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characteristics of the lower and upper plates, Tectonics, 25,. TC4006 .... up northeastern Canada (Quebec–Baffin Island segment of. THO, Figure 2) in the late ...
TECTONICS, VOL. 25, TC4006, doi:10.1029/2005TC001907, 2006

Trans-Hudson Orogen of North America and HimalayaKarakoram-Tibetan Orogen of Asia: Structural and thermal characteristics of the lower and upper plates Marc R. St-Onge,1 Michael P. Searle,2 and Natasha Wodicka1 Received 20 September 2005; revised 13 April 2006; accepted 21 April 2006; published 18 July 2006.

[1] The Trans-Hudson Orogen (THO) of North America and the Himalaya-Karakoram-Tibetan Orogen (HKTO) of Asia preserve a Paleoproterozoic and Cenozoic record, respectively, of continentcontinent collision that is notably similar in scale, duration and character. In THO, the tectonothermal evolution of the lower plate involves (1) early thinskinned thrusting and Barrovian metamorphism, (2) out-of-sequence thrusting and high-T metamorphism, and (3) fluid-localized reequilibration, anatexis, and leucogranite formation. The crustal evolution of the Indian lower plate in HKTO involves (1) early subduction of continental crust to ultrahigh pressure (UHP) eclogite depths, (2) regional Barrovian metamorphism, and (3) widespread high-T metamorphism, anatexis, and leucogranite formation. The shallow depths of the high-T metamorphism in HKTO are consistent with early to mid-Miocene ductile flow of an Indian lower plate midcrustal channel, from beneath the southern Tibetan Plateau to the Greater Himalaya. Melt weakening of the lower plate in THO is not observed at a similar scale probably due to the paucity of pelitic lithologies. Tectonothermal events in the upper plate of both orogens include precollisional accretion of crustal blocks, emplacement of Andean-type plutonic suites, and high-T metamorphism. Syncollisional to postcollisional events include emplacement of garnet-biotite-muscovite leucogranites, anatectic granites, and sporadic metamorphism (up to 90 Myr following the onset of collision in THO). Comparing the type and duration of tectonothermal events for THO and HKTO supports the notion of tectonic uniformitarianism for at least the later half of dated Earth history and highlights the complementary nature of the rock record in an older ‘‘exhumed’’ orogen compared to one undergoing present-day orogenesis. Citation: St-Onge, M. R., M. P. Searle, and N. Wodicka (2006), Trans-Hudson Orogen of North America and HimalayaKarakoram-Tibetan Orogen of Asia: Structural and thermal 1 2

Geological Survey of Canada, Ottawa, Ontario, Canada. Department of Earth Sciences, University of Oxford, Oxford, UK.

Copyright 2006 by the American Geophysical Union. 0278-7407/06/2005TC001907

characteristics of the lower and upper plates, Tectonics, 25, TC4006, doi:10.1029/2005TC001907.

1. Introduction [2] The question ‘‘how far back in the 4.4 Ga time span recorded by dated Earth materials does uniform tectonics apply’’ remains a source of much discussion in Earth sciences (for recent reviews and opinions, see de Wit [1998], Hamilton [1998, 2003], Murphy and Nance [1999], Bleeker [2002], McCall [2003], and Stern [2005]). Much of the current debate centers on the interpretation and dating of rock sequences, how these fit into geodynamic models utilized for the Earth’s evolution [e.g., Hamilton, 2002; Turcotte and Schubert, 2002], and how they pertain to the issues of mantle fractionation and differentiation, the growth of continental crust through geological time, the scale and mechanics of mantle convection, and the driving mechanism(s) for lithospheric plates [e.g., Anderson, 2000, 2002; Artemieva and Mooney, 2001; Grand, 2002; Bercovici, 2003; Kreemer et al., 2003; Stern, 2005]. [3] In this paper we take a geological, field researchbased approach to address the issue of secular changes in tectonic uniformitarianism for at least the later half of dated Earth history by reviewing the salient crustal features of the lower (subducted) and upper (nonsubducted) tectonic plates of two large collisional orogenic/mountain belts: the late Paleoproterozoic Trans-Hudson Orogen of North America and the Cenozoic Himalaya-Karakoram-Tibetan Orogen of Asia. The premise on which our approach is based is that if one can compare the type and duration of magmatic, deformation, and metamorphic events (including the amounts and time spans of crustal shortening, thickening, regional metamorphism, and magmatism) for two orogens of comparable size, magnitude, and constituent tectonic components, yet separated in time by over 1750 Myr, then it provides a basis for considering the principle of actualism for crustal tectonics at least as far back as the late Paleoproterozoic. This in turn allows the India-Asia collision zone, and the structural and thermal evolution of its Himalayan, Karakoram, and Tibet domains to be viewed as present-day analogues for some of the older orogenic belts in the geological record back to 2 Ga. Remaining differences in the accessible rock record of collisional orogens can then be evaluated in terms of contrasting depths of erosion, degree of thermal preservation, and original tectonostratigraphy, as we will demonstrate with the TransHudson and Himalaya-Karakoram-Tibetan orogens below.

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Figure 1. (a) Summary geological map of North America including Greenland. Map shows extent of Trans-Hudson Orogen and location of bounding Archean crustal blocks and cratons. Abbreviations are GO, Grenville Orogen; MP, Manitoba promontory; QP, Quebec promontory. (b) Location map for the Himalaya-Karakoram-Tibetan Orogen in Asia. Abbreviations are HKu, Hindu Kush mountain range; Kar, Karakoram mountain range; NB, Namche Barwa syntaxis; NP, Nanga Parbat syntaxis. Note similar scale to that of Figure 1a. This study puts flesh on the contention, which goes back at least a century to Sederholm [1932, 1934], that orogenic development has changed little since the early Proterozoic, and extends back in terms of Earth history similar conclusions reached for the Neoproterozoic Pan-African Mobile Belt of northwest Africa [Black et al., 1979; Caby et al., 1981].

2. Trans-Hudson Orogen (THO) [4] The Trans-Hudson Orogen (THO) [Hoffman, 1988; Lewry and Collerson, 1990] is a Himalayan-scale collisional orogenic belt that extends from the south central part of the North American continent to its northeastern edge (Figure 1a), where it is truncated by the younger Mesoproterozoic to Neoproterozoic Grenville Orogen. THO separates the lower plate Archean Superior craton from an upper plate collage of Archean crustal blocks that includes the

Wyoming craton to the west, the Hearne domain and Rae craton of the Churchill plate to the north, and the Nain or North Atlantic craton to the east (Figure 1a). The orogen is over 4600 km in strike length and, in certain segments, over 800 km across strike. The Manitoba promontory in the west (MP, Figure 1a) and the Quebec promontory in the northeast (QP, Figure 1a) mark the corners of the indenting Superior lower plate. Beyond North America, THO has been correlated with the Nagssugtoqidian Orogen of Greenland, the Kola-Karelia Orogen of the Baltic Shield, and possibly with the Transantarctic Orogen of Eastern Antarctica [Zhao et al., 2002]. [5] Recently, St-Onge and Searle [2004] and Searle and St-Onge [2004] have argued that the tectonic record of collision and indentation of the Superior craton into the collage of cratons and terranes (Churchill plate) that made up northeastern Canada (Quebec – Baffin Island segment of THO, Figure 2) in the late Paleoproterozoic (circa 1830–

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Figure 2. Simplified geological map of the Quebec-Baffin segment of Trans-Hudson Orogen. Principal tectonostratigraphic units and bounding crustal sutures are described in the text. Lines of section A-B and C-D-E for Figure 4 are shown. Abbreviations are BaS, Baffin suture; BeS, Bergeron suture; CB, Cumberland batholith; CF, Chioak Formation; CS, Cumberland Sound; HBG, Hoare Bay Group; L&B, Lake Harbour Group, Blandford Bay assemblage and crystalline basement (Ramsay River orthogneiss); NA, Narsajuaq arc; Ord, Ordovician cover; P&S, Parent and Spartan groups; P&C, Povungnituk and Chukotat groups; PG, Piling Group; SRS, Soper River suture; and WG, Watts Group. 1795 Ma) was similar to the record of collision and indentation of India into the collage of plates and terranes that made up central Asia in the Eocene (beginning at circa 50.6 Ma [Rowley et al., 2004]). 2.1. Crustal Components [6] An overview of the tectonostratigraphic framework of the Quebec-Baffin segment of THO (segment which is superbly exposed in northern Quebec and Baffin Island and

