Journal of Petrology Advance Access originally published online on November 3, 2006
Journal of Petrology 2007 48(2):271-302; doi:10.1093/petrology/egl060
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Polymetamorphic Evolution of the Trans-Hudson Orogen, Baffin Island, Canada: Integration of Petrological, Structural and Geochronological Data
Geological Survey of Canada, 601 Booth Street, Ottawa, Ontario, Canada, K1A 0E8
RECEIVED MAY 26, 2005; ACCEPTED SEPTEMBER 21, 2006
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONGEOLOGICAL BACKGROUNDTECTONIC EVENTS, DEFORMATION...METHODSRESULTSDISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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Supracrustal units metamorphosed at mid-crustal conditions within the Paleoproterozoic Trans-Hudson Orogen are preserved within an obliquely exposed continental collision zone on Baffin Island (Canada). Early granulite-facies assemblages yield thermobarometric data and phase diagram information that define a steep, compressive P–T path segment. These assemblages are bracketed between ca. 1849 and 1835 Ma, and are interpreted to result from (1) heat advection by an 1865 +4/–2 to 1848 ± 2 Ma Andean-type granitic batholith, and (2) a ca. 1845 Ma crustal thickening event associated with accretion of an intra-oceanic arc terrane. A subsequent regional metamorphic event is characterized by the growth of retrograde, upper amphibolite-facies assemblages that define a clockwise, decompressive P–T path. Mineral growth is bracketed between 1820 ± 1 and 1813 ± 2 Ma, and is localized within deformation zones associated with the 1820 +4/–3 to 1795 ± 2 Ma collision of the Rae and Superior cratons. The metamorphic history of Baffin Island supports a progressive change from plate-margin to intraplate processes within an evolving convergent orogen during the Paleoproterozoic that is similar to those documented in younger collisional belts.
KEY WORDS: polymetamorphism; geochronology; Paleoproterozoic; Trans-Hudson Orogen
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL BACKGROUNDTECTONIC EVENTS, DEFORMATION...METHODSRESULTSDISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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Building on the innovative work of England & Thompson (1984
), regional metamorphism in convergent orogens has been shown by coupled thermal–mechanical models to result from the complex interplay of several, generally causative, probably interlinked processes that include subduction-driven convergence and tectonic accretion or collision, crustal thickening and deformation, structural and/or erosional exhumation, underthrusting of materials rich in radiogenic heat-producing elements, and magmatic heat advection (e.g. Jamieson et al., 1996, 1998, 2002; Huerta et al., 1998, 1999; and references therein). Concurrently, field-based petrological–structural–thermochronological studies have successfully documented the fundamental polymetamorphic character of many higher-grade terrains within collisional orogenic belts ranging in age from Eocene to Neoproterozoic (e.g. Jamieson et al., 1995; Vannay & Hodges, 1996; Culshaw et al., 1997; Harrison et al., 1997; Moecher, 1999; Searle et al., 1999, 2003; Walker et al., 1999; Carr et al., 2000; Indares et al., 2000; Wodicka et al., 2000; Fraser et al., 2001; Johnson et al., 2003).
In this paper, we extend the field-based integrated petrological–structural–thermochronological study of convergent orogens to the Paleoproterozoic, and are able to relate key petrological aspects of polymetamorphic mineral assemblages to a well-constrained history of crustal deformation and evolution. The focus of our work is a part of southern Baffin Island (Nunavut, Canada) where an oblique crustal cross-section through a continental collision zone of Paleoproterozoic age is superbly preserved and exposed (St-Onge et al., 2002). The exceptional tectonic context afforded by the geology of southern Baffin Island allows for the internally consistent correlation of deformation fabrics, polymetamorphic assemblages, mineral chemistry, and age data. The correlation of these multiple datasets for southern Baffin Island, as presented in this paper, allows us to propose unambiguous links between thermal events or processes and consequent regional metamorphic associations.
We begin by providing the regional tectonic context for our work and describing the main crustal components of the Quebec–Baffin segment of the Himalayan-scale Trans-Hudson Orogen. The large amount of published work pertaining to this Paleoproterozoic collisional orogen (St-Onge et al., 1992, 1999a, 2001, 2002; Scott et al., 2003; and references therein) provides us with the context within which we can evaluate the significance of the detailed petrological and thermochronological data presented here and which is specific to the generally less well understood orogenic internal zone. Key structural and petrological observations outline the deformation and metamorphic histories of the southern Baffin Island area. Integrated thermobarometric, phase diagram, and thermochronological work documents the pressure–temperature–time–deformation history of the polymetamorphic internal zone of the Trans-Hudson Orogen on Baffin Island. These results distinguish two regionally significant metamorphic events (M1a, M2) and document the thermal evolution of crustal rocks in the hanging wall of a major suture (tectonic zone separating two crustal blocks of contrasting geological history and parentage or provenance) during accretion- and collision-related deformation. We conclude by considering the potential implications of the south Baffin Island results for the thermal evolution of convergent to collisional orogenic belts.
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GEOLOGICAL BACKGROUND |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL BACKGROUNDTECTONIC EVENTS, DEFORMATION...METHODSRESULTSDISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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The Trans-Hudson Orogen (Hoffman, 1988
; Lewry & 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, where it is truncated by the younger Meso- to Neoproterozoic Grenville Orogen (Fig. 1). The Trans-Hudson Orogen separates a partially subducted or underthrust collisional lower plate (the Superior craton) from a collisional upper-plate (assembled in the hanging wall of a subduction zone) collage of Archean crustal blocks that includes the Wyoming craton, the Hearne domain and Rae craton of the northern Churchill plate, and the North Atlantic craton of Labrador and Greenland (Fig. 1). The orogen is over 4600 km in strike-length and, in certain segments, over 800 km in strike-width. Beyond North America it has been correlated (Zhao et al., 2002) with the Nagssugtoqidian Orogen of Greenland, the Kola–Karelia Orogen of the Baltic Shield, and possibly with the Transantarctic Orogen of Eastern Antarctica.
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Recently, Searle & St-Onge (2004
) and St-Onge & Searle (2004) have argued that the tectonic record of the Trans-Hudson Orogen in the late Paleoproterozoic (ca. 1820–1795 Ma) was similar to the record of collision and indentation of India into SE Asia in the Eocene (beginning at ca. 50.6 Ma; Rowley et al., 2004). An overview of the tectonostratigraphic framework of the Quebec–Baffin segment of Trans-Hudson Orogen is given below from south to north (St-Onge et al., 2002; Scott et al., 2003, and references therein), to provide a regional context for the southern Baffin Island study area (Fig. 2).
