two locations for the
fracture analysis in Fig. 4. (b) Detailed structure of the Salt Wash graben. Equal angle stereonets show strike and dip of faults with lineations (triangles) indicating direction of fault
movement. N is the number of faults measured. Closed triangles correspond to northern fault lineations, while open triangles indicate southern fault lineations. (For interpretation
of the references to colour in this gure legend, the reader is referred to the web version of this article).
B. Dockrill, Z.K. Shipton / Journal of Structural Geology 32 (2010) 17681782 1770
created multiple structural terraces with varying dips (Fig. 2; see
also Fig. 7 in Vrolijk et al., 2005). Away from the central zone,
a single fault strand is present. In the subsurface, the down-dip
extent and geometry of the fault is uncertain, though drilling
records for two exploration wells (Amerada Hess 1 and 2) that
intersect the fault at depth suggest that the fault may offset rocks as
old as Pennsylvanian in age.
The dip of the main fault strands along the Little Grand Wash
fault varies but is generally steep, averaging 70
W/30
56
W/54
90
counterclockwise from
the fault planes and have lengths of decimeters to meters, whereas
the lengths of the linked fault segments are from meters to deca-
meters (Figs. 8 and 9). Most splay cracks are tensile fractures and do
not show any evidence of shearing although some splay cracks,
particularly those within fault linkages, show evidence of shearing.
These shear splay cracks are thought to be sheared tensile fractures
which have originated as tensile fractures, because the orientation
of the shear splay cracks is the same as those of the adjacent tensile
splay cracks.
In addition, the major FCBs were shown to have slipped under
a horizontal stress state with s1 normal to the fold axis and s3
parallel to the fold axis, determined by the multiple inverse method
using the fault displacement data obtained at this outcrop (Ishii and
Fukushima, 2006). Under this stress regime the above orientation
of the evolving splay cracks is in agreement with those generally
predicted by the numerical simulations of Lunn et al. (2008).
4.2. Fracture logging of the boreholes
Siliceous mudstones of the Wakkanai Formation were inter-
sected by borehole HDB-6 from 262 m to 620 m depth; by HDB-9
from 0 m to 520 m depth; by HDB-11 from 460 m to 1020 m depth
and by PB-V01 from 237 m to 520 m depth. Fractures in the core
from these sections were examined. At shallower depths in bore-
holes HDB-6, HDB-11 and PB-V01, the diatomaceous mudstones of
the Koetoi Formation were intersected. FCBs were commonly
intersected by the boreholes; examination of recovered core indi-
cated that for every 10 m long section of core there were between
0 and 44 FCBs with an average of 9.4 FCBs over 10 m long sections
(frequency: 044 per 10 m, avg. 9.4 per 10 m). Additionally, 44% of
the FCB population were oriented by detailed correlation between
core and BHTV imagery. Moreover, 14% of the population of FCBs
observed in core are major FCBs (frequency: 09 per 10 m, avg. 1.3
per 10 m), and 23% of the major FCB population could be oriented
by detailed correlation with the BHTV imagery. Strikes of all the
oriented FCBs are predominantly EW, oblique to the fold axes (i.e.
N40
), though tensile
strengths on core from nearly the same depths are 3.1 0.2 MPa
and 0.9 0.1 MPa respectively. Internal friction angles were also
obtained from ten core samples of the Wakkanai Formation in the
other boreholes (HDB-1 and HDB-2) (Niunoya and Matsui, 2005),
and the average angle of internal friction for the twelve core
samples from these boreholes (HDB-1, HDB-2 and HDB-6) is
26.1 2.9
(Fig. 11b).
5. Discussion
5.1. Relative age of formation of the major FCBs
The relative timing of fault formation can be inferred from their
crosscutting relationships with other structures. The major FCBs
displace bedding faults without exception. The bedding faults are
considered to have formed synchronously with exural folding
(Ishii and Fukushima, 2006). Therefore, the later major FCBs are
thought to have developed during and/or after uplift and erosion
following folding. This deduction is also supported by observations
from the horizontal outcrop where a major FCB displaces layered
opaline chert, which is considered by Iijima and Tada (1981) to have
formed during uplift and erosion. This implies the major FCBs were
growing during and/or after uplift and erosion.
5.2. Stress state suitable for propagation of splay cracks in the
Wakkanai Formation
What stress state would lead to the development of the abun-
dant splay cracks that are seen to propagate from the major FCBs?
Fig. 10. En echelon shear fractures (minor FCBs) evolving at and beyond the minor FCB tips, FCB1 and FCB2 tips in this gure, at much lower angles to the existing fault planes
(a core sample from 469.6 m depth in PB-V01). (a) an overview shot of the core sample (b) occurrence at and beyond the FCB1 tip, (c) hanging wall surface of the FCB1, (d)
occurrence at and beyond the FCB2 tip, and (e) footwall surfaces of the shear fractures developing at and beyond the FCB1 tip. Splay fracturing was not observed near the FCB1 and
FCB2 tips. Sense of displacement of the FCB1, FCB2 and shear fractures are the same; sinistral, predominantly strike-slip, although slickenlines and slickensteps of the shear fractures
indicate only very weak shearing.
E. Ishii et al. / Journal of Structural Geology 32 (2010) 17921805 1798
Most splay cracks propagating froma fault can be formed as tensile
fractures, particularly at the fault tips, by concentration of tensile
stress generated when a slip patch nucleates and propagates in
a fault (Martel and Pollard, 1989). But such tensile failure does not
necessarily occur at the fault tips under all initial stress states.
Previous numerical simulations suggest that shear failure can also
occur at the fault tips following fault slip, resulting in propagation
of shear fractures subparallel to the original fault plane or forma-
tion of shear fractures that propagated back into the compressive
quadrants (Shen and Stephansson, 1993; Bourne and Willemse,
2001; Willson et al., 2007; Lunn et al., 2008). Although an extension
of the initial fault in its own plane also has been previously simu-
lated by Du and Aydin (1993, 1995), tensile failure is suppressed in
their conceptual model of mechanical failure.
Following Bourne and Willemse (2001), the distance between
any prevailing stress state in a rock mass to either the tensile or the
shear failure conditions can be described by a stress quantity, X on
a Mohr diagram. In Fig. 12, X corresponds to the shortest distance
from the GrifthCoulomb failure envelope of rock strength to the
Mohr circle, the initial stress state. Brittle failure occurs when
the Mohr circle stresses reach the failure envelope, i.e. X 0, and
the mode of failure is determined by whichever failure condition is
met; either X
shear
(X
s
) 0 or X
tensile
(X
t
) 0. Furthermore, reduction
of effective normal stress and increase of shear stress near fault tips,
that are produced by fault slip (e.g., Martel and Pollard, 1989;
Bourne and Willemse, 2001), are important factors causing brittle
failure near fault tips.
When fault slip occurs under an initial stress state where X
s
is
smaller than X
t
(Fig. 13a, top), two cases can be assumed for failure
induced by reduction of effective normal stress and/or increase of
shear stress near the fault tip (Fig. 13b). In the rst case, shear
failure occurs before tensile failure can occur (Fig. 13b, top), and, at
least in the case of the Wakkanai Formation, en echelon shear
fractures can develop at and beyond the fault tips in the plane of
and thus parallel to the existing faults as shown in Fig. 10 and
Fig. 13c, top. Shear fractures that propagate back into the
compressive quadrants were not observed in the outcrop. Eventu-
ally, the en echelon shear fractures would coalesce into a fault
represented by fault rocks such as a major FCB. In the other case,
tensile failure occurs before shear failure can occur (Fig. 13b,
bottom) and splay cracks propagate from the fault tips by tensile
failure (Fig. 13c, bottom).
However, when fault slip occurs in the initial stress state where
X
s
is larger than X
t
(Fig. 13a, bottom), tensile failure will certainly
occur before shear failure due to reduction of effective normal
stress and increase of shear stress near the fault tip (Fig. 13b,
bottom), and splay cracks will propagate from the fault tips by
tensile failure. Therefore, the suitable stress state for propagation of
the splay cracks in the Wakkanai Formation is likely to be the case
where X
s
X
t
is positive, that is when X
s
>X
t
.
The above value of X
s
X
t
is written by the GrifthCoulomb
criterion as:
X
s
X
t
2S
T
cos f
i
0:5
1 sin f
i
s
r
m
where S
T
, f
i
and s
m
r
are the tensile strength, the angle of internal
friction, and the remote mean stress, respectively (Fig. 12). This
formula means that failure modes which occur at fault tips by the
fault slip depend on the tensile strength, the angle of internal
friction, and the remote mean stress at the time of fault slip, and, in
the case of rocks which do not signicantly vary in both the tensile
strength and the angle of internal friction, the lower remote mean
stress is suitable for propagation of tensile splay cracks. This
formula also means that higher values of the tensile strength
promote tensile failure near fault tips when fault slip occurs while
the lower values tend to induce shear failure, provided the remote
mean stresses are similar (and the angles of internal friction are
also). This can be consistent with the relationship between rock
strengths and failure modes suggested by eld observations of
siliceous shale units of the Monterey Formation of coastal California
(Gross, 1995; Dholakia et al., 1998; Eichhubl and Boles, 2000).
5.3. Estimation of X
s
X
t
distributions in the boreholes
In this study, the distributions of X
s
X
t
in the Wakkanai
Formation in each borehole were estimated using the above
formula, with tensile strengths and angles of internal friction
determined by the laboratory testing and the assumed remote
mean stresses. Tensile strengths were set as the strengths obtained
by Brazilian tests (0.83.2 MPa) at depths in 50100 m intervals in
each borehole (Fig. 14). Although the internal friction angles
Fig. 11. Histograms on rock mechanical properties of the Wakkanai Formation
measured by laboratory tests. (a) tensile strengths of core samples from HDB-6, HDB-9,
HDB-11 and PB-V01 examined in this study. (b) internal friction angles of core samples
from HDB-6 and boreholes HDB-1 and HDB-2 (Niunoya and Matsui, 2005).
Fig. 12. Proximity of stress state to brittle failure conditions is represented by the
smallest stress increment required to reach that stress state from either the shear
failure envelope, X
s
, or the tensile failure envelope, X
t
: see Bourne and Willemse
(2001); the cohesive strength, s
0
, is assumed to be equal to twice the tensile strength,
S
T
, following Brace (1960).
E. Ishii et al. / Journal of Structural Geology 32 (2010) 17921805 1799
obtained by the triaxial compression tests in the boreholes are only
from two locations in HDB-6, the average angle of internal friction
for the twelve core samples including core samples from the other
boreholes (HDB-1 and HDB-2) is 26.1 2.9
,
which are lower and higher than the average degree by 2s (5.8),
were represented as the lowangle case (Case 1) and the high angle
case (Case 2) respectively (Fig. 14). For the remote mean stresses,
the following was assumed: i) the principal remote stresses are
horizontal and vertical based on the results of the multiple inverse
Fig. 13. Modes of brittle failure near the fault tip induced by reduction of effective normal stress and/or increase of shear stress that are produced by fault slip in the Wakkanai
Formation. The failure modes are determined by the initial stress state. See Fig. 12 for the envelopes and circles.
Fig. 14. The estimated depth proles of X
s
X
t
estimated in each borehole. Results in Case 1 and Case 2 are based on a low friction angle (20.3
) respectively, and the results are almost the same. The lithostatic load estimated by density logging and the tensile strengths determined by laboratory testing are also shown.
E. Ishii et al. / Journal of Structural Geology 32 (2010) 17921805 1800
analyses using the fault slip data sampled in the horizontal outcrop,
ii) the vertical stress is lithostatic, iii) the pore pressure is hydro-
static since the present pore pressures of the sections examined in
the borehole are nearly under hydrostatic state (Kurikami, 2007;
Funaki et al., 2009) and further the elevated pore pressure, which
could have been raised during the early stage of folding following
burial diagenesis (e.g., Engelder, 1985), are inferred to be nearly
released by early fracturing predating the major FCBs development,
and iv) the remote mean stress is a vertical effective stress as the
sense of displacement of the major FCBs is predominantly strike-
slip (Fig. 6). The lithostatic load was calculated using data obtained
by density logging in each borehole (Fig. 14). Although the esti-
mation based on the above assumptions is accompanied by some
errors, the distribution of approximate X
s
X
t
could be determined.
The estimated depth proles of X
s
X
t
in Case 1 and Case 2 are
shown in Fig. 14. In this study, the domain where X
s
X
t
is positive
is dened as SDT (Suitable Domain for pervasive Tensile failure),
while, the domain where X
s
X
t
is negative is dened as UDT
(Unsuitable Domain for pervasive Tensile failure). The results using
upper and lower values for internal friction angle are almost the
same (Fig. 14). The SDT domains are estimated to be above 400 min
HDB-6, above 300 m in HDB-9, and above 450 m in PB-V01, while
the UDT domains were estimated to occur below these depths in
HDB-6, in HDB-9, and in PB-V01, respectively and below 500 m in
HDB-11 (the depth of the SDT domain in HDB-11 is unknown, but is
no deeper than 500 m). Thus, the SDT and UDT domains roughly
correspond to the rock mass above and below 400 m depth in each
borehole, respectively, if we take into consideration estimation
errors.
5.4. Comparison between the distribution of the SDT and UDT and
fractures in core
The presence of many splay cracks along a fault is an indication
of numerous slip events along a fault (Martel and Pollard, 1989).
Similarly, a fault associated with fault rocks such as breccia is likely
formed by several slip events. Hence, if tensile splay cracks are
easily developed by fault slip in an SDT, the number of tensile
fractures near major FCBs is expected to be higher in an SDT. Fig. 15
shows the frequencies (number per 10 m) of major FCBs and of
tensile fractures, the ratios of the frequencies of tensile fractures to
major FCBs, and the distributions, in terms of depth and bound-
aries, of the SDTand UDT in each borehole. Comparison of the ratios
of tensile fracture to major FCB frequencies and the SDT/UDT
Fig. 15. Frequency distributions of brittle fractures in the boreholes in the Wakkanai Formation. Frequency is expressed as a number per 10 m for FCBs, major FCBs, and tensile
fractures in each borehole. Also, tensile strengths, histograms showing the ratios of tensile fracture frequency to major FCB frequency, and the distribution of the SDT and UDT in
each borehole are provided.
E. Ishii et al. / Journal of Structural Geology 32 (2010) 17921805 1801
distributions, indicate that the higher ratios (e.g. 5) are restricted,
with one exception in HDB-9, to the SDT domain. The frequencies of
tensile fractures also are higher in the SDT than the UDT. This
evidence supports the above hypothesis. Although a considerable
number of tensile fractures are observed also in the range of 470
480 m depth in HDB-6 in the UDT, it could be the result of the
following: i) mechanical effects of slip on a considerable number of
FCBs observed at 440460 m depth (since tensile failure also can
occur by fault slip even if within a UDT as shown in Fig. 13), or ii)
error in estimation of the location of the SDT/UDT boundary
(perhaps several tens meters) due to the assumptions relevant to
the remote mean stresses at the time of fault slip.