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which has been the focus of several recent mapping and research projects by the Geological Survey of Canada) is given below from south to north [St-Onge et al., 2002; Scott et al., 2003, and references therein]. All age dates and age constraints for THO cited in subsequent sections of the paper are taken from Parrish [1989], Jackson et al. [1990], St-Onge et al. [1992, 2006], Machado et al. [1993], Scott and St-Onge [1995], Dunphy et al. [1995], Scott [1997, 1998, 1999], Wodicka and Scott [1997], Scott and Wodicka [1998], Wodicka et al. [2002a, 2002b], Bethune and Scammell [2003], D. J. Scott et al. (manuscript in preparation, 2006), and N. Wodicka et al. (manuscript in preparation, 2006). 2.1.1. Lower Plate Superior Craton and Northern Margin [7] In northern Quebec, the exposed lower plate Archean Superior craton (Figure 2) comprises dominantly felsic orthogneisses and plutonic units ranging in age between 3220+32/ 23 and 2737±2 Ma. Unconformably overlying the felsic basement is a suite of parautochthonous basal clastic sedimentary units, carbonatitic volcaniclastic rocks, and continental tholeiitic flood basalt and rhyolite (Povungnituk Group, Figure 2) associated with initial Paleoproterozoic rifting of the northern Superior craton (present-day coordinates). These have yielded ages between 2038+4/ 2 and 1959+3.1/ 2.7 Ma. Stratigraphically overlying the initial rift sedimentary and volcanic rocks is a younger succession of predominantly komatiitic to tholeiitic basalts (Chukotat Group, Figure 2) accumulated during renewed rifting along the northern continental margin and dated between 1887+39/ 11 and 1870±4 Ma. The ages of the younger volcanic succession indicate that 150 Myr elapsed between the onset of initial continental rifting and the subsequent rifting event [St-Onge et al., 2000a]. 2.1.2. Bergeron Suture Zone and Watts Group Ophiolite [8] A south verging tectonic boundary or crustal suture (Bergeron suture; Figures 2, 3, and 4) separates the northern Superior margin strata from allochthonous crustal elements of the Churchill plate to the north [St-Onge et al., 2001]. Associated with and sitting in the hanging wall of the Bergeron suture are the crustal components of an obducted Paleoproterozoic ophiolite (Watts Group, Figure 2, 3, and 4) dated as 1998±2 Ma. Rocks of the ophiolite represent two magmatic suites, each with a mantle source that is Nd isotopically distinct [Scott et al., 1992, 1999]. The older mid-ocean ridge basalt-like suite comprises >5 km of pillowed and massive volcanic rocks, sheeted dikes, dominantly gabbroic layered cumulate rocks, and rare plagiogranite. The younger suite includes >4 km of light rare earth element and large-ion lithophile element-enriched sheeted mafic dikes and layered ultramafic to mafic cumulate rocks that have a within-plate oceanic island chemical signature. Preservation of the ophiolite and the higher structural levels it represents within the Paleoproterozoic orogen is entirely a function of the late to postcollisional, crustal-scale, orogenparallel folding and orogen-perpendicular cross-folding which characterizes the southern, lower plate margin of THO in northern Quebec [Lucas and Byrne, 1992]. The offshore extension of the Bergeron suture, and others in

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Figure 3. Geological compilation of the southern margin of the Trans-Hudson Orogen in northern Quebec [after St-Onge et al., 2001] (with permission from Elsevier). The Bergeron suture separates the parautochthonous units of the northern Superior craton in its footwall from allochthonous ophiolitic, island arc, and Andean margin units in its hanging wall. Line of section A-B for Figure 4 is shown. Abbreviations are BeS, Bergeron suture, and CSB, Cape Smith Thrust Belt. THO that are described below, is based on the interpretation of geophysical data as presented by St-Onge et al. [2002]. 2.1.3. Narsajuaq Island Arc Terrane [9] North of the Bergeron suture and Watts Group ophiolite, a magmatic arc terrane forms the leading edge of the Churchill plate and includes (Figures 2, 3, and 4): fore-arc clastic deposits (Spartan Group), a dominantly volcanic sequence (Parent Group), and a dominantly plutonic assemblage (Narsajuaq arc). The volcanic and plutonic units of the arc terrane can be grouped into two temporally and petrologically distinct, suites [Dunphy and Ludden, 1998]. The older suite includes calc-alkaline layered diorite-tonalite gneiss and tholeiitic to calc-alkaline basaltic andesite to rhyolite, and is dated between 1863±2 and 1845±2 Ma. It is interpreted as an island arc assemblage built on Paleoproterozoic oceanic crust (Watts Group above) and a rifted sliver of Archean continental crust [The´riault et al., 2001]. The younger suite comprises crosscutting, gneissic to massive, monzodiorite to granite plutons. It is bracketed between 1842+5/ 3 and 1820+4/ 3 Ma and is interpreted as having been emplaced in an Andean-type

continental arc setting [Dunphy and Ludden, 1998; The´riault et al., 2001]. 2.1.4. Meta Incognita Microcontinent [10] On southern Baffin Island, a second south verging tectonic boundary or crustal suture (Soper River suture, Figures 2 and 4 [St-Onge et al., 2001]) separates the Narsajuaq arc terrane in its footwall from the Archean and Paleoproterozoic units of the Meta Incognita microcontinent in its hanging wall. The microcontinent (Figures 2 and 4) comprises: (1) a continental margin shelf succession (Lake Harbour Group) and its crystalline basement (Ramsay River orthogneiss), (2) an overlying foreland basin succession (Blandford Bay assemblage), and (3) an extensive suite of Andean margin-type quartz diorite to monzogranitic plutons (Cumberland batholith) that intrude both 1 and 2. The age of the shelf succession and foreland basin is bracketed between 1934±2 and circa 1880 Ma, and that of the underlying crystalline basement is bracketed between 3019±5 and 1950+6/ 4 Ma. Plutons of the Cumberland batholith have yielded ages between 1865+4/ 2 and 1848±2 Ma. At this point it remains unclear whether the microcontinent was

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Figure 4. Schematic, strike-perpendicular cross section of the Trans-Hudson Orogen in northern Quebec and southern Baffin Island [after St-Onge et al., 1999] (with permission from Elsevier). Lines of section A-B and C-D-E are shown on Figures 2 and 3. The cross section is of the collisional orogen following orogenparallel thick-skinned crustal folding [cf. Lucas and Byrne, 1992]. The position of the Moho is assumed. Abbreviations are BeS, Bergeron suture; CP, Churchill plate; SP, Superior plate; SRS, Soper River Suture. initially rifted from the Superior craton as suggested by StOnge et al. [2000a], whether it constitutes a rifted fragment of the northern Rae craton (described below), or whether it represents crust exotic with respect to both bounding cratons. 2.1.5. Upper Plate Rae Craton and Margin [11] In the central Baffin Island area (top part of island on Figure 2) the upper plate Rae craton comprises felsic orthogneisses, metamorphosed sedimentary and volcanic rocks, and younger felsic plutonic rocks ranging in age between 2868+13/ 12 and 2702±3 Ma. The southern margin of the craton is unconformably overlain by the Paleoproterozoic Piling Group (Figure 2), which comprises [Scott et al., 2003]: continental margin clastic and carbonate platform strata (younger than 2159±16 Ma), mafic intrusive, extrusive, and sedimentary rift units (younger than 1980±11 Ma), and foredeep turbidites (younger than 1915±8 Ma). Various felsic plutonic rocks, ranging in age from 1.90 to 1.85 Ga and including the northernmost component of the Cumberland batholith, intrude the south facing Piling Group. 2.2. Structural Evolution of the Lower Plate [12] Parautochthonous sedimentary and volcanic strata along the northern margin of the Superior craton are imbricated by D2a thrust faults (Table 1) above a regional basal de´collement [Lucas, 1989] within the Cape Smith Thrust Belt (Figures 3 and 4). Fault displacement was in a southerly direction, with thin-skinned imbrication and associated folding occurring in a piggyback sequence toward

the foreland. Cumulative displacement along the D 2a de´collement is estimated to be at least 100 km south of the Bergeron suture, although this must be viewed as an absolute minimum given that the greater part of the foreland thrust-fold belt in northern Quebec has been eroded away, which leads to the current extensive exposures of footwall crystalline basement (Figures 3 and 4). [13] D2a thrust deformation was initiated after 1870±4 Ma, the age of the youngest unit within the parautochthonous Superior craton cover sequence. D2a may be as young as circa 1830– 1815 Ma (age range overlapping with the M2a metamorphic event; see below). St-Onge et al. [2000b] have suggested that D2a deformation may record early terrane accretion (ophiolite obduction?) and/or subduction of the lower plate crystalline basement prior and leading up to the main (D2b) collisional event. [14] A distinct suite of late or ‘‘out-of-sequence’’ thrust faults that postdate the D2a structures, reimbricate the parautochthonous cover units of the Superior craton [Lucas, 1989]. These younger structures are thick skinned (involving both crystalline basement and Paleoproterozoic cover units) and are interpreted as D2b in age (Table 1). They are collisional in origin as they can be linked to accreted terrane boundary faults and the Bergeron suture [St-Onge et al., 2001], and are thus considered as the main structural manifestation of the Superior-Churchill collision within the lower plate. A minimum estimate on the cumulative amount of shortening (tectonic overlap between Churchill plate and Superior plate units) accommodated by the D2b

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Table 1. Principal Tectonothermal Events Identified in the Quebec-Baffin Segment of Trans-Hudson Orogen Event D0

M0 D1a

M1a

M1b M1c D2

M2 D2a

M2a

D2b

M2b M3 M3

Age Constraintsa

Description accretion of the Meta Incognita microcontinent to the southern margin of the upper plate Rae craton (closure of Baffin suture) low-pressure crustal metamorphism at higher structural levels within the upper plate accretion of the intra-oceanic Narsajuaq arc to the southern margin of the composite upper plate (closure of Soper River suture) regional granulite-facies metamorphism following emplacement of the Cumberland batholith in the composite upper plate and accretion of the intra-oceanic Narsajuaq arc localized thermal reequilibration of the upper plate associated with post-D1a crustal magmatism granulite-facies metamorphism of lower crustal levels within the upper plate Narsajuaq arc (Andean margin phase) collision between the composite upper plate Churchill domain and the northern margin of the lower plate Superior craton as documented by recumbent folds and steep-reverse faults in the upper plate retrograde amphibolite-facies metamorphism of upper plate domains associated with collisional D2 structures first phase of collision/terrane accretion to the northern margin of Superior craton as documented by early piggyback sequence thrust faults in the lower plate regional, kyanite-sillimanite grade metamorphism associated with early thin-skinned thrusting of cover units in the lower plate second phase of collision between the composite upper plate Churchill domain and the northern margin of the Superior craton as documented by out-of-sequence thrust faults in the lower plate amphibolite facies reequilibration of the lower plate Superior craton crystalline basement and cover during D2b postcollisional reequilibration of upper plate domains associated with the emplacement of leucogranites postcollisional reequilibration of lower plate domains associated with the emplacement of leucogranites

1883±5 – 1865+4/

2

Ma

circa 1877 – 1850 Ma 1845±2 – 1842+5/

3

Ma

circa 1849 – 1835 Ma

circa 1833 – 1829 Ma 1836±2 – 1825±3 Ma 1820+4/

3

– 1805±2 Ma

1820±1 – 1808±3 Ma ca. 1830 – 1815 Ma

1820+4/

3

– 1815±4 Ma

1815±4 – 1795±2 Ma

1815±4 – 1785±1 Ma 1797±2 – 1785±4 Ma ca. 1795±2 Ma

a

References for age constraints are provided in text.

structures along the southern margin of the orogen is 260 km (Figure 4) [St-Onge et al., 1999]. [15] D2b faults truncate the M2a metamorphic isograds within the Cape Smith Thrust Belt and thus must postdate circa 1820+4/ 3 –1815±4 Ma [Be´gin, 1992] (see below). They predate the age of emplacement of postkinematic syenite plugs and syenogranite dikes (1795±2 –1758.2±1.2 Ma).

