Petrogenesis of the Dunite Peak ophiolite, south-central Yukon, and the distinction between upper-plate and lower-plate settings: A new hypothesis for the late Paleozoic–early Mesozoic tectonic evolution of the Northern Cordillera

1 1 Petrogenesis of the Dunite Peak ophiolite, south-central Yukon and the 2 distinction of upper plate and lower plate settings: a new hypothesis for 3 the late Paleozoic – early Mesozoic tectonic evolution of the Northern 4 Cordillera 5 A.J. Parsons1*, A. Zagorevski2, J.J. Ryan1, W.C. McClelland3, C.R. van Staal1, M.J. Coleman4, M.L. 6 Golding1. 7 * andrew.parsons@earth.ox.ac.uk 8 1 Dept. Earth Sciences, University of Oxford, South Parks Road, Oxford, OX1 3AN, United Kingdom. 9 2 Geological Survey of Canada, 1500-605 Robson Street, Vancouver, BC, V6B5J, Canada. 10 3 Geological Survey of Canada, 601 Booth St., Ottawa, ON, K1A 0E8, Canada. 11 4 Dept. Earth and Environmental Sciences, 115 Trowbridge Hall, University of Iowa, Iowa City, IA, 52242, 12 USA. 13 5 14 Canada. Dept. Earth and Environmental Sciences, University of Ottawa, 120 University, Ottawa, ON. K1N 6N5, 15 16 ABSTRACT 17 Upper plate and lower plate settings within subduction zones have distinct geological signatures. Identifying 18 and discriminating between these settings is crucial to the study of accretionary orogens. We apply this 19 distinction to the Northern Cordillera in Yukon, British Columbia and Alaska, and focus on the 20 identification of upper plate and lower plate domains during the late Paleozoic to early Mesozoic evolution 21 of the allochthonous Yukon-Tanana terrane, the west Laurentian margin and the intervening Slide Mountain 22 Ocean. We present new data from the Dunite Peak ophiolite in south-central Yukon, previously interpreted 23 as ocean plate stratigraphy that was obducted from the subducting Slide Mountain Ocean (i.e. lower plate). 24 Whole-rock geochemical and Sm-Nd isotopic analyses, and U-Pb zircon geochronology indicate that the 25 Dunite Peak ophiolite formed in an intra-oceanic suprasubduction zone setting (i.e. upper plate) with 2 26 magmatism at 265 ± 4 Ma. We propose that the Dunite Peak ophiolite correlates with other mid-Permian 27 suprasubduction zone ophiolites of the Slide Mountain terrane, collectively defining the previously 28 unrecognized mid-Permian Dunite Peak intra-oceanic arc. This intra-oceanic arc was active from ~280 to 29 260 Ma, located within the Slide Mountain Ocean, between the Yukon-Tanana terrane and west Laurentia. 30 Existence of this arc is incompatible with previous models which proposed that accretion of the Yukon- 31 Tanana terrane to Laurentia was facilitated by Permian subduction of Slide Mountain Ocean beneath the 32 Yukon-Tanana terrane. Our results, combined with existing datasets suggest that during the mid- to late 33 Permian, Yukon-Tanana terrane subducted eastward beneath the Dunite Peak intra-oceanic arc. Subsequent 34 collision and accretion of the Yukon-Tanana – Dunite Peak composite terrane with Laurentia must have 35 occurred after the Middle Triassic. 36 37 3 38 1. INTRODUCTION 39 The understanding of subduction, accretion and collision processes in active and ancient settings has been 40 vastly improved by modern geochemical and geochronological constraints. Specifically, subduction zone 41 upper-plate settings and subduction zone lower-plate settings in accretionary orogens may be discriminated 42 using key geochemical indicators, especially with respect to the origin and timing of formation of ophiolites 43 (e.g. Wilson, 1989; Wakabayashi & Dilek, 2000; Pearce, 2008, 2014; Dilek & Furnes, 2014; McGoldrick et 44 al., 2017). In this study, we investigate the petrogenesis of Permian ophiolites associated with the Slide 45 Mountain terrane (SMT) of the Northern Cordillera accretionary orogen in Alaska, Yukon and British 46 Columbia (Figure 1), with particular focus on distinction between subduction zone upper-plate and 47 subduction zone lower-plate processes. This includes new field, geochemical and geochronological analyses 48 of the Dunite Peak ophiolite (de Keijzer et al., 2000; Parsons et al. 2017a, b) in south-central Yukon, which 49 has been assigned previously to the SMT (e.g. Colpron et al., 2016). Previous studies consider the SMT to 50 be the accreted remnants of the Slide Mountain Ocean, which they interpreted as a back-arc oceanic basin 51 between the western Laurentian margin and the allochthonous Yukon-Tanana terrane (YTT, a continental 52 island arc), during the Late Devonian to Permian (see review of Nelson et al., 2013). 53 In this study, we first present the concept of upper-plate and lower-plate distinctions in subduction zone 54 settings and their relevance for studies of accretionary orogens. This is followed by an introduction to 55 current tectonic models for the Northern Cordillera and the locally defined mid- to late Permian ‘Klondike 56 orogeny’ (260-252.2 Ma) as proposed by previous works (e.g. Mortensen, 1992a; Beranek & Mortensen, 57 2011; Nelson et al., 2006; 2013; Colpron et al., 2006a; 2007). We present new data from the Dunite Peak 58 ophiolite (Figure 1e), integrate it with existing data from the SMT, and argue that most mid-Permian 59 ophiolites assigned to the SMT are supra-subduction zone (SSZ) ophiolites, derived from the upper plate of 60 a previously unrecognized intra-oceanic arc. We define this arc as the Dunite Peak intra-oceanic arc. 61 Our findings are incompatible with current tectonic models for Northern Cordillera and require an 62 alternative explanation. We integrate our findings with existing datasets to present a new hypothesis for the 63 late Paleozoic to early Mesozoic evolution of the Northern Cordillera. This study not only bears significance 4 64 for our understanding of the Northern Cordillera, but also demonstrates the importance of identifying and 65 distinguishing between upper-plate and lower-plate components in any accretionary orogen, and how 66 misidentification of such components can result in markedly different tectonic interpretations. 67 1.1. Upper plate – lower plate distinction in accretionary orogens 68 Distinguishing subduction zone upper-plate processes/components from subduction zone lower-plate 69 processes/components is a complicated and yet essential part of any study of accretionary orogens (e.g. 70 Wakabayashi & Dilek, 2000; Zagorevski & van Staal, 2011; McGoldrick et al., 2017). Here, we define 71 accretion as the sequential addition of material from a subducting lower plate to an overriding upper plate or 72 vice versa, via underplating obduction or transform faulting. We define the term collision as entry of a lower 73 plate oceanic plateau, arc or continental lithosphere into a subduction zone. During the life of an active 74 subduction zone, it is possible to classify the structural components that interact with it into one of three 75 groups: (1) Upper-plate material of the active subduction zone. These include the arc, fore-arc and back-arc, 76 plus the arc substrate comprising continental lithosphere, oceanic lithosphere or fragments of the two. Where 77 oceanic lithosphere is present in the upper plate, SSZ ophiolites may also form in this setting (e.g. Dilek & 78 Furnes, 2014). (2) Subducted lower-plate material. This group accounts for almost all lower-plate oceanic 79 lithosphere that enters a subduction zone and is rarely preserved in the geological record as most of this 80 material subducts into the mantle (e.g. Dewey, 2003; Wakabayashi, 2017). At the point of cessation of 81 subduction, this lower-plate material may be preserved, juxtaposed to the upper plate via a suture. In such 82 cases, cessation of subduction typically corresponds to a collision, facilitated by the arrival of a lower-plate 83 arc, continent, or ocean plateau into the subduction zone (e.g. Cloos, 1993; Mann & Taira, 2004; Afonso & 84 Zlotnik, 2011; Brown et al., 2011). (3) Accreted lower-plate material. Material derived from the lower plate 85 that is accreted to the upper plate during subduction-accretion (e.g. the Franciscan complex, California; 86 Wakabayashi, 2017). For ancient subduction zones this is usually the only remaining record of the lower 87 plate and is commonly used to reconstruct the paleo-ocean plate stratigraphy of the lower plate (e.g. Isozaki 88 et al., 1990; Kusky et al., 2013; Wakabayashi, 2017). 5 89 Distinguishing between upper-plate and accreted lower-plate material may be challenging in any subduction 90 zone setting, but is particularly difficult when studying polyphase accretionary orogens such as the Central 91 Asian orogenic belt (e.g. Lin et al., 2018), the Appalachians (e.g. van Staal et al., 2009; Zagorevski & van 92 Staal, 2011) or the Northern Cordillera (e.g. Nelson et al., 2013). Such orogens preserve episodic collisions, 93 formation of composite terranes, and subduction polarity reversals, which result in upper- and lower-plate 94 material of older subduction zones (including SSZ ophiolites) residing in either the lower or upper plates of 95 a younger subduction zone (e.g. Lush’s Bight ophiolite, Newfoundland, Zagorevski & van Staal, 2011). In 96 such settings, the study of ophiolites and accretionary complexes can provide an invaluable record of 97 subduction-accretion and paleo-subduction polarity during multiple collisional and suturing events (e.g. 98 Dilek & Furnes, 2014; Zagorevski & van Staal, 2011; Wakabayashi, 2017). However, misidentification of 99 upper-plate, accreted lower-plate and subducted lower-plate material in an accretionary orogen can lead to 100 101 102 unlikely or implausible tectonic models, as is discussed below (section 2). 2. THE NORTHERN CORDILLERA AND THE KLONDIKE OROGENY (PREVIOUS WORK) 103 The North American Cordillera accretionary orogen (Figure 1) has served as a type example for accretionary 104 orogenesis since the establishment of the terrane concept (Coney et al., 1980). It has a complex tectonic 105 evolution that involved multiple extensional, collisional and magmatic events since the Proterozoic, 106 followed by as much as 1000-2000 km lateral displacement of accreted allochthonous and parautochthonous 107 units during the Cenozoic (see review of Nelson et al., 2013). This study focuses on the Paleozoic-Mesozoic 108 accretionary history of YTT, SMT and the Laurentian margin, which occurred prior to Mesozoic to Eocene 109 dextral translation of terranes along the Cordilleran margin (e.g. Gabrielse et al., 2006). 110 It is generally accepted that YTT is an allochthonous continental island arc terrane (Figure 1) built on a 111 substrate of peri-cratonic (Laurentian) continental crust (e.g. Colpron et al., 2006a; Piercey & Colpron, 112 2009; Nelson et al., 2013). The SMT (Figure 1) comprises Mississippian to early Permian supracrustal 113 igneous and sedimentary oceanic rocks and mid-Permian mafic-ultramafic ophiolites that are interpreted as 114 vestiges of accreted ocean floor stratigraphy of the Slide Mountain Ocean (see summaries by Nelson et al., 6 115 2006, 2013). Most models propose that YTT rifted from the west margin of the Laurentia during the Late 116 Devonian – Early Mississippian via back-arc rifting of the Yukon-Tanana arc to form the Slide Mountain 117 Ocean (e.g. Mortensen, 1992a; Nelson et al., 2006; 2013; Colpron et al., 2006a, 2007). These models then 118 propose a subduction polarity reversal across YTT from east-dipping subduction of the Panthalassa Ocean to 119 west-dipping subduction of the Slide Mountain Ocean, which led to re-accretion of YTT to Laurentia and 120 suturing of Slide Mountain Ocean in the late Permian (e.g. Beranek & Mortensen, 2011). 121 Beranek and Mortensen (2011) referred to late Permian (260-252.5 Ma) collision between YTT and 122 Laurentia as the ‘Klondike orogeny,’ although similar versions of this model had been proposed prior to 123 their work by others (e.g. Mortensen, 1992a; Dusel-Bacon et al., 1995; 2002; Nelson et al., 2006; Berman et 124 al., 2007; Colpron et al., 2007; Johnston et al., 2007). Most studies concerning the Paleozoic to early 125 Mesozoic evolution of YTT, SMT and Laurentia conducted since Beranek and Mortensen (2011) have been 126 interpreted within, and supportive of this framework for the Klondike orogeny (e.g. Nelson et al., 2013; 127 Petrie et al., 2015; 2016; Colpron et al., 2015; 2016; Golding et al., 2016a; Staples et al., 2016; Gilotti et al., 128 2017). Such hypotheses typically cite the following as supporting evidence: 129 (1) Similarities in lithostratigraphy and detrital zircon populations of pre-Late Devonian metasedimentary 130 basement of YTT and metasedimentary rocks of Laurentia are interpreted as indication that YTT initially 131 formed part of the Laurentian margin prior to rifting (e.g. Colpron et al., 2006a; Piercey & Colrpon, 2009). 132 (2) Occurrences of Mississippian to Permian arc plutonic and volcanic rocks within YTT that are not 133 recorded within Laurentia are interpreted as indications that YTT rifted from Laurentia during the Late 134 Devonian and then evolved independently, remaining allochthonous to Laurentia at least until the end of the 135 Permian (e.g. Mortensen, 1992a; Nelson et al., 2006; 2013). The presence of Mississippian and Permian to 136 Triassic metamorphism in YTT but its absence from Laurentia are also used as evidence for separation until 137 the Triassic (e.g. Berman et al., 2007; Staples et al., 2016). 138 (3) Outcrops of SMT located along the present-day YTT-Laurentia boundary are interpreted as Slide 139 Mountain ocean floor stratigraphy accreted to YTT and/or Laurentia, marking a suture that formed 7 140 following Permian subduction and closure of the Slide Mountain Ocean (e.g. Mortensen, 1992a; Murphy et 141 al., 2006). 142 (4) Mid-Permian blueschist and eclogite assumed to be metamorphosed slivers of Slide Mountain oceanic 143 crust (e.g. Creaser et al., 1997; 1999; Erdmer et al., 1998), and mid- to late Permian Sulphur Creek plutonic 144 suite and volcanic Klondike schist (collectively referred to as the Klondike assemblage), within YTT (e.g. 145 Mortensen, 1990; Colpron et al., 2006a) have been interpreted as a ‘paired metamorphic belt’ that recorded 146 subduction of Slide Mountain Ocean beneath YTT (e.g. Mortensen, 1992a; Nelson et al., 2006; Beranek & 147 Mortensen, 2011). Beranek and Mortensen (2011) split the Klondike cycle into an early subduction-related 148 phase of arc magmatism, deformation and metamorphism between 260 and 254 Ma and a later phase of 149 crustal-melt derived post-tectonic magmatism between 254 and 252 Ma. These studies bracket the closure of 150 Slide Mountain Ocean and the Klondike orogeny between 260 and 252.5 Ma (Beranek & Mortensen, 2011; 151 Nelson et al., 2013). 