Furthermore, the ratios of tensile fracture frequency to major
FCB frequency are near zero throughout the dened UDT, though
the frequency of major FCBs in the UDT (09 per 10 m; avg. 1.3 per
10 m) is similar to that in the SDT (09 per 10 m; avg. 1.4 per 10 m).
This observation implies that the principal mode of failure in a UDT
is shear. This implication is also supported by the fact that the
occurrences of en echelon shear fractures developing at and
beyond minor FCB tips in cores, i.e., those not associated with
splays, were observed more often in the UDT than the SDT. This is
based on both the frequency distribution of en echelon shear
fractures mentioned in Section 4.2, and the SDT and UDT domains
dened in Section 5.3.
Depths above 400 m in HDB-6 and the depths above 450 m in
PB-V01 where tensile strengths are higher appear to correspond to
the zones of high ratios of tensile fracture to major FCB frequencies
and high frequencies of tensile fractures. But the zones with high
ratios of tensile fracture to major FCB frequencies and high
frequencies of tensile fractures are also recognized above 300 m
depth in HDB-9, where tensile strengths are not higher. This shows
that the failure mode near fault tips during faulting in the Wak-
kanai Formation, depends not only on rock strength, but also the
remote mean stresses.
5.5. Growth model of the major FCBs
Concerning fault growth models in brittle sedimentary rocks,
models in sandstones have been developed in previous studies.
Two main ways that faults grow in sandstones are recognized
(Davatzes et al., 2003; Flodin and Aydin, 2004): i) deformation
band-based faulting; and ii) sheared-joint-based faulting. In the
rst model, faults grow by localization and amalgamation of indi-
vidual deformation bands to form deformation band zones with
subsequent coalescence of the zones and discontinuous slip
surfaces nucleated along deformation bands to form through going
deformation band-style faults (e.g. Aydin and Johnson, 1978;
Antonellini and Aydin, 1995; Shipton and Cowie, 2001). In the
second model, faults grow by linking with neighboring faults via
linkage structures such as splay cracks, which formed near the fault
tips by stress concentrations following slip nucleation on preex-
isting structures (e.g., Flodin and Aydin, 2004; Myers and Aydin,
2004). Although fault growth models in mudstones are seldom
known, Dholakia et al. (1998) suggested that faults grew by linking
via splay cracks produced by shearing along initial joints, i.e.
the sheared-joint-based faulting model, in siliceous shale units of
the Monterey Formation. In the case of the siliceous mudstone, the
Wakkanai Formation, it is inferred by this study that, during and/or
after uplift and erosion the major FCBs grewmainly by linking with
adjacent faults via numerous splay cracks, which formed by tensile
failure as shown in Figs. 8, 9 and Fig. 13c-bottom, in an SDT (i.e.
above roughly 400 m depth). Such growth mechanism in an SDT
domain is similar to the sheared-joint-based faulting model. In
contrast, in a UDT (i.e. belowroughly 400 m depth), the major FCBs
grew predominantly by shear failure, and could develop in direc-
tions parallel to the fault planes through coalescence of the en
echelon shear fractures as shown by Fig. 10, Fig. 13c, top and Fig. 16
though the developing direction would not necessarily be limited
to the above directions. A fault growth model in mudstones like the
mechanism in a UDT domain has not been reported, however, it
may be similar to the idealized fault growth model suggested by
Cowie and Scholz (1992).
5.6. Comparison between SDT and UDT and permeability of the
Wakkanai Formation
Based on the above, the linking of adjacent, major FCBs via
numerous tensile splay cracks is assumed to develop in an SDT.
Previous studies pointed out that such structures associated with
tensile fractures have important hydraulic properties. Eichhubl
et al. (2009) showed by using the distribution of fault-related
calcite cement at the Moab fault system in southern Utah as an
indicator of paleouid migration, that uid ow was focused
especially at locations of structural complexity such as fault inter-
sections, extensional steps, and fault-segment terminations, where
many tensile fractures developed. Dholakia et al. (1998) indicated
that increased hydrocarbon concentrations in the Monterey
Formation in the southern San Joaquin Valley and coastal California
are almost exclusively associated with faults which grew through
linkage via many tensile splay cracks. Khang et al. (2004) indicated
by numerical simulations that geometry of the structures linking
faults via tensile splay cracks plays an important role in uid ow
through the fracture network. Mazurek et al. (1998, 2003) and Lunn
et al. (2008) also imply that linkage by structures such as tensile
Fig. 16. A growth model of the major FCBs in the Wakkanai Formation during and/or
after uplift and erosion. When fault slip occurs in an SDT domain, splay cracks form
near the fault tips by tensile failure, resulting in faults strongly interconnected via
numerous tensile splay cracks. In contrast, when fault slip is in a UDT domain, shear
failure is predominant near fault tips and en echelon shear fractures could form at and
beyond the fault tips.
E. Ishii et al. / Journal of Structural Geology 32 (2010) 17921805 1802
splay cracks may represent preferential pathways for uid ow. In
consideration of the work of these authors, comparisons between
the distributions of SDT and UDT domains in each borehole and
variations in permeability in the Wakkanai Formation were done in
this study.
Fig. 17 shows the results of the comparisons in the each bore-
hole. The SDT domains generally have higher permeability (e.g.
hydraulic transmissivity: 10
5
m
2
/s) whereas the UDT domains
clearly show lower permeability (e.g. hydraulic transmissivity:
10
5
m
2
/s). However, some lower permeability sections are also
observed in the SDT domains in HDB-9 and PB-V01. These sections
have relatively lower frequencies of total number of FCBs and
tensile fractures and thus are not active in the main connected ow
paths in an SDT. This observation suggests that the hydrogeological
environment in an SDT differs from that in a UDT due to hydraulic
effects of the faults strongly interconnected via many tensile splay
cracks, and that these structures result in the restricted distribution
of the highly permeable sections to depths less than about 400 min
the Wakkanai Formation, as shown in Figs. 3 and 17.
Diagenetic alteration products can also provide some indication
of permeability of fractured rock masses (e.g., Milodowski et al.,
1998; Solumet al., 2005; Eichhubl et al., 2009). But in the Wakkanai
Formation, neither mineral precipitation on fractures nor alteration
along fractures has been observed in core samples though oxidized
zones along the major FCBs, which imply high permeability, are
observed at surface (Ishii and Fukushima, 2006). Moreover,
although the harder siliceous mudstones, which might be formed
due to additional silica cementation during uplift as suggested by
Iijima and Tada (1981), are observed in the subsurface, a relation-
ship between the harder siliceous mudstones and permeability of
the Wakkanai Formation has not been found. If the issue of whether
or not some diagenetic indicator for uid ow exists in the Wak-
kanai Formation is validated, further detailed X-ray analyses such
as SEM, XRD and XRF must be done on rock matrices and fracture
surfaces.
6. Conclusions
This paper discusses the growth mechanisms of the major FCBs
in the Wakkanai Formation by the following methods: i) geological
characterization by fracture mapping at an outcrop and by fracture
logging in several boreholes; ii) rock mechanical characterization
by laboratory tests on core samples for tensile strength and angle of
Fig. 17. Comparison of the SDT and UDT distributions and permeability in the Wakkanai Formation. Also, tensile strengths, the frequencies of total number of FCBs and tensile
fractures and the ratios of tensile fracture frequency to major FCB frequency in each borehole are shown.
E. Ishii et al. / Journal of Structural Geology 32 (2010) 17921805 1803
internal friction, and iii) theoretical analysis using the Grifth
Coulomb criterion. This paper suggests the following ideas:
1. The principal mode of failure, which occurs near fault tips by
fault slip in the Wakkanai Formation, depends not only on rock
strength, but also on the remote mean stresses.
2. During and/or after uplift and erosion the major FCBs grew
mainly by linking with adjacent major FCBs via numerous splay
cracks formed by tensile failure in an SDT (i.e. above roughly
400 mdepth), while, in a UDT (i.e. belowroughly 400 mdepth),
the major FCBs predominantly grow by shear failure, and could
develop in directions parallel to the fault planes through coa-
lescence of en echelon shear fractures though the direction of
development is not necessarily limited to the above directions.
3. The hydrogeological environment in an SDT differs fromthat in
a UDT due to hydraulic effects of the faults strongly inter-
connected via many tensile splay cracks. The preferential
development of the linking structures results in the restricted
distribution of the high permeability sections to less than about
400 m depth in the Wakkanai Formation.
Such conclusions are useful also for three dimensional geolog-
ical modeling and for groundwater ow simulations (e.g., for the
geological disposal of high level radioactive waste). However, in
order to conduct the more detailed modeling and simulations for
the Wakkanai Formation, it may be necessary to solve unanswered
questions, such as the comprehensive deformation history, and the
detailed hydraulic effects of the faults strongly interconnected via
many tensile splay cracks.
Acknowledgements
We thank G. McCrank and W.R. Alexander for English editing
and their helpful suggestions on this manuscript. We also express
our gratitude to H. Moir, J.G. Solum and Z. Shipton for their critical
review of the manuscript.
Appendix
The acronyms and their denition used in this study are as
shown in Table A1.
References
Andersson, J., Ahokas, H., Hudson, J.A., Koskinen, L., Luukkonen, A., Lo fman, J., Keto, V.,
Pitka nen, P., Mattila, J., Ikonen, A.T.K., Yla -Mella, M., 2007. Olkiluoto Site
Description 2006. POSIVA Technical Report POSIVA 2007-03, Olkiluoto, Finland.
Antonellini, M., Aydin, A., 1995. Effect of faulting on uid ow in porous sandstone:
geometry and spatial distribution. American Association of Petroleum Geologist
Bulletin 79, 642671.
Aydin, A., Johnson, A.M., 1978. Development of faults as zones of deformation
bands and as slip surfaces in sandstone. Pure and Applied Geophysics 116,
931942.
Bossart, P., Hermanson, J., Mazurek, M., 2001. Analysis of fracture network based on
the integration of structural and hydrogeological observations on different
scales. SKB Technical Report 01-21, SKB, Stockholm, Sweden.
Bourne, S.J., Willemse, E.J.M., 2001. Elastic stress control on the pattern of tensile
fracturing around a small fault network at Nash Point, UK. Journal of Structural
Geology 23, 17531770.
Brace, W.F., 1960. An extension of the Grifth theory of fracture to rocks. Journal of
Geophysical Research 65, 34773480.
Caine, J.S., Evans, J.P., Forster, C.B., 1996. Fault zone architecture and permeability
structure. Geology 24, 10251028.
Cowie, P.A., Scholz, C.H., 1992. Physical explanation for displacement-length rela-
tionship of faults using a post-yield fracture mechanics model. Journal of
Structural Geology 14, 11331148.
dAlessio, M.A., Martel, S.J., 2004. Fault terminations and barriers to fault growth.
Journal of Structural Geology 26, 18851896.
Davatzes, N.C., Aydin, A., Eichhubl, P., 2003. Overprinting faulting mechanisms
during the development of multiple fault sets, Chimney Rock fault array, Utah,
USA. Tectonophysics 363, 118.
Dholakia, S.K., Aydin, A., Pollard, D.D., Zoback, M.D., 1998. Fault-controlled hydro-
carbon pathways in the Monterey Formation, California. American Association
of Petroleum Geologist Bulletin 82, 15511574.
Du, Y., Aydin, A., 1993. The maximum distortional strain energy density criterion for
shear fracture propagation with applications to the growth paths of en e chelon
faults. Geophysical Research Letters 20, 10911094.
Du, Y., Aydin, A., 1995. Shear fracture patterns and connectivity at geometric
complexities along strike-slip faults. Journal of Geophysical Research 100,
1809318102.
Eichhubl, P., Boles, J.R., 2000. Focused uid ow along faults in the Monterey
Formation, coastal California. Geological Society of America Bulletin 112,
16671679.
Eichhubl, P., Davatzes, N.C., Becker, S.P., 2009. Structural and diagenetic control of
uid migration and cementation along the Moab fault, Utah. American Asso-
ciation of Petroleum Geologist Bulletin 93, 653681.
Engelder, T., 1985. Loading paths to joint propagation during a tectonic cycle: an
example from the Appalachian Plateau, U.S.A. Journal of Structural Geology 7,
459476.
Evans, J.P., Forster, C.B., Goddard, J.V., 1997. Permeability of fault-related rocks, and
implications for hydraulic structure of fault zones. Journal of Structural Geology
19, 13931404.
Flodin, E.A., Aydin, A., 2004. Evolution of a strike-slip fault network, Valley of Fire
State Park, southern Nevada. Geological Society of America Bulletin 116 (1/2),
4259.
Fukusawa, H., 1985. Late Neogene Formations in the Tempoku-Haboro region,
Hokkaido, Japan stratigraphic reinvestigation of the Wakkanai and Koetoi
Formations. The Journal of the Geological Society of Japan 91, 833849.
Funaki, H., Ishii, E., Tokiwa, T., 2009. Evaluation of the role of fracture as the major
water-conducting feature in Neogene sedimentary rocks. Journal of the Japan
Society of Engineering Geology 50, 239248.
Genter, A., Castaing, C., Dezayes, C., Tenzer, H., Traineau, H., Villemin, T., 1997.
Comparative analysis of direct (core) and indirect (borehole imaging tools)
collection of fracture data in the Hot Dry Rock Soultz reservoir (France). Journal
of Geophysical Research 102, 1541915431.
Gross, M.R., 1995. Fracture partitioning: failure mode as a function of lithology in
the Monterey Formation of coastal California. Geological Society of America
Bulletin 107, 779792.
Gutmanis, J.C., Lanyon, G.W., Wynn, T.J., Watson, C.R., 1998. Fluid ow in faults:
a study of fault hydrogeology in Triassic sandstone and Ordovician volcani-
clastic rocks at Sellaeld, north-west England. Proceedings of the Yorkshire
Geological Society 52, 159175.
Hiraga, N., Ishii, E., 2008. Mineral and Chemical Composition of Rock Core and
Surface Gas Composition in Horonobe Underground Research Laboratory
Project (Phase 1). JAEA Technical Report JAEA-Data/Code 2007-022. Japan
Atomic Energy Agency, Tokai-mura, Japan.
Iijima, A., Tada, R., 1981. Silica diagenesis of Neogene diatomaceous and volcani-
clastic sediments in northern Japan. Sedimentology 28, 185200.
Ikeda, Y., 2002. The origin and mechanism of active folding in Japan. Active Fault
Research 22, 6770.
Isaacs, C.M., 1982. Inuence of rock composition on kinematics of silica phase
changes in the Monterey Formation, Santa Barbara area, California. Geology 10,
304308.
Ishii, E., Fukushima, T., 2006. A case study of analysis of faults in Neogene siliceous
rocks. Journal of the Japan Society of Engineering Geology 47, 280291.
Ishii, E., Yasue, K., Tanaka, T., Tsukuwi, R., Matsuo, K., Sugiyama, K., Matsuo, S., 2006.
Three-dimensional distribution and hydrogeological properties of the Omagar
Fault in the Horonobe area, northern Hokkaido, Japan. The Journal of the
Geological Society of Japan 112, 301314.
Ishii, E., Hama, K., Kunimaru, T., Sato, H., 2007. Change in groundwater pH by
inltration of meteoric water into shallow part of marine deposits. The Journal
of the Geological Society of Japan 113, 4152.