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2.3. Metamorphic Evolution of the Lower Plate [16] Three phases of metamorphism characterize the thermal evolution of the lower plate. An M2a greenschistto amphibolite-facies regional metamorphism is associated with early thin-skinned thrusting of cover units along the north margin of the Superior craton (Table 1) [Be´gin, 1992]. Earlier M0, M1a, M1b, M1c thermal events (Table 1) are restricted to the orogenic upper plate (see section 2.5). M2a is a regional, Barrovian facies, kyanite-sillimanite-grade, medium temperature, and pressure metamorphism during which P-T conditions from foreland to hinterland along the exposed base of the Cape Smith Thrust Belt increased from 400°C and 6.3 kbar to 575°C and 9.1 kbar [St-Onge et al., 2000b]. M2a metamorphism is bracketed between 1820+4/ 3 and 1815±4 Ma and is interpreted as a consequence of the relaxation of isotherms in the tectonically thickened thrust belt [St-Onge and Lucas, 1991]. [ 17 ] Regionally extensive M 2b upper amphibolitefacies mineral domains within the lower plate basement broadly parallel the exposed basement-cover contact (Table 1) [St-Onge and Lucas, 1995]. Microstructures indicate that reequilibration of the footwall basement occurred during D2b with P-T conditions from foreland to hinterland along the exposed base of the thrust belt increasing from 585°C and 7.7 kbar to 720°C and 9.8 kbar [St-Onge et al., 2000b]. Growth of coronitic titanite, a mineral found only in reequilibrated basement units [St-Onge and Ijewliw 1996] occurred between circa 1814 and 1789 Ma, and growth of metamorphic monazite and zircon within overlying cover units is documented between 1815±4 and 1785±1 Ma. M2b is interpreted by St-Onge and Lucas [1995] to have resulted from syncollisional relaxation of crustal isotherms, with new mineral growth occurring within zones of fluid infiltration focused along the basement-cover contact. [18] A third phase of metamorphism characterizing the thermal evolution of the lower plate involves localized postcollisional M3 reequilibration to middle amphibolitefacies conditions (Table 1). The M3 metamorphism is associated with partial melting and the emplacement of tourmaline-garnet-muscovite leucogranites. P-T conditions range from 675°C and 7.0 kbar to 775°C and 8.9 kbar. M3 is dated at circa 1795±2 Ma in the collisional lower plate. 2.4. Lower Plate Magmatism [19] Leucogranites emplaced within the lower plate of THO contain quartz, K-feldspar and plagioclase with varying amounts of tourmaline, garnet, muscovite, and biotite, and minor amounts of sillimanite and cordierite. The leucogranites typically have a pegmatitic texture, are undeformed, and cut all fabrics in the host rock [St-Onge et al., 1992]. These granites, although not abundant, are widespread in the exposed supracrustal domains of the lower plate and are interpreted as derived from melting of lower Povungnituk Group distal pelites. At the mostly lower structural levels preserved within the eastern THO, the leucogranites occur primarily as dikes, and rarely as sills, with the age of emplacement ranging between 1795±2 and 1758±1.2 Ma.

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2.5. Upper Plate Crustal Accretion, Andean-Type Magmatism and Metamorphism (Precollisional) [20] A number of precollisional, deformation, magmatic, and metamorphic events of Paleoproterozoic age can be documented in the rock record of the upper (Churchill) plate of THO [see also St-Onge et al., 2006]. From north to south, the sequence of accretion, plutonism, and metamorphism includes (Table 1): (1) accretion of the Meta Incognita microcontinent to the southern Rae craton (D0) and consequent low-P crustal metamorphism at preserved high structural levels (M0), (2) accretion of the intraoceanic Narsajuaq arc to the southern margin of the Meta Incognita microcontinent (D1a), (3) regional lower granulite-facies metamorphism resulting from emplacement of the Cumberland batholith and the accretion of Narsajuaq arc (M1a), (4) a distinct thermal reequilibration event associated with postaccretion magmatism (M1b), and (5) a separate granulite-facies metamorphic event spatially and temporally related (see below) to the southern Andean phase of Narsajuaq arc (M1c). [21] Accretion of the Meta Incognita microcontinent to the southern margin of the Rae craton (and closing of the intervening Baffin suture) led to the formation of a D0 north verging, thin-skinned, thrust-fold belt along the north margin of THO on Baffin Island (northern thrust-fold belt, Figure 2). Thrust belt structures imbricate the shelf margin strata of the Piling Group (youngest unit dated at 1883±5 Ma) and predate emplacement of the 1865+4/ 2 – 1848±2 Ma stitching Cumberland batholith [Scott et al., 2003]. D0 fabrics and M0 assemblages are well preserved at high structural levels within the northern thrust-fold belt with cordierite-andalusite assemblages yielding P-T estimates of 550°C to 600°C and 3.0 kbar to 4.0 kbar [Allan and Pattison, 2003] between circa 1877 and 1850 Ma. [22] Continued convergence between the leading edge of the Churchill plate (i.e., the southern margin of the accreted Meta Incognita microcontinent) and crustal domains to the south led to D1a accretion of the Narsajuaq arc and formation of the Soper River suture (Figure 2). Arc accretion is constrained to have taken place after 1845±2 Ma, the age of the youngest gneissic unit associated with the intraoceanic phase of Narsajuaq arc and to predate the subsequent Andean margin phase of Narsajuaq arc for which the oldest plutonic unit is dated at 1842+5/ 3 Ma [St-Onge et al., 2006]. [23] Prograde M1a metamorphism within the southern Meta Incognita microcontinent is characterized by lower granulite-facies pelitic assemblages containing garnet-cordierite-sillimanite-biotite-plagioclase – K-feldspar – quartzgranitic pod. P-T conditions range from 790°C and 6.9 kbar to 845°C and 8.4 kbar at circa 1849 –1835 Ma [St-Onge et al., 2006]. M1a metamorphism thus immediately follows the emplacement of the Andean margin Cumberland batholith and is largely concurrent with accretion of the Narsajuaq arc along the southern margin of the Meta Incognita microcontinent. Concordant U-Pb ages from metamorphic monazite and zircon in siliciclastic and gneiss samples suggest that rocks from the Meta Incognita microcontinent experienced a subsequent M1b thermal perturbation at circa 1833 – 1829 Ma, which is interpreted as related to the emplacement

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of circa 1834 – 1831 Ma postaccretion felsic dikes and plugs [St-Onge et al., 2006]. [24] Metaplutonic rocks from the lower crustal levels of the Narsajuaq arc contain M1c mineral assemblages consistent with granulite-facies conditions. P-T estimates range from 800°C and 6.6 kbar to 900°C and 9.2 kbar [St-Onge et al., 2000b]. U-Pb dating of monazite and metamorphic zircon in supracrustal rocks of the arc suggest the M1c metamorphism culminated at circa 1836±2 – 1825±3 Ma. 2.6. Upper Plate Crustal Metamorphism and Magmatism (Syncollisional to Postcollisional) [25] D2 collisional deformation (outcrop to map-scale, south verging, recumbent folding and steep-reverse faulting of earlier fabrics and basement/cover units) is associated with M2 retrograde upper amphibolite-facies metamorphism of upper plate domains on Baffin Island and northern Quebec. D2 deformation is constrained between 1820+4/ 3 (age of the youngest dated upper plate Narsajuaq arc unit) and 1805±2 Ma (age of a postdeformation leucogranite dike) in the upper plate (Table 1). Zones of D2 deformation are typically characterized by the growth of retrograde micaceous M2 assemblages at the expense of M1a and M1c lower granulite-facies assemblages. P-T determinations range from 665 to 765°C and 5.0 to 7.7 kbar. New growth of zircon, monazite, and titanite in upper plate metasedimentary units indicates that M2 retrograde metamorphism occurred at circa 1820±1 – 1808±3 Ma (Table 1), with the source of the retrogressive fluid phases proposed by StOnge et al. [2000b] to be the underthrust and dehydrating supracrustal units of the orogenic lower plate (described above). [26] Syncollisional to postcollisional, two-mica, garnettourmaline leucogranite dikes, as well as postcollisional, anatectic, equigranular to K-feldspar megacrystic, monzogranite to syenogranite plutons [Dunphy and Ludden, 1998] are emplaced in the upper plate and range in age from 1805±2 to 1742.2±1.3 Ma. Localized postcollisional M3 reequilibration to middle amphibolite-facies conditions is associated with partial melting and emplacement of the twomica, garnet-tourmaline leucogranites. M3 is bracketed between 1797±2 Ma and 1785±4 Ma in the upper plate (Table 1). 2.7. Syncollisional Sedimentation [27] Flexure of the lower plate Superior craton as it underthrusts the southern margin of THO resulted in the development of a foreland basin that accumulated debris eroded from the folded and thrusted continental margin strata [Hoffman, 1988]. Although largely eroded away due to the current relatively deep structural levels exposed along the southern margin of the Paleoproterozoic orogen, remnants of the molasse basin are nevertheless preserved southwest of Ungava Bay as the Chioak Formation (Figure 2). This formation comprises largely syntectonic, immature, continental clastic sedimentary rocks (fluvial conglomerate and arkose) [Wares and Goutier, 1990] and occurs no less than 500 km south of the current southern

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Figure 5. Geological map of the Himalaya [after Searle and Szulc, 2005] (with permission from Elsevier). Abbreviations are K, Katmandu; MBT, Main Boundary Thrust; MCT, Main Central Thrust; MFT, Main Frontal Thrust; MMT, Main Mantle Thrust; NP, Nanga Parbat syntaxis; P, Peshawar; S, Spontang ophiolite; SS, Shyok Suture Zone; STD, South Tibetan Detachment.