152 (5) The Triassic Jones Lake Formation (Beranek et al., 2010) that overlies parts of Laurentia, SMT and YTT 153 in Yukon and British Columbia has been interpreted as a syn-orogenic sedimentary cover succession 154 initiated in the Early Triassic (e.g. Colpron et al., 2005; 2006; 2007; Beranek et al., 2010; Beranek & 155 Mortensen, 2011; Golding et al., 2016a). Mississippian to Permian detrital zircon populations from these 156 strata have been interpreted as indications of a YTT-derived sediment source, which formed an overlap 157 succession marking the final stages of collision, accretion and related subsidence between YTT and 158 Laurentia (Colpron et al., 2006; 2007; Beranek & Mortensen, 2011). 159 2.1. Unresolved issues, conflicting datasets and alternative models 160 Other datasets and models for the tectonic evolution of YTT, SMT and Laurentia disagree, or are hard to 161 reconcile with models of the Klondike orogeny. These are summarized below and are considered in light of 162 our own findings, following presentation of our geochemical and geochronological analyses from the Dunite 163 Peak ophiolite. These include: 8 164 (1) Mid-Permian eclogites previously interpreted as metamorphosed slices of SMT oceanic crust have since 165 been identified as mafic intrusions within YTT (Petrie et al., 2016). Furthermore, it has been demonstrated 166 that the YTT metasedimentary rocks which host these intrusions also record mid-Permian eclogite facies 167 metamorphism (Gilotti et al., 2017). As such, deformation, metamorphism and magmatism previously 168 attributed to the Klondike orogeny (e.g. Mortensen, 1990; 1992a; Berman et al., 2007; Beranek & 169 Mortensen 2011; Staples et al. 2016; Petrie et al., 2016; Gilotti et al., 2017) are exclusively recorded in 170 YTT, which forms the upper plate in these models. There is no evidence of contemporaneous lower-plate 171 (i.e. Laurentia) deformation or metamorphism to support these models (Berman et al., 2007; Staples et al., 172 2016). 173 (2) Existing models do not account for the presence of mid-Permian igneous rocks in SMT with 174 geochemical signatures that require formation in an SSZ setting (e.g. Nelson et al., 1993; Plint & Gordon, 175 1997; Dusel-Bacon & Cooper, 1999; Fallas et al., 1999; Colpron et al., 2005; 2006b; Dusel-Bacon et al., 176 2006). These data are incompatible with models which propose that Slide Mountain Ocean subducted 177 beneath YTT during the mid-Permian (e.g. Mortensen, 1992a; Murphy et al., 2006; Nelson et al., 2006; 178 2013; Johnston et al., 2007; Beranek & Mortensen 2011). 179 (3) Detrital zircons and conodont fauna do not validate, and in some cases are inconsistent with a model of 180 deposition of Triassic syn-orogenic clastic sediments in a regional-scale basin overlapping YTT, SMT and 181 Laurentia following Permian collision (e.g. Colrpon et al., 2007; Beranek et al., 2010; Beranek & 182 Mortensen, 2011). With the exception of three Mississippian zircons, detrital zircon populations of Early 183 Triassic strata on the west Laurentian margin in Yukon (cf. Beranek et al., 2010) may be entirely explained 184 by a Laurentian source (Archean, Proterozoic and pre-340 Ma Paleozoic zircons). Middle to Late Triassic 185 strata deposited on Laurentia in Yukon and northern British Columbia contain pre-340 Ma zircons that may 186 be explained by a Laurentian source or by a YTT source, plus Mississippian to Triassic zircons that require 187 either a YTT source (i.e. a Laurentian zircon population plus abundant Mississippian to Permian zircons), or 188 Stikinia/Quesnellia source (i.e. Mississippian, minor Permian and abundant Triassic zircons). In contrast, 189 most Middle Triassic strata overlying SMT in Yukon and northern British Columbia (e.g. Nelson & 9 190 Bradford, 1993; Murphy et al., 2002; Colpron et al., 2005) have unimodal/bimodal Permian-Triassic ± 191 Devonian – Early Mississippian detrital zircon populations (Beranek & Mortensen, 2011). This is suggestive 192 of sediment derived from a late Paleozoic – early Mesozoic arc such as Stikinia/Quesnellia, and is not 193 consistent with sediment derived from YTT ± Laurentia, which would also contain Pennsylvanian, 194 Devonian to Cambrian, Proterozoic and Archean zircons. In the Finlayson Lake district (Figure 1c), the Late 195 Triassic Bug Island Limestone, which overlies SMT, yielded a detrital zircon population of Laurentian 196 zircons plus five Mississippian zircons that may suggest a YTT source (Beranek, 2009). Conodont fauna 197 from this limestone includes the “exotic” Tethyan species Norigondolella hallstattensis (Orchard, 2006) that 198 has been reported from other allochthonous terranes (e.g. Wrangell terrane; Orchard, 1991), but is not 199 recorded from anywhere along the autochthonous Laurentian margin (Orchard, 2006; Beranek, 2009). 200 Additionally, the Bug Island Limestone overlies Middle Triassic clastic sedimentary rocks with 201 unimodal/bimodal zircon populations that are not suggestive of a Laurentian or YTT source (Beranek, 2009; 202 Beranek & Mortensen, 2011). It is therefore unlikely that these Middle to Late Triassic strata that overlie 203 parts of SMT were deposited in the same basin system as coeval strata deposited on Laurentia (Orchard, 204 2006). 205 In central to southern British Columbia, Anisian (Middle Triassic) to Rhaetian (Late Triassic) strata 206 overlying Laurentia are interpreted as foreland basin deposits derived from YTT and Laurentia following 207 their collision in the Permian (Golding et al., 2016a). However, detrital zircon populations in these strata can 208 be explained by a Laurentian source with the minor addition of sediment from YTT or Stikinia, to account 209 for the presence of four zircon grains with Mississippian ages and three zircon grains with Permian ages, 210 (Ferri, 2009; Ferri et al., 2010; Golding et al., 2016a). Additionally, Rhaetian clastic strata in central British 211 Columbia analysed by Golding et al. (2016b), are dominated by Late Triassic euhedral zircon grains that 212 suggest a volcanic source from Quesnellia. As such, these data indicate that Middle to Late Triassic 213 sedimentary strata in southern and central British Columbia can be explained coherently with sources 214 derived from Laurentia and Stikinia/Quesnellia, with no demand for a YTT sediment source (Golding et al., 215 2016a, b). As such, detrital zircon data from Triassic strata in Yukon, British Columbia and one sample in 216 Alaska do not indicate that YTT collided and accreted to Laurentia during the Permian. 10 217 (4) Paleomagnetic studies from across the Northern Cordillera suggest that between 90-70 Ma and 50 Ma, 218 YTT, SMT and parautochthonous parts of Laurentian were fixed together as they moved northwards to their 219 current position relative to stable North America (see reviews of Gabrielse et al., 2006 and Enkin, 2006). 220 Prior to this, paleomagnetic data from Middle Pennsylvanian – early Permian red cherts from SMT in the 221 Sylvester allochthon (Figure 1a) and at Sliding Mountain, in northern and central British Columbia, suggest 222 that these parts of SMT were deposited ~20° south of their current locations with respect to stable North 223 America (Richards et al., 1993). Secondary magnetization suggests that following deposition at equatorial 224 latitudes, these chert units migrated northwards with respect to Laurentia until the Early to Middle Jurassic. 225 Based on the occurrence of an Early Jurassic chemical remagnetization event (Cole et al., 1992) recorded 226 locally within the Cache Creek terrane (Figure 1), and on the intersection of the small circle for the 227 secondary magnetization pole in these chert units with the Middle to Late Jurassic portion of the apparent 228 polar wander path for Laurentia (their Figure 8), Richards et al. (1993) favoured an Early to Middle Jurassic 229 timing of obduction and accretion of these SMT units to Laurentia. This contradicts the Permian timing for 230 closure of Slide Mountain Ocean favoured by most authors. 231 (6) Of most striking contrast to the Klondike orogeny model of Beranek and Mortensen (2011) and others, 232 are the Cordilleran Ribbon Continent models for the Northern American Cordillera (Johnston, 2008; 233 Hildebrand, 2009). These models proposed that parts of the parautochthonous Laurentian margin, such as 234 the Cassiar platform (CA – Figure 1) in Yukon and British Columbia are actually exotic with respect to 235 Laurentia and collided with the intermontane terranes (YTT, Stikinia, Quesnellia) in the Late Triassic to 236 form a composite continental ribbon. Accretion of the ribbon continent to Laurentia is proposed to occur in 237 the Cretaceous (Johnston, 2008; Hildebrand, 2009). This model has received little support from the scientific 238 community (e.g. Monger, 2014; Colpron et al., 2015; Sigloch & Mihalynuk, 2017; Matthews et al., 2017; 239 2018), partly because the occurrence of a Cretaceous suture between ribbon continent and Laurentia, which 240 should be located inboard of the Cassiar platform, remains cryptic (e.g. Johnston, 2008; Monger, 2014). 241 3. STUDY AREA: THE DUNITE PEAK OPHIOLITE AND LOCAL GEOLOGY 11 242 The Dunite Peak ophiolite (DPO, Figure 1e and 2) forms klippen of mid-Permian mafic-ultramafic rocks 243 (267 ± 10 Ma, de Keijzer et al., 2000) previously assigned to SMT (Figure 2), overlying rocks previously 244 assigned to YTT (Parsons et al., 2017a,b). The location of the DPO study area (Figure 2) with respect to the 245 major terrane boundaries is indicated on Figure 1. The tectonic boundary between YTT and Laurentia 246 (Cassiar terrane, Figure 1), is inferred to lie 5 km east of the study area shown in Figure 2 (Colpron et al., 247 2016a). This boundary is structurally equivalent to the Middle Triassic to Early Jurassic ductile and post- 248 Jurassic brittle Tummel fault zone (Figure 1i, Colpron et al., 2005) and the post-Late Triassic Inconnu thrust 249 (Murphy et al., 2006). To the west of the study area, the N-S striking d’Abbadie fault system (Figure 2) had 250 an early normal sense of motion and subsequent post-96 Ma right-lateral strike-slip motion with ~4 km 251 dextral displacement that truncated the DPO along its length (de Keijzer et al., 1999; Colpron et al., 2017). 252 The dominant transposition foliation within the DPO and all underlying units is sub-parallel to 253 compositional layering. Mineral stretching lineations and shear-sense indicators in these units broadly 254 correspond to a NE-SW transport direction, most commonly with a top-SW shear sense. These fabrics are 255 folded by an open synform that plunges gently WSW (Figure 2). The bedrock geology in the Dunite Peak 256 region (Figure 2) may be split into three distinct units, listed from structurally lowest to highest: (1) A basal 257 marble unit; (2) a marine metasedimentary succession; and (3) the Dunite Peak ophiolite. Within each of 258 these, unit subdivisions are identified, as described below. Additional field observations from this area are 259 described by Parsons et al., 2017a, b and de Keijzer et al. (1999, 2000). 260 3.1. Basal marble unit 261 The structurally lowest unit (Figure 2) comprises coarse grained massively bedded marble, weathered 262 yellow-grey, ≥600 m thick, with rare garnet-kyanite bearing metapelitic and garnet-amphibole-epidote 263 bearing meta-igneous layers. This basal marble unit is identified as part of the Snowcap assemblage of YTT 264 based on the occurrence of Mississippian Simpson Range suite orthogneiss within this metasedimentary 265 unit, both north and south of the study area (de Keijzer et al., 2000; Westberg et al., 2009, 2010; Colpron et 266 al., 2016a). Based on its lithology, the basal marble unit probably forms the northern continuation of the 12 267 Scurvy Creek succession, a locally defined subdivision of the Snowcap assemblage which crops out 10-30 268 km south of our study area (Westberg et al., 2009; 2010; Colpron et al. 2016a). 269 3.2. Marine metasedimentary succession 270 The marine metasedimentary succession has a structural thickness of ~350-500 m (Figure 2) and is visually 271 distinct from the underlying basal marble unit. From bottom to top, this succession divides into four mapped 272 units (Figures 2 and 3a): (1) MMS1 – Micaceous and carbonaceous quartzite and carbonaceous shale (~100- 273 150 m thick); (2)MMS2 – white quartzite overlain by interbedded quartzite, marble and subordinate 274 carbonaceous calc-silicate layers (~100 m thick); (3) MMS3 – a unit of carbonaceous semipelite and shale, 275 which locally contains an upper portion of thinly interbedded carbonaceous shale/semipelite and subordinate 276 chert, quartzite (silt/sandstone) and metavolcanics (150-250 m thick); at one outcrop, pillow basalt was also 277 observed within this unit; and (4) MMS4 – dark grey, micritic, mylonitized marble (200-250 m thick) with 278 locally less deformed and fossiliferous zones containing deformed bryozoa and unidentified shelly fauna. 279 Mylonitic foliation within the marine metasedimentary succession increases in deformation intensity up- 280 section towards the base of the DPO crustal section. A metamorphic assemblage of biotite ± garnet ± 281 staurolite ± chlorite suggests that this unit is of a lower metamorphic grade than the basal marble unit. Based 282 on its lithology and structural relationship with the basal marble unit we speculate that the marine 283 metasedimentary succession may correlate with either the Nasina assemblage or Swift River Group of the 284 Finlayson assemblage (e.g. Colpron et al., 2006a; Nelson et al., 2006). 285 3.3. The Dunite Peak ophiolite (DPO) 286 The DPO has a variable structural thickness of ~450 m to >1650 m, and comprises a crustal section of mafic 287 to intermediate volcanic/volcaniclastic and plutonic rocks, structurally overlain by ultramafic rocks (Figures 288 2 and 3). 