Ishii, E., Yasue, K., Ohira, H., Furusawa, A., Hasegawa, T., Nakagawa, M., 2008.
Inception of anticline growth near the Omagari Fault, northern Hokkaido, Japan.
The Journal of the Geological Society of Japan 114, 286299.
Table A1
Acronyms and denitions.
Acronym Denition
FCB Fault crosscutting bedding planes at a high angle
Major FCB FCB composed of fault planes with associated
fault rocks such as breccia
Minor FCB FCB composed of fault planes with associated
slickenlines and/or slickensteps but not fault
rocks such as breccia
X Shortest distance from the GrifthCoulomb failure
envelope of rock strength to the Mohr circle
X
s
Shortest distance from the Coulomb failure envelope
of rock strength to the Mohr circle
X
t
Shortest distance from the Grifth failure envelope
of rock strength to the Mohr circle
SDT Suitable domain for pervasive tensile failure
UDT Unsuitable domain for pervasive tensile failure
E. Ishii et al. / Journal of Structural Geology 32 (2010) 17921805 1804
Japanese Industrial Standard Committee, 2000. Method of Test for Tensile Strength
of Rock JIS M 0303. Japanese Standards Association, Tokyo, Japan.
Khang, N.D., Watanabe, K., Saegusa, H., 2004. Fracture step structure: geometrical
characterization and effects on uid ow and breakthrough curve. Engineering
Geology 75, 107127.
Kovari, K., Tisa, A., Einstein, H.H., Franklin, J.A., 1983. Suggested methods for
determining the strength of rock materials in triaxial compression: revised
version. International Journal of Rock Mechanics and Mining Sciences 20,
283290.
Kulander, B.R., Dean, S.L., Ward Jr., B.J., 1990. Fractured core analysis: interpretation,
logging, and use of natural and induced fractures in core. In: AAPG Methods in
Exploration Series 8. AAPG, Oklahoma, USA.
Kurikami, H., 2007. Groundwater Flow Analysis in the Horonobe Underground
Research Laboratory Project: Recalculation based on the Investigation until
Fiscal Year 2005. JAEA Technical Report JAEA-Research 2007-036. Japan Atomic
Energy Agency, Tokai-mura, Japan.
Kurikami, H., Takeuchi, R., Yabuuchi, S., 2008. Scale effect and heterogeneity of
hydraulic conductivity of sedimentary rocks at Horonobe URL site. Physics and
Chemistry of the Earth Parts A/B/C 33 (Suppl. 1), S37S44.
Lunn, R.J., Willson, J.P., Shipton, Z.K., Moir, H., 2008. Simulating brittle fault growth
from linkage of preexisting structures. Journal of Geophysical Research 113,
B07403. doi:10.1029/2007JB005388.
Martel, S.J., Pollard, D.D., 1989. Mechanics of slip and fracture along small faults and
simple strike-slip fault zones in granitic rock. Journal of Geophysical Research
94, 94179428.
Mazurek, M., 1998. Geology of the crystalline basement of northern Switzerland
and derivation of geological input data for safety assessment models. NAGRA
Technical Report NTB 93-12, NAGRA, Wettingen, Switzerland.
Mazurek, M., 2000. Geological and hydraulic properties of water-conducting features
in crystalline rocks. In: Stober, I., Bucher, K. (Eds.), Hydrogeology of Crystalline
Rocks. Kluwer Academic Publishers, Dordrecht, Netherlands, pp. 326.
Mazurek, M., Bossart, P., Eliasson, T., 1996. Classication and characterization of
water-conducting features at A
spo
I: geological basis and model calibration. Journal of Contaminant Hydrology
61, 157174.
Milodowski, A.E., Gillespie, M.R., Naden, J., Fortey, N.J., Shepherd, T.J., Pearce, J.M.,
Metcalfe, R., 1998. The petrology and paragenesis of fracture mineralisation in
the Sellaeld area, west Cumbria. Proceedings of the Yorkshire Geological
Society 52, 215241.
Mitsui, K., Taguchi, K., 1977. Silica mineral diagenesis in Neogene tertiary shales in
the Tempoku district, Hokkaido, Japan. Journal of Sedimentary Petrology 47,
158167.
Murata, K.J., Nakata, J.K., 1974. Cristobalitic stage in the diagenesis of diatomaceous
shale. Science 184, 567568.
Myers, R., Aydin, A., 2004. The evolution of faults formed by shearing across joint
zones in sandstone. Journal of Structural Geology 26, 947966.
Niunoya, S., Matsui, H., 2005. The Investigation on Rock Mechanics in HDB-1
and HDB-2 Boreholes in order to Select the URL Area. JNC Technical Report
TN5400 2005-012. Japan Nuclear Cycle Development Institute, Tokai-mura,
Japan.
Nordqvist, R., Gustafsson, E., Andersson, P., Thur, P., 2008. Groundwater ow and
hydraulic gradients in fractures and fracture zones at Forsmark and Oskar-
shamn. SKB Technical Report R-08-103, SKB, Stockholm, Sweden.
O
hman, J., Niemi, A., Tsang, C.-F., 2005. Probabilistic estimation of fracture trans-
missivity from Wellbore hydraulic data accounting for depth-dependent
anisotropic rock stress. International Journal of Rock Mechanics and Mining
Sciences 42, 793804.
Ota, K., Abe, H., Yamaguchi, T., Kunimaru, T., Ishii, E., Kurikami, H., Tomura, G.,
Shibano, K., Hama, K., Matsui, H., Niizato, T., Takahashi, K., Niunoya, S., Ohara, H.,
Asamori, K., Morioka, H., Funaki, H., Shigeta, N., Fukushima, T., 2007. Horonoe
Underground Research Laboratory Project, synthesis of Phase I investigations
20012005. JAEA Technical Report JAEA-Research 2007-044. Japan Atomic
Energy Agency, Tokai-mura, Japan.
Petit, J.P., 1987. Criteria for the sense of movement on fault surfaces in brittle rocks.
Journal of Structural Geology 9, 597608.
Rhe n, I., Forsmark, T., Forssman, I., Zetterlund, M., 2006. Evaluation of hydro-
geological properties for Hydraulic Conductor Domain (HCD) and Hydraulic
Rock Domains (HRD). SKB Technical Report R-06-22, SKB, Stockholm, Sweden.
Segall, P., Pollard, D.D., 1983. Nucleation and growth of strike slip faults in granite.
Journal of Geophysical Research 88, 555568.
Shen, B., Stephansson, O., 1993. Numerical analysis of mixed Mode I and Mode II
fracture propagation. International Journal of Rock Mechanics and Mining
Sciences 30, 861867.
Sibson, R.H., 1977. Fault rocks and fault mechanisms. Journal of Geological Society
133, 191213.
Shipton, Z.K., Cowie, P.A., 2001. Damage zone and slip-surface evolution over mm to
km scales in high-porosity Navajo sandstone, Utah. Journal of Structural
Geology 23, 18251844.
Solum, J.G., van der Pluijim, B.A., Peacor, D.R., 2005. Neocrystallization, fabrics and
age of clay minerals from an exposure of the Moab Fault, Utah. Journal of
Structural Geology 27, 15631576.
Tada, R., Iijima, A., 1982. Petrology and diagenetic changes of Neogene siliceous
rocks in northern Japan. Journal of Sedimentary Petrology 53, 911930.
Twiss, R.J., Moores, E.M., 2007. Structural Geology, second ed. W.H. Freeman and
Company, New York.
Wei, D., Seno, T., 1998. Determination of the Amurian plate motion. In: Flower, M.,
Chung, S.L., Lo, C.H., Lee, T.Y. (Eds.), Mantle Dynamics and Plate Interactions in
East Asia. Geodynamics Series 27. American Geophysical Union, Washington,
D.C., USA, pp. 337346.
Wei, Z.Q., Egger, P., Descoeudres, F., 1995. Permeability predictions for jointed rock
masses. International Journal of Rock Mechanics and Mining Sciences 32, 251261.
Welch, M.J., Davies, R.K., Knipe, R.J., 2009. A dynamic model for fault nucleation and
propagation in a mechanically layered section. Tectonophysics. doi:10.1016/
j.tecto.2009.04.025.
Williams, J.H., Johnson, C.D., 2004. Acoustic and optical borehole-wall imaging
for fractured-rock aquifer studies. Journal of Applied Geophysics 55,
151159.
Willson, J.P., Lunn, R.J., Shipton, Z.K., 2007. Simulating spatial and temporal evolu-
tion of multiple wing cracks around faults in crystalline basement rocks. Journal
of Geophysical Research 112, B08408. doi:10.1029/2006JB004815.
Yabuuchi, S., Kunimaru, T., Ishii, E., Hatsuyama, Y., Ijiri, Y., Matsuoka, K., Ibara, T.,
Matsunami, S., Makino, A., 2008. Horonobe Underground Research Laboratory
Project, Overview of the Pilot Borehole Investigation of the Ventilation Shaft
(PB-V01): Hydrogeological Investigation. JAEA Technical Report JAEA-Data/Code
2008-026. Japan Atomic Energy Agency, Tokai-mura, Japan.
Yamaji, A., 2000. The multiple inverse method: a new technique to separate
stresses from heterogeneous fault-slip data. Journal of Structural Geology
22, 441452.
Yamamoto, H., 1979. The geologic structure and the sedimentary basin off northern
part of the Hokkaido Island. Journal of the Japanese Association of Petroleum
Technologists 44, 260267.
E. Ishii et al. / Journal of Structural Geology 32 (2010) 17921805 1805
Structural and petrophysical evolution of extensional fault zones in low-porosity,
poorly lithied sandstones of the Barreiras Formation, NE Brazil
F. Balsamo
a,
*
, F. Storti
a
, F. Salvini
a
, A.T. Silva
b
, C.C. Lima
b
a
Universita` degli Studi Roma Tre, Dipartimento di Scienze Geologiche, 00146 Roma, Italy
b
Cenpes, Petrobras, Rio de Janeiro, Brazil
a r t i c l e i n f o
Article history:
Received 5 February 2009
Received in revised form
2 October 2009
Accepted 5 October 2009
Available online 6 November 2009
Keywords:
Grain size reduction
Fractal distribution
Porosity
Cataclasis
Dilatancy
Fault-zone permeability
a b s t r a c t
We describe the structural and petrophysical evolution of extensional fault zones developed in low
porosity, poorly lithied, quartz-dominated sandstones from the Mio-Pliocene continental Barreiras
Formation, NE Brazil. We studied eight fault zones developed as sands were lithied. Fault displacement
ranges from a few centimetres to w50 m. A diagnostic feature of the studied fault zones is the lack of
deformation bands, which typically develop in high porosity sand(stone)s. Structural and microstructural
analyses, grain size and shape analyses, porosity and pore size analyses, and laboratory and in situ
permeability measurements show relationships between deformation processes and hydrologic prop-
erties. Undeformed rocks are very poorly sorted, medium- to ne-grained, clay-rich sandstones with an
average intergranular porosity of about 3%. Sandstones in damage zones record non-destructive dilatant
granular ow and formation of opening-mode intergranular extensional fractures, which increase
porosity, pore connectivity and permeability. Deformation in fault cores evolved from particulate ow to
compactional cataclastic ow, with progressive grain size reduction increasing the amount of silt- and
clay-size fractions. Porosity was dramatically reduced to an average value of 0.2% and permeability is
generally lower than the related protoliths. All this evidence highlights a conduit/barrier behaviour of the
studied fault zones, which signicantly differs from the sealing behaviour of deformation band fault
zones commonly observed in high-porosity sandstones.
2009 Elsevier Ltd. All rights reserved.
1. Introduction
Porosity and its evolution through time are an important factor
controlling what kind of mesoscopic deformation structures
develop during rock failure (e.g. Dunn et al., 1973; Vernik et al.,
1993; Kwon et al., 2005; Fossen et al., 2007). In the last four
decades, considerable attention has been devoted to the under-
standing of fault zone evolution in high-porosity (>1015%), lithi-
ed to loose granular material both in the eld (Aydin, 1978; Aydin
and Johnson, 1978; Pittman, 1981; Lucas and Moore, 1986; Anto-
nellini and Aydin, 1994; Fowles and Burley, 1994; Fossen and
Hesthammer, 1997; Heynekamp et al., 1999; Cashman and Cash-
man, 2000; Shipton and Cowie, 2001; Rawling and Goodwin, 2003;
Flodin et al., 2003, 2005; Johansen et al., 2005; Minor and Hudson,
2006) and in experimental studies (Borg et al., 1960; Mandl et al.,
1977; Mene ndez et al., 1996; Wong et al., 1997; Zhu and Wong,
1997; Mair et al., 2000). This because of their important inuence
on uid ow in hydrocarbon reservoirs and groundwater aquifers
(e.g. Haneberg, 1995; Walsh et al., 1998; Heynekamp et al., 1999;
Aydin, 2000; Fisher and Knipe, 2001; Rawling et al., 2001;
Manzocchi et al., 2002; Nelson et al., 2009). Deformation in high-
porosity granular materials occurs by development of small
displacement deformation structures comprehensively referred to
as deformation bands, which evolve into zones of deformation
bands and slip surfaces with increasing offset (e.g. Aydin and
Johnson, 1978; Fowles and Burley, 1994; Shipton and Cowie, 2001;
Fossen et al., 2007). A typical result of deformation band faulting in
high-porosity sandstones is that their extensive development in
fault damage zones may reduce fault transmissibility, thus
providing an effective barrier to uid ow (e.g. Antonellini et al.,
1994, 1999; Sigda et al., 1999; Rotevatn et al., 2007). This hydraulic
behaviour differs from the typical conduit behaviour of fault
damage zones in low-porosity fully lithied rocks, where defor-
mation is dominated by opening-mode fracturing (e.g. Caine et al.,
1996; Billi et al., 2003; Kim et al., 2004).
The lower threshold porosity limit for deformation band
development is at about 1015% (e.g. Dunn et al., 1973; Flodin et al.,
2003; Wong et al., 1997). Below this threshold limit, shear strength
* Corresponding author. Tel.: 39 0657338049; fax: 39 0657338201.
E-mail address: balsamo@uniroma3.it (F. Balsamo).
Contents lists available at ScienceDirect
Journal of Structural Geology
j ournal homepage: www. el sevi er. com/ l ocat e/ j sg
0191-8141/$ see front matter 2009 Elsevier Ltd. All rights reserved.
doi:10.1016/j.jsg.2009.10.010
Journal of Structural Geology 32 (2010) 18061826
becomes a fundamental parameter controlling deformation
mechanisms. Joints and slip surfaces are expected to develop in
fully lithied sandstones (e.g. Johansen et al., 2005; Fossen et al.,
2007). On the other hand, deformation in low-porosity poorly
lithied sand(stone)s is still poorly understood. In this paper, we
attempt to bridge the gap by describing the structural and petro-
physical evolution of extensional fault zones developed in low-
porosity, poorly lithied quartz-dominated sandstones of the Bar-
reiras Formation, NE Brazil. The relative compositional maturity
and homogeneity of the Barreiras sandstones allow us to discount
the effects of clay smearing and tectonic mixing of strongly
different sedimentary units within fault zones (e.g. Antonellini and
Aydin, 1994; Gibson, 1998; Heynekamp et al., 1999; Caine and
Minor, 2009). Results of structural, microstructural, grain size, grain
shape, and porosity analyses and permeability measurements are
described with the aim of (1) inferring the deformation mecha-
nisms that governed the evolution of these extensional fault zones;
(2) proposing an evolutionary model of grain size, grain shape and
porosity changes during extensional faulting; and (3) assessing the
inuence of faulting on uid ow by establishing a relationship
between fault-related permeability variations and fault displace-
ment. The latter provides a useful tool for predicting the expected
permeability and transmissibility of sub-seismic and seismic fault
zones in sand-dominated clastic reservoirs.