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Lower-plate cratonic margin
In northern Quebec, the exposed lower-plate Archean Superior craton (Fig. 2) comprises dominantly felsic orthogneisses and plutonic units ranging in age between 3220 + 32/–23 and 2737 ± 2 Ma (Fig. 3; Parrish, 1989
, unpublished data; St-Onge et al., 1992; Scott & St-Onge, 1995). Unconformably overlying the felsic basement is a suite of parautochthonous sedimentary strata, continental tholeiitic flood basalt and rhyolite (Povungnituk Group, Figs 2 and 4a) associated with initial Paleoproterozoic rifting of the northern Superior craton and yielding ages between 2038 + 4/–2 and 1959 + 3·1/–2·7 Ma (Fig. 3; Parrish, 1989; Machado et al., 1993). Overlying the initial rift sedimentary and volcanic rocks is a younger succession of predominantly Mg-rich komatiitic to tholeiitic basalts (Chukotat Group, Figs 2 and 4a) accumulated during renewed rifting along the northern continental margin (St-Onge et al., 2000a) and dated between 1887 + 37/–11 and 1870 ± 4 Ma (Fig. 3; Wodicka et al., 2002a; R. R. Parrish, unpublished data).
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Ophiolitic and magmatic arc assemblages
A major tectonic boundary or crustal suture (Bergeron suture, Figs 2 and 4e; St-Onge et al., 2001
), which accommodated displacements to the south, separates the Superior margin strata (described above) from allochthonous northern crustal elements. Accreted units in the hanging wall of the Bergeron suture include (Figs 2 and 4): (1) an ophiolite (Watts Group) dated to 1998 ± 2 Ma (Fig. 3; Parrish, 1989); (2) a fore-arc clastic sequence (Spartan Group), which is undated; and (3) a magmatic arc terrane (Parent Group and Narsajuaq arc). The arc terrane is characterized by two temporally and petrologically distinct suites: an older intraoceanic suite and a younger Andean margin suite (Dunphy & Ludden, 1998). The older suite includes calc-alkaline diorite–tonalite gneiss, basaltic andesite, and rhyolite, and is dated between 1863 ± 2 and 1845 ± 2 Ma (Fig. 3; St-Onge et al., 1992; Machado et al., 1993; R. R. Parrish, unpublished data). The younger suite comprises crosscutting monzodiorite to granite plutons and is bracketed between 1842 +5/–3 and 1820 +4/–3 Ma (Fig. 3; Parrish, 1989, unpublished data; Machado et al., 1993; Scott, 1997; Scott & Wodicka, 1998).
Microcontinent and granite batholith
On southern Baffin Island, a second south-verging crustal suture (Soper River suture, Figs 2 and 4c; St-Onge et al., 2001) separates the Narsajuaq arc in its footwall from Archean and Paleoproterozoic units of the Meta Incognita microcontinent (St-Onge et al., 2000a) in its hanging wall. The microcontinent comprises (Figs 2 and 4): (1) a clastic–carbonate shelf succession (Lake Harbour Group) and its crystalline basement; (2) an overlying foreland basin succession (Blandford Bay assemblage); (3) an extensive suite of quartz diorite to monzogranitic plutons (Cumberland batholith) that intrude both the platformal and foreland basin rocks and underlying crystalline basement. The age of the shelf succession and foreland basin is bracketed between 1934 ± 2 and ca. 1880 Ma (Fig. 3; Scott, 1997, unpublished data; Scott et al., 2002). Orthogneisses from the stratigraphic basement to the Lake Harbour Group have yielded Archean ages between 3019 ± 5 and 2784 ± 9 Ma (Scott, 1998, 1999) and a Paleoproterozoic age of 1950 + 6/–4 Ma (Scott & Wodicka, 1998), whereas plutons of the Cumberland batholith have yielded ages between 1865 + 4/–2 and 1848 ± 2 Ma (Jackson et al., 1990; Wodicka & Scott, 1997; Scott & Wodicka, 1998; Scott, 1999) (Fig. 3).
Upper-plate cratonic margin
In the central Baffin Island area (top part of island shown in Fig. 2) felsic orthogneisses, metamorphosed supracrustal rocks, and younger felsic plutonic rocks range in age between 2868 + 13/–12 and 2702 ± 3 Ma (Fig. 3; Jackson, 2000; Wodicka et al., 2002b; Bethune & Scammel, 2003; D. J. Scott, unpublished data; N. Wodicka, unpublished data) and belong to the upper-plate Rae craton (Jackson, 1969, 2000; Bethune & Scammell, 2003; Scott et al., 2003). The southern margin of the craton is unconformably overlain by the Paleoproterozoic Piling Group (Scott et al., 2003; and references therein; Figs 2 and 4a), which comprises rift, shelf margin, and foredeep strata ranging in age between 2159 ± 16 Ma and 1897 + 7/–4 Ma (Fig. 3; Wodicka et al., 2002b; N. Wodicka, unpublished data). Various felsic plutonic rocks, ranging in age from 1·90 to 1·85 Ga (Fig. 3) and including the northernmost component of the Cumberland batholith (Wodicka et al., 2002b), intrude the Piling Group. Accretion of the Meta Incognita microcontinent to the southern margin of the Rae craton is proposed by St-Onge et al. (2005a) to have occurred across the Baffin suture (BaS; Fig. 3) between ca. 1880 and 1865 Ma, with the trace of the early suture tentatively situated at the head of Cumberland Sound (Fig. 2).
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TECTONIC EVENTS, DEFORMATION FABRICS, AND METAMORPHIC ASSEMBLAGES, SOUTHERN BAFFIN ISLAND |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL BACKGROUNDTECTONIC EVENTS, DEFORMATION...METHODSRESULTSDISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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Integrated tectonostratigraphic, structural and petrological fieldwork in the southern Baffin Island region has led to the recognition of a number of regional tectonothermal events of Paleoproterozoic age. These are schematically illustrated in Fig. 4 and an overview of the salient features of each event, and consequent development of deformation fabrics and metamorphic assemblages as documented in the study area are presented below.
D0: accretion of Meta Incognita microcontinent
Map-scale D0 tectonic contacts, which imbricate basement and cover units within the study area (Fig. 5), post-date deposition of the supracrustal units of the Lake Harbour Group and pre-date emplacement of the Cumberland batholith (St-Onge et al., 1999b). The repetition of units by D0 structures is suggestive of thrust imbrication, but the direction, geometry, and magnitude of early displacements in southern Baffin Island are not known. The D0 structures are interpreted as resulting from the accretion of the Meta-Incognita microcontinent to the southern Rae margin (Fig. 4a) and closing of the Baffin suture (BaS; Fig. 3). In the study area, mesoscopic and microscopic deformation fabrics and mineral assemblages potentially associated with the D0 structures are not preserved, because of the pervasive recrystallization and foliation development that occurred during the subsequent prograde D1a–M1a event (see below).