289 Crustal Section (DPc) 290 The DPO crustal section (DPc, Figure 2) is an internally deformed thrust stack of metamorphosed basaltic to 291 andesitic volcanics and volcaniclastics, gabbro and leucogabbro with a variable thickness of ~300-650 m 292 (Figures 2 and 3). The degree of metamorphism and deformation in these rocks is such that it is often not 13 293 possible to distinguish between volcanic, volcaniclastic and hypabyssal lithologies. Henceforth, we use the 294 term ‘greenstone’ to collectively refer to the metamorphosed products of these rock types. Generally, the 295 lithologies of this unit (Figure 4) transition from a lower portion of fine to medium grained (≤1 mm) 296 greenstone to an upper coarser grained portion of interlayered greenstone, gabbro and leucogabbro. 297 The base of the DPO crustal section is dominated by light green, very fine grained, homogeneous and 298 locally fissile greenstone with rare pelitic and silicic horizons and sheared pillows (Figure 4a-c). These rocks 299 probably represent volcanic and volcaniclastic successions erupted/deposited in a subaqueous environment 300 with intermittent chert and argillite deposition. Up-section, the DPO crustal section is dominated by gabbro 301 and coarser grained pyroxene- ± plagioclase-phyric greenstone that probably represent a mix of plutonic and 302 vent-proximal volcanic and hypabyssal successions (Figure 4d-e). In the upper portion of the crustal section, 303 pelitic and silicic horizons are not observed, and gabbro and leucogabbro commonly intrude greenstone 304 (Figure 4e-f). Variability in grain size and in relative proportions of leucocratic to melanocratic minerals 305 within the gabbro and leucogabbro define compositional layering sub-parallel to foliation with locally 306 undulating and discordant boundaries corresponding to original intrusive relationships. The top of the crustal 307 section consists of medium to coarse grained (~0.5-2 mm) cumulate to adcumulate gabbro, leucogabbro and 308 plagiogranite and subordinate fine grained greenstone layers (Figure 4f). Leucogabbro contains rare rafts of 309 greenstone. 310 The DPO crustal section is pervasively sheared with a mylonitic fabric at its lower and upper contacts. 311 Thrust imbrication of this unit incorporates thrust-bounded slivers of the marine metasedimentary succession 312 (MMS3, Figure 2). A peak metamorphic mineral assemblage of hornblende-actinolite + epidote-clinozoisite 313 + albite-oligoclase is indicative of upper greenschist to lower amphibolite facies conditions. Peak 314 metamorphic conditions do not appear to vary between vertical structural positions or between 315 metavolcanic/volcaniclastic units and metaplutonic units. As such, the observed metamorphism of the DPO 316 probably corresponds to burial and heating during or after emplacement, rather than metasomatism during 317 ophiolite formation, which would display increasing metamorphic grade from supracrustal to plutonic 318 sections of an ophiolite. 319 Ultramafic Section (DPu) 14 320 The ultramafic section of the DPO is the structurally highest unit (DPu, Figures 2 and 3) and forms isolated 321 klippen emplaced onto the DPO crustal section (DPc) with variable thicknesses of ~150 m to >1000 m 322 (Figures 2-3). The lowermost ~5-50 m of these klippen (DPu1, Figure 2) are ubiquitously defined by 323 sheared, locally cataclastic blue-green and dark-blue-black ‘fish-scale’ serpentinite (Figure 3b). This 324 transitions up-section into orange-weathered, variably serpentinized harzburgite and dunite with subordinate 325 lherzolite (DPu2, Figure 2). Pyroxene-rich layers define compositional layering within these rocks, which 326 vary in orientation across the klippen. Planar intrusions of pyroxenite and gabbro cross-cut these ultramafic 327 rocks (DPu3, Figures 2 and 3b). Dunites contain aggregates of undeformed olivine and serpentine- 328 psuedomorphs of olivine with interlocking, adcumulate textures and no interstitial material between grains. 329 Harzburgites contain orthopyroxene phenocrysts with oikocrysts of olivine and clinopyroxene, plus chromite 330 with orthopyroxene overgrowths that may be indicative of exsolution cumulate residuum. The basal 331 serpentinite of the DPO ultramafic section (DPu1) displays S-C fabrics with C-planes parallel to mylonitic 332 foliation in underlying DPO crustal section. S-planes show no consistent transport direction. Above this, 333 thrust faulting and chaotic folding decrease in intensity up-section until they are absent from the orange- 334 weathered upper portion of planar-stratified ultramafic layers (DPu2). 335 4. U-Pb ZIRCON GEOCHRONOLOGY 336 Zircon U-Pb isotopic analysis was conducted for geochronology on sample 16RAY-AP074A1; a coarse 337 grained gabbro intruded into greenstone from the top of the DPO crustal section (Figure 2). This gabbro 338 does not cross-cut the deformation fabric in the greenstone. Analyses were conducted by secondary ion 339 microprobe spectrometry (SIMS) at the U.S. Geological Survey-Stanford University SHRIMP-RG (sensitive 340 high resolution ion microprobe-reverse geometry) facility. Analytical procedures are described in detail in 341 Supplementary Materials 01. The full U-Pb isotopic dataset is presented in Supplementary Materials 02. 342 Sample 16RAY-AP074A1 yielded euhedral to subhedral zircon with oscillatory zoning and some thin 343 cathodoluminescence (CL) dark, high-U rims that may reflect resorption (Figure 5a). Spot analyses (Figure 344 5a) conducted on 12 zircon grains produce a simple cluster of grain ages overlapping concordia and, 345 excluding the oldest analysis, define a weighted mean 206 Pb/238U age of 265 ± 4 Ma (MSWD = 0.9; Figure 15 346 5b). The oldest analysis (283 ± 4 Ma) was obtained from a euhedral, CL-bright, low-U (50 ppm) grain core 347 domain (spot 4.1, Figure 5a) and inclusion of this grain yields a mean age of 268 ± 6 Ma (MSWD = 2.7). 348 Whilst analytical scatter cannot be ruled out, textural evidence suggests that the core domain represents a 349 xenocryst that records an earlier magmatic history. Chondrite-normalized trace element signatures are 350 depleted in LREEs relative to HREEs and have little variation between each grain (Figure 5c). Together, the 351 trace element signatures, grain morphology and spread of individual grain ages suggest that the calculated 352 mean age reflects the timing of crystallization of the magmatic protolith. Ti-in-zircon thermometry (Ferry & 353 Watson, 2007) yields a mean zircon crystallization temperature of 761±21 °C (1σ error). 354 5. GEOCHEMISTRY 355 The degree of metamorphism and deformation recorded by the DPO makes it difficult to identify and 356 distinguish between specific igneous protoliths using field observations and optical microscopy alone. We 357 utilize whole-rock major and trace element and Sm-Nd isotopic analyses to provide additional constraints on 358 protoliths and igneous petrogeneses of different rock types in the DPO. Whole-rock geochemical analyses of 359 52 samples were conducted by Activation Laboratories Ltd. (Ancaster, ON). Sm-Nd isotopic analysis of 7 360 samples was conducted at Carleton University (Ottawa, ON). Analytical procedures for both analyses are 361 described in Supplementary Materials 01. 362 5.1. Immobile element fingerprinting 363 For interpretation of the whole-rock element analysis, only immobile rare earth elements (REEs) and high 364 field strength elements (HFSEs) are considered, in order to minimize the effects of element mobility during 365 metamorphism (e.g. Pearce, 1996; 2014). REE-HSFE profiles (Figure 6) and normalized element 366 concentrations and ratios (Table 1) are given relative to N-MORB (Sun & McDonough, 1989, plus V and Sc 367 from Klein, 2004) for Groups 1-4, defined below and relative to primitive mantle (Sun & McDonough, 368 1989, plus V and Sc from McDonough & Frey, 1989) for Group 5, defined below. Immobile element ratios 369 Nb/Y and Zr/Ti provide proxies for Na2O+K2O (total alkalinity) and SiO2, respectively (Cann, 1970; Pearce, 370 1996) (Figure 7a). Fractionation of immobile elements Y, La and Nb provides a means to classify crustal 371 samples in terms of magmatic setting, due to the variable sensitivity of these elements to degree of partial 16 372 melting/fractionation, composition of melt source and subduction zone interaction (Figure 7b) (Cabanis & 373 Lecolle 1989). Relative concentrations of Th, Nb, and Yb, provide a means to distinguish suprasubduction 374 zone (SSZ) settings from intra-plate and mid-ocean spreading ridge settings (Figure 7c) (Pearce, 2008, 375 2014). Th and Nb behave similarly in the mantle (Th/Nb ≈ 1), except during subduction where Nb is 376 retained in the subducting slab, whilst Th is released from the subducting slab into the mantle wedge. This 377 results in high positive Th/Nb ratios in mantle wedge-derived magmas (Pearce, 2008). The concentration of 378 Yb, relative to Th and Nb provides an indication of degree of fractionation, which may be used to 379 distinguish intra-oceanic arc from continental arc settings and mid-ocean ridge from intra-plate settings 380 (Pearce, 2008; 2014). 381 From the 52 samples analyzed, five geochemically distinguishable groups are defined. Groups 1 to 3 contain 382 mafic to intermediate samples from the DPO that are distinguished from each other on geochemical bases, 383 irrespective of lithology, following the identification of three distinct geochemical signatures. Group 4 384 contains a single sample of pillow basalt from the marine metasedimentary succession (MMS3). Group 5 385 contains ultramafic samples from the DPO. Locations of samples are given in Figure 2 with group-specific 386 symbology. Seven fine-grained greenstone samples from the DPO were selected for Sm-Nd isotopic analysis 387 prior to distinction of Groups 1 to 3. Of these seven samples, four derive from Group 1 and three derive 388 from Group 3. The results of Sm-Nd isotopic analyses are reported alongside the whole-rock analyses of 389 these groups. 390 Group 1 (DPO mafic to intermediate rocks) 391 Group 1 (n =18) displays LREE enriched spectra and large Th/Nb anomalies (Figure 6a), and may be 392 divided into two subgroups based on the relative concentrations of Ti, V and Sc (Figure 6a). Group 1a has 393 less enrichment ([La/Yb]N = 2.87, LaN = 1.08, [Th/Nb]N = 14.67) than Group 1b ([La/Yb]N = 4.68, LaN = 394 4.21, [Th/Nb]N = 27.95). Group 1a displays minor negative and a single strong positive Ti anomaly. V and 395 Sc are enriched relative to other HREEs and Ti. These samples display minor negative anomalies in Zr-Hf 396 (Figure 6a). In contrast, Group 1b displays strong negative Ti, V and Sc anomalies, but no Zr-Hf anomalies 397 (Figure 6a). Small positive Eu anomalies are common in both subgroups. Sm-Nd isotopic analysis of four 17 398 samples from Group 1 yielded ƐNd(t=265) values of +7.2, +7.4, +7.5 (Group 1a) and +7.4 (Group 1b) (Table 399 1). 400 Group 2 (DPO mafic to intermediate rocks) 401 Group 2 (n = 11, Figure 6b) displays moderately LREE-enriched to flat spectra ([La/Yb]N = 2.91). Element 402 concentrations are depleted relative to N-MORB by up to a factor of 10 (LaN = 0.85, YbN = 0.32). Group 2 403 samples have large Th/Nb ratios ([Th/Nb]N = 17.43), negative Ti anomalies and strong positive V and Sc 404 anomalies relative to HREEs and Ti (Table 1). Negative Zr-Hf anomalies and positive Eu anomalies are 405 sometimes displayed (Figure 6b). 406 Group 3 (DPO mafic to intermediate rocks) 407 Group 3 (n = 6, Figure 6b) displays flat spectra ([La/Yb]N = 1.41) with N-MORB-comparable 408 concentrations (LaN = 1.49, YbN = 1.07). All samples have positive Th/Nb ratios ([Th/Nb]N = 5.29) and 409 negligible to small positive V anomalies, negative Sc anomalies, but no significant Ti or Eu anomalies. 410 Three samples from Group 3 yielded ƐNd(t=265) values of +8.3, +8.1 and +9.0 (Table 1). 411 Group 4 (pillow basalt) 412 Group 4 (n = 1, Figure 6c) has an LREE-enriched spectrum ([La/Yb]N = 2.14) with N-MORB-comparable 413 concentrations of HREEs (LaN = 1.42, YbN = 0.66). Unlike Groups 1-3, Group 4 has no Nb anomaly 414 ([Th/Nb]N = 1.55), nor does it have Ti, V, Sc or Eu anomalies (Table 1). 415 Group 5 (DPO ultramafic rocks) 416 Group 5 (n = 16, Figure 6d) comprises samples from the DPO ultramafic section (DPu). CIPW normative 417 major element concentrations of Group 5 samples mostly plot within the harzburgite (6 samples) and dunite 418 (6 samples) fields. A minority plot within lherzolite (3 samples) and pyroxenite (1 sample) fields (Figure 419 7d). Two reference compositions for depleted-MORB mantle (DMM – Salters & Stracke, 2004; Workman 420 & Hart 2005) are also included on the REE-HFSE diagram (Figure 6d). Group 5 is depleted in HREEs 421 relative to primitive mantle and DMM (YbN = 0.21). LREEs are enriched relative to HREEs and DMM (LaN 18 422 = 0.50, [La/Yb]N = 2.37). Several samples have positive Th anomalies equal to, or greater than primitive 423 values (ThN = 1.06), and most have Nb concentrations below the analytical detection limit. One sample had 424 detectable amounts of both Th and Nb, producing [Th/Nb]N = 14.26. Positive V and Sc and Zr anomalies are 425 common. 426 5.2. Interpretation of geochemistry results 427 Group 1 comprises greenstone and gabbro from the DPO crustal section (Figure 6a). Differences between 428 Groups 1a and 1b probably reflect crystal fractionation processes. Group 1a plots within subalkaline basaltic 429 to basaltic andesite island arc tholeiite (IAT) fields and probably represents plutonic/hypabyssal IATs 430 (Figure 7a-b). These samples have positive V, Eu, and Sc anomalies, plus one positive Ti anomaly, and 431 negative Zr-Hf anomalies. These are generally attributed to crystallization and retention of magnetite/rutile 432 (Ti enrichment) and clinopyroxene (Sc enrichment) and extraction of an intermediate-acid melt (Zr-Hf 433 depletion) (e.g. Pearce, 1996). Positive V anomalies are generally attributed to hydrous mantle melting (e.g. 434 Pearce, 2014). Positive Eu anomalies are most easily explained by the presence of cumulate plagioclase (e.g. 435 Rollinson, 1993). Positive Th/Nb ratios and negative Nb anomalies suggest a SSZ setting (Figure 7c). 