2. Analytical methods
Structural analysis was used to constrain the mesoscale archi-
tecture and kinematics of the studied extensional fault zones. Where
offset markers were available, stratigraphic separations were
measured in the eld and then converted into true fault displace-
ment values by using fault kinematics (Butler and Bell, 1989). Fault
core thicknesses were measured to determine whether or not there
was a predictive statistical relationship that could be used for esti-
mating fault displacement from fault core width when direct
measurements were not possible (e.g. Walsh and Watterson, 1988).
Undeformed sandstones and rocks in damage zones and fault
cores were sampledat eachstudied eldsite. About 0.5 kgof material
was collected for each sample. After complete disaggregation and
chemical removal of Fe-oxides in the laboratory, grain size analysis of
about 60gof granular material was completedbycombiningstandard
sieve and laser diffraction analyses in order to account for coarser
(gravel- and sand-size) and ner (silt- and clay-size) fractions,
respectively (Selley, 2000). A total of 44 samples were analysed
including 17 undeformed sandstones, 16 fault core rocks, and 11
damage zone sandstones. Results of grain size data were plotted as
frequency distribution curves using the Phi scale arrangement
(Krumbein, 1934, 1938). Grain size distributions were quantitatively
described in terms of the following statistical parameters: (a) the
mean size Phi
m
(a measure of the average size of the curve); (b) the
standard deviation So (a measure of the size spread around the mean
value, or sorting); (c) the skewness Sk (a measure of the curve
symmetry around the mean value, or preferential spread to one side
of the mean); and (d) the kurtosis (the degree of concentration of the
grain sizes relative to the mean value) (Inman, 1952; Grifths, 1952).
These parameters were obtainedby mathematical methods (Folkand
Ward, 1957; Krumbein and Pettijohn, 1938) employing the entire
sample grain size populations (McManus, 1988). Grain size distribu-
tions were also transformed into equivalent particle numbers (e.g.
Storti et al., 2003) assuming spherical particles with a density of 2.65
g/cm
3
and plotted against the equivalent size classes in bilogarithmic
diagrams to obtain their fractal dimensions (D) as the slope of the
best-t lines (e.g. Blenkinsop, 1991).
Additional non-destructive sampling was carried out in the
three structural domains (undeformed sands, damage zones and
fault cores) for blue-dyed epoxying and thin section analysis.
Microstructural analyses were carried out with a standard petro-
graphic microscope connected to a digital photo camera.
Computer-based image analysis techniques (e.g. Francus, 1998;
Heilbronner and Keulen, 2006) were applied to thin section images
using the Optimas-6.5
vh
vx
cos q sin q
(1)
where h(x,t) is the thickness of the current, x is the alongslope
position, m is the viscosity and k is the effective permeability of the
uid as it migrates through the rock, where we account for
the effects of relative permeability in a very simple fashion with the
single permeability parameter. Most of the interest in this work is
in modelling gas or supercritical uid dispersion and we model the
motion through the rock in terms of an effective permeability. This
has been shown to give good leading order predictions compared to
the full two-phase ow relations for such buoyancy driven ows
(cf. Nordbotten and Celia, 2006; Hesse et al., 2006). The rate of loss
of uid from the current to the overlying low permeability layer,
through unit length of the boundary, depends on the permeability
k
b
, the thickness b of the seal layer, and the thickness of the current
(cf. Pritchard et al., 2001) according to the relation
Loss
k
b
gDrh bcos q
bm
(2)
These equations are then combined with the relation for the
conservation of mass, which in the invading ow has the form
f1 s
w
1 R
vh
vt
v
vx
hu Loss for
vh
vt
> 0 (3)
since as the invading gas advances into the formation, there is
a fraction s
w
of the pore space which remains saturated in the
original uid, and gas then dissolves into this uid, representing an
effective additional pore volume fs
w
R for the injected uid. Here
R denotes the mass fraction of gas dissolved in the original uid
(which occupies the fraction s
w
of the pore volume) multiplied by
the density of the original uid and divided by the density of the
free gas phase. Also, in this expression f denotes the porosity of the
rock.
In contrast, in any part of the ow where the depth of the
current decreases with time, then as the buoyant uid vacates the
pore space and is displaced with water, there will be some residual
gas trapped which occupies a fraction s
g
of the pore spaces. Here,
for simplicity, we model this as being a constant (cf. Barenblatt,
1996; Hesse et al., 2006, 2008) and so the conservation of mass
takes the form
f
1 s
w
s
g
vh
vt
v
vx
hu Loss for
vh
vt
< 0 (4)
In order to solve for the motion of the current we require some
boundary conditions. First, it follows that the nose of the current
propagates at the rate
dx
dt
u
f1 s
w
Rs
w
(5)
while we assume that the source ux, at x 0, gradually wanes, at
a rate
Qt Q
o
expt=s
kDrgsin q
m
1 cot q
vh0; t
vx
h0; t
(6)
Fig. 1. Cartoon of the ow geometry for the present problem.
A.W. Woods, S. Norris / Journal of Structural Geology 32 (2010) 18271833 1828
where t is the e-folding time over which the source ux decays. It is
this waning source ux, coupled with the dynamics of continuous
leakage of uid through the overlying seal rock, and the capillary
retention at the tail of the current, which provides the newanalytic
results of this paper.
From eq. (6), we deduce that there is no drainage if the source
ux Q
o
is smaller than a critical value
Q
o
<
h
c
gDrsin q
m
Qcrit (7)
3. Approximations and analytical solutions
With this system of equations, we can now develop a series of
solutions for the motion of the plume of gas along the inclined low
permeabilitylayer. Thesesolutions areuseful for exposingsomeof the
key controls on the distance travelled along the layer, and also how
the current partitions between that component which is retained in
the original layer and that component which migrates through the
low permeability partial seal layer and higher into the formation.
Before developing solutions for the motion, there are some
simplications which we can introduce which simplify the analysis,
and allow for an approximate analytical solution. First, as the
current spreads out and disperses into a relatively long and thin
ow, of lateral scale L and depth H say, such that L >Hcot q, then
the alongslope component of gravity, proportional to sin q, domi-
nates the force associated with the cross-slope component of
gravity which acts on variations in the alongslope depth of the
current and is proportional to cos q vh=vx . In this limit, h cot q <L,
the dynamical term proportional to cos qv=vxhvh=vx,which arises
in the rst term on the right had side of eqs. (3) and (4), as may be
inferred by combining these eqns with eq. (1), can be neglected.
Indeed, we demonstrate that our analytical solutions are consistent
with this approximation.
Also, in the limit that the critical current depth required for
draining, h
c
, satises h
c
>b, the depth of overlying lowpermeability
layer, then the numerator (h b) in the loss term can be approxi-
mated by h (>h
c
) since the loss only arises if the ow is sufciently
deep to overcome the capillary entry pressure (cf. Woods and Far-
cas, 2009a).
It is also convenient to introduce the scaling for the speed
U
kDrgsin q
mf
1 s
w
s
g
; (8)
the dimensionless ratio of the speed of the front and the tail of the
current, as given by
l
1 s
w
1 R
1 s
w
s
g
(9)
and the inverse of the time-scale for the draining ux
b U
k
b
cos q
kbsin q
(10)
With a waning source, these approximations lead to the governing
equations
l
vh
vt
U
vh
vx
bh for h > h
c
and
vh
vt
> 0 (11)
l
vh
vt
U
vh
vx
for h < h
c
and
vh
vt
> 0 (12)
and
vh
vt
U
vh
vx
bh for h > h
c
and
vh
vt
< 0 (13)
vh
vt
U
vh
vx
for h < h
c
and
vh
vt
< 0 (14)
with the boundary conditions that at the source,
1 s
w
s
g
Uh0; t Q
o
expt=s (15)
and that at the nose of the current, x x
n
(t),
hx
n
; t 0 and
dx
n
dt
U
l
(16)
It may be seen from the denition of l that l >1, and so the
advection speed of the current in the region in which the current is
invading new rock, U/l, is slower than the advection speed of the
current in the region in which it is receding from the rock, U. This
means that the nose of the current in which the depth decreases
from a maximum to zero, occurs across a localized region whose
detail depends on the cross-slope component of gravity. In this
simplied model, this is represented by a localised front at x x
n
(t).
The structure of the current behind this front depends on the
source ow rate compared to the draining rate, and we now
consider a range of cases in turn.
3.1. Small supply ux with no leakage current
If Q<Q(crit), then there is no drainage into the overlying layer,
and as the current moves forward along the boundary, the source
ux gradually wanes, leading to a waning plume. In this case, with
the draining term neglected, the solution of the equation for the
current depth, eq. (14), can be written in the form
hx; t h0; 0exp
t
s
x
Us
(17)
It follows that surfaces of constant depth advance forward at
a speed given by U (eg see Fig. 2 below). This is faster than the speed
of the leading front, U/l, and so the leading edge of the current
gradually becomes shallower with time. This is a result of the loss of
uid through capillary retention as the depth of the current at
a given point in space behind the leading front gradually decreases
in time.
Indeed, by direct substitution, it follows that the depth of the
current at the leading edge, x
n
Ut/l, has the form
distance
time
x=Ut /
dx/dt=U
Depth remains
constant along
Leading edge
Fig. 2. illustrates the evolution of the ow in terms of the motion of surfaces of
constant depth in the xt plane.
A.W. Woods, S. Norris / Journal of Structural Geology 32 (2010) 18271833 1829
hUt=l; t h0; 0exp
l 1t
ls
(18)
We can also calculate the area of the zone of the rock which is
invaded by the current. This can be used to estimate the volume of
the residual uid which is trapped in the pore space once the
current recedes, although we note that, in time, this trapped uid
may dissolve into the water, and be carried off by any hydrological
ow. The solution above for the shape of the current as a function of
distance and time illustrates that at each point in space, the current
is deepest on rst arriving at that point. The current rst arrives at
each point x after a time lx/U (Fig. 2), and so the maximum depth of
the current at a distance x from the source, h
max
(x), is
h
max
x h0; 0exp
xl 1
Us
(19)
This curve describes the locus of the zone in which there may be
some residual plume uid once the plume has drained and moved
on, and hence in which there may be a possible source of
contaminant in a subsequent hydrological ow. In Fig. 3 below, we
illustrate the envelope of the zone contaminated with gas and
compare this with the instantaneous proles of the buoyant plume
at different times.
3.2. Larger source ux and drainage
With a larger source ux, Q>Q(crit), then initially there will be
some drainage into the overlying layer, with h >h
c
. In the region of
the current where h >h
c
the solution may be written in the form,
hx; t h0; 0exp
t
s
x
Us
1 bs
(20)
and the rate of propagation of surfaces of constant depth is now
faster than in the case with no draining, as the uid leaks off
through the overlying layer. We will now see that these new
solutions are very different fromthe case with no draining (sect 3.1)
3.2.1. Slow draining or rapid decay of the source ux
In the case 1 >bs, the depth of the current h increases with
distance from the source at a given time since the draining of the
uid is slow compared to the decay of the source and hence the
uid at the source has the smallest ux and so is shallowest; as
a consequence, the depth rst decreases to value h h
c
at x 0
when t t
c
, as given by (see Fig. 4 and 5)
t
c
s ln
h
c
h0; 0
(21)
Subsequently, as the ux continues to wane, the depth of the
current near the source decreases to values h <h
c
(Fig. 5) and the
location of the point at which the depth has value h
c
migrates away
from the source and has position x x
c
(t). In the region 0 <x <x
c
,
the current does not drain since h <h
c
. As the zone in which h <h
c
advances outwards from the source, the leading part of this region,
where h h
c
has position given by (Fig. 4)
x
c
t
Ut t
c
1 bs
(22)
Meanwhile the leading edge of the current has position x
n
Ut/l
(cf. eq. (16)), and so the zone in which draining occurs advances
progressively further from the source. Eventually, the front
x x
c
(t),which represents the closest point to the source at which
the depth has value h
c
and hence can drain, reaches the leading
edge of the ow. This occurs at time
t t
d
t
c
1
bs
l
1
l
1
(23)
Subsequently, there is no more draining anywhere in the ow
(Fig. 4). For times t >t
c
the closest point to the source at which the
depth has value h
c
is given by the front x x
b
(t) where
X
b
t Ut t
c
(24)
This lags behind the front x x
c
, and between these fronts, in
the region x
b
<x <x
c
, the depth has the constant value h
c
but there
is no draining. In the near source region, 0 <x <x
b
, in which the
current depth h <h
c
, the plume has shape (cf. eq. (17) and Fig. 4)
hx; t h0; 0exp
t
s
x
Us
(25)
In this near source region, the depth increases with distance
from the source, and reaches the critical depth h h
c
at the point
x
b
(t). Eventually, at time t
b
, the nearest point to the source at which
the current increases to depth h
c
, as given by x x
b
, reaches the
leading edge of the current, so that x
b
(t) x
n
(t). This occurs at time
t t
b
given by
t
b
lt
c
=l 1 (26)
0
0.2
0.4
0.6
0.8
1
1.2
0.00 0.50 1.00 1.50
d
i
m
e
n
s
i
o
n
l
e
s
s
h
e
i
g
h
t
dimensionless distance from source
time=0.5 1.0 1.5
Fig. 3. Comparison of the envelope of the zone invaded by the injected uid and the
instantaneous shape of the injected uid plume at times 0.5s, 1.0s and 1.5s.
Fig. 4. (x,t) plot to illustrate the evolution of the fronts x x
b
, x
c
and x
n
as they evolve
with time in the current. At times earlier than t
c
the current is deeper than h
c,
but for
times greater than t
b
the whole current is thinner than the critical depth h
c.
For
intermediate times, the near source region, x <x
b
, is shallower than h
c
while the more
distal parts of the current are either of depth h
c
, for x
b
<x <x
c
or of depth greater than
h
c
for x
n
>x >x
c
.
A.W. Woods, S. Norris / Journal of Structural Geology 32 (2010) 18271833 1830
Subsequently the current is described by relation (17), and is
everywhere shallower than h
c
. This sequence of ow regimes is
illustrated in the Fig. 4 and 5 shown below.