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D1a: accretion of Narsajuaq arc
Continued convergence between the southern margin of the accreted Meta Incognita microcontinent and crustal domains to the south led to D1a accretion of the Narsajuaq arc (Fig. 4c) and formation of the Soper River suture (Fig. 2). Closing of the suture (SRS; Fig. 3) is bracketed between 1845 ± 2 Ma, the age of the youngest unit associated with the intra-oceanic phase of Narsajuaq arc (Dunphy & Ludden, 1998
) and 1842 + 5/–3 Ma (Scott, 1997), the age of the oldest plutonic unit of the Andean-margin phase of the Narsajuaq arc.
The Soper River suture marks the structural base of the Meta Incognita microcontinent. Deformation in the suture's footwall includes: (1) structural repetitions and truncations of distinct tectonostratigraphic units yielding an overall SW-verging ramp–flat fault geometry (Scott et al., 1997); (2) development of ribbon mylonites; and (3) transposition of crosscutting intrusive units into parallelism with the suture. Deformation in the hanging wall is both extensive and penetrative, with D1a manifest as a regional syn-metamorphic foliation (S1a; see below) that affects all pre- to syn-accretion units within the microcontinent terrane.
M1a: prograde granulite-facies metamorphism, Meta Incognita microcontinent
The M1a metamorphism of southern Baffin Island is characterized by pelitic and psammitic units of the Lake Harbour Group generally containing subsets of the granulite-facies assemblage garnet–cordierite–sillimanite–biotite–plagioclase–K-feldspar–quartz–melt (here and elsewhere interpreted from granitic leucosome) (Fig. 6a) and mafic units containing subsets of the assemblage garnet–orthopyroxene–clinopyroxene–hornblende–biotite–plagioclase (Fig. 7a).
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At the macroscopic and microscopic scale, the distribution of these assemblages is strongly domainal and associated with the development of a pervasive millimetre- to centimetre-scale compositional foliation (S1a; Fig. 6a). In metapelitic units, S1a is defined by layers of coarse (2 cm x 1 cm) subhedral garnet, abundant (2–3 mm) euhedral sillimanite, (1–2 mm) biotite, and (4 mm x 1·5 mm) anhedral cordierite (plus quartz and plagioclase), alternating with layers of dominantly plagioclase, K-feldspar, and quartz. The S1a foliation is further enhanced by the alignment of biotite, sillimanite, and elongate grains of garnet and cordierite (Fig. 6a). Both garnet and cordierite are strongly poikiloblastic, containing abundant sub-millimetre scale inclusions of sillimanite generally aligned parallel to the dominant S1a foliation. Overall, garnet and sillimanite dominate the S1a foliation, suggesting that these phases were principal products of the M1a event, whereas cordierite, although almost always present, is notably less abundant. In the samples selected for this study, all of these minerals are unaltered and found in mutual contact, thus appearing to be in textural equilibrium, with the exception of S1a cordierite (Fig. 7b). This mineral generally presents a marked recessive or embayed morphology with adjacent garnet, biotite and/or sillimanite within the S1a foliation, which suggests a (late?) stage of disequilibria during the M1a thermal event.
Within the study area, the granulite-facies foliation is generally shallow-dipping (<45°), invariably parallel to lithological contacts between supracrustal units (St-Onge et al., 1999b) and, in metasedimentary units, is interpreted as a metamorphic enhancement of primary bedding (S0). Mafic units are generally massive, although locally S1a is defined by the alternating distribution of dominantly ferromagnesian-rich layers (comprising 1–1·5 cm scale garnet porphyroblasts, and granoblastic millimetre-scale orthopyroxene, clinopyroxene, hornblende, and biotite grains; Fig. 7a) and layers comprising dominantly plagioclase. The alignment of hornblende and biotite also highlights the S1a foliation when present. Garnet porphyroblasts are generally poikiloblastic and contain inclusions of all the other S1a phases (plagioclase, orthopyroxene, clinopyroxene, hornblende, and biotite), suggesting equilibrium growth during M1a.
M1b: high-T thermal perturbation, Meta Incognita microcontinent
Concordant U–Pb ages from metamorphic monazite and zircon in siliciclastic and gneiss samples (Scott & Gauthier, 1996; Wodicka & Scott, 1997; Scott & Wodicka, 1998; this study) suggest that rocks from the study area experienced a thermal perturbation at ca. 1833–1829 Ma. As discussed below, these ages may represent a distinct thermal event (here termed M1b) related to emplacement of ca. 1834–1831 Ma felsic dykes and plugs. However, other than the growth of monazite and zircon, as well as orthopyroxene in some of the more mafic plugs, no other metamorphic mineral phase or regional foliation development is documented for M1b. Whether M1b is a distinct thermal pulse or part of a protracted M1 metamorphic event is discussed further after presentation of the geochronological data from the southern Baffin Island study area.
D2: collision of the northern (Churchill) plate and the Superior craton
The D2 collision of the northern (Churchill) plate (i.e. Rae craton plus accreted microcontinent, arc, and ophiolitic units described above) with the Superior craton involved: (1) closure of the Bergeron suture (Figs 2 and 4e) and formation of a foreland thrust–fold belt within the lower-plate Superior domain (Fig. 4e; Lucas, 1989; St-Onge et al., 2001); (2) imbrication of tectonostratigraphic units on southern Baffin Island including reactivation and further shortening across the Soper River suture; and (3) D2–M2 modification of pre-existing fabrics in the study area (see below). Terminal collision across the Bergeron suture (BeS; Fig. 3) is bracketed between 1820 +4/–3 Ma, the age of the youngest dated plutonic unit in the hanging wall of the suture (Scott & Wodicka, 1998) and 1795 ± 2 Ma, the age of a post-collisional syenogranitic dyke that crosscuts the suture on southern Baffin Island (Wodicka & Scott, 1997).
Within the hanging wall of the Bergeron suture on southern Baffin Island, principal modifications to pre-existing fabrics during D2 include the development of a more finely layered S2 foliation in metasedimentary units (see below) and the transposition of angular compositional discordances in metaplutonic rocks. Local reactivation of the Soper River suture is manifest by outcrop- to map-scale recumbent folding of earlier (S1a) foliations (St-Onge et al., 1999b).
M2: retrograde amphibolite-facies metamorphism, Meta Incognita microcontinent
D2 collisional deformation is associated with prograde greenschist- to amphibolite-facies metamorphism of the lower-plate Superior cratonic margin and retrograde upper amphibolite-facies metamorphism of the granulites of southern Baffin Island (St-Onge et al., 2000b).