436 Group 1b plots within basaltic-andesite to andesite IAT–calc-alkaline (CA) fields (Figure 7a-b) and 437 represents a more fractionated component of IAT to CA hypabyssal magmatism or volcanism. This is 438 suggested by negative Ti anomalies and the absence of Zr-Hf anomalies; Ti is depleted in moderately 439 evolved magmas due to early fractionation of Ti-bearing phases in basaltic magmas, whereas Zr and Hf 440 remain incompatible during fractionation of intermediate melts. Negative Sc anomalies can be attributed to 441 early clinopyroxene fractionation in basaltic magmas. Negative V anomalies may be attributed to increased 442 compatibility of V during crystal fractionation of intermediate melts (e.g. Pearce, 1996). Positive Th/Nb 443 ratios and negative Nb anomalies suggest a SSZ setting (Figure 7c). ƐNd values from Groups 1a and 1b 444 ranging between +7.2 and +7.5, indicate that these magmas did not interact with continental crust (e.g. 445 Rollinson, 1993). 446 Group 2 comprises layered gabbro, leucogabbro and subordinate greenstone from the DPO crustal section 447 (Figure 6b) with subalkaline basaltic to basaltic andesite compositions (Figure 7a). These data form an array 19 448 from the intersection between IAT-CA and CA fields to the intersection between IAT, N-MORB and back- 449 arc basin basalt (BABB) fields, mostly plotting in the IAT field (Figure 7b). Based on the same 450 interpretations as provided for Group 1a (see above), the depleted nature of Group 2 including Zr-Hf 451 negative anomalies and V-Sc positive anomalies suggests that these rocks represent crustal cumulates 452 derived from hydrous mantle melting. Variations in Ti and V most likely reflect variable proportions of 453 cumulate phase (Ti depleted) versus melt phase (Ti enriched) ± later addition of Ti-bearing fluids, as 454 suggested by the presence of titanite veins in some samples. Positive Th/Nb ratios and negative Nb 455 anomalies suggest a SSZ setting (Figure 7c). 456 Group 3 comprises gabbro intrusions from the DPO ultramafic section and greenstone from the DPO crustal 457 section (Figure 6b) with subalkaline basaltic N-MORB-comparable compositions (Figure 7a-b) plus Nb and 458 V anomalies suggesting subduction zone interaction (Figure 7c) and hydrous mantle melting (e.g. Pearce, 459 1996; 2014). Minor negative Sc anomalies probably reflect earlier crystallization of clinopyroxene (e.g. 460 Pearce, 1996). Group 3 is interpreted as primary BABB melts derived directly from depleted mantle, with 461 little or no subsequent crystal fraction or melt extraction. Positive Th/Nb ratios and negative Nb anomalies 462 suggest a SSZ setting (Figure 7c). ƐNd values from Group 3 ranging between +8.8 and +9.0, indicate that it 463 has not interacted with continental crust (ƐNd of depleted mantle = 9.0-9.6, Kimura et al., 2017). 464 Group 4 is a pillow basalt sample from the marine metasedimentary succession (Figure 6c), and has a 465 geochemical signature comparable to subalkaline E-MORB (Figure 7a-c). The stratigraphic association of 466 Group 4 with the marine metasedimentary succession and absence of E-MORB signatures from the DPO 467 (one Group 1a sample excepted) suggest that Group 4 does not derive from the same SSZ setting as Groups 468 1-3. 469 Group 5 comprises ultramafic rocks from the DPO (Figure 6d) that displays lithological, petrographic (see 470 Section 3.3) and geochemical characteristics typical of lower crustal or mantle cumulates, rather than mantle 471 tectonites. It is dominated by dunite and harzburgite with cumulate and adcumulate textures. Enrichment in 472 LREEs relative to HREEs and large V/Ti ratios suggest that Group 5 derived from partial melting of the 473 mantle under hydrous conditions. Positive Sc anomalies probably reflect the presence of cumulate 20 474 clinopyroxene. Positive Zr anomalies may correspond to enrichment from slab-derived fluids (e.g. Bizimis 475 et al., 1999) or recycling of lower crustal material during crustal foundering (e.g. Arndt & Goldstein, 1989). 476 Strong depletion in Nb (all but one Group 5 sample have Nb concentrations below the detection limit), and 477 enrichment in Th and the presence of BABB intrusions (Group 3) within these rocks suggests that the DPO 478 ultramafic section formed in the same SSZ setting as Groups 1-3. 479 6. PETROGENESIS OF THE DUNITE PEAK OPHIOLITE 480 Large Th/Nb ratios, large positive ƐNd values, enriched concentrations of LREEs and N-MORB-comparable 481 concentrations of HREEs indicate that the DPO is an SSZ ophiolite (Figure 7b-c) comprising a dismembered 482 crustal section and lower crustal/lithospheric mantle cumulates of an intra-oceanic arc system. U-Pb zircon 483 geochronometry from a Group 2 IAT gabbro (this study, Figures 2 and 5) and a leucocratic granitic dike (de 484 Keijzer et al., 2000, Figure 2) within the DPO crustal section yielded respective ages of 265 ± 4 Ma and 267 485 ± 10 Ma. The ages and geochemical signatures of these rocks suggest that constructional volcanism was 486 occurring along an intra-oceanic arc that resided somewhere within the ocean separating YTT and Laurentia 487 during the mid-Permian (currently modelled as the Slide Mountain Ocean). The single zircon xenocryst age 488 (283 ± 4 Ma) from sample 16RAY-AP074A1 (this study), may derive from the Group 1 IAT greenstone into 489 which it intrudes (e.g. Figure 4e-f), corresponding to the earliest record of magmatism along this intra- 490 oceanic arc. Comparison of volcanic glass and whole rock REE-HFSE signatures of basalt, basaltic andesite 491 and andesite samples collected from four active intra-oceanic arc settings in the SW Pacific Ocean (Figure 492 8) with geochemical signatures from the Dunite Peak ophiolite supports our interpretation that the Dunite 493 Peak ophiolite derived from an intra-oceanic arc system (Figure 8b, e) in which arc (IAT/CA) and 494 subordinate back-arc (BABB) magmatic assemblages were generated co-spatially (Figure 8c, f). 495 6.1. Correlative SSZ ophiolites in the Northern Cordillera 496 The DPO is comparable to other mid-Permian ophiolites currently assigned to SMT that comprise sections 497 of oceanic crust and lithospheric mantle with mafic to intermediate SSZ geochemical signatures (e.g. 498 Mortensen, 1990; Nelson & Bradford, 1993; Dusel-Bacon et al., 2006; Petrie et al., 2015). These 499 comparable sections are as follows: 21 500 In Division II of the Sylvester allochthon, north British Columbia (Figure 1a), the Quartzrock Creek gabbro, 501 Cassiar block, Zus Mountain block and Blue Mountain block form ophiolitic sections of coarse grained 502 cumulate gabbro and leucogabbro, layered clinopyroxene and dunite and tectonized harzburgite (Gabrielse 503 et al., 1993; Nelson, 1993; Nelson & Bradford, 1993). U-Pb zircon geochronology from a layered 504 leucogabbro in the Zus Mountain block yielded an age of 268.6 +6.8/-3.4 Ma (Figure 1a) (Gabrielse et al., 505 1993). Geochemical signatures from Division II basalts (Nelson, 1993) are mostly N-MORB; however a 506 minority of these samples have arc signatures (IAT or CA) requiring an SSZ setting. 507 In south-central Yukon, the St Cyr klippe and Tower Peak assemblage (Figure 1d) comprise harzburgite, 508 gabbro, leucogabbro and basaltic to andesitic metavolcanic rocks currently assigned to SMT (Fallas et al., 509 1998; 1999; Petrie et al., 2015). Andesitic greenschists from the Tower Peak assemblage have IAT to CA 510 geochemical signatures (Fallas et al., 1999) indicative of a SSZ setting. In the Finlayson Lake district 511 (Figure 1c) most occurrences of SMT have N-MORB, E-MORB and OIB geochemical signatures. However, 512 localized occurrences of SMT greenstone, gabbro and leucogabbro have signatures ranging between 513 IAT/CA, BABB and N-MORB (Plint & Gordon, 1997; Murphy et al., 2006; Piercey et al., 2012) suggestive 514 of a SSZ setting. Locally, in the Finlayson Lake district (Figure 1c), gabbro and leucogabbro are commonly 515 observed intruding greenstone (Plint & Gordon, 1997), as is observed in the DPO. U-Pb zircon 516 geochronology of a leucogabbro, and of a plagiogranite within a serpentinized shear zone (Figure 1c-d), 517 respectively yielded crystallization ages of 273.4 ± 1.4 Ma (Murphy et al., 2006) and 274.3 ± 0.5 Ma 518 (Mortensen, 1992b). 519 In the Glenlyon region, central Yukon, the Tummel fault zone (Figure 1i) forms the boundary between YTT 520 and Laurentia, comprising fault slices of SMT serpentinized harzburgite mantle tectonite, gabbro, 521 greenstone and associated marine sedimentary rocks (Colpron et al., 2005, 2006b). Greenstones have 522 basaltic to andesitic compositions and geochemical signatures that mostly range between N-MORB and IAT 523 (Colpron et al., 2005, 2006b), suggesting an SSZ origin. To the northwest, the Ragged Lake klippe (Figure 524 1j) comprises serpentinite and gabbro, structurally overlying Laurentia with an IAT/CA geochemical 525 signature (Colpron et al., 2005, 2006b). U-Pb zircon geochronology from an IAT andesitic greenstone in the 22 526 Tummel fault zone and from the IAT/CA gabbro in the Ragged Lake klippe respectively yielded 527 crystallization ages of 267.8 ± 1.5 Ma and 260.3 ± 0.8 Ma (Figure 1, Colpron et al., 2005, 2006b). 528 In west-central Yukon and east-central Alaska, ophiolitic klippen of SMT and the Seventymile terrane 529 (Alaskan correlative of SMT) structurally overlie YTT, comprising pyroxene-phyric greenstone, cumulate 530 gabbro, harzburgite, dunite and subordinate clinopyroxenite (Figure 1o, Foster & Keith, 1974; Foster et al., 531 1994; Mortensen, 1990; Dusel-Bacon et al., 2006). Notable ophiolite occurrences include Clinton Creek 532 (Mortensen, 1990; Dusel-Bacon et al., 2006), Mount Sorensen, Salcha River and American Creek (Foster et 533 al., 1994; Dusel-Bacon et al., 2006). Geochemical signatures of greenstones in the Eagle quadrangle 534 (Alaska) and Clinton Creek/Dawson area (Yukon) (Figure 1o) range between N-MORB, BABB and CA 535 (Dusel-Bacon et al., 2006), suggestive of a SSZ setting. Metaharzburgite in the Eagle quadrangle has a 536 similar Nb depletion and Th enrichment to our Group 5 DPO samples that is suggestive of a SSZ setting 537 (Dusel-Bacon et al., 2013). Most notably, the Wolf Mountain klippe is dominated by greenstone with 538 basaltic to andesitic and IAT to CA geochemical signatures (Dusel-Bacon & Cooper, 1999; Dusel-Bacon et 539 al., 2006) that closely resemble IAT/CA greenstone from the DPO. 540 7. THE DUNITE PEAK INTRA-OCEANIC ARC 541 As outlined above, mid-Permian SSZ ophiolites with mean crystallization ages of ~275-260 Ma (including 542 the DPO), are distributed along the structural boundary between YTT and Laurentia over a distance of 543 ~1000 km between northern British Columbia, Yukon and eastern Alaska (Figure 1). Based on their similar 544 lithological, structural, geochemical and geochronological characteristics, we propose that collectively, they 545 represent tectonic slivers of a regionally extensive intra-oceanic extensional arc system that was active 546 within the Slide Mountain Ocean from 275 and 260 Ma. If our oldest zircon dated from IAT gabbro sample 547 16RAY-AP074A1 (283 ± 4 Ma) was inherited from the greenstone layer also of IAT composition (sample 548 16RAY-AP074A2, Table 1), into which it intrudes, then magmatism along this intra-oceanic arc occurred as 549 early as ~280 Ma. This is the first recognition of this arc system, which we define here as the Dunite Peak 550 intra-oceanic arc. 23 551 The DPO and correlative ophiolitic sections at the Tummel fault zone and Ragged Lake, St Cyr and Wolf 552 Mountain klippen (Figure 1) are dominated by basaltic to andesitic arc-magmatic rocks (IAT/CA), which we 553 interpret as constructional volcanic centers along the Dunite Peak intra-oceanic arc. Correlative ophiolitic 554 sections in the Sylvester allochthon (Figure 1a) and Finlayson Lake district (Figure 1c) have been 555 interpreted as remnants of a fossil transform fault system (Nelson & Bradford, 1993; Murphy et al., 2006; 556 Piercy et al., 2012) or ocean core complexes and syn-volcanic extensional faults (Ryan et al., 2015), and 557 yield geochemical signatures indicative of extensional magmatism in an SSZ setting (BABB, N-MORB ± E- 558 MORB). We interpret these sections as extensional magmatic zones along the Dunite Peak intra-oceanic arc. 559 The distribution of constructive volcanic centers (e.g. DPO, Wolf Mountain klippe) and extensional 560 magmatic zones (e.g. Zus Mountain block, Sylvester allochthon) suggests that the Dunite Peak intra-oceanic 561 arc may represent a disorganized extensional intra-oceanic arc (e.g. Weissel et al., 1981; Tamaki, 1985; 562 Wysoczanski et al., 2010). These systems are characterized by constructional volcanic centers surrounded 563 by localized zones of disorganized extensional faulting and associated magmatism. Extensional magmatism, 564 typically considered indicative of back-arc spreading (BABB) may occur along strike, adjacent to, and 565 contemporaneous with constructive arc-related magmatism (IAT/CA). Modern examples of disorganized 566 extensional intra-oceanic arcs include the Kermadec arc – Havre trough system (Wysoczanski et al., 2006; 567 2010; Smith et al., 2009), and the Lau basin (Keller et al., 2008; Sleeper & Martinez, 2014; Sleeper et al., 568 2016) in the SW Pacific Ocean (Figure 8) and the Calabrian arc in the Mediterranean Sea (Florio et al., 569 2011). In the geological record, comparable disorganized extensional intra-oceanic arc systems have been 570 inferred for the ancient settings of the Permian Nahlin ophiolite of the Cache Creek terrane in northern 571 British Columbia (McGoldrick et al., 2017) and the Middle Ordovician Annieopsquotch Accretionary Tract 572 of the Newfoundland Appalachians (Zagorevski et al., 2015). The development of the Dunite Peak intra- 573 oceanic arc as a disorganized extensional intra-oceanic arc provides a plausible explanation for the relatively 574 small volume of identifiable remnants of this arc and its elusiveness in the geological record. 