3.2.2. Fast draining or slow decay of the source
In the case that 1 <bs, and with Q>Q
c
, the current is initially
deeper than the critical value for draining, h >h
c
, and the ow
initially advances with prole
hx; t h0; 0exp
t
s
x
us
1 bs
(27)
in the region 0 <x <Ut/l (Fig. 6). In this case, at a given time, the
current becomes shallower with distance from the source, as
a result of the draining occurring more rapidly than the rate of
decay of the source, so that the owfurther fromthe source has less
ux than that at the source (Fig. 7, dashed lines). As a result, the
leading front of the current eventually reaches the critical depth at
which draining ceases, h h
c
. This occurs when
t t
d
t
c
1
bs
l
1
l
1
(28)
Subsequently, the leading edge of the current continues forward
with depth h h
c
while the closest point to the source at which the
depth of the current h h
c
, as given by x x
c
, migrates backwards
towards the source according to the relationship (cf. eq. (22) and
Fig. 6)
x
c
U
t t
c
1 bs
(29)
This front eventually reaches the source when t t
c
. (Fig. 6).
Subsequently, for t >t
c
, the depth of the current at the source
decreases to values h <h
c
and is given by the original solution (17)
in the region 0 <x <x
b
(t). From this solution, it follows that the
point nearest to the source at which the current depth equals the
critical depth h
c
has position
x
b
Ut t
c
(30)
This front eventually catches up with the leading edge of the
current, which advances at the rate
X
n
Ut/l, at the time given by (cf. eq. (27) and Fig. 6)
t t
b
l
t
c
l 1
(31)
Subsequently, the whole ow evolves according to the simple
non-draining solution (17) (see Fig. 7, dotted line).
3.3. Fraction which drains
In general the fraction of the ow which drains depends on the
capillary pressure, which suppresses the draining, the source ow
rate, the ratio of the draining time to the decay time of the source,
bs, and also the residual saturation of the water and the gas at the
advancing and receding fronts, as expressed by l.
In general the expression is complex to calculate, but is found
by comparing the volume input at the source with the volume
which remains in the formation, with the difference represent-
ing the fraction which has drained. There is however a useful
limit when h
c
<h(0) in which case, to leading order, the fraction
retained in the original layer may be found by integration of eq.
(20) evaluated at t lx/u. This leads to the result that the frac-
tion of the source uid which remains trapped in the formation,
F, is given by
F
l 1
l 1 bs
(32)
where we note that l >1 (eq. (9)). This expression effectively
compares the process of capillary trapping at the nose and tail of
the ow with the drainage through the upper boundary. It illus-
trates that if the draining time 1/b is short compared to the decay
time of the source, t, then F will become relatively small, and much
of the injected uid can drain away, whereas if the draining time is
comparable to or longer than the decay time of the source, then
much of the injected uid remains in the original layer. As the
capillary pressure increases, this further restricts the fraction of the
ow which drains, and so the above expression provides a lower
bound on the fraction of the ow which remains trapped in the
original layer of the formation.
We illustrate the variation of F with bs for a series of represen-
tative values of l (0.05, 0.1 and 0.15) in Fig. 8 below.
Fig. 5. Illustration of the evolution of the current with time. Each prole corresponds
to a vertical line (ie constant time) in the (x,t) plane of Fig. 4; in this case, the proles at
t/4 and t/2 lie in the range (t
d
<t <t
b
), while the proles at 3t/4 and t correspond to
times greater than t
b
.
Fig. 6. Illustration of the structure of the current on an xt plot. The gure shows how
the various transition points in the current evolve with time. For t <t
d
the ow is
everywhere deeper than h
c
. For t
d
<t <t
c
the near source region is deeper than h
c
while the distal part of the ow, x >x
c
, has constant depth h
c
. For t >t
c
the near source
region has become shallower than h
c
, while for x >x
b
the ow has constant depth.
At late times, t >t
b
the ow is everywhere shallower than h
c.
A.W. Woods, S. Norris / Journal of Structural Geology 32 (2010) 18271833 1831
4. Draining through faults
In comparison with the above results in which the draining
occurs through the upper boundary of the formation, we now
consider the case in which uid leaks off through a localised fault
which cuts across the layers. Typically faults are narrow compared
to the length scale of the ow, but provide a higher permeability
route to the surface. If the fault connects the ow in the lower
owing layer to a layer of high permeability above the seal layer,
then the ux through this fault will have the form (cf. Pritchard,
2007)
Q
fault
k
f
Drg cos q hw
mb
Uh (33)
where w is the width of the fault, and b the vertical thickness of the
fault, across which the gas pressure acts to drive the ow through
the fault. k
f
is the permeability of the fault, and h is the current
thickness just upstream of the fault. Here we assume that h >b so
that the hydrostatic pressure driving the ow through the fault is
associated with the buoyancy of the plume of injected uid in the
lower owing layer.
The fault ux Q
fault
represents a discontinuity in the ux of gas
along the layer. Since the alongslope ux scales as Uh, it follows that
the drainage through the fault dominates the ux along the
formation if U>U, and in this case, there will be no ux beyond the
fault, which for convenience we assume is located at x x
f
.
In this case, the fraction which remains in the formation is given
by the fraction which is trapped by capillary retention upstream of
the fault
F
1 exp
x
f
l 1
Us
(34)
Here, the critical balance is between the distance the fault lies
away from the source and the distance that would be travelled by
the plume over the time required for the source to decay, Us.
5. Application
It is useful to examine the implications of the model for a typical
example of the ow in a layered permeable rock. In the case of
a geological waste repository, there may be a ux of buoyant gas
with a decay time of order 300 years, and an initial ux of 10
5
m
2
/s
per unit length of the repository. If there is a layer of rock of
permeability 10
15
m
2
bounded above by a layer of permeability
10
17
m
2
, then with a porosity of 0.1 and a layer inclination of 10
o
,
the along layer velocity scale U has value of order 10
8
m/s with
uid of viscosity 10
4
Pa s. If the overlying seal layer has thickness
of order 1 m, then b has value of order 10
10
s
1
and so bs w1,
suggesting that the draining and the decay of the source occur over
approximately comparable times. If the capillary entry pressure to
the overlying layer is small, then the fraction of the owretained in
the layer is given by relation (32), within the simplied framework
of this model, and this has value of about F w0.10.2.
In the case of a rapidly decaying source or a current with slow
drainage rate, 1 <bs, then the drainage ux through the seal layer
F
D
(x,t), per unit length along the current, which is supplied to
points higher in the formation, is given by
F
D
x; t bh0; 0exp
t
s
x
Us
1 bs
1 bs
< x < Ut=l if t
c
< t < t
d
35
While in the case of a slowly decaying source or current with
high drainage rate, bs >1, the drainage ux, again given by the
same expression as in (35), is always located near to the source,
with drainage in the region 0 <x <Ut/l if t <t
d
and drainage in the
region 0 <x <U(t t
c
)/(1 bs) when t
d
<t <t
c
. Subsequently there
is no draining.
With a capillary entry pressure corresponding to a depth of
order 1 m, then with the above values for U, b and s it follows that
t
c
wt at which time the current has travelled a distance of order
100 m. The draining zone then evolves away from the source in the
slow draining case, until time t
d
w(1.11.2) t
c
. Similarly in the fast
draining case, the current will propagate about 100 m from the
source, while the draining persists, with the illustrative parameters
given in the example above. The plume will then cease to drain and
will migrate along the original layer as a thin, elongate ow. For
smaller capillary entry pressure, the ow may drain for times cor-
responding to several multiples of t, and hence the draining region
may extend several hundred metres alongslope.
In a different situation of CO
2
sequestration, the injection period
may only be of order 30 years. If the injectivity of the formation
D
i
m
e
n
s
i
o
n
l
e
s
s
D
e
p
t
h
e c n a t s i D s s e l n o i s n e m i D
Fig. 7. Illustration of the case in which for early times, t <t
d
, there is a draining zone
near the source (long and short dashed lines) and that as the current loses mass
through draining it reaches the critical depth at the nose of the ow, x x
n
, when t t
c
.
For t
c
>t >t
d
the location of the point closest to the source at which the depth equals
the critical value h
c
progressively migrates back to the origin which it reaches at t t
c
(dot-dashed line) . For t >t
c
, the depth is smaller than h
c
in the region x <x
b
, while the
more distal part of the ow, x >x
b
, has constant depth h
c
(dotted line). Eventually, for
t >t
b,
the whole ow is shallower than h
c.
0
0.2
0.4
0.6
0.8
1
0 0.2 0. 4 0.6 0. 8 1
f
r
a
c
t
i
o
n
r
e
t
a
i
n
e
d
i
n
t
h
e
o
r
i
g
i
n
a
l
l
a
y
e
r
Time of Decay of Source / Draining Time
0.15
0.05
Fig. 8. Illustration of the variation of the fraction of the ow which remains trapped in
the original layer of the formation as a function of the time of decay of the source
compared to the draining time across the thin partial seal layer.
A.W. Woods, S. Norris / Journal of Structural Geology 32 (2010) 18271833 1832
becomes impeded with time, the continuing injection ux may
then decay with time, and the present model may give a guide to
the ow. Initially, the ux per unit length injected into a long
horizontal well may again be of order 10
5
m
2
/s, and if the
formation has similar properties to the example above, this would
correspond to the case bs w0.1 which represents a rapidly decay-
ing source compared to the drainage rate. Now the current would
continue draining for a period t
d
w5s which is about 150 years,
in which time it would propagate about 50 m from the source. We
note that in the case of a maintained steady ux, the drainage
dynamics are somewhat different, as described by Woods and
Farcas (2009).
From both these idealized examples, we see that with a decay-
ing source, the injected uid may rapidly spread alongslope and
hence thin out, thereby limiting the fraction of the ow which
drains into the overlying formation compared to the fraction which
becomes capillary trapped in the original owing layer.
6. Summary
Using a simplied approach, we have identied and modeled
some of the controls on the migration of buoyant uid through
a layered permeable rock issuing from a waning source of buoyant
uid. We have focussed on the dynamics of the current in a single
layer of the formation, examining the balance between leakage
from the layer, and lateral spreading of the current along that layer.
In modelling the leakage, we have accounted for the capillary
entry pressure into an overlying seal layer, and shown that this
leads to a localized region of leakage which evolves in time. As the
current wanes, the capillary entry pressure suppresses further
leakage and the remainder of the current migrates through the
original layer. Capillary trapping of the uid in the original layer
leads to a continual loss of uid from the ow, and as the plume
disperses, it is eventually trapped within this layer. The balance
between the fraction of the ow which is trapped in the original
layer, and the fraction which leaks into the overlying layer depends
on the ratio of draining time through the overlying layer compared
to the decay time of the source. With a rapidly decaying source,
most of the uid remains trapped in the original layer, whereas
with a slowly decaying source, much of the uid is able to leak into
higher parts of the geological formation.
We have also shown that if the current reaches a fracture, then
a signicant part of this current may be diverted through the
fracture and then migrate higher into the formation. The critical
controlling parameter in this case is the ratio of the distance of the
fracture from the source to the product of the Darcy ux and the
decay time of the source. The further the fracture from the source
the greater the fraction of the ow which is sequestered in the
original layer in which the buoyant uid is injected.
References
Barenblatt, G.I., 1996. Dimensional Analysis, Self-Similarity and Intermediate
Asymptotics. CUP.
Bear, J., 1972. Dynamics of Flow in Porous Media. Elsevier, pp. 1746.
Bickle, M., Chadwick, A., Huppert, H.E., Hallworth, M., Lyle, S., 2007. Modelling
carbon dioxide accumulation at Sleipner: implications for underground carbon
storage. Earth Planet. Sci. Lett. 255, 164176.
Farcas, A., Woods, A.W., 2009a. The effect of drainage on the capillary retention of
CO
2
in a layered permeable rock. J Fluid Mech. 618, 349359.
Farcas, A., Woodsm A.W. On the steady drainage of a gravity current, from a point
source, up a sloping layered permeable rock, sub-judice. J Fluid Mech. in press.
Hesse, M.A., Tchelepi, H.A., Orr, F.M. Jr., 2006. Scaling analysis of the migration of
CO
2
in aquifers. In: SPE 102796, Presented at the 2006 SPE Annual Technical
Conference and Exhibition, San Antonio, TX, 2427 Sept 2006.
Hesse, M.A., Orr Jr., F.M., Tchelepi, H.A., 2008. Gravity currents with residual trap-
ping. J Fluid Mech. 351, 3560.
Huppert, H.E., Woods, A.W., 1995. Gravity-driven ows in porous layers. J. Fluid
Mech. 292, 5569.
Kharaka, Y.K., Cole, D.R., Hovorka, S.D., Gunter, W.D., Knauss, K.G., Freifeld, B.M.,
2006. Gas water rock interactions in Frio formation following CO
2
injection:
implications for the storage of greenhouse gases in sedimentary basins.
Geology 34 (7), 577580.
Mitchell, V., Woods, A.W., 2006. Gravity driven ow in conned aquifers. J. Fluid
Mech. 566, 345355.
Nordbotten, J.M., Celia, M.A., 2006. Similarity solutions for uid injection into
conned aquifers. J. Fluid Mech. 561, 307327.
Obi, E.-O.I., Blunt, M.J., 2006. Streamline-based simulation of carbon dioxide storage
in a North Sea aquifer. Water Resour. Res. 42, W03414. doi:10.1029/
2004WR003347.
Pritchard, D., Woods, A.W., Hogg, A.J., 2001. On the slow draining of a gravity
current moving through a layered permeable medium. J. Fluid Mech. 444,
2347.
Pritchard, D., 2007. Gravity currents over fractured substrates in a porous medium.
J Fluid Mech. 584, 415431.
Vella, D., Huppert, H.E., 2007. Gravity currents in a porous medium at an inclined
plane. J Fluid Mech. 555, 353362.
Woods, A.W., Farcas, A., 2009. Capillary entry pressure and the leakage of gravity
currents through a sloping layered permeable rock. J Fluid Mech. 618,
361379.
A.W. Woods, S. Norris / Journal of Structural Geology 32 (2010) 18271833 1833
Clay smear in normal fault zones The effect of multilayers and clay cementation
in water-saturated model experiments
J. Schmatz
a,
*
, P.J. Vrolijk
b
, J.L. Urai
a
a
Structural Geology, Tectonics and Geomechanics, Geological Institute, RWTH Aachen University, Lochnerstrasse 4-20, 52056 Aachen, Germany
b
ExxonMobil Upstream Research Co., P.O. Box 2189, Houston, TX, USA
a r t i c l e i n f o
Article history:
Received 29 April 2009
Received in revised form
10 December 2009
Accepted 13 December 2009
Available online 4 January 2010
Keywords:
Clay smear
Fault seal
Mechanical layering
Competence contrast
Experimental model
a b s t r a c t
We studied the evolution of fault zones in water-saturated model experiments consisting of sand and
clay layers above a normal fault dipping 70
(see
also Mandl, 2000; Van der Zee, 2002; Ferrill and Morris, 2003;
Adam et al., 2005). Water-saturated models allowed the deforma-
tion of wet clay and cohesionless sand together in one model
(Schmatz et al., 2010). The basement fault moved at 40 mm/h to
a maximum offset of 60 mm. The models were run between two
glass plates lubricated to minimize edge effects. At this deformation
rate, the thick clay layers were sheared under undrained condi-
tions, whereas pore pressures likely remained hydrostatic inside
the ne-grained sand. The resulting material properties were
characterized by a series of standard geotechnical measurements.