(erosional) margin of the foreland thrust-fold belt in northern Quebec.

are interpreted as the (deeper) geological extension of the Lhasa Block of South Tibet [Searle, 1991]. 3.1. Crustal Components

3. Himalaya-Karakoram-Tibetan Orogen (HKTO) [28] The collision of the Indian plate with Asia approximately 50 Myr ago resulted in the formation of the Himalayan and Karakoram –Hindu Kush mountain ranges and the Tibetan Plateau (Figure 1b), arguably the largest continental collision in the last 450 Myr of Earth history. The Himalayan mountain range (shaded, Figure 1b) is approximately 2500 km long and stretches from the Afghanistan-Pakistan border in the west, along the Ladakh, Zanskar, Spiti, Himachal, Garhwal ranges of northern India, and along the Nepal, Sikkim, Bhutan and Aranachal Pradesh regions to western Yunnan and northern Myanmar (Burma) in the east. The Nanga Parbat syntaxis in the far northwest (NP, Figure 1b) and the Namche Barwa syntaxis in the far northeast (NB, Figure 1b) mark the corners of the indenting Indian lower plate. The Tibetan Plateau is the largest area (3500  1500 km2) of high elevation (average 5023 meters above sea level) and thick crust (65 – 85 km) in the world. It has low relief (generally less than 1 km) and is extremely flat in the internally drained interior of the plateau. The Karakoram mountain range along the borders of Pakistan and India with Tadjikistan and southwest China

[29] Many consider the India-Asia collision zone as the type continent-continent collisional orogenic belt and consequently its relatively well constrained tectonic evolution has been extensively reviewed and synthesized along several segments of the orogen over the past number of years [e.g., Gansser, 1964; Molnar and Tapponnier, 1975; Alle`gre et al., 1984; Coward and Ries, 1986; Malinconico and Lillie, 1989; Dewey et al., 1989; Searle, 1991; Treloar and Searle, 1993; Chamberlain and Zeitler, 1996; Le Fort, 1996; Searle, 1996; MacFarlane et al., 1999; Khan et al., 2000; DiPietro and Pogue, 2004]. For the purposes of this paper, a brief overview of the principal tectonic domains of the Himalaya-Karakoram-Tibetan Orogen (HKTO) will therefore suffice and is given below from south to north [see also Searle, 2006] with the aim of comparing and contrasting with similar domains in the THO. Readers are referred to the papers and syntheses listed above (and references therein) for more detailed descriptions and documentation of individual tectonic domains within the Asian orogen. [30] The major tectonic zones and principal crustal faults within the Himalaya are shown on the map of Figure 5. Figure 6 is a north-south cross section across the western

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Figure 6. Cross section of the western Himalaya (Kishtwar-Zanskar region, northwest India) and central Karakoram (northern Pakistan) [after Searle, 1991]. Abbreviations are GHS, Greater Himalayan sequence; MBT, Main Boundary Thrust; MCT, Main Central Thrust; MHT, Main Himalayan Thrust; MKT, Main Karakoram Thrust; ZSZ, Zanskar Shear Zone (South Tibetan Detachment).

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Himalaya and the central Karakoram of northeastern Pakistan and northwestern India showing the major features of each domain. All age dates and age constraints for the HKTO cited in subsequent sections of the paper are taken from Searle et al. [1999a], Walker et al. [1999], Fraser et al. [2001], Hildebrand et al. [2001], Grujic et al. [2002], Leech et al. [2005], Searle and Szulc [2005], Searle [2006], and references therein. 3.1.1. Lower Plate Indian Craton and Northern Margin [31] The lower plate Indian craton (Figure 1b) comprises dominantly charnockite, orthogneiss, migmatite, granite, metapelite, and quartzite units of Archean age that are variably overlain by Proterozoic paragneisses and volcanic rocks or intruded by granites [Naqvi and Rogers, 1987]. In central western India, the Precambrian basement is overlain by the circa 65 Ma Deccan Trap tholeiitic lavas. Along the northern margin of the craton, the crystalline basement is overlain by thin early Paleozoic cover sedimentary rocks. A flexural foreland basin, the Siwalik molasse basin, formed along the southern margin of the Himalaya, that accumulated much of the erosional detritus of the evolving and uplifting Himalaya to the north. [32] Prior to collision with Asia, the Indian craton was attached to southern Africa, Madagascar, the Seychelles, and Antarctica as part of the Gondwana supercontinent during the early Mesozoic. India-Madagascar broke from the African plate and rifted away approximately 155 Myr ago [de Wit et al., 1988; Reeves and de Wit, 2000; de Wit, 2003]. By marine magnetic anomaly 34 (80 – 89 Ma) India had separated from Madagascar and begun its rapid northward drift across the Indian Ocean as the mid-Indian ridges began a period of rapid extrusion and ocean floor spreading. At magnetic anomaly 22 (circa 53 Ma) India made first contact with the Asian crust [see Ali and Aitchison, 2005], roughly at equatorial latitude, and drift rates decreased rapidly from 25 –20 to 5 cm/yr. Since 50.6 Ma [Rowley et al., 2004], India has indented northward into Asia and some 2000 km of crustal shortening has been taken up within the Himalaya-Tibet region since the early Eocene. GPS shows that the Indian plate is continuing to move north relative to the stable Siberian platform [Bilham et al., 1997]. Older mountain ranges along the northern margin of Tibet (Pamir, Tien Shan, Kun Lun, Altyn Tagh) have been reactivated and uplifted further as a result of the Indian plate collision and indentation. Uplift may extend even further to the north as far as the Altay ranges of Mongolia. 3.1.2. Himalayan Mountain Ranges [33] The Himalayan mountain ranges resulted from the collision of a lower plate passive continental margin (Indian plate) with an active upper plate Andean-type continental margin (Lhasa Block and Karakoram terrane) characterized by the Trans-Himalayan calc-alkaline batholith (Figure 5) and associated calc-alkaline volcanics. The southern margin of the Asian plate may have had a crustal thickness and elevation similar to the present-day Andes, prior to the Indian plate collision [England and Searle, 1986]. [34] One of the most striking aspects of the Himalayas is the lateral continuity of the mountain belt’s major tectonic elements. Since Medlicott and Blanford [1879], Heim and

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Gansser [1939], and Gansser [1964] the Himalayas are classically divided into four zones that can be followed for more than 2500 km along the length of the belt (Figures 5 and 6): [35] 1. The Sub-Himalaya form the foothills of the Himalayan Range and are dominantly composed of Miocene to Pleistocene molassic sediments derived from erosion of the active mountain belt. The molasse map units (Murree and Siwalik formations), as well as underlying Paleocene sandstones and marine limestones, are internally folded and imbricated. The Main Frontal Thrust (MFT, Figure 5) is the presently active southernmost thrust of the Himalaya that cuts the Quaternary alluvium deposited by the rivers (Ganges, Indus, Brahmaputra) coming from the collisional mountain ranges to the north. [36] 2. The Lesser Himalaya, are mainly formed by Neoproterozoic clastic and carbonate units, Paleozoic limestone, shale and sandstone, and Cambro-Ordovician granites. The low-grade sedimentary rocks are thrust over the Siwalik molasse basin along the Main Boundary Thrust (MBT, Figure 5). Lesser Himalaya rocks also appear in tectonic windows (e.g., Kishtwar and Kulu windows, Figure 5) within the High Himalaya domain (below). [37] 3. The Greater Himalaya (or High Himalaya) Sequence (GHS) form the backbone of the Himalayan Orogen and include most of the area of highest topographic relief. Protolith rocks include Neoproterozoic and early Paleozoic dominantly pelitic sedimentary rocks, circa 500 Ma Cambro-Ordovician orthogneisses, and Paleozoic – early Mesozoic sedimentary rocks of the North Indian passive margin sequence. The entire GHS has been affected by Tertiary Himalayan regional metamorphism up to kyanite and sillimannite-grade, which peaked during the period 32 –20 Ma. Widespread anatexis was responsible for the formation of the tourmaline-bearing Himalayan leucogranites and the melting event may have triggered the ductile extrusion of a hot midcrustal channel [e.g., Beaumont et al., 2001; Grujic et al., 2002; Jamieson et al., 2004; Searle and Szulc, 2005] bordered by a crustal-scale thrust fault, the Main Central Thrust (MCT, Figure 5), at the base and a low-angle normal fault, the South Tibetan Detachment (STD, Figure 5) (also known as the North Himalayan Normal Fault), along the top of the extruding channel. [38] 4. The Tibetan-Tethys Zone (or Tethys Himalaya), forms the mainly sedimentary upper crust of the Indian plate passive continental margin with some ophiolitic thrust sheets rarely preserved at the highest structural levels. An almost complete stratigraphic record ranging from Neoproterozoic to Eocene is preserved within the sedimentary rocks of the Tibetan-Tethys Zone. The transition between the generally low-grade sedimentary units of the Tethys Himalaya and the underlying low- to high-grade rocks of the Greater Himalayan Sequence is marked by the South Tibetan Detachment. Intense folding and thrusting across the entire Tethyan zone has resulted in a minimum crustal shortening of 150 km. 3.1.3. Indus-Yarlung Tsangpo Suture and Tethyan Ophiolites [39] The Indus-Yarlung Tsangpo (or simply Indus) Suture Zone (Figures 5 and 6) marks the boundary of the Indian