24 575 8. INTEGRATION WITH EXISTING DATASETS: A NEW HYPOTHESIS FOR THE LATE 576 PALEOZOIC – EARLY MESOZOIC TECTONIC EVOLUTION OF THE NORTHERN 577 CORDILLERA 578 Slide Mountain Ocean could not have formed the lower plate of a westward-dipping subduction zone 579 beneath YTT (e.g. Mortensen, 1992a; Nelson et al., 2006; Beranek and Mortensen, 2011), at the same time 580 that it generated SSZ ophiolites with upper-plate intra-oceanic arc geochemical compositions (as required by 581 geochemical and geochronological constraints). This would require a more complicated double subduction 582 zone model that would still fail to adequately explain the variable duration and timing of YTT eclogitization, 583 SSZ ophiolite formation, Klondike cycle magmatic rocks and mid-Permian to Middle Triassic exhumation 584 of YTT. We propose an alternative model: structural juxtaposition of mid-Permian SSZ ophiolites and mid- 585 Permian YTT eclogites along the eastern margin of YTT (Figure 1) (e.g. Fallas et al., 1998; Petrie et al., 586 2015, 2016; Gilotti et al., 2017) by eastward subduction and collision of YTT (lower plate) beneath the 587 Dunite Peak intra-oceanic arc (upper plate) (Figure 9). In the following sections, we present a new 588 hypothesis for the Klondike orogeny (Figure 9) in the context of our analyses of the Dunite Peak ophiolite, 589 combined with existing datasets from across the Northern Cordillera. This includes datasets that were 590 previously inconsistent or hard to reconcile with previous models (see Section 2.1). Our proposed model is 591 supported by an extensive geochronology dataset (total of 110 published ages), displayed in Figure 10 (see 592 Supplementary Materials 05 for details), to constrain the timing of specific events. Note that in our model, 593 use of the term ‘Klondike orogen’ refers to the accreted mass (i.e. composite terrane) of YTT and the Dunite 594 Peak intra-oceanic arc. 595 8.1. Mid- to late Permian arc-continent collision: the Klondike orogeny 596 ~280-260 Ma: Dunite Peak intra-oceanic arc magmatism and subduction and collision of YTT 597 continental crust 598 Subduction related magmatism within the Dunite Peak intra-oceanic arc (Figure 1) is bracketed between 599 ~280 and 260 Ma (Figures 9a & 10, this study; Mortensen, 1992b; Gabrielse et al., 1993; de Keijzer et al., 600 2000; Colpron et al., 2005; 2006b; Murphy et al., 2006). Along the length of this intra-oceanic arc, 25 601 lithological and geochemical variations between ophiolites (see Sections 6.1 and 7) probably reflect 602 development of constructive volcanic centers and localized zones of upper-plate extension and associated 603 magmatism, representative of a disorganized extensional intra-oceanic arc such as the present day Kermadec 604 arc – Havre trough system (e.g. Wysoczanski et al., 2010). It is expected that most of the lower plate oceanic 605 lithosphere subducted into the mantle during that period, although some lower plate material may have 606 accreted to the Dunite Peak intra-oceanic arc (blue accretionary prism, Figure 9a). Early Permian and older 607 occurrences of SMT that contain chert, argillite and pillow basalt with mid-ocean ridge and/or within-plate 608 geochemical signatures such as parts of the Campbell Range basalts in the Finlayson Lake district (Figure 609 1c, Murphy et al., 2006; Piercy et al., 2012) may represent accreted remnants of this oceanic lower plate. 610 Subduction and collision of YTT continental crust beneath the Dunite Peak intra-oceanic arc began as early 611 as ~275 Ma (Figure 10, Gilotti et al., 2017) and continued, coeval with magmatism in the Dunite Peak intra- 612 oceanic arc to ~260 Ma (Figures 1e, h, i, g & 10, Creaser et al., 1997; Fallas et al., 1998; Godwin-Bell, 613 1998; Petrie et al., 2016; Gilotti et al., 2017). Many of the recorded eclogite occurrence from YTT are 614 described as eclogite pods hosted within amphibolite, greenschist or blueschist facies rocks (e.g. Erdmer et 615 al., 1998; Fallas et al., 1998; Petrie et al., 2015). Where blueschist is preserved, such as in the Faro, Ross 616 River and St Cyr regions (Figure 1d, g-h), muscovite cooling ages of ~263 to 235 Ma (Erdmer et al., 1998; 617 Fallas et al., 1998) suggest that these rocks may represent lower plate material from an accretionary complex 618 that formed above the subduction zone between YTT and the Dunite Peak intra-oceanic arc (e.g. blue shaded 619 accretionary complex, Figure 9a). However, most exposures of SMT and/or the Dunite Peak intra-oceanic 620 are structurally interleaved with slices of YTT and/or Laurentia (e.g. Mortensen, 1990; Fallas et al., 1998; 621 Murphy et al., 2006; Petrie et al., 2015) suggesting that the original accretionary complex and subduction 622 interface between YTT and the Dunite Peak intra-oceanic arc has been overprinted or structurally displaced 623 from its original configuration. The arrangement of ultramafic over mafic rocks in the DPO and the apparent 624 absence of a metamorphic discontinuity between the DPO and the underlying YTT also suggest that 625 emplacement structures relating to the original subduction zone and accretionary complex between YTT and 626 the Dunite Peak intra-oceanic arc have been structurally overprinted and/or displaced. This probably 26 627 occurred during subsequent regional metamorphism and deformation record in YTT during the latest 628 Triassic to Early Jurassic and Late Jurassic to Early Cretaceous (e.g. Staples et al., 2016). 629 Amphibolite facies metamorphism of YTT metasedimentary units (Snowcap assemblage) in Canada 630 recorded at ~260 Ma (Berman et al., 2007), 259.3 ± 3.4 Ma (Villeneuve et al., 2003) and 263.6 ± 3.4 Ma 631 (Staples, 2014) probably relates to crustal thickening during the final stages of subduction and collision of 632 YTT (Figures 1 & 10). Subduction and eclogitization of YTT and associated magmatism in the Dunite Peak 633 intra-oceanic arc ceased sometime between 265 and 260 Ma (Figure 10). Based on the ages listed here, we 634 define the earliest record of subduction and collision of YTT (~275 Ma) as the start of the Klondike orogeny 635 (Figure 10). 636 The coeval generation of SSZ ophiolites in the Dunite Peak intra-oceanic arc and eclogite facies 637 metamorphism of YTT suggests collision was diachronous. This may imply that the subducting margin of 638 YTT had an irregular shape with promontories and re-entrants and/or that collision of YTT with respect to 639 the intra-oceanic arc was oblique (e.g. Cawood & Suhr, 1992). Similarly, we have no constraint on the 640 length of the Dunite Peak intra-oceanic arc, and so it is possible that YTT collided only with a portion of the 641 arc and that adjacent to the subducting YTT, normal oceanic subduction and associated intra-oceanic arc 642 magmatism was maintained. Present-day collision and subduction of the Australian continent adjacent to 643 subduction of oceanic lithosphere beneath the Solomon – New Hebrides intra-oceanic arc is a modern 644 example of this process (e.g. Hall, 2002). Because of these potential and likely complexities and the lack of 645 further constraints, we make no attempt at estimating the width of oceanic crust subducted beneath the 646 Dunite Peak intra-oceanic arc. 647 265-252 Ma: Slab break-off, subduction polarity reversal, orogenic collapse and associated magmatism 648 The period between 265 and 260 Ma (Figure 10) marks; (1) the latest occurrence of eclogite facies 649 metamorphism of YTT, and magmatism in the Dunite Peak intra-oceanic arc (see above); (2) the earliest 650 record of exhumation and cooling of YTT blueschist and eclogite (Wanless, et al., 1978; Erdmer & 651 Armstrong, 1988; Erdmer et al., 1998; Fallas et al., 1998); (3) the emplacement of orogenic peridotite within 652 YTT (Canil et al., 2003; Johnston et al., 2007); and (4) the beginning of the Klondike magmatic cycle within 27 653 YTT (e.g. Colpron et al., 2006a; Beranek & Mortensen, 2011). We interpret the temporal overlap of these 654 events between 265 and 260 Ma (Figure 10) as a record of cessation of subduction of YTT continental 655 lithosphere accompanied or soon followed by orogenic collapse and associated magmatism (Figure 9b). 656 The Buffalo Pitts peridotite (Figure 1m) has been interpreted as megaboudin of orogenic mantle peridotite 657 emplaced into metasedimentary rocks of YTT during lithospheric extension of the terrane (Canil et al., 658 2003; Johnston et al., 2007). U-Pb zircon crystallization ages of 261.5 ± 2.3 Ma from an associated 659 leucogabbro, and 262.3 ± 0.4 Ma from an emplacement-related migmatite derived from the YTT 660 metasedimentary rock which hosts the peridotite (Figure 10), mark the timing of peridotite emplacement 661 during extension and exhumation of YTT (Canil et al., 2003; Johnston et al., 2007). This occurred 662 contemporaneously with exhumation of blueschsits and eclogites within the east margin of YTT (Figure 10) 663 (e.g. Erdmer et al., 1998; Fallas et al., 1998). 664 The Klondike magmatic cycle is recorded by granitic plutons (Figure 1) and associated volcanics 665 (collectively referred to as the Klondike assemblage) within YTT between ~265 and 252 Ma (Figure 10) in 666 Yukon, northern British Columbia and eastern Alaska (e.g. Harms, 1985; Gabrielse et al., 1993; Nelson & 667 Friedman, 2004; Liverton et al., 2005; Murphy et al., 2006; Colpron et al., 2006a; Dusel-Bacon et al., 2006; 668 Nelson et al., 2006 and Beranek & Mortensen, 2011). Beranek and Mortensen (2011) argued for syn- 669 tectonic subduction-related magmatism, deformation and metamorphism between 260 and 254 Ma and post- 670 tectonic magmatism between 254 and 252 Ma. However, the reported occurrences of deformed and 671 undeformed Klondike magmatic assemblages collated in Figure 10 do not show a distinct phase of post- 672 tectonic magmatism. Instead, metamorphism and deformation of YTT metasedimentary units is recorded 673 between ~260 and ~252 Ma (Figures 1 & 10, Fallas, 1998; Villeneuve et al., 2003; Berman et al., 2007; 674 Staples, 2014), contemporaneous with Klondike cycle magmatism and exhumation of YTT (Figures 9c & 675 10, Wanless et al., 1978; Htoon, 1981; Erdmer & Armstrong, 1988; Hunt & Roddick, 1988; 1993; Oliver, 676 1996; Erdmer et al., 1998; Fallas et al., 1998; Breitsprecher & Mortensen 2004; Joyce et al., 2015). Within 677 our proposed tectonic framework, the Klondike magmatic cycle is interpreted as either, (1) exhumation- 678 driven and/or asthenospheric flow-driven crustal melting in response to orogenic collapse (i.e. extension) 28 679 after slab break-off (Figure 9b-c) (e.g. Dewey, 1988; Huw Davies & von Blackenburg, 1995; Atherton & 680 Ghani, 2002; Brown et al., 2011; Li et al., 2016), (2) arc-derived magmatism generated in response to 681 initiation of westward subduction beneath the Klondike orogen following slab break-off and subduction 682 polarity reversal (e.g. Stern, 2004); or; (3) a combination of both (1) and (2). All three interpretations fit with 683 geochemical and geochronological constraints from plutonic components of the Klondike magmatic cycle 684 (e.g. Ruks et al., 2006; Beranek & Mortensen, 2011). We also note that magmatism in response to eastward 685 subduction of Panthalassa Ocean beneath YTT cannot be ruled out. 686 We propose that the events described above occurred in response to buoyancy-driven slab break-off from 687 the subducted margin of YTT continental crust (e.g. Cloos, 1993; Afonso & Zlotnik, 2011) followed by 688 initiation of westward subduction and roll-back of the ocean basin that lay east of the Dunite Peak intra- 689 oceanic arc (e.g. Dewey, 1988; Huw Davies & von Blackenburg, 1995; Atherton & Ghani, 2002; Duretz et 690 al., 2010; Brown et al., 2011). Obduction of SSZ ophiolites of the Dunite Peak intra-oceanic arc on to YTT 691 probably occurred during or soon after these events, whilst they were young (<10 M.yr.) and hot (e.g. 692 Dewey, 2003). The period between 265 and 260 Ma marks the starting point for these events, triggered by 693 the cessation of subduction of YTT (Figure 10), but we note that the relative timing and duration of slab 694 break-off and subduction polarity reversal is poorly constrained and may have taken longer than five million 695 years (e.g. Brown et al., 2011). The latest occurrence of metamorphism of YTT coeval with exhumation and 696 Klondike magmatism marks the final stage of the Klondike orogeny. Thus, we bracket the duration of 697 Klondike orogeny between ~275 and ~252 Ma (Figure 10). This outlined sequence of events is comparable 698 to the evolution of other ancient arc-continent collisions, such as the Grampian, Taconic and Kamchatka arc- 699 continent orogens (e.g. Dewey, 2005; Boutelier & Chemenda, 2011; Brown et al., 2011 and references 700 therein). We also note that the duration of the Klondike orogeny (~20-25 M.yr) is comparable to the 701 duration of other arc-continent collisions (~20-50 M.yr.), as opposed to the longer duration of continent- 702 continent collisions (≥50 M.yr.), which would be expected for a collision between YTT and Laurentia (e.g. 703 Friedrich et al., 1999; Dewey, 2005; van Staal et al., 2007; Chew et al., 2010; Brown et al., 2011). 704 9. REMAINING PROBLEMS AND FUTURE WORK 29 705 9.1. Accretion of the Klondike orogen to Laurentia 706 Our study argues that YTT did not collide with Laurentia during the Permian. Previous studies proposed that 707 deposition of the Jones Lake Formation on YTT, Laurentia and SMT marked the final stages of collision, 708 accretion and related subsidence between YTT and Laurentia, following their collision in the Permian 709 (Colpron et al., 2006; 2007; Beranek & Mortensen, 2011). However, close inspection of variations in detrital 710 zircon populations and conodont fauna within the Jones Lake Formation (see Section 2.1 for a review of this 711 data) indicate that these data are at best, consistent with a post-Middle Triassic collision but do not validate, 712 and in some cases are inconsistent with a model of deposition in a regional Triassic basin overlapping YTT, 713 SMT and Laurentia following Permian collision (e.g. Colrpon et al., 2007; Beranek et al., 2010; Beranek & 714 Mortensen, 2011). We also note that the widespread Late Triassic to Middle Jurassic plutons and Permian 715 and Jurassic high-grade metamorphism recorded within YTT are absent from Laurentia in Yukon (e.g. 716 Berman et al., 2007; Staples et al., 2014; Colpron et al., 2016b). 