Full details of the methods for a free-surface boundary condition
are given in Schmatz et al. (2010). Here we summarize the most
important aspects and procedures that result from additional
boundary conditions and experiment designs.
2.2. Boundary conditions
We used 2 cm thick aluminum (r 2400 kg m
3
) top plates,
pre-cut at the same 70
in the direction of
the hanging wall. Almost immediately after 2 mm displacement on
the basement fault, deformation is also localized at the base of the
fault between the top plates, propagating downwards. With
progressive deformation these two zones interact, the precursor
fault becomes inactive, and deformation is localized in a stable,
planar deformation zone in the plane of the basement fault.
Localization is thus much more rapid than in the experiment with
a free top surface (displacement of w7
mm).
3.1.2. Sand and clay
Adding a 30 mm thick layer of rm clay to the model (Fig. 5,
Table 1, experiment N-9-f) has a major impact on the fault-zone
evolution. The experiment with a free surface is comparable to Figs.
9 and 11 of Schmatz et al. (2010). Here the fault-zone evolution is
complex. The steep precursor fault fails to propagate across the clay
layer, which instead deforms by bending. Fractures initiate at the
claysand interface at locations of high curvature and extension.
With increasing deformation the layer breaks into fragments of
various sizes which then rotate and become reworked into a ductile
clay gouge. The ow of sand grains along clay asperities is associ-
ated with asperity abrasion and mixing of sand and clay. Secondary
faults propagate fromfractures in the clay across the sand. A stable,
distinct planar fault plane fails to form for offset >30 mm (Fig. 5b).
a
b
Fig. 2. (a) Sketch of the underwater experimental apparatus with: 1) water-saturated
apparatus (inside), 2) waterproof container (outside), 3) rigid basement bottom with
drain perforation, 4) motor-driven fault offset, 5) basement fault dipping 70
, 6)
movable glass plate (inside), 7) movable glass plate (outside). (b) Sketch of setup
showing alternating layers of sand and clay overlying stiff basal block, covered with
pre-cut aluminum top plates. Red area indicates kinematically favored position of the
fault. (For interpretation of the references to colour in this gure legend, the reader is
referred to the web version of this article.)
J. Schmatz et al. / Journal of Structural Geology 32 (2010) 18341849 1836
In the same experiment with pre-cut top plates, both fault-zone
geometry and clay-gouge evolution are markedly different (same
vertical stress at the top of the clay layer; Fig. 6, Table 1, experiment
N-8-f-tp). Analogous to the experiment with sand only and top
plates, a precursor fault with a curvature towards the hanging wall
initiates at the basement fault tip, followed by a second fault that
initiates at the tip of the fault in the top plates and propagates
downwards. Bending and macroscopic brittle failure are absent.
The two fault segments link up across the clay layer and are
progressively straightened. At a displacement of 10 mm, a nearly
planar fault forms with a continuous, w2 mm thick clay gouge.
3.2. Effect of clay strength
3.2.1. Low competence contrast
Experiment R-2-s-tp contains a 40 mmthick sand layer between
two thin layers of soft clay (Fig. 7, Table 1); the clay beds compact
into normal clay before the experiment ends (Schmatz et al., 2010).
The competence contrast between sand and clay is thus low. The
rst fault initiates at the basement fault tip, followed by a second
fault initiating at the fault tip at the top plates. Belowthe lower clay
layer, the lower fault branched, with the branch on the footwall side
propagating through the clay. At the upper clay layer, there was
a mismatch between the two deformation bands arriving from
below and above, and the fault zone was initially segmented. The
nal, kinematically favored fault zone thus has a number of inactive
branches and is wider at the level of the clay layers with continuous
clay smear between the fault strands (Fig. 7b; cf. Fig. 3b of Van der
Zee et al., 2003).
Experiment C2-4-s-tp (Fig. 8, Table 1) has a similar initial stra-
tigraphy as the one described above, but the sand between the two
clay layers is 50 mm thick. To increase the competence contrast
slightly, 0.25 wt.% Portland cement was added to the soft-clay
mixture (this increased the strength only slightly from 0.2 kPa for
Table 1
Overview of experiment series.
Series New ID # Clay
layers
Thickness Layer setup (clay) Clay
type
Top
plates
#
Images
Velocity
[mm/min]
Water
content [%]
Hardening
[h]
Cement
[wt.%]
Figure SGR
diagram
Adam et al., 2005 P-1 0 150 133 0.7 4
1) Sand only P-2-tp 0 140 133 0.4 5, 19
P-3-tp 0 140 0.4
2) Changing two
parameters:
clay strength and layer
thickness
S-1-s 1 10 5/10/140 Soft 224 0.2 52 2
S-2-s 1 3 5/3/140 Soft 268 0.2 52 19
S-3-f 1 3 5/10/140 Firm 276 0.2 27
S-4-n 1 30 5/30/140 Normal 290 0.2 44 21
S-5-n 1 10 5/10/140 Normal 300 0.2 40 23
S-6-n 1 3 5/3/140 Normal 298 0.2 44 24
S-7-s 1 30 10/30/140 Soft 310 0.2 50 9
Schmatz et al., 2010 S-8-f 1 30 5/30/140 Firm 296 0.2 28 34
S-9-f 1 3 5/3/140 Firm 302 0.2 28 31
3) Changing four
parameters:
clay strength, layer
thickness,
number of layers, layer
orientation
N-1-s-tp 2 10 55/10/10/10/55 Soft 181 0.4 52 3
N-2-s-tp 1 30 55/30/55 Soft 183 0.4 52 12, 19 5
N-3-s-tp 1 30 30 Soft 185 0.4 52 13
N-4-n-tp 1 3 3 Normal 179 0.4 48 25
N-5-s-tp 1 3 3 Soft 146 0.4 52 17
N-6-s-tp 1 3 60/3/77 Soft 193 0.4 53 16
N-7-s-tp 2 3 50/3/6/3/78 Soft 182 0.4 52 10
N-8-f-tp 1 30 50/30/60 Firm 186 0.4 37 7 28
N-9-f- 1 30 50/30/60 Firm 189 0.4 34 6 29
N-10-n-tp 2 3 50/3/6/3/78 Normal 191 0.4 l45/u47 20
N-11-n-tp 1 10 50/10/80 Normal 191 0.4 44 22
4) Changing clay layer
distance
D-1-s-tp 2 3 40/3/30/3/64 Soft 187 0.4 53 8
D-2-s-tp 2 3 40/3/40/3/54 Soft 215 0.4 53 6
D-3-s-tp 2 3 40/3/50/3/44 Soft 203 0.4 53 18
5) Properties of cement C1-1-s-tp 1 2 60/2/78 Soft 195 0.4 50 72 10
C1-2-s-tp 1 3 60/3/77 Soft 100 0.4 53 72 1 35
C1-3-s-tp 1 3 60/3/77 Soft 158 0.4 52 72 5
C1-4-s-tp 1 5 60/5/75 Soft 181 0.4 52 48 10
C1-5-s-tp 1 3 60/3/77 Soft 126 0.4 51 22 10
C1-6-s-tp 1 3 60/3/77 Soft 157 0.4 53 2 10 10, 16 33
6) Reproduction R-1-s-tp 2 3 40/3/40/3/54 Soft 184 0.4 55 11
R-2-s-tp 2 3 40/3/40/3/54 Soft 181 0.4 52 8 7
R-3-s-tp 2 3 40/3/40/3/54 Soft 180 0.4 50 12
R-4-s-tp 2 3 40/3/40/3/54 Soft 191 0.4 51 14
R-5-s-tp 2 3 40/3/40/3/54 Soft 189 0.4 53 15
7) Cement content C2-1-s-tp 1 3 50/3/87 Soft 175 0.4 48 22 1 36
C2-2-s-tp 2 3 50/3/50/3/34 Soft 173 0.4 49 23 1 32
C2-3-s-tp 2 3 50/3/50/3/34 Soft 183 0.4 49 22 2 11 38
C2-4-s-tp 2 3 50/3/50/3/34 Soft 184 0.4 49 22 0 9 27
C2-5-s-tp 2 3 50/3/50/3/35 Soft 178 0.4 48 22 2 37
C2-6-s-tp 2 3 50/3/50/3/34 Soft 188 0.4 50 24 0 26
8) Multilayer M-1-s-tp 4 8 30/8/12/8/12/8/
12/8/30
Soft 186 0.4 44 13, 17 4
M-2-s-tp 4 Various 35/20/10/3/10/
20/10/2/40
Soft 185 0.4 50 14 1
9) Clay 2 50 20/50/5/50/40 Soft 163 0.4 48
10) Others K-1-f-tp 1 3 60/3/77 Firm 178 0.4 37 30
K-2-f-tp 1 3 50/3/87 Firm 133 0.4 34
K-3-f-tp 1 3 80/3/57 Firm 162 0.4 32
J. Schmatz et al. / Journal of Structural Geology 32 (2010) 18341849 1837
Fig. 3. (a) Image sequence showing experiment P-1 with 150 mm thick sand layer and
free top surface (see inset). The three stages record offset along the basement fault of
0.4, 15, and 22 mm. (b) Image sequence showing the corresponding PIV overlay to the
image sequence in (a) with the velocity vector eld displayed in the foreground and
the contour plot of incremental strain e
xy
in the background. Modied from Adam et al.
(2005). (c) Legend for insets in Figs. 318.
Fig. 4. (a) Image sequence showing experiment P-2-tp with a 140 mm thick sand
package and top plates (see inset). Fault offset is 2, 4, and 16 mm. Red lines indicate
location of rigid blocks at bottom and top. (b) Image sequence showing the corre-
sponding PIV overlay to image sequence in (a) with the velocity vector eld in the
foreground and a contour plot of the z-component of the incremental rotation eld in
the background. (For interpretation of the references to colour in this gure legend, the
reader is referred to the web version of this article.)
J. Schmatz et al. / Journal of Structural Geology 32 (2010) 18341849 1838
Fig. 5. (a) Image sequence showing experiment N-9-f with a 30 mm thick rm clay
layer in the center and a free top surface (see inset). Basement fault offset is 3, 6, and 35
mm. (b) Image sequence showing the corresponding PIV overlay to image sequence (a)
with the velocity vector eld in the foreground and a contour plot of the z-component
of the incremental rotation eld in the background.
Fig. 6. (a) Image sequence showing experiment N-8-f-tp with a 30 mm thick rm clay
layer in the center of the model and top plates (see inset). Basement fault offset is 4, 7
and 10 mm. (b) Image sequence showing the corresponding PIV overlay to image
sequence in (a) with the velocity vector eld in the foreground and a contour plot of
the z-component of the incremental rotation eld in the background.
J. Schmatz et al. / Journal of Structural Geology 32 (2010) 18341849 1839
Fig. 7. (a) Image sequence showing experiment R-2-s-tp with two 3 mm thick soft-
clay layers and top plates (see inset). Basement fault offset is 4, 6 and 20 mm. Red lines
indicate location of rigid blocks at bottom and top. (b) Image sequence showing the
corresponding PIV overlay to the image sequence in (a) with the velocity vector eld in
the foreground and a contour plot of the z-component of the incremental rotation eld
in the background. (For interpretation of the references to colour in this gure legend,
the reader is referred to the web version of this article.)
Fig. 8. (a) Image sequence showing experiment C2-4-s-tp with two 3 mm thick
cemented clay layers and top plates (see inset). Basement fault offset is 5, 9 and 36
mm. Red lines indicate location of rigid blocks at bottom and top. (b) Image sequence
showing the corresponding PIV overlay to the image sequence in (a) with the velocity
vector eld in the foreground and a contour plot of the z-component of the incre-
mental rotation eld in the background. (For interpretation of the references to colour
in this gure legend, the reader is referred to the web version of this article.)
J. Schmatz et al. / Journal of Structural Geology 32 (2010) 18341849 1840
soft clay to 0.25 kPa for soft clay with 0.25 wt.% cement). We
observe two precursor faults initiating fromthe basement and from
the top. The lower fault initiates with a steep dip, but, after inter-
secting the lower clay layer, it curves towards the footwall. The
upper fault evolves in its kinematically favored plane. Both faults
are active throughout the experiment, with a progressively thin-
ning but stable fault lens between the two active fault strands
(Fig. 8b). Continuous clay smear forms everywhere in the model.
The two experiments described above are therefore similar initially,
but the second experiment allows simultaneous displacement on
two fault surfaces whereas the rst experiment constrained
displacement to one fault strand.
3.2.2. High competence contrast
3.2.2.1. Strong clay. Experiment C1-6-s-tp (Fig. 9, Table 1) contains
a single, 3 mmthick clay layer 60 mmabove the basement. The clay
contains 10 wt.% cement and is cured only partly, so its strength is
approximately the same as the completely cured, 2 wt.% cemented
Fig. 9. Image sequence showing experiment C1-6-s-tp with one 3 mm thick cemented clay layer and top plates (see inset). Basement fault offset is 8, 16 and 28 mm. Red lines
indicate location of rigid blocks at bottom and top. (For interpretation of the references to colour in this gure legend, the reader is referred to the web version of this article.)
Fig. 10. Image sequence showing experiment C2-3-s-tp with two cemented (1.5 wt.%) clay layers and top plates (see inset). Basement fault offset is 5, 25, and 45 mm. Red lines
indicate location of rigid blocks at bottom and top. (Transparent, reddish box in center of photos is a camera reection). (For interpretation of the references to colour in this gure
legend, the reader is referred to the web version of this article.)
J. Schmatz et al. / Journal of Structural Geology 32 (2010) 18341849 1841
clay. Deformation involves monoclinal bending of the clay followed
by fracturing and block rotation. The shear zones in the sand form
a wide lens with continuous activity on multiple fault strands. The
interior of the lens is highly deformed, in contrast to the previously
described experiments.
Experiment C2-3-s-tp (Fig. 10, Table 1) has two layers of 3 mm
thick clay. The clay was fully cured but with only 1.5 wt.% cement.
As in the previous experiment, initial deformation causes brittle
failure in both layers. Progressive deformation leads not only to
displacement and rotation of the fragments, but also to progressive
erosion of the fragment edges, forming a ductile clay gouge with
clay clasts and coremantle structure around the old fragments.
3.2.2.2. Weak clay. Experiment N-2-s-tp contains one 30 mmthick,
soft-clay layer (Fig. 11, Table 1). A near-vertical precursor fault
initiates from the basement fault tip. After 2 mm displacement on
the basement fault, deformation localizes at the tip of the fault
between the top plates and propagates downward. With progres-
sive deformation these two zones interact and overlap, forming
a restraining relay zone across the clay layer (Fig. 11a, center). Initial
Fig. 11. (a) Image sequence showing experiment N-2-s-tp with one 30 mm thick soft-
clay layer (see inset) and top plates. Fault offset is 4, 6 and 34 mm. Approximate fault
location is traced with red lines. (b) Image sequence showing the corresponding PIV
overlay to the image sequence in (a) with the contour plot of incremental rotation eld
and the velocity vector eld. (For interpretation of the references to colour in this
gure legend, the reader is referred to the web version of this article.)