Zones of D2 deformation within psammitic and pelitic units in the study area are marked by the pervasive growth of S2 biotite–sillimanite–quartz ± garnet ± cordierite at the expense of the S1a garnet–cordierite assemblages described above. Abundant, small, S2 biotite (0·3–0·5 mm) and sillimanite (0·1–0·5 mm) laths define a distinct, penetrative, schistose (finely layered), sub-millimetre-scale foliation (Fig. 6b). Sparse, small, sub-equant (0·5–1·0 mm) euhedral to slightly ellipsoid subhedral grains of garnet are generally aligned with and overgrow the biotite–sillimanite foliation (Fig. 6b). Based on the relationship with the S2 foliation, these grains are interpreted as M2 in age. Their size and shape, and the fact that they overgrow the S2 foliation, make them readily distinguishable from the larger, elongate and more voluminous S1a garnets (Fig. 6a). Within zones of D2 deformation, the S1a garnets (when present) are wrapped by (and thus pre-date) the S2 foliation and display embayed, corroded outlines. Cordierite is either absent, or present as sub-millimetre grains aligned in the S2 foliation. As with garnet, the size and relationship of the S2 cordierite grains to the foliation allows S2 cordierite to be distinguished from S1a cordierite. In some thin sections, the S2 garnets are further mantled by a sub-millimetre rim of cordierite, suggesting late-stage replacement. New growth and recrystallization of zircon and monazite in metasedimentary units indicates that M2 retrograde metamorphism occurred at ca. 1820 ± 1 to 1813 ± 2 Ma (Wodicka & Scott, 1997; Scott et al., 2002; this study).
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METHODS |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL BACKGROUNDTECTONIC EVENTS, DEFORMATION...METHODSRESULTSDISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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A suite of 25 samples was collected (Fig. 5) for an integrated structural–petrological–geochronological study of supracrustal rocks from the Meta Incognita microcontinent. Of these, 14 samples were utilized for textural observations, thermobarometric data, and phase diagram construction, whereas 12 samples were employed for U–Pb geochronological analysis of zircon, monazite, and titanite.
Thermobarometry
Samples of pelitic and mafic units from the Lake Harbour Group containing S1a granulite-facies assemblages or S2 amphibolite-facies assemblages (Table 1) appropriate for thermobarometric work were selected for analysis. Sample locations are given in Fig. 5. The spot-analyses were obtained at the Geological Survey of Canada (GSC) using a Cameca microbeam electron microprobe equipped with four wavelength-dispersive spectrometers and automated to produce simultaneous multi-element analyses and data reduction (Pouchou & Pichoir, 1984). Operating conditions were 15 kV accelerating voltage, specimen current of 10–30 nA and a counting time of 30 s. Standards used were appropriate minerals, oxides, and metals. A 10 µm beam spot was used for biotite and plagioclase and a more focused beam (ca. 3 µm) for garnet, cordierite, orthopyroxene, and clinopyroxene.
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All ferromagnesian minerals and plagioclase grains were analysed along rim–core–rim traverses at regular intervals (0·02–0·25 mm depending on size, yielding on average 5–10 spot-analyses per traverse) to check for compositional zoning or heterogeneities. For each mineral grain, the identification of relatively homogeneous core and rim domains was followed by 5–7 spot-analyses per domain. These analyses were then averaged to provide a statistically robust dataset for the thermobarometric calculations. For any given sample, detailed analytical work focused on a part of the thin section (ca. 2–5 mm radius), which was judged to be optimal in terms of mutual mineral textural equilibrium and lack of alteration (with the exception of cordierite; see below), and contained within a single S1a compositional layer or S2 domain. Reconnaissance microprobing of mineral phases outside the areas of detailed analysis confirmed the compositional homogeneity of the mineral phases within any given compositional layer or domain for this suite of rocks. Within each area of detailed analytical work, the analysed minerals were selected so as to be separated from each other by one or two grains (distance of ca. 1 mm) that did not participate in the equilibrium of interest (for example, quartz or plagioclase in the case of analysed ferromagnesian minerals), to avoid potential local retrograde re-equilibration between minerals actually in contact (St-Onge, 1984
, and references therein). Representative analyses (each an average of 5–7 spot-analyses) of orthopyroxene, clinopyroxene, garnet, cordierite, plagioclase, and biotite for the 12 thermobarometric samples are given in Tables 2–6.
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Compositional zoning (see Spear, 1993
; and references therein) is absent from most mineral phases analysed with the exception of garnet. In S1a samples D515 and S502 (Table 3) garnet is characterized by minor prograde compositional change from core to rim (increase in pyrope content, decrease in almandine and spessartine contents). In S2 samples D518, J501 and N516, and samples M500, P501 and S501, which contain S1a garnet cores and S2 garnet rims (Table 3), garnet is characterized by a retrograde compositional change between core and rim (increase in calculated almandine and spessartine contents, decrease in pyrope content).
Pressure-temperature (P–T) estimates were calculated using the TWEEQU software (Berman, 1991) that incorporates the thermodynamic data of Berman (1988) together with revised data for garnet, cordierite, and biotite (Berman & Aranovich, 1996; TWQ version 2.02a; http://www.gis.nrcan.gc.ca/twq.html). The following independent set of equilibria were utilized for the P–T calculations:
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Equilibria (3), (4), (7), and (9) serve as barometers and equilibria (1), (2), (5), (6), and (8) serve as thermometers. Absolute errors on the thermobarometric results are considered to be ca. ± 50°C and 1 kbar [see Essene, 1989; Berman, 1991; see also St-Onge & Ijewliw (1996) for discussion]. For the samples discussed in the text, we also report one standard deviation errors for the TWEEQU equilibria intersections, as computed by Berman's (1991) INTERSX program. The INTERSX standard deviations allow an assessment of the overall consistency of the results, but do not represent statistically valid uncertainties (Berman, 1991). Finally, given that a same set of equilibria is used for the thermobarometric calculations of all the pelitic assemblages [equilibria (1)–(4) above] and another same set of equilibria for the mafic assemblages [equilibria (1), (5)–(9) above], the calculations of relative changes in recorded P–T conditions between and within either pelitic or mafic samples (see below) should be reliable, as relative thermobarometry with the same set of equilibria depends dominantly on changes in the compositional part of ln 
terms.