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plate versus Asian plate supracrustal units and the site of collision between the Indian passive continental margin with the Asian Andean-type destructive plate margin. The suture zone is marked by a discontinuous line of Tethyan ophiolites that mostly show suprasubduction zone affinity, deep-sea sedimentary rocks, rare blueschists, me´langes, and postcollision conglomerates and molasse sediments. [40] Some ophiolite complexes have been obducted southward onto the passive margin sedimentary rocks of the northern Indian plate. The largest of these is the Spongtang Ophiolite of Ladakh (S, Figure 5), which has a complete lherzolite-harzburgite-dunite upper mantle sequence and a crustal sequence comprising gabbroic and ultramafic cumulates, isotropic gabbros, sheeted dikes and pillow lavas with radiolarian cherts and deep-water pelagic sediments. The ophiolite tectonically overlies an accretionary complex composed of Permian to Cretaceous Tethyan deep-water sediments, ocean island-type alkali basalts, and me´lange [Corfield et al., 1999, 2001]. U-Pb zircon dating has shown that the Spontang ophiolite is a Jurassic (177±1 Ma) mid-ocean ridge basalt ophiolite overlain by a late Cretaceous andesite-dacite arc (88±5 Ma) [Pedersen et al., 2001]. The age of obduction is constrained as beginning during the late Cretaceous. It may have continued through to the Paleocene but was certainly completed before the closure of Tethys (50.6 Ma) and the beginning of crustal thickening and metamorphism along the Himalaya. 3.1.4. Kohistan Arc Terrane [41] In the western Himalaya of northern Pakistan, the Kohistan terrane (Figure 5) is sandwiched between the lower Indian plate (Himalaya) to the south and the Asian upper plate (Karakoram terrane) to the north. It is widely regarded as representing a large-scale intraoceanic island arc complex, which is obliquely exposed from a deep crustal root level in the south to a composite volcanic edifice level in the north [Treloar et al., 1996; Khan et al., 1997; Searle et al., 1999a]. The intraoceanic Kohistan island arc was a Cretaceous volcanic arc within Tethys, which first collided with Asia along the Shyok (or northern) Suture Zone (SS, Figure 5) during the latest Cretaceous, and later collided with India along the Indus suture zone (or Main Mantle Thrust; MMT, Figure 5) during the early Eocene [Searle et al., 1987, 1999a]. The Kohistan arc terrane thus evolved from an intraoceanic volcanic arc into an Andean-type magmatic arc (part of the Trans-Himalayan batholith) dominated by intrusion of large-scale granodioritic magmas following collision with Asia [Petterson and Windley, 1991]. 3.1.5. Tibetan Plateau [42] The Tibetan Plateau comprises a number of upper plate (Asian) crustal blocks, which are separated by major suture zones and terrane boundaries (Figure 7) that predate the India-Asia collision itself. These upper plate Asian accretionary suture zones are progressively younger from north to south and include the Kun Lun (late Triassic?), the Jinsha and Bangong-Nujiang (late Cretaceous), and finally the Indus-Yarlung Tsangpo (Eocene) collisional suture to the south. The age of the intervening crustal bocks (Figure 7) decreases from Precambrian to Triassic south

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of the Tsaidam Basin, Precambrian to Cretaceous for the Qiangtang block, to Precambrian to Eocene for the Lhasa block clearly documenting a north to south trend in the crustal accretion and growth of the upper plate. 3.1.6. Karakoram Terrane [43] The Karakoram terrane in northern Pakistan (Figures 5 and 6) is equivalent to the Lhasa block of south Tibet (Figure 7), and together they formed the southern margin of the upper Asian plate prior to first accretion of the Kohistan island arc terrane during the late Cretaceous and then collision with India in the early Eocene [Searle et al., 1999a]. Although the Karakoram and the Tibetan Plateau have a similar average elevation of around 5.02 km above mean sea level, the geological exposures are very different [Searle, 1991]. Tibet is a high, relatively flat uplifted plateau which has not been deeply eroded and only reveals upper crustal sedimentary and volcanic rocks, whereas the Karakoram terrane has large topographic differences (7 –8.7 km mountain tops to 2.5– 3 km valleys) and has undergone enormous amounts of deformation and erosion (up to 35 km of erosion since 35 Myr ago). 3.2. Structural Evolution of the Lower Plate [44] Southward verging thrusting is present across most of the Himalaya with large-scale thrusts generally propagating structurally down section with time, from the 50– 40 Ma deformation along the Indus-Yarlung Tsangpo suture to the mid-Miocene deformation in the Greater Himalaya to the younger Tertiary thrusting in the Lesser Himalaya (Figure 6). Several phases of thrusting and folding occur in each zone with late Tertiary north vergent backthrusting along the Indus-Yarlung Tsangpo suture and northern part of the North Himalaya (Figure 6). Total crustal shortening across the Himalaya could be up to 1000 km. Approximately 150 –300 km of shortening has been documented across the folded and thrusted Tethyan Himalaya in Ladakh [Corfield and Searle, 2000]. Amounts of shortening in the high-grade metamorphic rocks of the Greater Himalaya are difficult to ascertain due to the high degree of ductile strain, internal partial melting and crustal anatexis. [45] The structural evolution of the lower plate is dominated during the Miocene by the ductile flow of an Indian plate midcrustal channel structure from beneath southern Tibet south to the Greater Himalaya domain [Beaumont et al., 2001; Vannay and Grasemann, 2001; Grujic et al., 2002; Searle et al., 2003; Jamieson et al., 2004; Searle and Szulc, 2005]. The base of the Greater Himalayan Sequence crustal channel is the Main Central Thrust (MCT) zone, a crustal-scale low-angle thrust fault characterized by a zone of inverted and telescoped metamorphic isograds, and a ductile shear zone above, propagating down section with time to a brittle thrust fault [Stephenson et al., 2000, 2001]. The MCT acted as the lower boundary of the southward extruding layer of Indian middle crust and was moving synchronously with the South Tibetan Detachment (STD) low-angle normal fault at the top of the channel [Searle and Rex, 1989; Hodges et al., 1993, 1996]. [46] The South Tibetan Detachment (STD) system of low-angle normal faults bounds the extruding midcrustal

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Figure 7. Map of the Tibetan Plateau showing the approximate boundaries of the major suture zones, crustal blocks and strike-slip faults [after Searle, 2006]. Shaded area in the south is the Indian plate rocks of the Himalaya. Abbreviations are MBT, Main Boundary Thrust; MCT, Main Central Thrust; SS, Shyok suture; STD, South Tibetan Detachment. channel on the north (top) side. The STD is characterized by a right way up, but telescoped, series of metamorphic isograds from sillimanite + K-feldspar down to garnetbiotite. The shear zone propagates up section with time, from a ductile shear zone to a low-angle (north-side-down) brittle normal fault. Footwall/hanging wall offsets along the STD may be 100 km or more in the area of Mt. Everest [Searle et al., 2003]. 3.3. Metamorphic Evolution of the Lower Plate [47] Three phases of metamorphism characterize the thermal evolution of the HKTO lower plate. The earliest metamorphism is recorded by coesite-bearing eclogites along the northern margin of the Indian plate in the Kaghan area of northern Pakistan [O’Brien et al., 2001; Ernst, 2005] and in the Tso Morari dome in Ladakh [Sachan et al., 2001]. These rocks have basaltic protoliths within late Paleozoic –early Mesozoic carbonates that were subducted to depths of 100 km at 27– 29 kbar and 700– 750°C. Timing of peak ultrahigh-pressure (UHP) metamorphism at 49 –46 Ma for Kaghan and 53 Ma for Tso Morari [Leech et al., 2005] is almost the same as the timing of initial continent-continent collision (50.6 Ma), so it is debated as to whether these eclogites resulted from the final stage of

the ophiolite obduction process, as for example in Oman [Searle et al., 2004], or the initial stage of the continentcontinent collision process. [48] The second phase of metamorphism recorded by cover units of the lower plate is the main Himalayan (M1) metamorphic event, which characterizes the Greater Himalaya domain. M1 is a regional, Barrovian facies, kyanitesillimanite-grade, medium-temperature and medium-pressure metamorphism. Early kyanite-grade metamorphism peaked at 550– 680°C and 10– 12 kbar, corresponding to depths of 36– 45 km [Searle et al., 1992b, 1999b; Walker et al., 2001]. Timing is constrained as circa 35 –32 Ma from U-Pb dating of monazites and Sm-Nd garnet ages [Vance and Harris, 1999; Walker et al., 1999; Simpson et al., 2000]. [49] Subsequent (M2) sillimanite-grade metamorphism peaked at 650– 770°C and 3.7 –4.5 kbar, corresponding to depths of 14– 18 km [Searle et al., 1992b, 1999b, 2003]. Timing of peak sillimanite-grade metamorphism from U-Pb dating of monazites growing in equilibrium with garnetsillimanite ± cordierite was between 24 and 17 Ma [Noble and Searle, 1995; Simpson et al., 2000]. Sillimanite metamorphism was concomitant with partial melting that produced widespread migmatites within the Greater Himalaya