Fig. 12. (a) Image sequence showing experiment experiment M-1-s-tp with four 8 mm
thick soft-clay layers and top plates (see inset). Fault offset is 2 and 11 mm. Red lines
indicate location of rigid blocks at bottom and top. Fig. 17 shows nal stage of this
experiment. (b) Image sequence showing the corresponding PIV overlay to the image
sequence with the contour plot of incremental rotation eld and the velocity vector
eld. (For interpretation of the references to colour in this gure legend, the reader is
referred to the web version of this article.)
J. Schmatz et al. / Journal of Structural Geology 32 (2010) 18341849 1842
deformation in the clay layer is in this restraining zone (cf. Egholm
et al., 2008). With increasing displacement, the top precursor fault
becomes inactive, and deformation switches to a new strand
forming a restraining relay zone relative to the fault below the clay
layer (Fig. 11b, bottom). At the same time, the fault below the clay
layer rotates clockwise, producing a gentle curving geometry at the
lower sandclay contact that is similar to normal-drag folds. This
fault evolution produces a stable, planar deformation zone with an
unusually thick clay gouge.
3.2.3. Effect of multiple layers and layer thickness
Experiment M-1-s-tp has four 12 mm thick soft-clay layers
alternating with 8 mm thick sand layers (Fig. 12, Table 1). The
fault-zone evolution in this experiment is similar to the reference
experiment with only sand (Fig. 4). The initially steep precursor
faults form a restraining zone across the sandclay sequence with
minor fault segmentation. Deformation in this sequence evolves
into the kinematically favored plane. Each clay layer forms
a continuous clay smear in the layered gouge, but the sand
between the clay layers thins and in some cases becomes bou-
dinaged (Fig. 12). A similar setup with four soft clay layers but
variable layer thickness (Fig. 13, Table 1, experiment M-2-s-tp)
illustrates the effect of layer thickness on fault-zone evolution.
Here, the total thickness of the clay-rich interval is more than in
the previous experiment (75 vs. 62 mm), and the claysand ratio is
also higher (3:2 vs. 2:3). The two initially steeply dipping
precursor faults connect in a zone of diffuse deformation with
a shallower dip than the initial segments. Localization of defor-
mation occurs on both sides of this zone with a less deformed lens
in between. With progressive deformation the sheared sand in the
fault zone becomes discontinuous as individual, sheared clay
layers coalesce.
4. Discussion
4.1. Effect of boundary conditions on fault-zone evolution
Faults in nature form under a variety of local boundary condi-
tions, but in many cases those boundary conditions might be less
than those imposed here (Schmatz et al., 2010). In this study we
explored normal fault development in weak materials layered
between two rigid layers faulted with the same 70
f=2 froms
0
1
. The major principal effective stress s
0
1
is the vertical overburden stress s
0
v
in a basin under tectonic
extension, which can be converted to a depth of faulting assuming
a lithostatic gradient for the basin.
In the design of the experiments, we assume an effective
lithostatic gradient of 10 MPa/km and a friction angle of 30
to
calculate the normal stress to be applied so that failure occurs
approximately at a given burial depth. The actual depth of faulting
is back-calculated after the experiments from the measured
stresses at peak condition.
3.4. Test description
The experiments are described in Table 2. After mounting and
saturation inside the ring shear cell, the samples are rst consoli-
dated to nominal vertical stresses. A small cycling load (100 cycles,
5 kN amplitude) is used during normal loading to ensure good
contact between the sample and the knives in the ring shear device.
After consolidation, the samples are sheared under constant
effective normal stress with a prescribed rotation velocity of 1 deg/
5 min for a total length of 90 degrees (220 mmshear displacement).
The velocity has been chosen such that the estimated pore pressure
build-up at mid-height of the sample is negligible.
The permeability between two opposite openings across clay
smear and sandsand juxtaposition is measured during normal
loading and every 1530 degrees of shear rotation by imposing
a ow rate Q at the injection point and a constant pressure at the
outlet, and recording the pressure difference DP between inlet and
outlet after steady state conditions are reached. Note that shearing
is stopped during permeability measurements. Typically for
measuring permeability across the clay smear, two opposite
openings at the center of the fault throw are used, while the
Table 1
Mineralogy of the Troll clay fromX-ray diffraction analyses (fromLunne et al., 2007).
Mineral % Clay fraction (<2 mm) % Sand fraction (>60 mm)
Illite 54
Kaolinite 19
Smectite 13
Chlorite 6 10
Quartz 2 65
Feldspar 2 17
Calcite Minor
Pyrite 8
Table 2
Laboratory program for experiments dedicated to clay smear.
Test Effective normal
stress (MPa)
Over-consolidation Burial depth
a
(m)
RT15 3.5 Normally consolidated 500
RT16 0.7 Normally consolidated 100
RT17 7 Normally consolidated 1000
RT18 10.5 Normally consolidated 1500
RT19 3.5 Normally consolidated 500
RT20 3.510.5 Overconsolidated max 1500
a
Estimated before testing.
Table 3
Strength data from ring shear experiments. The depth of burial at time of faulting is
estimated from peak stress and a pore pressure gradient of 10 MPa/km.
Test Effective normal
stress (MPa)
Peak shear
stress (MPa)
Depth of burial at
faulting (m)
RT15 3.5 2.20 570
RT16 0.7 0.79 149
RT17 7 3.90 1090
RT18 10.5 5.32 1582
RT19 3.5 2.19 569
RT20 10.53.5 2.46 596
1
This condition could differ in other experiments.
F. Cuisiat, E. Skurtveit / Journal of Structural Geology 32 (2010) 18501863 1853
remaining openings are closed. The interpreted coefcient of
permeability is an average value of the sample containing unde-
formed sand, shear band, and clay smear. The permeability of the
sample can be back-calculated using Darcys law if one assumes
that the ow is one dimensional and that the involved ow cross
section is known, through the equation:
k
Qm
A
DP
H
(1)
where:
Q constant ow rate across sample (m
3
/s)
DP pressure difference from inlet to outlet (Pa)
k permeability (m
2
)
m uid viscosity (10
3
Pa s, water)
H sample height (m)
A assumed ow area (m
2
).
In practise the permeability is expressed in mD using
10
15
m
2
z 1mD. A constant ow rate Q equal to 2.3 10
7
m
3
/s is
used resulting in typical values of pressure difference equal to
58 10
3
Pa at steady state. Steady state owconditions are achieved
after two to three minutes.
The main uncertainty when using Darcys law relates to the size
of the owarea. In the following, we will assume that the owarea
is equal to 1/24th of the ring area for the interpretation of the
permeability experiments. Given that this assumption may only be
an approximation in reality, 3D nite difference modelling was
conducted of the ow conditions during a permeability test in the
ring shear apparatus. The modelling shows that the ow area may
be closer to the area of the clay smeared zone, depending on the
contrast between clay and sand permeability.
In tests RT15RT18, three segments of clay 120
apart fromeach
other are used. In tests RT19 and RT20 5 segments are used, 3 of
them being only separated 15
1
1
k
sand
k
smear
t
smear
t
tot
t
smear
t
tot
(2)
where:
k
smear
is the permeability of the clay smear
k
sand
is the permeability of the sand
t
tot
is the total height of the sample
t
smear
is the thickness of the clay smear.
Fig. 12. Sketch showing formation of sand wedge during faulting in claysand layered sediments, and observed cross-sections at end of shearing in test RT19. Shearing at 3.5 MPa
effective normal stress (estimated burial depth at the time of faulting of 500 m). All sections are within 5 cm along the shear direction. The sand wedge thins in the direction of
rotation of the lower ring. Photograph width approximately equal to 2.54 cm.
F. Cuisiat, E. Skurtveit / Journal of Structural Geology 32 (2010) 18501863 1859
The permeability ratio k
avg
/k
sand
is plotted in Fig. 15 for different
values of k
sand
/k
smear
, i.e. ratio between sand and clay smear
permeability. The results from the ring shear tests are plotted on
the gure, based on the measured permeability at end of shearing
(90
/
)
Linear Displacement (mm)
RT15 -3.5 MPa
RT19 -3.5 MPa
RT20 -3.5 MPa
RT15 -500 m
RT19 -500 m (CS)
RT20 -500m (CS/OC)
Fig. 13. Normalised shear stress versus linear displacement for tests RT15, RT19
(composite smear CS) and RT20 (composite smear and over-consolidated CS/OC). All
tests are sheared under 3.5 MPa effective normal stress.
0
500
1000
1500
2000
2500
0 50 100 150 200 250
Linear Displacement (mm)
P
e
r
m
e
a
b
i
l
i
t
y
a
c
r
o
s
s
s
m
e
a
r
(
m
D
)
RT19 - 1 clay segment
RT19 - 3 clay
segments
RT19 - 1/2 clay
segment
RT20 - 1 clay segment
RT20 - 1/2 clay segment
O
v
e
r
c
o
n
s
o
l
i
d
a
t
e
d
Fig. 14. Permeability across clay smear from single clay segment, half segment and
three segments versus linear displacement in tests RT19 and RT20 (over-consolidated).
Both tests are sheared under 3.5 MPa effective normal stress.
0.001
0.01
0.1
1
10
0 0.02 0.04 0.06 0.08 0.1
k
o
i
t
a
r
y
t
i
l
i
b
a
e
m
r
e
P
e
c
n
e
u
q
e
s
k
/
d
n
a
s
Normalised thickness of clay smear t
smear
/t
total
ksand/ksmear=10
ksand/ksmear=100
ksand/ksmear=1000
RT18
RT17
RT15
RT16
Fig. 15. Average permeability of sandclay smear layer sequence (harmonic average)
versus ratio of clay smear thickness to total thickness for different ratios of sand
permeability k
sand
/clay smear permeability k
smear
. The ring shear tests are plotted with
black lled symbols. The thickness of the clay smear is estimated after dismounting the
samples (Fig. 8).
Table 4
Shale Smear Factor SSF, Clay Smear Potential CSP and Shale Gouge Ratio SGR
calculated at mid-point along slipped interval from ring shear experiments on clay
smear. Tests RT19, RT15 and RT20 are sheared under 3.5 MPa effective normal stress,
but RT20 is over-consolidated. Tests RT16, RT17 and RT18 are sheared under 0.7, 7
and 10.5 MPa effective normal stress, respectively. The corresponding (estimated)
burial depths at the time of faulting are: 100 m (RT16), 500 m (RT15, RT19, RT20),
1000 m (RT17), 1500 m (RT18).
RT19, RT20 RT15RT20 RT19, RT20
Throw (mm) 219.5 219.5 219.5
Clay segment
thickness (mm)
15.5 31 31
Clay volume
factor (%)
100 100 100
Number of
clay segments
layer 1 layer 3 layers
Shale Smear
Factor (SSF)
14.2 7.1 2.4
Clay Smear
Potential (CSP)
1.1 4.4 39.4
Shale Gouge
Ratio (SGR)
7.1 14.1 42.4
Thickness of
smear (mm)
0.12 23 24 up to 7
Continuity of smear Discontinuous (RT19)
Continuous (RT20)
Continuous
except at high
burial depth (RT18)
Continuous
In bold gures: probable discontinuous smear predicted from empirical
relationships.
F. Cuisiat, E. Skurtveit / Journal of Structural Geology 32 (2010) 18501863 1860
observed for test RT19 (half-segment width) in agreement with SSF
and SGR estimations. For one clay segment, the values of SSF and
SGR are close to the limit for breach of the smear (SSF 7 and
SGR 14). Continuous smear is observed in all experiments,
although a small puncture is observed at high burial depth
(1500 m, RT18).
The previous parameters do not provide quantitative relation-
ship between smear development, rock properties, stress condi-
tions and uid ow properties. Takahashi (2003) for instance
showed through laboratory experiments that a critical SSF for seal
breach depends on the effective normal stress applied to the fault
during smearing. Our experiments also indicate that the burial
depth (or normal stress) might control the continuity of the smear.
Nevertheless, our experiments are in agreement with eld
evidence showing a rst order correlation between SGR and seal
capacity (Naruk, 2008).
The results of the ring shear experiments are also compared to
the empirical relationship proposed by Sperrevik et al. (2002)
which relates fault rock permeability to maximum burial depth,
depth at the time of faulting and clay content in fault rock. In Fig. 16,
the empirical relationship is plotted for two values of maximum
burial depth (same as depth at the time of faulting in this case)
equal to 500 and 1500 m together with the permeability data from
ring shear tests with one clay segment. The ring shear test results
show a rst order agreement with the empirical relationship of
Sperrevik et al. (2002) based on data of nearly 100 normal faults
from clastic reservoirs in the North Sea (Fig. 16).
5.4. Fault permeability reduction with increasing burial depth
Despite uncertainties related to the interpretation of ow
measurements (in particular the owarea), the test results suggest
that clay smear and shear displacements are the main permeability
reduction mechanism for faults formed at low burial depths
(Fig. 17). Sandsand juxtaposition shear is dominated by grain
rolling causing only minor permeability reduction. At greater
depths grain crushing along the clay smear zone and in the sand
sand juxtaposition areas contributes as much to permeability
reduction as the formation of a clay smear (Fig. 17). The transition
between the two scenarios is likely to occur at around 500 m burial
depth. At this depth, some grain crushing is observed in the
experiments (e.g. test RT15), but not as developed as for experi-
ments conducted at 1000 mand 1500 mburial. In general the burial
depth at time of faulting has a larger inuence on the estimated
permeability change than the clay segment thickness and over-
consolidation ratio, within the range of parameter variation
investigated in the experiments.
6. Conclusions
In this paper, we have presented the results from experimental
work carried out with a high stress ring shear apparatus to inves-
tigate mechanisms of clay smear along faults in uncemented sedi-
ments at various burial depths, and its impact on uid ow
properties. The experiments consist of shearing a ring of sand with
embedded clay segments, thereby simulating faulting through
a layered sandclay sequence. Baskarp sand No 15 and a natural
glacio-marine clay from the Troll eld are used for testing. Visual
inspection of the samples after testing, analyses of thin sections and
permeability measurements across the shear zone are used to
describe clay smear continuity, thickness and permeability.
Deformation processes such as grain reorientation, clay smear
and cataclasis are identied from the tests. The complexity of the
shear zone is observed to increase with greater burial depth at time
of faulting. The experiments suggest that at shallow burial depth in
clay-rich sediments, clay smear is the most efcient mechanism for
permeability reduction. At this depth, sandsand juxtaposition
shear is dominated by grain rolling causing only minor perme-
ability reduction. At greater burial depths, permeability reduction is
dominated by grain crushing. Measurements of permeability both
across clay smear and sandsand juxtaposition yield similar values.