Phase diagram calculations
The program DOMINO (De Capitani, 1994; http://titan.minpet.unibas.ch/minpet/groups/theriak/theruser.html, August, 2003 version) and the thermodynamic dataset of Holland & Powell (1998) were used to calculate phase diagrams in the expanded pelitic model system MnO–CaO–Na2O–K2O–FeO–MgO–Al2O3–SiO2–H2O (MnCNKFMASH; e.g. Spear, 1993; Tinkham et al., 2003). The phases considered in the construction of the phase diagrams were orthopyroxene, garnet, cordierite, sillimanite, kyanite, andalusite, K-feldspar, plagioclase, quartz, biotite, muscovite, chlorite, staurolite, melt, and H2O. Details and sources of the activity–composition models employed in this study have been given by Tinkham et al. (2003), with the biotite and chlorite models reparameterized for use with DOMINO (De Capitani, 2004, unpublished data). The H2O content in the system was reduced approximately to the point that the assemblage was just saturated immediately subsolidus at 6 kbar, following the procedure reported by White et al. (2001).
The bulk-rock composition of a large (12 kg) unaltered sample of Lake Harbour Group metapelite (N016; Fig. 5) containing an S1a garnet–cordierite–sillimanite–biotite–plagioclase–K-feldspar–quartz–melt assemblage, and characterized by regular millimetre-scale compositional layering, was obtained from X-ray fluorescence (XRF) analysis as reported by Thériault et al. (2001). Whole-rock XRF data for sample N016 are considered a reasonable proxy for bulk composition during M1a regional metamorphism and porphyroblast growth in the study area based on the large size of the sample and the millimetre scale of the S1a compositional layering, the limited documentation of growth zoning in S1a minerals, and the absence of S2 minerals and foliation, which suggests a lack of subsequent M2 retrograde re-equilibration of this sample.
Another large (9 kg) sample of Lake Harbour Group metapelite (D024; Fig. 5) containing an S2 sillimanite–biotitez–plagioclase–quartz–K-feldspar–garnet–cordierite–melt assemblage, and exhibiting a well-developed, penetrative, sub-millimetre-scale foliation, was also analysed by XRF to determine its bulk-rock composition (Thériault et al., 2001). We are aware that the growth of overprinting assemblages and the imposition of chemical zoning in porphyroblasts during subsequent metamorphic events (i.e. M2), may render such data inappropriate for the calculation of phase diagrams because of the possibility that porphyroblast cores may no longer be in equilibrium with surrounding matrix minerals. However, the strong development of a penetrative S2 schistose foliation in sample D024, which by analogy with other samples in this study (e.g. D518, D527, E501, J501, and N516) points to an exclusively M2 origin, suggests that the whole-rock XRF data can be considered to approximate the effective bulk composition during the overprinting metamorphic event (e.g. Stüwe, 1997; Marmo et al., 2002). In other words, based on the textural criteria available, D024 is interpreted to have completely re-equilibrated during the younger M2 metamorphic event.
U–Pb geochronology
Eight samples of metasedimentary and metaplutonic rocks from the Lake Harbour Group and four samples of felsic intrusive rocks were studied for U–Pb geochronology using isotope dilution thermal ionization mass spectrometry (ID-TIMS) at the GSC. Sample locations are given in Fig. 5. Preliminary results for three of the eight metasedimentary and metaplutonic samples (N056, D094B, N312) were presented by Wodicka & Scott (1997), Scott & Wodicka (1998) and Davis et al. (1998). Zircon, monazite, and titanite were separated by standard crushing, grinding, hydrodynamic, and heavy liquid techniques. Purification of the heavy mineral concentrates was carried out using a Frantz LB-1 isodynamic separator. All zircon fractions and some monazite and titanite fractions were air abraded following the method of Krogh (1982). The analytical methods for U–Pb analysis of zircon and monazite have been summarized by Parrish et al. (1987) and those of titanite in Davis et al. (1997). Treatment of analytical errors follows that outlined by Roddick (1987), with regression analysis modified after York (1969). For samples whose analyses are concordant or nearly concordant within error, the cited age is the weighted average of the 207Pb/206Pb ages (or 207Pb/235U ages in the case of reversely discordant monazite analyses), calculated using Isoplot version 3.00 of Ludwig (2003). Decay constants used are those recommended by Steiger & Jäger (1977).
Back-scattered electron (BSE) images were obtained from selected grains of monazite to evaluate the dispersion of ages obtained from two samples (N056 and D332A). Imaging was carried out on a Cambridge Instruments S360 scanning electron microprobe also located at the GSC. Operating conditions were 20 kV accelerating voltage and 2–5 nA beam current. Most images reveal internal compositional variations that reflect differences in the average atomic number of monazite. Based on several monazite studies (e.g. Zhu et al., 1997; Stern & Sanborn, 1998; Foster et al., 2002), these differences largely reflect differences in the concentration of rare earth elements (REE), U, and especially Th.
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RESULTS |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL BACKGROUNDTECTONIC EVENTS, DEFORMATION...METHODSRESULTSDISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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Thermobarometry
S1a assemblages
Four of the analysed samples contain exclusively S1a mineral assemblages, with two samples characterized by a well-developed compositional foliation (D515, S502), two by a weakly developed planar foliation (N503A, N503B), and all four devoid of overprinting S2 assemblages and modification by S2 fabrics (Table 1). The two well-layered samples are pelitic (D515, S502) whereas the two foliated samples are mafic (N503A, N503B). All four yield relatively high T and P estimates (Table 7) with average determinations of 812°C and 7·5 kbar for mineral core compositions and 810°C and 8·3 kbar for mineral rim compositions. If one includes the P–T estimates based on S1a mineral core compositions from three samples (M500, P501, S501; Table 7) that contain overprinting S2 (rim) assemblages, the average mineral core P–T determinations for S1a remain virtually unchanged at 806°C and 7·6 kbar.
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S2 assemblages
Five analysed samples contain exclusively S2 mineral assemblages that define (biotite, sillimanite, plagioclase) and overgrow (garnet) a well-developed schistose foliation (Fig. 6b; Table 1). The five samples (D518, D527, E501, J501, N516) yield lower T and P estimates (Table 7) when compared with the P–T determinations for S1a (cores and rims) presented above, with average S2 determinations of 735°C and 6·5 kbar for mineral core compositions and 710°C and 6·0 kbar for mineral rim compositions. In contrast to the thermobarometric results for S1a, the S2 data document a slight decrease in both P and T from mineral core to mineral rim analyses.
Phase diagrams
To further delimit potential segments of P–T paths as derived from core–rim thermobarometry and to substantiate the apparently contrasting trends in recorded P–T evolution for the M1a and M2 events (see above), two phase diagrams were constructed for the pelitic rocks of the study area (Figs 8–11) utilizing the XRF data presented by Thériault et al. (2001). As the samples in that study were completely consumed by the whole-rock chemical analysis, subsequent direct thermobarometric work was not an option and consequently some degree of mismatch might be expected when comparing predicted compositional mineral trends with those observed in the analysed thermobarometric samples.