717 The present day lithospheric structure of the Northern Cordillera suggests that YTT forms an overriding 718 thrust sheet emplaced on top of Laurentia (e.g. Gordey, 2002; Cook et al., 2004; Calvert et al., 2017). To 719 satisfy this constraint, we suggest that after the Klondike orogeny, oceanic lithosphere that lay between 720 Laurentia and the Dunite Peak intra-oceanic arc was subducted westward beneath the Klondike orogen 721 (Figure 9d). We propose that the Jones Lake Formation represents accreted remnant ocean plate stratigraphy 722 (magenta accretionary complex, Figure 9c-d) and/or forearc/passive margin basin sediments (e.g. Gordey, 723 2013) deposited in the basin between the Klondike orogen and Laurentia (e.g. Ingersole, 1988; Ingersole et 724 al., 2003), which closed sometime after the Middle Triassic (e.g. Hansen et al., 1991; Stevens et al., 1996; 725 Plint & Gordon, 1997; Hansen and Dusel-Bacon 1998; Gordey, 2002, 2013). This is consistent with the 726 occurrence of unimodal/bimodal Permian-Triassic ± Devonian – Early Mississippian detrital zircon 727 populations (Beranek & Mortensen, 2011) and “exotic” Tethyan conodont species (Orchard, 2006) in 728 Middle to Late Triassic strata that overlie parts of SMT Yukon and northern British Columbia. Additionally, 729 paleomagnetic analysis of red chert units of SMT (Richards et al., 1993), suggested that parts of SMT 730 remained allochthonous with respect to Laurentia until the Early Jurassic (see Section 2.1.). We propose that 30 731 these red chert units derived from the basin between the Dunite Peak intra-oceanic arc and the Laurentian 732 margin. 733 The exact timing of collision between Laurentia and the Klondike orogen is unclear. Early Jurassic 734 amphibolite facies metamorphism followed by Early to Middle Jurassic exhumation recorded within YTT in 735 Canada and eastern Alaska may record collision between Laurentia and the Klondike orogen and its 736 subsequent collapse (e.g. Dusel-Bacon et al., 2002; Berman et al., 2007; Joyce et al., 2015; Morneau, 2017; 737 Morneau et al., 2017). However, Jurassic metamorphism and exhumation within YTT have also been 738 interpreted as a record of collision and accretion between YTT and Stikinia-Quesnellia (e.g. Colpron et al., 739 2015; Clark, 2017). Collision between YTT-Stikinia and the Insular terranes (Alexander-Wrangellia, Figure 740 1) is hypothesized during the Early to Middle Jurassic (Monger, 2014) or the Late Jurassic to Cretaceous 741 (Sigloch & Mihalynuk, 2017). The timing of collision and accretion and the polarity of subduction zones 742 responsible for accretion between the Intermontane terranes (YTT-Stikinia-Quesnellia), the Insular terranes 743 and Laurentia is still debated (e.g. Johnston, 2008; Hildebrand, 2009; Monger, 2014; Sigloch & Mihalynuk, 744 2017). 745 The distribution of Early Cretaceous plutonic suites emplaced within both YTT and Laurentia (e.g. Colpron 746 et al., 2016b) provides a latest constraint on this collision and accretion, suggesting that YTT had accreted to 747 Laurentia by ~120 Ma. Emplacement of Early Cretaceous plutons into YTT and Laurentia (e.g. Colpron et 748 al., 2016b) was contemporaneous with amphibolite facies metamorphism of the underlying Laurentian 749 margin (e.g. Gibson et al., 2008; Moynihan, 2013; Staples et al., 2016; Ryan et al., 2017). This has led some 750 authors to propose that the allochthonous terranes first accreted with each other before accreting to Laurentia 751 sometime during the Cretaceous (e.g. Johnston, 2008; Hildebrand, 2009). 752 Integration of our study with existing datasets indicates that that collision and accretion between the 753 Klondike orogen and Laurentia occurred after the Middle Triassic (e.g. Hansen et al., 1991; Richards et al., 754 1993; Stevens et al., 1996; Plint & Gordon, 1997; Hansen and Dusel-Bacon 1998; Gordey, 2002, 2013), and 755 perhaps as late as post-Jurassic (e.g. Johnston, 2008; Hildebrand, 2009). However, the variability of timing 756 estimates and subduction zone polarity models hypothesized by previous studies (described above), 31 757 demonstrates that the timing, order and sequence of collisions between the allochthonous terranes and 758 Laurentia that occurred after the Klondike orogeny requires further investigation. The disruption and/or 759 displacement of original emplacement structures associated with the Permian subduction zone beneath the 760 Dunite Peak intra-oceanic arc occurred during one or more of the subsequent collisional events listed above. 761 9.2. Re-assessment of the Slide Mountain terrane 762 Within our model we make a clear argument that the DPO and other correlative SSZ ophiolites derive from 763 the upper plate of the Dunite Peak intra-oceanic arc. Less certain, is the identification of ocean plate 764 stratigraphy (OPS) derived from the lower plate that subducted beneath this intra-oceanic arc or from the 765 basin between the Dunite Peak intra-oceanic arc and Laurentia. We suggest that early Permian and older 766 occurrences of SMT may represent accreted remnants of OPS from the subducting oceanic plate that was 767 attached to the east continental margin of YTT. These comprise supracrustal sections of basinal sedimentary 768 rocks and associated mafic volcanics with mid-ocean ridge and/or within-plate geochemical signatures, such 769 as parts of the Campbell Range basalts in the Finlayson Lake district (Figure 1c, Murphy et al., 2006; Piercy 770 et al., 2012). As our model predicts that the basin between the Dunite Peak intra-oceanic arc and Laurentia 771 closed following the Klondike orogeny, OPS attached to the passive margin of Laurentia should be capped 772 by Triassic and younger sediments that are possibly represented by the Jones Lake Formation (see Section 773 9.1.). 774 Our model also predicts structural separation between remnants of mid-Permian SSZ ophiolites, 775 Mississippian to Permian OPS attached to YTT, and Mississippian to Triassic or younger OPS from the 776 basin between the Dunite Peak intra-oceanic arc and Laurentia. In our study area, the single outcrop of 777 pillow basalt from the marine metasedimentary succession with an E-MORB geochemical signature (Group 778 4) may be one such example of this lower-plate accreted OPS. Elsewhere, a possible example of structurally 779 distinct elements of SMT is presented in the Finlayson Lake district (Figure 1c) where SMT containing N- 780 MORB, and BABB geochemical signatures appears to be structurally separated from SMT containing N- 781 MORB, E-MORB and OIB (Murphy et al., 2006, their Figure 16). 32 782 As our study is the first to predict these structurally distinct elements of OPS within SMT, it is difficult to 783 identify these elements of SMT from previous work without reliable age and geochemical constraints. 784 Furthermore, it is likely that subsequent deformation during the Jurassic and Cretaceous (e.g. Staples et al., 785 2016) overprinted or removed original emplacement-related structural relationships. Further work should be 786 conducted to robustly identify and distinguish between our three predicted elements of SMT. Based on their 787 distinct age, lithology and structural separation from other parts of SMT, we suggest that all upper-plate 788 assemblages associated with the Dunite Peak intra-oceanic arc be recognized as a distinct terrane (e.g. 789 Coney et al., 1981; Ryan et al., 2015). 790 9.3. Spatial and temporal extent of the Dunite Peak intra-oceanic arc 791 We encourage further work to investigate its spatial and temporal extent of the Dunite Peak intra-oceanic 792 arc. Our model requires that following the Klondike orogeny, westward subduction of the ocean basin 793 between the Klondike orogen and Laurentia occurred so that they could collide sometime after the Middle 794 Triassic (see Section 9.1.). A potential caveat of this hypothesis is the relatively low abundance of Triassic 795 arc-magmatic rocks within YTT (e.g., the Stikine suite, Colpron et al., 2016b). Although we can only 796 speculate, a potential solution is that the magmatic response to westward-subduction of the ocean basin 797 between the Klondike orogen and Laurentia may be recorded by the Triassic Quesnel intra-oceanic arc (see 798 review of Nelson et al., 2013). This would imply that the Quesnel intra-oceanic arc formed on or close to the 799 eastern margin of the Klondike orogen, as a successor arc to the Dunite Peak intra-oceanic arc. Other 800 potentially misidentified correlatives of the Dunite Peak intra-oceanic arc may include greenstone from the 801 Klinkit intra-oceanic arc, which is currently considered part of YTT (e.g. Colpron et al., 2006; Nelson et al., 802 2013), and yet yielded a U-Pb igneous age of 281 ± 2 Ma and juvenile (ƐNd = +6.7 to +7.4) IAT geochemical 803 signatures (Roots et al., 2002; Simard et al., 2002). Similarly, recent study of the Permian Nahlin ophiolite, 804 currently assigned to the Cache Creek terrane, also calls for the presence of a previously unrecognized 805 Permian-Triassic intra-oceanic arc that is allochthonous to the intermontane terranes (McGoldrick et al., 806 2017). Based on our recognition of the Dunite Peak intra-oceanic arc and its similarity with other intra- 807 oceanic arcs in the Northern Cordillera, plus our prediction of two structurally confined OPSs within SMT 33 808 that are distinct from each other and the Dunite Peak intra-oceanic arc, we suggest that the current 809 tectonostratigraphic framework of terranes and assemblages in the Northern Cordillera be re-evaluated. 810 10. CONCLUSIONS 811 This study demonstrates the importance of identifying and distinguishing between upper-plate and lower- 812 plate components when attempting to understand the tectonic evolution of an accretionary orogen. We have 813 applied this concept to our study of the Dunite Peak ophiolite in south-central Yukon and other previously 814 studied mid-Permian ophiolites in Alaska, Yukon and British Columbia. Our findings indicate that a new 815 explanation is required for the late Paleozoic to early Mesozoic tectonic interaction between of the Yukon- 816 Tanana terrane, Slide Mountain terrane and Laurentian margin of the Northern Cordillera. 817 The Dunite Peak ophiolite (DPO) forms klippen of mafic-ultramafic strata structurally emplaced over 818 metasedimentary rocks of Yukon-Tanana terrane (YTT). Field structural, geochemical and geochronological 819 analyses conducted on the DPO and underlying metasedimentary strata yield the following conclusions: 820 1) Whole-rock geochemical and Sm-Nd isotopic analyses of mafic-ultramafic assemblages from the 821 DPO identify 5 distinct geochemical groups. Groups 1-3 correspond to arc (Groups 1-2 = IAT to CA) and 822 back-arc (Group 3 = BABB) magmatic components of a SSZ ophiolite formed in the upper plate of an intra- 823 oceanic arc (ƐNd = +7.2 to +9.0). Group 5 corresponds to highly depleted mantle/lower crustal ultramafic 824 cumulates, formed in a SSZ setting. The geochemical signatures from these groups are comparable to those 825 derived from modern day intra-oceanic arcs, such as the Izu-Bonin, Mariana and Kermadec arcs and Lau 826 basin, in the SW Pacific Ocean. Group 4 corresponds to E-MORB pillow basalt with no record of interaction 827 with a subduction zone and may therefore by tectonically unrelated to the DPO. 828 829 2) U-Pb zircon geochronology of an IAT gabbro from the DPO yielded a mean igneous crystallization age of 265 ± 4 Ma. 830 3) Geochemical and geochronological constraints correlate the DPO with other mid-Permian (275-260 831 Ma) ophiolites previously assigned to Slide Mountain terrane (SMT). These correlatives include Quartzrock 832 Creek gabbro, Cassiar block, Zus Mountain block and Blue Mountain block (northern British Columbia, 833 Sylvester allochthon), St Cyr klippe and Tower Peak assemblage and the Finlayson Lake greenstones and 34 834 gabbros (south-central Yukon), Tummel fault zone and Ragged Lake klippe (central Yukon), Clinton Creek 835 ophiolite (west Yukon) and the Wolf Mountain klippe (east Alaska) (Figure 1). Together, these ophiolites 836 represent the dismembered upper-plate remnants of a regionally extensive intra-oceanic arc active between 837 YTT and Laurentia during the mid-Permian (~280-260 Ma). We name this arc the Dunite Peak intra- 838 oceanic arc. 839 4) We propose that the Klondike orogeny records deformation, metamorphism and magmatism 840 associated with mid-Permian eastward subduction and collision of YTT beneath the Dunite Peak intra- 841 oceanic arc. Subduction of YTT and associated intra-oceanic arc magmatism terminated at ~265-260 Ma. 842 This was accompanied/followed by slab break-off, orogenic collapse and associated magmatism, and the 843 initiation of westward subduction (present coordinates) beneath the east margin of the Klondike orogen 844 (composite of YTT and the Dunite Peak intra-oceanic arc) between ~265 and ~252 Ma. During this time, 845 the Klondike magmatic cycle occurred in response to either: (1) exhumation-driven crustal melting; (2) 846 newly initiated westward subduction; or (3) a combination (1) and (2). 847 5) Accretion of the Klondike orogen with Laurentia occurred after the Permian, and probably after the 848 Middle Triassic. This is supported by reassessment of detrital zircon populations and conodont fauna with 849 Triassic sedimentary rocks on YTT, SMT and Laurentia, and the distribution of Triassic to Early Jurassic 850 metamorphism, magmatism and exhumation within YTT that is not record by Laurentia. 851 6) Mid-Permian SSZ ophiolites of the Dunite Peak intra-oceanic arc are distinct from older supracrustal 852 sections of SMT that formed in an intra-plate/mid-ocean spreading ridge setting. The definition of SMT 853 should be modified to formally recognize the distinction between mid-Permian SSZ ophiolites from older 854 SMT sections. Other parts of SMT may be subdivided into accreted ocean plate stratigraphy derived from 855 (1) the oceanic lower plate attached to YTT that subducted beneath the Dunite Peak intra-oceanic arc and (2) 856 the ocean basin which formed between the Dunite peak intra-oceanic arc and Laurentia. The latter should 857 contain younger strata (post-Permian) than the former. This younger lower plate material from the basin 858 between the Dunite Peak intra-oceanic arc and Laurentia may include the Triassic Jones Lake Formation. 