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Sequence and the anatectic event resulting in crystallization of the Himalayan leucogranites (below). [50] The youngest metamorphism within the orogenic lower plate is recorded in the cores of the Nanga Parbat and Namche Barwa syntaxes (Figure 1b), where Plio-Pleistocene sillimanite-cordierite – K-feldspar-grade gneisses show partial melting and formation of garnettourmaline leucogranites less than 5 Myr old and cordierite-bearing melt seams that crystallized from a 5 kbar melt, less than 1 Myr ago [Zeitler et al., 1993]. 3.4. Lower Plate Magmatism [51] Himalayan leucogranites contain quartz, K-feldspar and plagioclase with varying amounts of tourmaline, garnet, muscovite, and biotite, and minor amounts of sillimanite, cordierite, and apatite. The leucogranites are entirely derived from melted crustal pelitic rocks, probably Neoproterozoic shales (Haimanta Group) and result from dehydration melting of muscovite in the presence of fluids [Harris and Massey, 1994; Prince et al., 2001]. The leucogranites were emplaced between circa 25 and 16 Ma [see Searle et al., 1999b, and references therein] as sills along the top of the Greater Himalayan Sequence midcrustal channel and beneath the passive roof fault zone of the STD normal fault [e.g., Searle et al., 1993, 1997, 2003]. 3.5. Upper Plate (Precollisional) Andean-Type Magmatism [52] The Karakoram batholith includes a series of precollisional granodioritic-tonalitic granitoids subsequently deformed and metamorphosed to amphibolite-facies orthogneisses (e.g., K2 gneiss, Muztagh Tower gneiss, Hushe gneiss) [Searle et al., 1989, 1990; Crawford and Searle, 1992], as well as a massive precollisional I-type biotite and hornblende quartz diorite, granodiorite, and granite batholith dated as mid-Cretaceous (Hunza Plutonic Unit) [Fraser et al., 2001]. Specific age determinations on the orthogneisses range from circa 150 to 95 Ma [Treloar et al., 1989; Searle et al., 1990; Searle, 1991; Fraser et al., 2001; Hildebrand et al., 2001] and predate both the KohistanAsia and India-Asia accretion/collision events. All the precollisional magmatic units display similar chemical and isotopic characteristics, which suggest a uniform source within the mantle wedge above the subducting slab. Taken as a set, these Andean-type plutonic units bear witness to the long-lived nature of the southern upper plate margin of Asia, with most magmatic units likely formed above precollisional northward directed subduction zones from the Late Jurassic to the Late Cretaceous prior to accretion of the Kohistan Arc at circa 85– 75 Ma [Searle et al., 1999a]. [53] The Trans-Himalayan batholith [Singh and Jain, 2003] lies immediately north of the Indus Tsangpo Suture Zone (Figure 5) and occurs as a long linear belt of granitegranodiorite-diorite plutons, extending about 2500 km in strike length and 20– 80 km across strike. The batholith is interpreted as resulting from continued Andean-type magmatism along the southern margin of Eurasia, following accretion of the Kohistan island arc, and related to continued subduction of north dipping Tethyan oceanic crust [Raz

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and Honeger, 1989]. The Trans-Himalayan plutons predate the India-Asia collision with available U-Pb zircon ages on the Ladakh-Gangdese batholith ranging from 102±2 Ma to 49.8±0.8 Ma [Scharer et al., 1984; Weinberg and Dunlap, 2000]. 3.6. Upper Plate (Postcollisional) Crustal Magmatism [54] A large postcollisional S-type pluton, the Baltoro plutonic unit in the eastern Karakoram Range comprises biotite monzogranite to garnet-muscovite-biotite leucogranite. The pluton is dated by U-Pb method at 21±0.5 Ma and interpreted as a product of crustal melting with some input of heat from the upper mantle, based on the simultaneous intrusion of minor lamporphyric dikes across the northern Karakoram and western Tibet [Parrish and Tirrul, 1989; Searle et al., 1989, 1992a]. Leucogranites from the broader Hindu Kush and Karakoram mountain region range in age from 48.5±0.8 Ma to 9.3±0.1 Ma [Fraser et al., 2001; Hildebrand et al., 2001]. 3.7. Upper Plate Metamorphism [55] Multiple episodes of high-grade regional Barrovian metamorphism occur along the southern Karakoram and Hindu Kush ranges, spanning circa 130 Ma [Searle and Tirrul, 1991; Fraser et al., 2001; Hildebrand et al., 2001]. U-Pb monazite ages from staurolite-grade schists and sillimanite-grade gneisses predating both the Kohistan-Asia and India-Asia collisions peak at circa 135– 126, 106 –102, and 82.9±6.1 –61.9±4.7 Ma [Fraser et al., 2001; Hildebrand et al., 2001]. Metamorphism postdating the two collision events is dated at 44, 28, 16, and 5 Ma [Fraser et al., 2001]. The youngest dated metamorphism is a high-temperature, sillimanite-grade partial melting event in the core of a metamorphic core complex along the southern Karakoram (Dassu gneiss at 5.4±0.2 Ma [Fraser et al., 2001]) a metamorphism that is probably still ongoing today at depth. 3.8. Syncollisional Sedimentation [56] Flexure of the Indian plate as it underthrusts the Himalaya resulted in the development of an extensive foreland basin that has accumulated debris eroded from the rising Himalaya since the mid-Tertiary. The Siwalik molasse basin has a linear depocenter immediately south of the active Main Boundary Thrust (Figure 5). Sediments are transported along river systems into the foreland basin and the Ganges-Bhramaputra rivers drain the basin out to the Bay of Bengal. Sediments eroded from the Himalaya have been deposited in the Bengal fan as far south as offshore Sri Lanka, 1800 km south the Himalayan mountain front [Me´tivier et al., 1999].

4. Tectonothermal Events in THO and HKTO [57] Both THO and the HKTO show protracted precollisional and postcollisional deformation, magmatic and metamorphic histories, which can be effectively considered and compared within a broad collisional upper plate versus lower plate crustal context. In sections 4.1 and 4.2, we

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Figure 8. Summary of available U-Pb ages for the main magmatic, deformation, and metamorphic events within the upper (Churchill) plate Meta Incognita microcontinent (Baffin Island), Narsajuaq arc (Baffin Island and Quebec), and lower (Superior) plate (Baffin Island and Quebec) of Trans-Hudson Orogen. See text for sources of data and discussion. Abbreviations are BaS, Baffin suture; BeS, Bergeron suture; SRS, Soper River suture. review and illustrate the timing constraints for tectonothermal events in THO (Figure 8) and the HKTO (Figure 9) with the aim of comparing the type and duration of upper versus lower plate events in both orogens (Figure 10). 4.1. Timing of Deformation, Magmatism, and Metamorphism in THO [58] In the THO all precollisional deformation, magmatic, and metamorphic events are restricted to the upper plate, as

illustrated for the Churchill plate (Meta Incognita microcontinent) and Narsajuaq arc domains of the orogen on Figure 8. On Baffin Island, early accretion events within the Churchill plate (Table 1) are constrained at 1883±5 – 1865+4/ 2 Ma (D0 accretion of the Meta Incognita microcontinent to the southern Rae craton and closure of the Baffin suture) and circa 1845±2 – 1842+5/ 3 Ma (D1a accretion of the intraoceanic Narsajuaq arc to the southern margin of the Meta Incognita microcontinent and closure of the Soper River suture). Periods of precollision Andean-type magmatism

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Figure 9. Summary of available U-Pb and Th/Pb ages for the main magmatic and metamorphic events documented in the Karakoram – Hindu Kush terrane of the upper (Asian) plate, and age ranges for the thermal, magmatic, and deformation events in the Kohistan arc and the lower (Indian) plate of the Himalaya-Karakoram-Tibetan Orogen. See text for sources of data and discussion. are bracketed, from north to south, at circa 1.9 Ga (south margin of the Rae craton, not shown on Figure 8), 1865+4/ 2 – 1848±2 Ma (emplacement of the Cumberland batholith, Figure 8), and 1842+5/ 3 – 1820+4/ 3 Ma (Andean phase of Narsajuaq arc, Figure 8). Regional metamorphism in the upper plate (Table 1) peaked at circa 1877–1850 Ma (M 0 low-P crustal metamorphism along the southern margin of the Rae craton, not shown on Figure 8), circa 1849– 1833±2 Ma (M1a regional lower granulite-facies metamorphic event resulting from accretion of Narsajuaq arc and

emplacement of the Cumberland batholith, Figure 8), circa 1833–1829 Ma (M1b thermal reequilibration associated with postaccretion magmatism, Figure 8), and 1836±2 – 1825±3 Ma (M1c granulite-facies metamorphism related to the Andean phase of Narsajuaq arc, Figure 8). [59] In contrast, syncollisional deformation and metamorphism are recorded by both upper and lower plates of THO (Figure 8 and Table 1). Early syncollisional deformation of the Superior craton cover units (lower plate) is imprecisely bracketed between circa 1830 and 1815 Ma (D2a piggyback