The observed thickness of clay smear is more sensitive to the
thickness of the sheared clay layers than other parameters tested
within the limited test program. Shearing of multiple clay layers (3
layers) produces a composite clay smear 23 times thicker than for
a single clay layer, whereas when reducing the clay layer thickness
to one half of the reference layer, a thin and discontinuous clay
smear is produced. The permeability across the clay smear is found
to increase as the thickness of the clay source decreases for single
clay layers, but the permeability for composite smear is more
complex and a clear trend is not found from the only two tests
performed. Over-consolidation of the sample prior to shearing has
no signicant inuence on the thickness and continuity of the clay
smear produced, but a reduction in permeability both for clay
smears and initial sand is found when compared to a normal
1.E-08
1.E-06
1.E-04
1.E-02
1.E+00
1.E+02
1.E+04
1.E+06
1.E+08
0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1
)
d
m
(
y
t
i
l
i
b
a
e
m
r
e
P
Clay content in fault Vclay
1500 m
500 m
RT16 - 0.7 MPa
RT15 - 3.5 MPa
RT17 - 7 MPa
RT18 - 10.5 MPa
Fig. 16. Comparison of permeability data from ring shear tests (one clay segment) with
empirical relationship proposed by Sperrevik et al. (2002). The clay content in the fault
rock is taken equal to the SGR for the ring shear data. The corresponding (estimated)
burial depths at the time of faulting are: 100 m (RT16), 500 m (RT15), 1000 m (RT17),
1500 m (RT18).
Deep burial
depth
(~1500 m)
Shallow
burial depth
(~100m)
Bef ore
shearing
220 mm
Permeability
(mD)
~2500
~1200
~800
~1700
~10
~100
~500
110 mm
na
SGR 28 14
Shear displacement
Fig. 17. Synthesis of permeability trends from ring shear experiments. Values indicate
permeability in mD for clay smear (blue) and sandsand juxtaposition (yellow) or
a combination (yellow/blue). No ow measurement is available in the sandsand
juxtaposition for loweffective normal stress (i.e. low burial depth) at end of shear. Ring
shear displacements have been converted to shale gouge ratio (SGR). (For interpreta-
tion of the references to colour in this gure legend, the reader is referred to the web
version of this article.)
F. Cuisiat, E. Skurtveit / Journal of Structural Geology 32 (2010) 18501863 1861
consolidated test. This implies that the initial density of the
material is important for the permeability measurements.
The results from the experiments show many similarities with
eld or outcrop natural observations. Next to clay smearing, drag or
injection of clay sand along the fault plane also occurs. Due to
mixing of clay and sand into the fault core, the ow properties of
the fault are expected to be anisotropic with higher permeability
along the shear direction (Bense and Van Balen, 2004). However,
more experimental data should be collected in order to develop
better trend lines and predictive models of fault properties. Field
and outcrop observations could be used actively to constrain the
experimental conditions to use in future tests. Future research will
complement the existing dataset as well as address formation of
phyllosilicate framework fault rocks in unclean sand with varying
clay content, clay mineralogy, burial depth, and fault throw. More
tests are also needed for improving the understanding of defor-
mation bands in multiple clay layers and the effect of composite
smear on fault permeability. Other types of experimental set-ups
such as a biaxial plane strain apparatus (Rykkelid and Skurtveit,
2008) may also be used to reproduce more accurately the stress and
strain conditions during basin extension and fault propagation
through layered sandclay sequences. Ultimately, through more
extensive databases, models can be implemented into reservoir
simulators to capture better the impact of faults on oil recovery
(Jolley et al., 2007).
Acknowledgements
The work presented in this paper was carried out as part of the
PROFUSE project (2007) with support from Total, BP and Sta-
toilHydro. Comments from Yves Leroy and an anonymous reviewer
are greatly acknowledged.
References
Annunziatellis, A., Beaubien, S.E., Bigi, S., Ciotoli, G., Coltella, M., Lombardi, S., 2008.
Gas migration along fault systems and through the vadose zone in the Latera
caldera (central Italy): implications for CO
2
geological storage. International
Journal of Greenhouse Gas Control 2, 353372.
Antonellini, M., Aydin, A., Pollard, D., 1994. Microstructure of deformation bands in
porous sandstones at Arches National Park, Utah. Journal of Structural Geology
16, 941959.
Bense, V., Van den Berg, E., Van Balen, R., 2003. Deformation mechanisms and
hydraulic properties of fault zones in unconsolidated sediments; the Roer
Valley rift system, The Netherlands. Hydrogeology Journal 11, 319332.
Bense, V.F., Van Balen, R., 2004. The effect of fault relay and clay smearing on
groundwater owpatterns in the lower rhine embayment. Basin Research
Volume 16, 397411.
Bishop, A.W., Green, G.E., Garga, V.K., Andersen, A., Browns, J.D., 1971. A new ring
shear apparatus and its application to the measurement of residual strength.
Geotechnique 21, 273328.
Bjo rlykke, K., Ho eg, K., 1997. Effects of burial diagenesis on stresses, compaction and
uid ow in sedimentary basins. Marine and Petroleum Geology 14, 267276.
Bolton, A., Maltman, A., 1998. Fluid-ow pathways in actively deforming sediments:
the role of pore uid pressures and volume change. Marine and Petroleum
Geology 15, 281297.
Bouvier, J.D., Kaars-Sijpesteijn, C.H., Kluesner, D.F., Onyejekwe, C.C., Van der Pal, R.C.,
1989. Three dimensional seismic interpretation and fault sealing investigations,
Nun River eld, Nigeria. AAPG Bulletin 73, 13971414.
Caine, J.S., Evans, J.P., Forster, C.B., 1996. Fault zone architecture and permeability
structure. Geology 24, 10251028.
Cardozo, N., Cuisiat, F., 2008. Fault propagation folding modeling with FLAC. In:
Proceedings of the 1st International FLAC/DEM Symposium on Numerical
Modeling, 2527 August 2008, Minneapolis, USA.
Chuhan, F.A., Kjeldstad, A., Bjorlykke, K., Hoeg, K., 2002. Porosity loss in sand by
grain crushing experimental evidence and relevance to reservoir quality.
Marine and Petroleum Geology 19, 3953.
Clausen, J.A., Gabrielsen, R.H., 2002. Parameters that control the development of
clay smear at low stress states: an experimental study using ring shear appa-
ratus. Journal of Structural Geology 24, 15691586.
Crawford, B.R., Myers, E.D., Woronov, A., Faulkner, D.R., Rutter, E.H., 2002. Porosity
Permeability Relationships in Clay-bearing Fault Gouge. Society of Petroleum
Engineers Inc. SPE/ISRM 78214, Rock Mechanics Conference, October 2002.
Cuisiat, F., Skurtveit, E., Cleave, R., 2007. Fault seal prediction in unconsolidated
sediments with a novel experimental apparatus. In: Proceedings of the 7th
ISRM Congress on Rock Mechanics, Lisboa, Portugal.
Doughty, P.T., 2003. Clay smear seals and fault sealing potential of an exhumed
growth fault, Rio grande rift, new Mexico. American Association of Petroleum
Geologists Bulletin 87, 427444.
Egholm, D.L., Clausen, O.R., Sandiford, M., Kristensen, M.B., Korstgard, J.A., 2008. The
mechanics of clay smearing along faults. Geology 36, 787790.
Faerseth, R.B., 2006. Shale smear along large faults: continuity of smear and the
fault seal capacity. Journal of the Geological Society 163, 741751.
Fisher, Q.J., Knipe, R.J., 1998. Fault sealing processes in siliciclastic sediments. In:
Jones, G., Fisher, Q., Knipe, R. (Eds.), Faults, Fault Sealing and Fluid Flow in
Hydrocarbon Reservoirs. Geological Society of London, special publication, vol.
147, pp. 117134.
Fisher, Q.J., Knipe, R.J., 2001. The permeability of faults within siliciclastic petroleum
reservoirs of the north sea and Norwegian continental shelf. Marine and
Petroleum Geology 18, 10631081.
Gudehus, G., Karcher, C., 2007. Hypoplastic simulations of normal faults without
and with clay smears. Journal of Structural Geology 29, 530540.
Hickman, S., Sibson, R.H., Bruhn, R., 1995. Introduction to special section:
mechanical involvement of uids in faulting. Journal of Geophysical Research
100, 1283112840.
Hungr, O., Morgenstern, N.R., 1984. High velocity ring shear tests on sand. Geo-
technique 34, 415421.
Jolley, S.J., Dijk, H., Lamens, J.H., Fisher, Q.J., Manzocchi, T., Eikmans, H., Huang, Y.,
2007. Faulting and fault sealing in production simulation models: Brent Prov-
ince, northern North Sea. Petroleum Geoscience 13, 321340.
Karakouzian, M., Hudyma, N., 2002. A new apparatus for analog modeling of clay
smears. Journal of Structural Geology 24, 905912.
Koledoye, B.A., Aydin, A., May, E., 2003. A new process-based methodology for
analysis of shale smear along normal faults in the Niger Delta. American
Association of Petroleum Geologists Bulletin 87, 445463.
Kvaale, T., 2002. Et stadium av reservoarsand i ringskjrapparat for simulere
oppfrselen i forkastningssoner. Master thesis, University of Oslo.
Lambe, T.W., Whitman, R.V., 1979. Soil Mechanics. John Wiley, New York.
Lehner, F.K., Pilaar, W.F., 1997. The emplacement of clay smears in synsedimentary
normal faults: inferences from eld observations near Frechen, Germany. In:
Mller-Pedersen, P., Koestler, A.G. (Eds.), Hydrocarbon Seals: Importance for
Exploration and Production. Norwegian Petroleum Society, Special Publication,
vol. 7, pp. 3950.
Lindsay, N.G., Murphy, F.C., Walsh, J.J., Waterson, J., 1993. Outcrop studies of shale
smears on fault surfaces. Special Publications International Association of
Sedimentologist 15, 113123.
Lunne, T., Long, M., Uzielli, M., 2007. Characterisation and engineering properties of
troll clay. In: Tan, Phoon, Hight, Leroueil (Eds.), Characterisation and Engi-
neering Properties of Natural Soils.
Lupini, J.F., Skinner, A.E., Vaughan, P.R., 1981. The drained residual strength of
cohesive soils. Geotechnique 31, 181213.
Mandl, G., 2000. Faulting in Brittle Rocks. Springer Verlag, Berlin, 434 p.
Mandl, G., de Jong, L.N.J., Maltha, A., 1977. Shear zones in granular material. An exper-
imental studyof their structureandmechanical genesis. RockMechanics9, 95144.
Manzocchi, T., Walsh, J.J., Nell, P., Yielding, G., 1999. Fault transmissibility multipliers
for ow simulation models. Petroleum Geoscience 5, 5363.
Naruk S.J., 2008. Empirical calibrations of proven sealing faults. Joint Meeting of the
Geological Society of America, Soil Science Society of America, American Society
of Agronomy, Crop Science Society of America, Gulf Coast Association of
Geological Societies with the Gulf Coast Section of SEPM.
Nowacki, E.H.F., 1967. Anvendelse av ringskjrapparatet for studier av omrrte
leirersegenskaper ved store deformasjoner. Det store eksamensarbeidet i
Geoteknikk og fundamenteringslre. NTH, Trondheim.
Olson, R.E., 1974. Shearing strength of kaolinite, illite and montmorillonite. Journal
of the Geotechnical Division, ASCE 100 (GT11), 12151299.
Rawling, G.C., Goodwin, L.B., 2006. Structural record of the mechanical evolution of
mixed zones in faulted poorly lithied sediments, Rio Grande rift, New Mexico,
USA. Journal of Structural Geology 28, 16231639.
Rutqvist, J., Birkholzer, J., Cappa, F., Tsang, C.-F., 2007. Estimating maximum
sustainable injection pressure during geological sequestration of CO2
using
coupled uid ow and geomechanical fault-slip analysis. Energy Conversion
and Management 48, 17981807.
Rutter, E.H., Maddock, R.H., White, S.H., 1986. Comparative microstructures of
natural and experimentally produce clay-bearing fault gouges. Pure and
Applied Geophysics 124, 330.
Rykkelid, E., Skurtveit, E., 2008. Experimental work on unconsolidated sand: the
effect of burial depth at time of deformation. In: Proceedings of the 42nd US
Rock Mechanics Symposium and 2nd US-Canada Rock Mechanics Symposium,
Paper No 08-236, San Francisco, June 29July 2.
Sandbkken, G., Berre, T., Lacasse, S., 1986. In: Yong, R.N., Townsend, F.C. (Eds.),
Oedometer Testing at the Norwegian Geotechnical Institute. Consolidation of
Soils: Testing and Evaluation, ASTM STP 892. American Society for Testing and
Materials, Philadelphia, pp. 329353.
Sassa, K., Gonghui, W., Fukuoka, H., 2003. Performing undrained shear test on
saturated sands in a new intelligent type ring shear apparatus. Geotechnical
Testing Journal 26, 19.
Sibson, R.H., 2000. Fluid involvement in normal faulting. Journal of Geodynamics
29, 469499.
F. Cuisiat, E. Skurtveit / Journal of Structural Geology 32 (2010) 18501863 1862
Sperrevik, S., Frseth, R., Gabrielsen, R., 2000. Experiments on clay smear formation
along faults. Petroleum Geoscience 6, 113123.
Sperrevik, S., Gillespie, P.A., Fisher, Q.J., Halvorsen, T., Knipe, R.J., 2002. Empirical
estimation of fault rock properties. In: Koestler, A.G., Hunsdale, R. (Eds.),
Hydrocarbon Seal Quantication. Elsevier, Amsterdam, pp. 109125.
Takahashi, M., 2003. Permeability change during experimental fault smearing.
Journal of Geophysical Research 108, 2235.
Takizawa, S., Kamai, T., Matsukura, Y., 2005. Fluid pathways in the shearing zones of
kaolin subjected to direct shear tests. Engineering Geology 78, 135142.
Tika, T.E., Vaughan, P.R., Lemos, L.J., 1996. Fast shearing of pre-existing shear zones
in soil. Geotechnique 46, 197233.
Torabi, A., Braathen, A., Cuisiat, F., Fossen, H., 2007. Shear zones in porous sand:
insights from ring-shear experiments and naturally deformed sandstones.
Tectonophysics 437, 3750.
van der Zee, W., Urai, J.L., 2005. Processes of normal fault evolution in a siliciclastic
sequence: a case study from miri, sarawak, Malaysia. Journal of Structural
Geology 27, 22812300.
Weber, K.J., Mandl, G., Pilaar, W.F., Lehner, F., Precious, R.G., 1978. The role of faults in
hydrocarbon migration and trapping in Nigeriangrowth fault structures. In: 10th
Annual Offshore Technology Conference Proceedings, vol. 4, pp. 26432653.
Wibberley, C.A.J., Yielding, G., Di Toro, G., 2008. Recent Advances in the Under-
standing of Fault Zone Internal Structure: A Review. In: Geological Society,
London, Special Publications, vol. 299 533.
Wibberley, C.A.J., Shimamoto, T., 2003. Internal structure and permeability of major
strike-slip fault zones: the median tectonic line in mie prefecture, southwest
Japan. Journal of Structural Geology 25, 5978.
Yielding, G., Freeman, B., Needham, T., 1997. Quantitative fault seal prediction.
American Association of Petroleum Geologists Bulletin 81, 897917.
F. Cuisiat, E. Skurtveit / Journal of Structural Geology 32 (2010) 18501863 1863