859 The Dunite Peak intra-oceanic arc should be recognized as a distinct terrane and that the current 35 860 tectonostratigraphic framework of terranes and assemblages in the Northern Cordillera should be re- 861 evaluated. 862 Acknowledgements 863 This research was has received funding from the Geological Survey of Canada, GEM-II Cordillera project and the 864 European Research Council (ERC) under the European Union’s Horizon 2020 research and innovation programme 865 (grant agreement 639003 “DEEP TIME”). We thank Brad Singer (science editor), John Waldron (ass. editor), John 866 Wakabayashi, Cynthia Dusel-Bacon and Steve Johnston for constructive reviews. JoAnne Nelson (BCGS) is thanked 867 for helpful discussion during initial write-up. Dejan Milidragovic (BCGS) is thanked for guidance during initial 868 interpretations of geochemical data. We thank Shuangquan Zhang (Carlton University) for the Sm-Nd isotopic 869 analyses presented in this study. We thank Capital Helicopters Inc. and the Yukon Geological Survey for logistical 870 support during fieldwork. 871 872 36 873 References 874 Afonso, J.C., Zlotnik, S. 2011. The subductability of continentallithosphere: the before and after story. In: Brown, D. 875 & Ryan, P.D. (eds.) Arc-Continent Collision. Berlin, Heidelberg: Springer Berlin Heidelberg, p. 53-86. 876 Arndt, N.T., Goldstein, S.L. (1989). An open boundary between lower continental crust and mantle: its role in crust 877 878 879 formation and crustal recycling. Tectonophysics, 161, 201-212. Atherton, M.P., Ghani, A.A. (2002). Slab breakoff: a model for Caledonian, Late Granite syn-collisional magmatism in the orthotectonic (metamorphic) zone of Scotland and Donegal, Ireland. 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(eds.) Arc-Continent Collision. 320 Berlin, Heidelberg: Springer Berlin Heidelberg, p. 341-371. 321 Zagorevski, A., McNicoll, V.J., Rogers, N., van Hees, G.H. (2015). Middle Ordovician disorganized arc rifting in the 322 peri-Laurentian Newfoundland Appalachians: implications for evolution of intra-oceanic arc systems. Journal 323 of the Geological Society. 173, 76-93. 324 52 325 Figure 1. Terrane map of the Northern Cordillera. Modified from Colpron et al. (2007). Study area (Dunite Peak 326 ophiolite, DPO) outlined by black box (Figure 2). Inset map (bottom right corner) shows location of Figure 1 with 327 respect to North America. Geochronology localities discussed in this article and synthesized into our new hypothesis 328 are labelled with letters (a) to (o). Age ranges are given for localities with multiple geochronology analyses and 329 correspond to the oldest age plus error and the youngest minus error. Klondike magmatic assemblage (Ass) and 330 correlative units drawn from Cui et al. (2015) and Colpron et al. (2016a). (a-o) Superscript numbers next to ages 331 correspond to geochronology data sources (discussed in text): (1) Gabrielse et al., 1993, (2) Erdmer et al., 1998, (3) 332 Murphy et al. 2006, (4) Mortensen, 1992b, (5) Fallas et al., 1998, (6) Petrie et al., 2015, (7) Petrie et al., 2016, (8) 333 Gilotti et al., 2017, (9) de Keijzer et al., 2000, (10) Creaser et al., 1997, (11) Erdmer & Armstrong, 1988, (12) 334 Godwin-Bell, 1998, (13) Philippot et al. 2001, (14) Colpron et al., 2005, (15) Oliver, 1996, (16) Joyce et al., 2015, 335 (17) Johnston et al., 2007, (18) Colpron et al., 2006b, (19) Breitsprecher & Mortensen 2004, (20) Staples, 2014, (21) 336 Berman et al., 2007, (22) Htoon, 1981. 337 Figure 2. Geological map of the Dunite Peak ophiolite (DPO) with locations of geochronology and geochemistry 338 samples and Figure 3 photo locations. Geochemistry symbols correspond to geochemical groups in Figures 6-7. 339 Figure 3. (a-b) Photographs of the eastern klippe of the Dunite Peak ophiolite (DPO) with unit boundaries overlain. 340 See Figure 2 and text for map unit descriptions. See Figure 2 for photo locations. 341 Figure 4. Field images of the Dunite Peak ophiolite (DPO) crustal section. (a) Greenstone volcaniclastic rocks (Group 342 1 geochemistry) with subordinate chert horizons, lower DPO crustal section. (b) Sheared pillow basalt and greenstone 343 (Group 1a geochemistry) with subordinate sedimentary layers, lower DPO crustal section. (c) Typical lithological 344 representation of volcanic/volcaniclastic greenstone. (d) Typical lithological representation of gabbro (Group 2 345 geochemistry, upper DPO crustal section). (e) Gabbro (Group 2 geochemistry) intruding fragmental volcaniclastic 346 greenstone, upper DPO crustal section. (f) Gabbro (Group 2 geochemistry) intruding greenstone (Group 1 347 geochemistry), upper DPO crustal section. 348 Figure 5. U-Pb geochronology, sample16RAY-AP074A1 (a) Representative cathodoluminescence (CL) images of 349 zircon grains analyzed and (b) Tera–Wasserburg and (c) chondrite-normalized rare earth element plots of sensitive 350 high-resolution ion microprobe reverse-geometry (SHRIMP-RG) U/Pb and trace element data from sample 16RAY- 351 AP074A1. U/Pb data plotted as 1σ error ellipses uncorrected for common Pb. Black ellipses are used in calculating 352 concordia ages. Weighted mean age uncertainty is reported at the 95% confidence level. 53 353 Figure 6. Rare earth element – high field strength element (REE-HFSE) spider diagrams. (a-c) Groups 1-4 normalized 354 against N-MORB (Sun & McDonough, 1989). N-MORB V and Sc concentrations based on Klein (2004). (d) Group 5 355 normalized against primitive mantle (Sun & McDonough, 1989). Primitive mantle V and Sc concentrations based on 356 McDonough and Frey (1989). Standard REE-HFSE profiles for depleted-MORB mantle (DMM) are displayed for 357 comparison (Salters & Stracke, 2004; Workman & Hart 2005). 1358 Figure 7. Geochemical-tectonomorphic discrimination diagrams. (a) Nb/Y (total alkalinity proxy) vs. Zr/Ti (SiO2 359 proxy), based on Pearce (1996). (b) Y-La-Nb ternary discrimination diagram based on Cabanis & Lecolle (1989). (c) 360 Suprasubduction/spreading ridge/within-plate magmatism discrimination diagram (Nb/Yb vs. Th/Yb) based on Pearce 1361 (2008, 2014). Black squares show typical values for the magma types indicated by adjacent labels. Vertical arrows 1362 represent subduction zone enrichment. (d) Geochemical classification of Group 5 ultramafic samples, based on CIPW 1363 normative major element concentrations. 364 Figure 8. Geochemical signatures of basalt, basaltic andesite and andesite samples collected from active intra- 1365 oceanic arcs in the SW Pacific Ocean compared with data collected from the Dunite Peak ophiolite. (a-f) Volcanic 1366 glass and whole rock rare earth element – high field strength element (REE-HFSE) concentrations from the Izu-Bonin 1367 arc (red shaded area – e.g. Tollstrup et al., 2010; Ishizuka et al., 2014), Mariana arc (orange shaded area – e.g. 1368 Pearce et al., 2005; Tamura et al., 2014), Kermadec arc (blue shaded area – e.g. Wysoczanski et al., 2006; 1369 Smith et al., 2009) and Lau Basin (green shaded area – e.g. Keller et al., 2008; Bézos et al., 2009), drawn from 1370 datasets of n samples. Data are split into arc samples (a-c) and back-arc/trough/ridge samples (d-e). Note that only 1371 data collected from back-arc/trough/ridge settings are available from the Lau Basin (green shaded area). Geochemical 1372 data from the Dunite Peak ophiolite are delineated with black to grey lined, unshaded areas (g-i) and overlain on SW 1373 Pacific data sets for comparison (a-f). Spider diagrams are normalized against N-MORB (Sun & McDonough, 1989). 1374 N-MORB V and Sc concentrations based on Klein (2004). Explanation and abbreviations for geochemical- 1375 tectonomorphic discrimination diagrams (b-c), (e-f), and (h-i) are same as in Figure 7b-c. Data complied and extracted 1376 from GEOROC database (Sarbas & Nohl, 2008). Full dataset and sources (total of 57 sources) are presented in 1377 Supplementary Materials 04. 1378 Figure 9. Cartoon summary of our new model for the Paleozoic–Mesozoic tectonic evolution of the Northern 1379 Cordillera, see text for discussion. Yukon-Tanana terrane crust – dark green; YTT eclogite – purple; Laurentian crust 1380 – blue: Laurentian margin accretionary complex – magenta; Dunite Peak intra-oceanic arc crust – light grey; Dunite 54 1381 Peak arc accretionary complex – dark blue; oceanic crust – dark grey; lithospheric mantle – light green; volcanic 1382 centres – orange triangles. Not drawn to scale. 1383 Figure 10. Synthesis of isotope geochronology data from Yukon-Tanana terrane (YTT) and the Dunite Peak intra- 1384 oceanic arc. Sample numbers (x-axis) correspond to published geochronology data entries in Supplementary Materials 1385 05. Time scale (y-axis) based on Geological Society of America timescale. Data points are divided into specific 1386 tectonic processes, labeled in top header, with a color bar to highlight period of activity (e.g. duration of Dunite Peak 1387 intra-oceanic arc magmatism indicated by orange bar). Tectonic interpretations of data summarized on right panel. 1388 Hatched zone at 265-260 Ma corresponds to cessation of subduction and collision of YTT with the Dunite Peak intra- 1389 oceanic arc and the initiation of slab break-off, subduction polarity reversal and associated magmatism. Hatched zone 1390 at post-220 Ma marks possible timing of collision and accretion between YTT and Laurentia. See text for discussion. Figure 1 140°W 21 0 (l) 264 ± 3 Ma 20 255 ± 3 Ma 20 239 ± 3 Ma 20 AL AS KA (n) 239 ± 7 Ma KS NORTHW EST TERRITORIES (k) 217 ± 1 Ma19 (j) 260 ± 1 Ma18 (h) 270-245 Ma 13 SM WR CG NAb (g) 255 ± 8 Ma12 AX Cordilleran Terranes Outboard YT (b) 207 ± 1 Ma2 Geochronology Localities (a) Zus Mountain, Blue Dome (Sylvester allocthon)* (b) Klatsa (c) Finlayson Lake* (d) St Cyr / Quiet Lake* (e) Dunite Peak* (f) Last Peak (g) Ross River (h) Faro (i) Tummel fault zone* (j) Ragged Lake klippe* (k) Little Kalzas Lake (l) McQuesten (m) Buffalo Pitts (n) Stewart River (o) Clinton Creek / Eagle Quadrangle* * Denotes mid-Permian supra- NAc (a) 269 ± 7 Ma1 CC CG CA Ancestral N America CA - Cassiar NAb - NA basinal NAp - NA platform NAc - NA craton & cover 62°N subduction zone ophiolite locality Y U KO N TE R R ITOR B RIT Y IS H C O LU M B IA WR - Wrangellia AX - Alexander KS - Kluane,Windy,Coast CC - Cache Creek ST - Stikinia QN - Quesnellia YT - Yukon-Tanana SM - Slide Mountain Klondike Ass (YT) (c) 273 ± 1 Ma3 CA YT 200 CC Insular Intermontane (c) 274 ± 0.5 Ma4 DPO (Fig. 2) (d) 275-261 Ma 5,6,7,8 (d) 265-223 Ma 5 YA - Yakutat CG - Chugach 58°N (g) 275-237 Ma2,11 (f) 269 ± 2 Ma 10 (f) 267 ± 7 Ma 2 KS (e) 267 ± 10 Ma 9 WR YA (h) 263-257 Ma 2 (i) 215 ± 4 Ma 16 YT ST WR NAp (i) 243-208 Ma 15 (i) 268 ± 2 Ma14 (m) ~262 Ma 17 100 km NAp (n) ~260 Ma 21 62°N 124°W (o) 245 ± 8 Ma 22 YT NAb 132°W NAp CG AX YT NAc ST Slide Mountain Terrane - U/Pb age Buffalo Pitts peridotite - U/Pb age Figure 1 Permian Yukon-Tanana terrane eclogite facies metamorphism - U/Pb & Sm/Nd age Permian-Triassic Yukon-Tanana terrane amphibolite facies metamorphism - U/Pb age KS& Ar/Ar age Exhumation of Yukon-Tanana terrane eclogite/blueschist - K/Ar QN Exhumation of Yukon-Tanana terrane (other than eclogite/blueschist)- K/Ar & Ar/Ar age (a) Letters corresponds to geochronology locality 1 Superscript numbers correspond to data source (see figure caption) 132°W 124°W 58°N 134.04° W 134° W MMS3 133.96° W 1700 Locality U-Pb zircon age 61.66° N Field station 140 0 DPc 1800 00 Group 5 (Pyrxnt.) 2 0 1600 3 4 Kilometers 00 BM 15 00 00 1800 14 1500 1300 1700 0 14 1500 1 N 170 00 1700 1700 00 13 00 16 MMS4 00 Group 5 (Lherz.) 18 00 MMS3 Group 5 (Dun./Harz.) Group 2 (IAT, plutonic) 0 0.5 16 13 Group 4 (E-MORB) Group 1b (IAT-CA) Group 3 (BABB) 130 00 MMS3 Group 1a (IAT/IAT-CA) 180 1500 0 14 DPu 61.64° N Superscript number refers to data source: 1 - This study 2 - De Keijzer et al. 2000 15 MMS3 133.88° W Geochemical Samples (Figs. 6 & 7) 265 ± 4 Ma1 Figure photo, with viewing direction 133.92° W 1800 61.66° N 134.08° W BM 265 ± 4 Ma1 MMS2 61.64° N 134.12° W 18 00 134.16° W Figure 2 1700 BM 140 0 FAULT SYST MMS3 00 1300 14 00 MMS2 DPc MMS1 MMS3 DPc MMS3 DPu1 0 18 00 1400 18 170 00 0 0 DPc MMS3 1900 0 1000 DPu2 0 19 Fig. 3b 0 DPu2 19 00 1800 1900 0 180 MMS3 BM 2000 DPu2 MMS3 DPc DUNITE PEAK DPu2 DPu DPc 267 ± 10 Ma2 DPc MMS2 MMS1 00 16 DPc MMS2 0 MMS1 12 1300 00 BM 0 90 Marine metasedimentary succession DPu1 - Serpentinite, sheared and imbricated 00 14 0 MMS2 - White quartzite, interbedded quartzite & marble, subordinate carbonaceous calc-silicate layers MMS1 MMS1 - Micaceous/carbonaceous quartzite and carbonaceous shale 1100 1200 MMS2 10 Axial trace, synform 0 DPc 1400 DPc - Volcanic/volcaniclastic greenstone, subordinate shale, chert (lower portion), hypabyssal greenstone, gabbro, leucogabbro, sheared and imbricated 134.08° W 134.04° W Basal marble BM - Marble, subordinate siliciclastics, pelite, mafic CB 13 intrusions, garnet-epidote-amphibolite facies 00 134° W Normal fault, observed Normal fault, constrained Normal fault, inferred 133.96° W Geographic Symbols 1400 1100 Dunite Peak 1500 0 Crustal Ultramafic section section DPu1 MMS3 - Carbonaceous shale/semipelite, subordinate chert, greenstone 0 10 61.56° N DPu2 DPu2 - Dunite, harzburgite,subordinate lherzolite, variably serpentinized MMS3 Strike-slip fault, observed Strike-slip fault, constrained 00 Strike-slip13 fault, inferred Contact, observed Contact, inferred Contact, constrained 0 134.12° W DPu3 - Gabbro/pyroxenite dike/sill MMS4 - Marble/limestone, fossiliferous, mylonitized 120 K DPu3 1100 Dunite Peak ophiolite (DPO) BM Geologic Boundaires MMS4 0 Cretaceous Granite - Biotite granite (Last Peak Granite, 95.7 ± 0.6 Ma, Gallagher, 1999) 00 15 K 120 0 80 Geologic Map Units 61.58° N 900 61.58° N 0 MMS3 0 17 MMS3 130 110 0 00 Fig. 3a MMS3 DPu 10 61.6° N 170 0 DPu2 0 18 61.6° N 18 00 1900 DPc MMS3 Contour interval - 100 m 1500 Thrust fault, observed Thrust fault, constrained Thrust faiult, inferred River or stream 133.92° W 133.88° W Waterbody 61.56° N D’ABBADIE 00 16 DPu2 DPc DPu 1800 DPu2 1400 1500 DPu1 61.62° N 170 0 11 61.62° N MMS3 12 MMS3 1500 EM DPc Figure (a) 3 Looking NW Dunite Peak (2020 m) DP u1 DPu2 DPc MMS3 MMS2 MMS1 (b) DP Dunite Peak (2020 m) Looking W DPu3 u1 DPu2 DPc DPu2 Figure 4 (a) Sedimentary layer Greenstone - volcaniclastic (b) Greenstone Chert Sheared pillows Greenstone - volcaniclastic Sheared pillows Sheared pillows Chert Sheared pillows Greenstone - volcaniclastic