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Figure 10. Comparative time chart for the Trans-Hudson and Himalaya-Karakoram-Tibetan orogens recalibrated to the onset of continent-continent collision (circa 1830 Ma for Trans-Hudson Orogen and circa 50 Ma for the Himalaya-Karakoram-Tibetan Orogen). Age ranges of principal deformation (D) and magmatic and metamorphic (M) events for the upper and lower plates are shown. See text for sources of data and discussion.

sequence thrusting) followed by out-of-sequence thrusting (D2b) at 1815±4 – 1795±2 Ma. Regional, Barrovian-facies (M2a) metamorphism of the lower plate is bracketed between 1820+4/ 3 and 1815±4 Ma. High-T (M2b) reequilibration of exposed basement units occurred between circa 1814 and 1789 Ma, and M2b reequilibration of cover units is documented between 1815±4 and 1785±1 Ma. Within the

upper plate, D2 collisional deformation is constrained between 1820+4/ 3 and 1805±2 Ma and associated M2 metamorphism is constrained between circa 1820±1 and 1808±3 Ma. [60] Late to postcollisional leucogranite dikes, anatectic granitic plutons, and retrograde M3 metamorphism in TransHudson Orogen is bracketed between 1795±2 and 1742.2±1.3

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Ma for the lower Superior plate, 1805±2 and 1742.2±1.3 Ma for Narsajuaq arc, and 1797±2 and 1760±2 Ma for the Meta Incognita microcontinent (Churchill plate) (Figure 8). 4.2. Timing of Deformation, Magmatism, and Metamorphism in HKTO [61] In the HKTO, the bulk of the precollisional deformation, magmatic and metamorphic events are restricted to the upper Asian plate, as illustrated for the Karakoram–Hindu Kush domain on Figure 9. Periods of precollision Andean-type magmatism are bracketed between circa 150 and 95 Ma (Hushe gneiss, K2 gneiss, Tirich Mir pegmatite and granite, Hunza Plutonic Unit) and predate both the Kohistan-Asia and India-Asia accretion/collision events. Regional metamorphism peaked at circa 135 – 126 Ma and 106 – 102 Ma (garnetstaurolite-grade, eastern Hindu Kush), and 82.9±6.1 – 61.9±4.7 Ma (sillimanite–K-feldspar migmatite, Karakoram). [62] In contrast, syncollisional to late collisional deformation and metamorphism events are recorded in both the upper plate Karakoram – Hindu Kush domain and the lower Indian plate (Figure 9). Earliest UHP coesite eclogite metamorphism of the lower plate is bracketed between 53 and 46 Ma and follows deep northward subduction of thinned Indian continental crust. Regional, Barrovian facies (M1) kyanite-grade metamorphism characterizes the early evolution of the Greater Himalaya, is bracketed between 35 and 32 Ma, and occurs in response to tectonic thickening of Indian supracrustal units during the early stage of continentcontinent collision. Subsequent, widespread high-T but lowP sillimanite-grade metamorphism (M2) occurred between 24 and 17 Ma, again within the Greater Himalayan domain, and is associated with partial melting and formation of Himalayan leucogranites (crystallization ages of circa 25– 16 Ma). The very shallow depths of the M2 metamorphism and melting (3.7 – 4.5 kbar) are consistent with a model of mid-Miocene ductile flow of an Indian plate midcrustal channel southward from beneath southern Tibet to the Greater Himalaya. The orogenic channel is bounded by two crustal-scale shear zones, the MCT, with its zone of inverted metamorphic isograds along the base, and the STD system of low-angle normal faults along the top. [63] Large postcollisional S-type plutons (e.g., Baltoro granite) and leucogranites were emplaced into the postcollisional upper plate, with the leucogranites in the Hindu Kush and Karakoram mountain region ranging in age between 48.5±0.8 and 9.3±0.1 Ma. Sillimanite, kyanite and staurolite-grade metamorphism is dated at 44, 28, 16, and 5 Ma (Figure 9). [64] Possibly the youngest metamorphism within the collisional orogen is recorded in the cores of the Nanga Parbat and Namche Barwa syntaxes (Figure 1b), where sillimanitecordierite-K-feldspar-grade gneisses show partial melting and formation of leucogranites less than 5 Myr old and cordierite-bearing melt seams less than 1 Myr old (Figure 9).

5. Summary and Comparison of THO and HKTO [65] We are struck by the first-order similarities that can be documented for the structural and thermal evolution of

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the lower and upper orogenic plates of the eastern THO and the HKTO (see Figure 10). We would suggest that as established, these interorogenic similarities point to firstorder tectonic processes and boundary conditions that are important and common to at least these two mountain belts (and perhaps most large collisional orogens), effectively bridging (or bracketing) no less than 1.88 Ga (43%) of dated Earth history [Wyche et al., 2004]. While we recognize that no two collisional orogens are identical, that most mountain belts are unique, and that a given collisional orogen can exhibit important structural and petrological changes along strike, we would suggest that the similarities in length and timescales, as well as principal tectonic components outlined below for the THO and HKTO point to a broad similar tectonic context for these two orogens during the Paleoproterozoic and the Cenozoic. 5.1. Length-Scale, Timescale, and Principal Tectonic Components [66] The first similarities that can be drawn are those of orogenic length and timescales. Figure 1a shows that THO is no less than 4600 km in strike length (between Mesoproterozoic cutoffs), with 2500 km separating the western Manitoba promontory from the eastern Quebec promontory as measured along the northern Superior craton boundary. The greater Alpine-Himalayan system is 6600 km in strike length with 2500 km separating the Nanga Parbat syntaxis in the far northwest from the Namche Barwa syntaxis in the far northeast (Figure 1b). Promontories and syntaxes mark the corners of the indenting lower plate in both orogens. [67] In order to compare the timing, duration, and lower versus upper plate setting of tectonothermal events in both THO and HKTO, the age data shown on Figure 8 for THO and Figure 9 for HKTO have been integrated into a common temporal framework in Figure 10. Age constraints were converted into millions of years before the onset of collision (BOC) or millions of years after the onset of collision (AOC) for both orogens. Approximate ages of 1830 Ma and 50 Ma were used for the ‘‘onset of collision’’ datum in the THO and in the HKTO, respectively. Figure 10 thus shows the timing and duration of principal tectonothermal events in the upper and lower orogenic plates as recalibrated to the onset of collision in the THO (black bars) and the HKTO (gray bars). The constraints/duration of principal accretion events predating the onset of continentcontinent collision are shown as horizontal shaded domains for both orogens. [68] As shown on Figure 10, convergent, precollisional, upper plate deformation (crustal accretion), magmatism and metamorphism spans a period of over circa 50 Ma in the Baffin Island – northern Quebec segment of THO. Furthermore, syncollisional to postcollisional magmatic and thermal events are documented for another circa 90 Ma following the onset of the Churchill-Superior collision at circa 1830 Ma. In the HTKO, the magmatic units document crustal accretion and growth of the upper plate over a 100 Myr time period preceding the onset of the India-

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Asia collision at circa 50 Ma. Subsequent upper and lower plate magmatism and metamorphism is documented to less than 1 Myr ago and is almost certainly continuing today (Figure 10). [69] The coherent, indenting, lower plate in THO is the Superior craton, which comprises a precollisional rigid Eoarchean to Neoarchean crystalline basement overlain along the northern side (present-day coordinates) by continental margin cover strata of Paleoproterozoic age. In the HKTO, the lower Indian plate also includes a precollisional, rigid, crystalline basement (Mesoarchean to Mesoproterozoic) overlain along the north side by continental margin sedimentary rocks that range in age from Neoproterozoic to Paleocene. [70] In contrast, the upper plate of THO is the segmented, multiply reworked Churchill plate, which at the onset of collision with the Superior plate consisted of a collage of cratons and accreted terranes bound by dominantly Paleoproterozoic (pre-1.83 Ga) crustal structures (Figure 2). In the HKTO, prior to collision with the Indian plate, the Asian plate was also a segmented, nonrigid upper plate comprising a collage of crustal blocks, microplates, and terranes bound by crustal structures ranging in age from late Triassic (?) to Eocene in the Tibetan Plateau area (Figure 7). 5.2. Structural Evolution of the Lower Plate in THO and HKTO [71] The onset of collision in THO, beginning at circa 1830 Ma (Figures 8 and 10), led to the formation of a lower plate foreland thrust fold belt, of which the erosional remnant is the Cape Smith belt of northern Quebec (Figure 3). The Paleoproterozoic thrust-fold belt is characterized by early, piggyback sequence, southward verging thrusting (D2a) followed by out-of-sequence thrusting (D2b, also southward verging) (Figure 8). The deep erosion levels exposed in THO clearly show that early thinskinned deformation gave way to thick-skinned (basement involved) thrusting as the collision progressed. Minimum crustal shortening accommodated within the foreland thrust-fold belt, as preserved at the present erosion surface, is on the order of 400 km. Erosion of the imbricated continental margin strata led to deposition of molasse units, which are preserved up to 500 km south of the Cape Smith belt. [72] In the HKTO, the onset of collision at circa 50 Ma (Figures 9 and 10) led to southward verging thrusting across most of the Himalaya with large-scale thrusts in general propagating structurally down section with time (Eocene deformation along the Indus-Yarlung Tsangpo suture, midMiocene deformation in the Greater Himalaya, Tertiary thrusting in the Lesser Himalaya). Late Tertiary north vergent backthrusting is documented along the Indus suture and northern part of the North Himalaya. Considerable outof-sequence thrusting and thrust reactivation also occurs throughout the Himalaya. Total crustal shortening across the Himalaya could be up to 1000 km. Sediments eroded from the Himalaya are found up to 1800 km south the mountain front and the thickness of sediments derived from