S1a assemblages
The phase diagram based on the bulk-rock composition of sample N016 (Fig. 5) and for which saturation of the H2O content in the system occurs just below solidus, is shown in Fig. 8. The bulk-rock composition for N016, which contained only an S1a granulite-facies mineral assemblage and an associated, well-developed compositional foliation, restricts the garnet–cordierite–biotite–sillimanite–plagioclase–K-feldspar–quartz–melt assemblage to a specific stability field, which is highlighted with a dashed line. The average S1a (core) P–T determination derived from the thermobarometric results presented above is plotted as point ‘C’ in Fig. 8 and is in good agreement with the outlined stability field for this assemblage. In contrast, the average S1a (rim) P–T determination (point ‘R’ in Fig. 8) lies above the outlined stability field within the higher-P garnet–biotite–sillimanite–plagioclase–quartz–melt (cordierite absent) field. As previously noted, all phases defining the S1a assemblage appear to be in mutual textural equilibrium, with the exception of cordierite (Fig. 7b). Thus the increase in P during M1a as recorded by the thermobarometric results in Table 7, and the inferred consequent instability of cordierite as shown in the phase diagram of Fig. 8, appear to be corroborated texturally by the recessive and embayed aspect of S1a cordierite in thin section as previously described. The increase in P during M1a is also consistent with the minor increase in the grossular content of garnet from core to rim in analysed S1a samples D515 and S502 (Table 3).
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The essentially isothermal compressive P–T path segment for M1a as defined by core to rim thermobarometry (Fig. 8) can be corroborated by the isopleths of Mg/(Mg + Fe) and XGrs calculated for sample N016 and shown for the relevant garnet-bearing portion of the phase diagram in Fig. 9. The core to rim segment of the path (continuous-line arrow) can be seen to follow at a small angle the Mg/(Mg + Fe) isopleths for garnet (Fig. 9a), which agrees with the small to negligible increase in the Mg number from analysed S1a garnet cores to rims (Table 3). Likewise, the core to rim P–T path segment crosses a field characterized by a slight increase in XGrs, as documented by the isopleths of Fig. 9b and reflected in the S1a garnet analyses of Table 3. Finally, the absence of kyanite in S1a assemblages within the project area suggests that the subsequent cooling path remained within the sillimanite stability field and approximated the P–T path segment shown with a dashed line in Fig. 9.
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S2 assemblages
The phase diagram based on the bulk-rock composition of sample D024 (Fig. 5), and for which saturation of the H2O content in the system occurs just below solidus, is shown in Fig. 10. The bulk-rock composition for D024, which containes an S2 biotite–sillimanite–cordierite–garnet–quartz–melt mineral assemblage and a pervasive, well-developed schistose foliation, limits the S2 assemblage to the stability field that is highlighted with a dashed line in Fig. 10. The thermobarometric conditions determined with S2 core and rim mineral analyses (points ‘C’ and ‘R’ in Fig. 10) plot just outside the stability field determined from the bulk-rock composition, suggesting either that sillimanite may have equilibrated early along the P–T path for M2 (shown in Fig. 10 with a dashed P–T path line segment) or that the bulk-rock composition for sample D024 does not adequately approximate the effective bulk composition (e.g. Stüwe, 1997
; Marmo et al., 2002) for the Lake Harbour Group pelites during M2. The abundant fine-grained sillimanite defines the S2 foliation in all thin sections examined and thus appears to be in textural equilibrium with adjacent S2 biotite, cordierite, and plagioclase grains (Fig. 6b). In contrast, sillimanite is never observed in direct contact with the S2 garnet, the two mineral phases generally being separated by intervening plagioclase, cordierite, or biotite grains. The latter observation suggests that sillimanite may have indeed equilibrated early along the M2 path and become a relict phase prior to growth of the small, sub-equant S2 garnets.
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The M2 P–T path segment derived by thermobarometry is corroborated by the isopleths of Mg/(Mg + Fe), but less so by XGrs, as calculated for sample D024 and shown for the relevant garnet-bearing portion of the phase diagram in Fig. 11. The core to rim segment of the path (continuous-line arrow) can be seen to cross at a high angle the Mg/(Mg + Fe) isopleths for garnet (Fig. 11a), which agrees with the decrease in the Mg number for most analysed S2 garnet cores to rims (Table 3). The P–T path segment also crosses at high angles the XGrs isopleths for D024 (Fig. 11b), suggestive of an increase in Ca content from core to rim. Such a compositional trend is documented in only one S2 sample (N516, Table 3) with most other S2 garnet grains actually showing a decrease in Ca from core to rim. This apparent contradiction in predicted vs actual grossular compositional trends for S2 garnets is interpreted to reflect the variable nature of the metasomatic processes associated with the retrograde M2 metamorphic event, making our assumption of a representative M2 effective bulk composition largely invalid, at least in terms of Ca content. Alternatively, the bulk-rock composition derived from sample D024 may simply not be representative of the analysed S2 pelite samples in the study area.
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U–Pb geochronology
Twelve samples of Lake Harbour Group metasedimentary rocks and associated felsic intrusive units from the study area were analysed for U–Pb geochronology. Details pertaining to the morphology, colour, and chemistry of the analysed mineral phases are given in Electronic Appendix 1 (http://www.petrology.oxfordjournals.org). The results presented below, and summarized in Table 8, document the timing and duration of the M1a and M2 metamorphic events, the associated formation of the S1a and S2 deformation fabrics, and further characterize the polymetamorphic evolution of the study area.