Greenstone 5 cm 20 cm (c) (d) astic anicl lc al vo Gree nst one - ent fragm (e) (f) act sive t con u Intr Greenstone Intrusive contact Gabbro 5 cm Gabbro 5 cm Figure 5 a. b. 0.064 265 ± 4 Ma (MSWD= 0.9) 2.1 207 1.1 Pb/ 206 Pb 2.1 300 280 0.048 21 5.1 4.1 240 260 238 U/ 206 Pb 26 10.1 9.1 zircon/chondrite c. 10 3 10 2 10 1 10 0 10 -1 -2 10 100 µm La Ce Pr Nd Pm Sm Eu Gd Tb Dy Ho Er Tm Yb Lu 100 Figure 6 100 (a) Group 1 Group 1: Greenstone + gabbro, Dunite Peak ophiolite crustal section 10 Group 2: layered gabbro + leucogabbro + subordinate greenstone, Dunite Peak ophiolite crustal section Group 3: gabbro intrusions from the ultramafic section + greenstone from the crustal section, Dunite Peak ophiolite Rock/N-MORB Rock/N-MORB 10 (b) Groups 2 & 3 1 1 .1 .1 Group 1a (n = 9) Group 1b (n = 9) .01 Nb Th Ce La Sm Nd Hf Zr εNd = +7.2 to +7.5 Ti Eu Dy Gd Er Y Lu Yb Sc εNd = +8.1 to +9.0 Group 2 (n = 11) Group 3 (n = 6) .01 V Nb Th Ce La Sm Nd Hf Zr Ti Eu Dy Gd Er Y Lu Yb Sc V 100 100 Group 4: Pillow basalt, Marine metasedimentary succession (MMS3) (c) Group 4 (d) Group 5 Dunite Peak ophiolite ultramafic rocks Dunite/Harzburgite (n = 10) Lherzolite (n = 3) Pyroxenite (n = 1) DMM standard composition 10 Rock/N-MORB Rock/Primitive Mantle 10 1 1 .1 .1 Group 4 (n = 1) .01 Nb Th Ce La Sm Nd Hf Zr Ti Eu Dy Gd Er Y Lu Yb Sc V .01 Nb Th Ce La Sm Nd Hf Zr Ti Eu Dy Gd Er Y Lu Yb Figure (a) 7 (b) Y/15 alk. rhyolite 1 IAT rhyolite + dacite IAT-CA tephriphonolite trachyandesite + andes. Zr/Ti Group 1a (n = 9) Group 1b (n = 9) trach. .1 N-MORB phonolite * BABB * Group 2 (n = 11) E-MORB Group 3 (n = 6) basaltic andesite Calc-alkali basalt 1 Nb/Y 10 100 Nb/8 Ol (d) (c) Alkaline Intercontinental rifts La/10 Dunite SS OIB 1 ite hrl We Ha rz b urg ite Z 10 .1 Group 5 foidite alk. bas. .01 Continental. .01 Group 4 (n = 1) Th/Yb Lherzolite E-MORB PERIDOTITE PYROXENITE Ol-Websterite N-MORB SS Z .1 .01 .1 1 Nb/Yb 10 100 Opx Ortho- pyroxenite Websterite Clino- Cpx pyroxenite Dunite/Harzburgite (n = 12) Lherzolite (n = 3) Pyroxenite (n = 1) Izu-Bonin arc Mariana arc Kermadec arc Lau basin n (a-c) = 566 n (d-f) = 43 n (a-c) = 131 n (d-f) = 96 n (a-c) = 136 n (d-f) = 46 n (d-e) = 282 Y/15 10 (b) Group 1 (IAT/CA) (a) (c) IAT-CA Th/Yb 1 * SS Z .1 (d) (e) Z SS Th/Yb * E-MORB SS Z .1 N-MORB Calc-alkali .01 10 La/10 Group 1 (IAT/CA) (h) Z IAT-CA Th/Yb Rock/N-MORB IAT N-MORB * E-MORB Th Ce La Sm Nd Hf Zr Ti Eu Dy Gd Er Y Lu Yb Sc V Calc-alkali .01 .1 1 Nb/Yb 10 100 La/10 Continental. N-MORB SS Z .1 Nb * BABB E-MORB .1 .01 Nb/8 Y/15 OIB 1 1 Alkaline Intercontinental rifts (i) SS Group 2 (IAT/CA cumulates) Group 3 (BABB) * BABB E-MORB .1 10 N-MORB IAT IAT-CA (g) Nb/8 Y/15 OIB 1 1 Alkaline Intercontinental rifts (f) Group 3 (BABB) 10 .01 100 Dunite Peak ophiolite Calc-alkali La/10 Group 2 (IAT/CA cumulates) Rock/N-MORB Back-arc / Trough / Ridge N-MORB .01 10 Group 1 (IAT/CA) * BABB E-MORB E-MORB .1 .01 100 N-MORB OIB 1 Rock/N-MORB Arc 10 IAT SS Z Group 2 (IAT/CA cumulates) Group 3 (BABB) Continental. 100 Continental. Figure 8 Alkaline Intercontinental rifts Nb/8 (a) Mid Figure 9 Permian (280-260 Ma) Dunite Peak intra-oceanic arc (light grey unit) - Magmatism recorded by the Dunite Peak intra-oceanic arc between ~275 Accretionary complex of and 260 Ma and perhaps as early lower plate material (blue) as ~280 Ma - Eclogite facies metamorphism of Yukon-Tanana terrane recorded Yukon-Tanana between ~275 and 260 Ma terrane - Late Permian and older parts of Slide Mountain terrane may represent accreted lower plate material from the subducted oceanic lithosphere (dark blue) - Width of subducted ocean basin is unknown - St Cyr blueschist may represent accretionary complex of lower plate material (dark blue) Extensional zone Volcanic centers Width of basin between Dunite Peak intra-oceanic arc and Laurentian margin is unknown Laurentia Generation of surpasubduction zone ophiolites Eclogite facies metamorphism of Yukon-Tanana continental crust (purple) Earliest record of eclogite facies metamorphism of Yukon-Tanana terrane marks start of the Klondike orogeny. Coeval Dunite Peak magmatism and eclogite facies metamorphism of Yukon-Tanana terrane suggests Yukon-Tanana had an irregular shape with promontories and re-entrants and/or collided obliquely with the Dunite Peak intra-oceanic arc. Co-spatial and coeval development of constructional volcanic centres (e.g. Dunite Peak, Wolf Mountain klippe) and zones of extensional volcanism (e.g. Sylvester allochthon, Finlayson Lake district) is comparable to disorganized extensional intra-oceanic arcs such as the Kermedac arc - Havre trough system (b) Mid/late Permian (265-260 Ma) Orogenic collapse and associated magmatism (Klondike magmatic cycle) Suprasubduction zone ophiolites obducted onto Yukon-Tanana Width of basin between Dunite Peak - Cessation of Dunite Peak magmatism and subduction of Yukon-Tanana terrane by 260 Ma Yukon-Tanana - Exhumation of Yukon-Tanana terrane terrane blueschists and eclogites initiated at 265 Ma - Emplacement of Buffalo Pitts orogenic peridotite within Yukon-Tanana terrane at ~262 Ma - Main phase of Klondike magmatic cycle between 265 and 252 Ma Extension of Yukon-Tanana terrane accompanied by emplacement of the Buffalo Pitts orogenic peridotite intra-oceanic arc and Laurentian margin is unknown Laurentia Slab break-off Initiation of westward subduction of ocean between Klondike orogen and North American cratonic margin. Roll-back promotes extension. Buoyancy-driven slab break-off from the east margin of Yukon-Tanana terrane was followed by extension, exhumation and magmatism associated with orogenic collapse and initiation and roll-back of westward subduction beneath the Klondike orogen (c) Late Permian (260-252 Ma) Orogenic collapse and associated magmatism (Klondike magmatic cycle) - Continued exhumation of Yukon-Tanana terrane blueschists and eclogites between 265 and 256 Ma Yukon-Tanana - Klondike magmatic cycle continued until 252 Ma terrane - Amphibolite facies metamorphism of Yukon-Tanana terrane recorded between ~260 and 252 Ma Exhumation of YTT eclogite & blueschist Accretionary complex of lower plate material (magenta) Laurentia Latest occurrence of Klondike magmatism and metamorphism of Yukon-Tanana terrane coeval with exhumation marks the final stages of the Klondike orogeny, which ended by ~252 Ma (d) Post-Middle Triassic Accretionary complex of lower plate material may contain Triassic Jones Lake Fm. (magenta) Yukon-Tanana terrane Laurentia Westward subduction leads to collision and accretion of Klondike orogen to Laurentia - Collision and accretion of Klondike orogen to Laurentian margin facilitated by westward subduction of the ocean basin between the Klondike orogen and Laurentia - Triassic clastic deposits of the Jones Lake Formation (magenta) may represent obducted lower plate sediments from the basin between the Klondike orogen and Laurentia - Detrital zircons from the Jones Lake Formation suggest the Klondike orogen and Laurentian margin were in close proximity by the Late Triassic - Late Triassic - Jurassic magmatism and Early Jurassic metamorphism recorded by Yukon-Tanana terrane are absent from Laurentia in Yukon - Early Jurassic metamorphism and exhumation of Yukon-Tanana terrane records an Early Jurassic collisional event - Early Cretaceous plutonic suites in Yukon-Tanana terrane and Laurentia indicate collision and accretion of Klondike orogen to Laurentian margin occurred by that time. - Collision and accretion of Klondike orogen with Laurentian margin occurred after the Middle Triassic and by the Early Cretaceous Eclogite facies metamorphism Sub-eclogite facies metam. Klondike magmatic cycle YTT cooling/exhumation Undeformed by Permian fabrics Deformed by Permian fabrics or no information given YTT eclogite and blueschist facies crust YTT crust Buffalo Pitts orogenic peridotite 209 210 LATE NORIAN TRIASSIC 228 Tectonic interpreation See text for discussion 220-206 Ma Metamorphism of YTT Late Triassic Early Jurassic RHAETIAN Dunite Peak intra-oceanic arc magmatism Possible timing of YTT-NAC Accretion Figure 10 201 Exhumation of YTT 220 CARNIAN 230 MIDDLE LADINIAN 241 EARLY 247 250 240 ANISIAN Exhumation of YTT eclogites and blueschists 265-220 Ma 270-223 Ma 237 Exhumation of YTT (other than eclogites and blueschists) Buffalo Pitts peridotite emaplced during extension of YTT Orogenic collapse and subduction polarity reversal OLENEKIAN 250 INDUAN 252 269 WORDIAN 279 270 Dunite Peak intra-oceanic arc magmatism Subduction of YTT continental crust KUNGURIAN 280 Cisuralian 272 PERMIAN ROADIAN 265-260 Ma 265 275-260 Ma Guadalupian 260 CAPITANIAN ~262 Ma 260 Klondike magmatic cycle ~280-260 Ma WUCHIAPINGIAN 265-252 Ma 254 Klondike orogeny (275-252 Ma) Lopingian CHANGHSINGIAN Geochronology isotope and mineral system K-Ar Hornblende U-Pb Zircon ARTINSKIAN 290 290 SAKMARIAN 296 U-Pb Titanite K-Ar Muscovite U-Pb Monazite K-Ar Biotite Sm-Nd Ar-Ar Hornblende Rb-Sr Ar-Ar Muscovite Lu-Hf Ar-Ar Biotite Sample numbers correspond to geochronology sample data entry presented in Supplementary Materials including data source, sample description and analytical notes ASSELIAN 299 10 20 30 40 50 60 SAMPLE NUMBER 70 80 90 100 110 Table 1 Sample Group Lithology Map Unit LaN(1) Mean LaN(1) YbN(1) Mean YbN(1) [La/Yb]N(1) Mean [La/Yb]N(1) [Th/Nb]N(1) Mean [Th/Nb]N(1) [V/Ti]N(2,3) Mean [V/Ti]N(2,3) Ɛ Nd(4) 17RAY-AP034A1 1a Greenstone DP1 0.81 1.08 0.33 0.37 2.44 2.87 11.65 14.67 2.21 2.38 17RAY-AP043B1 1a Greenstone DP1 0.64 0.20 3.27 1.99 15RAY-JR074A 1a Greenstone DP1 1.20 0.41 2.93 6.47 2.36 7.2 15RAY-JR076A 1a Greenstone DP1 0.88 0.32 2.71 9.71 2.65 16RAY-AP074A2 1a Greenstone DP1 1.28 0.47 2.73 14.56 2.06 7.4 16RAY-AP075A2 1a Greenstone DP1 0.92 0.38 2.42 2.36 2.86 7.5 16RAY-AP076B1 1a Greenstone DP1 2.05 0.50 4.06 21.03 2.30 16RAY-AP088A1 1a Greenstone DP1 0.94 0.32 2.91 36.89 2.73 16RAY-JR111A2 1a Greenstone DP1 0.98 0.41 2.38 2.26 17RAY-AP035A2 1b Greenstone DP1 4.32 4.21 0.99 0.90 4.38 4.68 17.93 27.95 0.54 0.93 17RAY-AP040A1 1b Greenstone DP1 4.12 0.90 4.60 15.78 0.61 16RAY-AP076A1 1b Greenstone DP1 4.00 0.73 5.50 19.00 1.15 16RAY-AP087B2 1b Greenstone DP1 4.64 0.98 4.72 7.66 0.84 7.4 16RAY-AP104A1 1b Greenstone DP1 4.04 0.83 4.87 15.05 1.03 16RAY-AP159A1 1b Greenstone DP1 5.20 1.04 5.00 20.10 0.65 16RAY-AP181A1 1b Gabbro DP1 4.20 1.07 3.94 114.56 0.98 16RAY-AP209A1 1b Greenstone DP1 2.50 0.60 4.20 20.63 1.43 16RAY-AP234A1 1b Greenstone DP1 4.84 0.98 4.94 20.87 1.11 17RAY-AP041A1 2 Greenstone DP1 1.30 0.85 0.46 0.32 2.82 2.91 12.82 17.43 2.08 2.54 17RAY-AP043A1 2 Gabbro DP1 0.67 0.27 2.53 8.74 2.83 17RAY-AP043C1 2 Gabbro DP1 0.52 0.59 0.87 1.56 17RAY-AP044A1 2 Gabbro (cumulate) DP1 2.50 0.30 8.27 5.11 2.82 16RAY-AP074A1 2 Gabbro DP1 0.92 0.37 2.46 32.04 2.46 16RAY-AP078C1 2 Gabbro DP1 0.89 0.15 5.79 33.66 4.22 16RAY-AP095A1 2 Gabbro DP1 0.60 0.24 2.52 14.56 3.19 16RAY-AP120B1 2 Leucogabbro DP1 0.22 0.20 1.12 1.85 16RAY-AP195A1 2 Gabbro DP1 0.74 0.32 2.32 15.05 1.56 16RAY-AP202A1 2 Greenstone DP1 0.64 0.47 1.37 17.48 2.08 16RAY-AP212B1 2 Gabbro DP1 0.38 0.20 1.90 3.24 17RAY-AP041B1 3 Greenstone DP1 0.85 1.49 0.62 1.07 1.37 1.41 16.50 5.29 1.58 1.20 17RAY-AP053A1 3 Gabbro (cumulate) DP2c 1.92 1.35 1.42 2.62 1.06 16RAY-AP121B2 3 Greenstone DP1 1.76 1.04 1.69 3.52 1.17 16RAY-AP130C3 3 Gabbro (cumulate) DP2c 1.54 1.31 1.17 1.43 0.84 9.0 16RAY-AP179A1 3 Greenstone DP1 0.99 0.65 1.53 4.85 1.51 8.3 16RAY-AP199A1 3 Gabbro (cumulate) DP2c 1.88 1.46 1.29 2.84 1.03 8.8 16RAY-AP079A1 4 Pillow Basalt MMS3 1.42 1.42 0.66 0.66 2.14 2.14 1.55 1.55 1.43 1.43 15RAY-JR075A 5 Harzburgite DP2b 0.50 0.10 0.21 2.37 14.26 5.87 16RAY-AP088D1 5 Dunite DP2b 0.64 0.18 3.51 2.04 16RAY-AP093A1 5 Dunite DP2b 0.58 0.24 2.39 4.00 16RAY-AP094A1 5 Dunite DP2b 0.04 8.59 16RAY-AP099A2 5 Harzburgite DP2b 0.73 0.08 8.97 11.75 16RAY-AP101A1 5 Lherzolite DP2b 0.35 0.43 0.82 2.80 16RAY-AP113A2 5 Harzburgite DP2b 0.14 8.46 16RAY-AP123A2 5 Dunite DP2b 0.10 0.06 1.67 7.40 16RAY-AP151A1 5 Harzburgite DP2b 1.98 0.57 3.49 14.26 0.64 16RAY-AP172A2 5 Lherzolite DP2b 0.15 0.20 0.72 2.42 16RAY-AP172A1 5 Lherzolite DP2b 0.70 0.53 1.32 2.02 16RAY-AP174A2 5 Dunite DP2b 0.07 0.06 1.20 4.26 16RAY-AP178B1 5 Pyroxenite DP2b 0.13 0.10 1.29 26.82 16RAY-AP196D1 5 Dunite DP2a 0.51 0.22 2.28 1.57 16RAY-AP205A1 5 Harzburgite DP2a 0.42 0.37 1.16 2.47 16RAY-AP206A1 5 Harzburgite DP2a 0.20 0.10 2.01 2.80 Footnotes Full data set and QC presented in Supplementary Materials 02 -' indicates no data, due to concentratons outside of detection limits 1) Normalized against standard N-MORB composition (Sun & McDonough, 1989), except Group 5, which is normalized against standard primative mantle composition (Sun & McDonough, 1989) 2) Ti normalized against standard N-MORB composition (Sun & McDonough, 1989), except Group 5, which is normalized against primative mantle composition (McDonough & Frey, 1989) 3) V normalized against standard N-MORB composition (Klein, 2004), except Group 5, which is normalized against primative mantle composition (McDonough & Frey, 1989) 4) ƐNd calculation based on a model age of 265 Ma Latitude Longitude 61.63506654 61.59242971 61.59007387 61.59540921 61.63129486 61.59011971 61.5836507 61.63073 61.58143204 61.63232538 61.58670054 61.5836507 61.63234 61.62589037 61.59755302 61.60605985 61.62199554 61.58258804 61.58804837 61.59242971 61.59242971 61.59493771 61.63095953 61.59874752 61.62962702 61.59001219 61.59460302 61.58652204 61.61442871 61.58804837 61.61659902 61.59276702 61.60996203 61.60182918 61.59690452 61.62613451 61.60353936 61.63073 61.62613 61.62751169 61.62295119 61.6222022 61.61678803 61.60240819 61.62396118 61.59975069 61.59975069 61.59814669 61.61149835 61.59532635 61.59274121 61.59335037 -134.0680117 -134.0639359 -134.0590501 -134.0717427 -133.9679625 -134.0590276 -134.0376631 -133.96776 -134.0379607 -134.0755242 -134.0502722 -134.0376631 -133.97231 -134.0045107 -133.9073774 -133.9014004 -134.0470874 -134.0346879 -134.0538744 -134.0639359 -134.0639359 -134.0843975 -133.9672315 -133.9266214 -133.9452507 -133.9541316 -133.9347754 -134.0379319 -134.050282 -134.0538744 -133.9118072 -133.9526361 -133.9634552 -133.9070494 -133.9366562 -133.8870414 -133.9715166 -133.96776 -133.96498 -133.945937 -133.9497852 -133.993315 -133.994688 -133.9659106 -133.902718 -133.9312397 -133.9312397 -133.9385149 -133.9274497 -133.9356914 -134.0641057 -134.064108