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the Himalaya reaches a maximum of 20 km off the Ganges delta [Curray, 2005]. 5.3. Thermal Evolution of the Lower Plate in THO and HKTO [73] In THO (Figure 10), subduction of the Archean crystalline basement and tectonic thickening of the Paleoproterozoic cover led to (1) regional, Barrovian facies, kyanite-sillimanite metamorphism (M2a) 10– 15 Myr after the onset of collision (AOC), (2) extensive high-T reequilibration of basement and cover units (M2b) at upper amphibolite facies conditions 15 – 45 Ma AOC, and (3) postcollisional (M3) anatexis and formation of leucogranites 35– 70 Ma AOC. [74] In the HKTO (Figure 10), initial continent-continent collision may have resulted in coesite-eclogite metamorphism circa 3 –4 Ma immediately preceding and following the onset of collision. Within the Greater Himalaya domain, a regional, Barrovian facies, kyanite-sillimanite-grade metamorphism (M1) peaked circa 15– 18 Ma AOC. A subsequent (M 2 ) low-P sillimanite-grade metamorphism concomitant with partial melting peaked 26– 33 Ma AOC. The leucogranites were emplaced between circa 25 and 34 Ma AOC. The youngest metamorphism within the orogenic lower plate is recorded in the cores of the Nanga Parbat and Namche Barwa syntaxes with high-T anatexis dated at 45 Ma and 49 Ma AOC 5.4. Structural, Magmatic, and Thermal Evolution of the Upper Plate in THO and HKTO [75] Both the THO and the HKTO rock records document a precollisional, north-to-south trend in the crustal accretion and growth of the upper plate (Figure 10). In THO early accretion events are dated, from north to south, at 53– 35 Ma BOC (accretion of Meta Incognita microcontinent and closure of the Baffin suture), circa 15 Ma BOC (accretion of Narsajuaq arc and closure of the Soper River suture), and 0 Ma (onset of Churchill/Superior continent-continent collision). Periods of Andean-type magmatism are bracketed at circa 70 Ma BOC (not shown on Figure 10), 35 –18 Ma BOC (Cumberland batholith), and 12 Ma BOC to 10 Ma AOC (younger suite of Narsajuaq arc). Regional metamorphism peaked at circa 47– 20 Ma BOC (M0), circa 19– 3 Ma BOC (M1a), circa 3 Ma BOC to 1 Ma AOC (M1b), and 6 Ma BOC to 5 Ma AOC (M1c). [76] In the HKTO, a north to south trend in the crustal accretion and growth of the upper plate is indicated by the age of the crustal sutures (circa 150, 50– 25, and 0 Ma BOC) contained within the Tibetan Plateau (Figure 7), and by the periods of precollision Andean-type magmatism dated at circa 100 Ma BOC for the Hushe gneiss (not shown on Figure 10), and including the Ladakh-Gangdese granites of the Trans-Himalayan batholith (52 Ma BOC to 1 Ma AOC) and the arc volcanics of the Kohistan terrane (28 – 4 Ma BOC) (Figure 10). Consequent regional metamorphism is constrained to circa 85– 76 Ma BOC (not shown on Figure 10), 56– 52 Ma BOC (Barrovian metamorphism), and 33– 12 Ma BOC (sillimanite migmatite) in

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the Hindu Kush and Karakoram areas of the upper plate (Figure 10). [77] Two-mica, garnet-tourmaline leucogranite dikes, and large anatectic plutons were emplaced in the upper plate 25 –88 Ma AOC in THO and 1.5– 41 Ma AOC in the HKTO. Upper plate crustal metamorphism is constrained between 10– 22 Ma AOC (M2) and 33– 45 Ma AOC (M3) in THO and at 6, 22, 34, and 45 Ma AOC in the HKTO. 5.5. Differences in the Orogenic Record [78] While both THO and HKTO show broadly similar structural and thermal responses to the underthrusting of lower plate material (first oceanic and then continental), which is followed by the consequent tectonic thickening of continental margin units (Figures 8 and 9), differences in the rock record are recognized and discussed below: [79] 1. Early eclogite facies metamorphism is not documented in the eastern segment of THO as it is in the Kaghan area of northern Pakistan and in the Tso Morari dome in Ladakh. While notable, given the size and relatively well studied nature of the HKTO, these coesite-bearing eclogite localities are few and far between. Indeed it is quite likely that due to thorough and complete retrograde reaction, many older collisional mountain belts (those in the Paleoproterozoic included), which may have been subjected to early UHP conditions of metamorphism simply fail to retain any evidence of such conditions. In contrast, most collisional orogenic belts of Phanerozoic age seem to preserve some evidence of UHP conditions, at least locally [cf. Ernst, 2005]. [80] 2. Although the thermal evolution of the lower plate of both THO and the HKTO is characterized by the development of regional Barrovian facies mineral zones in response to the construction of a collisional mountain belt (M2a in THO; M1 in the HKTO), the subsequent high-T culmination (M2b in THO; M2 in the HKTO) is diagnostically low-P in the HKTO versus medium- to high-P in THO. This observation, coupled with the absence of MCT zone-type inverted metamorphic isograds and the absence of structures analogous to the STD zone, leads us to suggest that in THO a midcrustal channel flow structure was not involved in the thermal evolution of the lower plate south of the Bergeron suture. This may, in large part, be a reflection of the lack of voluminous anatectic melt generation in THO during M2b and a consequent insufficient amount of melt weakening [Jamieson et al., 2004] having occurred within the lower plate of THO versus that of the HKTO. We would suggest that in THO this is a consequence of the paucity of minimum-melt pelites contained in the Paleoproterozoic continental margin units of the northern Superior craton. [81] 3. The thermal evolution of the upper plate of both THO and the HKTO is punctuated by the emplacement of Andean margin-type magmatic suites and the consequent development of regional metamorphic domains. An apparent difference is the granulite-facies conditions attained in THO [St-Onge et al., 2006] versus the lower staurolitekyanite-sillimanite metamorphic temperatures prevalent in the upper plate Karakoram-Hindu Kush domain of THKO [Fraser et al., 2001; Hildebrand et al., 2001]. Considering

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the present extensive exposure of cratonic basement rocks in the upper plate of THO (Wyoming, Hearne, Rae, and North Atlantic cratons, Figure 1a) versus the limited exposure of crystalline basement units in the upper plate of the HKTO (Figure 1b) we would ascribe the apparent difference in attained metamorphic conditions in large part to a difference in the structural level of the present erosion surface in both orogens (deeper erosion levels in THO, shallower erosion levels in the HKTO). [82] 4. Restricted in THO and extensive in the HKTO, the preservation of synorogenic molasse units is also principally a function of the structural level of the erosion surface in a given collisional belt. As highlighted above, the large areas underlain by lower plate crystalline basement external and within the collisional belt in the older orogen (Figure 3) point to relatively deep postcollisional levels of erosion. In contrast, the extent of the Siwalik molasse basin in India points to ongoing flexure of the Indian plate and consequent preservation of accumulating synorogenic sedimentary units.

6. Conclusions [83] We have compared the type and duration of magmatic, deformation, and metamorphic events within the lower and upper collisional plates of two orogens of comparable size, magnitude, and constituent tectonic components, that are separated in time by over 1750 Myr. On the basis of these comparisons, we conclude that the many first-order, structural and thermal similarities that can be documented for the Paleoproterozoic THO and the Cenozoic HKTO support the notion of tectonic uniformitarianism for at least the later half of dated Earth history. Cognizant that collisional orogens can be highly diverse, this nevertheless suggests that the India-Asia collision zone, and the structural and thermal evolution of its Himalayan, Karakoram and Tibetan domains can be viewed as first-order younger analogues for ancient collisional orogenic belts back to 2 Ga. Remaining differences in the accessible orogenic rock record can then possibly be evaluated in terms of contrasting depths of erosion, degree of thermal preservation, and original tectonostratigraphic components, as was done in this paper for the eastern THO. [84] The comparison of the structural and thermal evolution of the eastern THO and the HKTO, and the similarity of plate tectonic processes that are implied, may owe their relevance to the fact that THO postdates the period of geological history when the early Earth is thought by some [cf. Hamilton, 2003] to have sufficiently cooled to allow the horizontal motion of plates to (gradually?) replace assumed vertically dominated processes of a youthful Earth. Hamilton [2003] places this transition at 2.0 Ga by which would be consistent with the observations presented here. However it must be emphasized that our data offer no lower constraint on when such a transition might have occurred. [85] Finally, reliable apparent polar wander paths have not yet been developed for the lower plate Superior craton or the upper Churchill plate in THO to test their relative movement in the critical 1890– 1795 Ma time period. Only

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when such paths become available will it be possible to use paleomagnetism to establish the flight path of the Superior craton within the Himalayan comparative model for THO presented here.

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[86] Acknowledgments. Discussions with David Corrigan and Dave Scott (GSC) are gratefully acknowledged. Helpful and thorough reviews of earlier versions of this paper were provided by Cees van Staal and Eric de Kemp (GSC), Associate Editor Robert Miller, and journal reviewers Paul Hoffman (Harvard University) and Maarten de Wit (University of Cape Town). This is GSC contribution 2005310.

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M. R. St-Onge and N. Wodicka, Geological Survey of Canada, 601 Booth St., Ottawa, Ontario, Canada K1A 0E8. ([email protected]; [email protected]) M. P. Searle, Department of Earth Sciences, University of Oxford, Parks Road, Oxford OX1 3PR, UK. ([email protected])