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White monzogranites, Lake Harbour Group
Siliciclastic rocks of the Lake Harbour Group are typically associated with white monzogranite occurring either as discrete, metre-thick tabular bodies or as concordant to sub-concordant layers or irregular pods less than 0·5 m thick and generally aligned parallel to, and wrapped by, the dominant S1a compositional foliation in the metasedimentary host. The monzogranite contains garnet with or without minor biotite, cordierite, and sillimanite. The aluminous mineral assemblage, the field association with the Lake Harbour Group, and the relationship to the dominant S1a compositional foliation suggest a derivation by partial melting during M1a regional metamorphism (e.g. St-Onge et al., 1996
; Scott, 1997). Both zircon and monazite were recovered from two samples of white monzogranite (N056, D094B; Fig. 5) making these rocks suitable for documenting the timing of M1a metamorphism and anatexis. Sample N056 is from a sheet of garnet-bearing anatectic monzogranite that is concordant to a well-developed S1a compositional foliation in semipelite. The monzogranite lies 500 m north of the trace of a D2 thrust (Fig. 5) and contains a strong internal S2 foliation defined by quartz ribbons and millimetre-scale polycrystalline feldspar bands that wrap around large feldspar and S1a garnet porphyroclasts. U–Pb analyses of five unabraded, single-grain fractions of monazite (MA–ME, Fig. 12a) are concordant to near-concordant (0·0–0·3%), but their 207Pb/206Pb ages exhibit a range of over 20 Myr between 1841 ± 2 and 1820 ± 2 Ma (Electronic Appendix 1). BSE imaging reveals complex internal structures for euhedral to round monazite grains, with cores showing faint sector, oscillatory, or patchy zoning overgrown by darker homogeneous or faintly zoned rims (Fig. 12b–d). In some cases, irregular embayment boundaries truncate zoning patterns in the core domains without change in the overall shape of the crystal (Fig. 12c and d). In contrast, anhedral fragments tend to be homogeneous (Fig. 12e). These textures, combined with a distinct Th/U ratio–age correlation (Electronic Appendix 1), suggest that the spread in 207Pb/206Pb ages exhibited by monazite in sample N056 results largely from recrystallization of older grains combined with younger metamorphic (over)growth, although the possibility that other processes (e.g. high-temperature Pb loss) were also involved cannot entirely be ruled out. Based on these observations and the petrofabric context for sample N056, the 1841 ± 2 Ma age for the oldest monazite fraction (MB, Fig. 12a) is interpreted as recording the minimum crystallization age of the monzogranite during M1a anatexis (Table 8), with new monazite growth and total recrystallization of M1a grains having occurred at 1820 ± 1 Ma (weighted average 207Pb/206Pb age of analyses MC and MD, Table 7; Fig. 12a), possibly in response to fluid influx during the younger M2 metamorphic event (Table 8).
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. (b–e) Back-scattered electron (BSE) images of selected monazite crystals from sample N056. (b) Euhedral equant monazite crystal with faint sector zoning overgrown by a concentrically zoned, dark round rim. The bright spot in the centre left of the grain is monazite. (c) Round monazite crystal with embayed oscillatory-zoned central domain. (d) Subhedral monazite crystal showing an irregular, patchy central domain resorbed by inward-directed fronts. (e) Simple homogeneous anhedral monazite fragment with sharp edges. (f) U–Pb concordia diagram for zircon (white ellipses) and monazite (vertically ruled ellipses) from white monzogranite sample D094B. Error ellipses represent 2Two zircon fractions in N056 (ZB and ZC, Fig. 12a) yield the oldest 207Pb/206Pb ages of 1836 ± 2 and 1830 ± 2 Ma, respectively, and have low Th/U ratios (Electronic Appendix 1), typical of zircon that has crystallized from hydrous crustal fluids generated during high-grade metamorphism (e.g. Krogh & Davis, 1973
). In contrast, two other fractions (ZA and ZD, Fig. 12a) yield distinctly higher Th/U ratios (Table 7) and younger 207Pb/206Pb ages of 1808 ± 2 and 1801 ± 6 Ma. The observed differences in chemistry and 207Pb/206Pb ages suggest at least two episodes of zircon growth. The age of M1a cannot be determined from the zircon data alone, but the monzogranite is likely to be older than the oldest discordant analysis (ZB) with a 207Pb/206Pb age of 1836 ± 2 Ma, consistent with the monazite data from this sample. Regressing the two youngest data points ZA and ZD together with concordant monazite data point MC yields an upper intercept age of 1821 + 11/–5 Ma [mean square of weighted deviates (MSWD) = 0·12], indistinguishable from the monazite weighted mean age of 1820 ± 1 Ma for the M2 event.
Sample D094B (Fig. 5) is a foliated (S1a + S2) anatectic monzogranite sheet containing both monazite and zircon. Wodicka & Scott (1997) originally analysed three single-grain monazite fractions (MA to MC, Fig. 12f) that gave a weighted mean 207Pb/206Pb age of 1838 ± 2 Ma. Four additional monazite analyses (MD to MG, Fig. 12f), conducted during subsequent U–Th–Pb studies (Davis et al., 1998; Stern & Sanborn, 1998), showed evidence for isotopic heterogeneity, with fraction MG having a significantly older U/Pb and Pb/Pb age (Electronic Appendix 1). Stern & Sanborn (1998) reported an age estimate of 1836 ± 1 Ma for this sample, based on the non-weighted mean age (1 standard error) of analyses MA to MF. Three multi-grain zircon fractions (ZA–ZC, Fig. 12f) from this same rock show a greater spread in 207Pb/206Pb ages and more extensive discordance (0·3–0·7%) than monazite (Electronic Appendix 1), similar to that observed in sample N056. A regression of the most precise collinear zircon and monazite data (fractions ZA, MA, MB, MC, ME, and MF, Fig. 12f) gives an upper intercept of 1837 + 2/–1 Ma (MSWD = 1·35). The age determination overlaps both the 1838 ± 2 Ma U–Pb age reported by Wodicka & Scott (1997) and the 1836 ± 1 Ma U–Th–Pb age estimate of Stern & Sanborn (1998). Consequently, 1837 + 2/–1 Ma is taken as the best estimate for the crystallization age of the white monzogranite and M1a partial melting at this locality (Table 8). The age for the younger zircon component (ZB and ZC, Fig. 12f) is interpreted to be equal to, or older than the 207Pb/206Pb age of the least discordant fraction ZB, which is 1818 ± 2 Ma (Table 8). This result corresponds closely to the age determined for M2 monazite and zircon in sample N056 described above.
Pelitic, semipelitic, and psammitic rocks, Lake Harbour Group
Sample S502 (Fig. 5) is a garnet–biotite–sillimanite–plagioclase–K-feldspar–quartz–melt-bearing metapelite characterized by a well-developed S1a compositional foliation that has yielded core and rim P–T estimates of 7·5 kbar and 814°C, and 8·1 kbar and 847°C, respectively (see above; Table 7). Five single-grain monazite analyses (MA–ME, Fig. 13a) define two distinct populations based on their morphology, 207Pb/206Pb ages and Th/U ratios (Electronic Appendix 1). The oldest population (MA, MB, MC, ME, Fig. 13a) has slightly discordant (0·1–0·4%) 207Pb/206Pb ages in the range 1849 ± 2 to 1844 ± 2 Ma. The second population is defined by a single prismatic crystal (MD) that is somewhat reversely discordant with a 207Pb/235U age of 1830 ± 2 Ma. Whereas the older monazite group records the time of M1a metamorphism at ca. 1849–1844 Ma (Table 8), the younger age of ca. 1830 Ma is attributed to a subsequent disturbance distinctly post-M1a but pre-M2 (discussed below).
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Sample N311 (Fig. 5) is a foliated (S1a + S2) grey psammite containing large (8 mm) S1a garnet poikiloblasts. Four, unabraded, single-grain fractions (MA–MD, Fig. 13b) of monazite yield near-concordant (0·1–0·3%) 207Pb/206Pb ages ranging from 1838 ± 2 to 1828 ± 2 Ma (Electronic Appendix 1). We interpret 1838 ± 2 Ma as the minimum age for monazite growth during M1a metamorphism (Table 8), and the dispersion to younger ages in the other grains as indicative of partial recrystallization and (or) younger metamorphic overgrowth [as documented in samples N056 and S502 (above), and D332A (below)].
Sample D513B (Fig. 5) is a foliated (S1a), grey biotite-garnet-bearing psammite. Three abraded monazite analyses (MA–MC, Fig. 13c) yield overlapping discordant (0·2–0·3%) 207Pb/206Pb ages of 1839 ± 2 to 1835 ± 2 Ma (Electronic Appendix 1). Based on similarities in colour, morphology, and chemistry (Electronic Appendix 1), we conclude that the analysed monazite grains define a single population with a minimum age of ca. 1839–1835 Ma for M1a (Table 8).
Sample D332A (Fig. 5) is a garnet–biotite–melt-bearing semipelite. Garnet occurs either as large, subhedral S1a porphyroblasts with inclusions of quartz and opaque minerals or as small inclusion-free, euhedral S2 grains. Four analyses, each consisting of a single unabraded monazite crystal (MA–ME, Fig. 13d), resulted in a range of concordant to near-concordant ages from 1845 ± 2 to 1813 ± 2 Ma (Electronic Appendix 1). BSE imaging documents fine oscillatory zoning in monazite grains with partial euhedral faces (Fig. 13e), suggestive of igneous growth, and patchy zoning in rounded grains (Fig. 13f). In addition, inner domains on a number of grains are truncated by homogeneous overgrowths of variable thickness (Fig. 13e and f). From these features, we interpret the oldest age, 1845 ± 2 Ma, to be the minimum age for M1a monazite growth, and 1813 ± 2 Ma as the time of extensive recrystallization or new monazite growth during M2 metamorphism (Table 8). The intermediate ages obtained for monazite fractions MA and ME (Electronic Appendix 1) might reflect partial recrystallization (or monazite overgrowth) during M2.
Carbonate and mafic rocks, Lake Harbour Group
Two suitable samples of titanite-bearing rocks (D020 and D513A) were collected in the study area (Fig. 5). Sample D020 is a coarse-grained, diopside-bearing siliceous marble. Three of four analyses of titanite (TC–TE, Fig. 14a) fit a discordia line with an upper intercept age of 1790 ± 1 Ma (MSWD = 0·96). We interpret this age as corresponding to complete isotopic resetting of older titanite grains (or new mineral growth) during a post-M2 thermal or fluid event (discussed below).
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Sample D513A is a well-foliated (S1a) metagabbro taken from the same outcrop as sample D513B (see above). Three of four multigrain fractions of titanite (TA–TD) analysed are normally discordant (0·5–0·7%) to concordant with 207Pb/206Pb ages of 1786 ± 7, 1784 ± 5, and 1778 ± 4 Ma (Fig. 14b). The other fraction (TD) is reversely discordant (–1·2%) and has an older 207Pb/206Pb age of 1796 ± 5 Ma, but plots on the same discordia line as the other three analyses. Regression through all four analyses yields an upper intercept age of 1785 ± 4 Ma (MSWD = 1·18) (Electronic Appendix 1). This age is distinctly younger than >1839–1835 Ma monazite separated from the adjacent semipelite D513B and is taken to indicate either the time of new titanite growth or isotopic resetting as a result of a post-M2 thermal event (see below).
Felsic intrusive rocks, Meta Incognita microcontinent
A number of monzogranite to syenogranite dykes and syenite to monzonite plugs are parallel or discordant to the principal deformation fabrics that characterize the metamorphic rocks of the study area. Previous work on these intrusive units revealed at least three phases of pre- to post-D2 dyke and plug emplacement: at ca. 1840, 1795–1784, and 1760 Ma (Scott, 1997; Table 8). Four additional samples of dykes and plugs were collected to further document the timing of magmatism concurrent with, and postdating, D1a deformation and M1a metamorphism in the study area.
Sample D534 (Fig. 5) is from an undeformed syenogranite pegmatite dyke oriented parallel to the S1a compositional foliation in host psammitic rocks. The undeformed nature of the dyke is suggestive of a late S1a emplacement. Data for five of the six analyses of zircon and monazite (MA–ZB, excluding fraction ZC, Fig. 14c) are collinear, and regression gives an upper intercept age of 1844 + 9/–3 Ma (MSWD = 0·16). This age is interpreted as dating the emplacement of the pegmatite and suggests that formation of the regional S1a foliation was waning by ca. 1844 Ma (Table 8).
Sample D532 (Fig. 5) is a monzogranite pegmatite that clearly cross-cuts the main S1a compositional layering in host psammitic rocks. Analysis of three single- and multigrain fractions of zircon (ZA–ZC, Fig. 14d) and three single-grain fractions of monazite (MA–MC, Fig. 14d), all abraded, are in good agreement and suggest an emplacement age of 1833 ± 1 Ma on the basis of their weighted mean 207Pb/206Pb age (MSWD = 0·89; probability of fit = 0·49) (Electronic Appendix 1); 1833 Ma represents a minimum age for development of the S1a compositional foliation in the study area (Table 8).
Sample D533 (Fig. 5) is an undeformed syenite plug intrusive into foliated (S1a) orthopyroxene–biotite monzogranite of the Cumberland batholith. Five analyses (ZA–ZE, Fig. 14e) of zircon display only minor discordance (0·2–0·4%) with 207Pb/206Pb ages ranging between 1834 ± 2 and 1823 ± 2 Ma (Electronic Appendix 1). A weighted mean age of fractions ZA, ZB, and ZE (Fig. 14e) yields 1831 ± 3 Ma (MSWD = 1·8; probability of fit = 0·16). This age dates the emplacement of the syenite and is a minimum age for S1a formation at this locality.
Sample N312 (Fig. 5) is from a cross-cutting, undeformed body of orthopyroxene-bearing monzonite (sample N312, Fig. 5) emplaced into the well-foliated (S1a) orthogneissic basement to the Lake Harbour Group. Five analyses of strongly abraded zircon fragments (ZA1-ZA3, ZD, ZE, Fig. 14f) are –0·3 to 0·4% discordant and yield the oldest 207Pb/206Pb ages, between 1834 ± 2 and 1829 ± 2 Ma (Electronic Appendix 1). Three of the five analyses (ZA1, ZA2, and ZE, Fig. 14f) yield a weighted mean 207Pb/206Pb age of 1834 ± 1 Ma (MSWD = 0·024; probability of fit = 0·98). The remaining two analyses, including a tip (ZA3, Fig. 14f), give a slightly younger age of ca. 1830 Ma. One analysis of small subequant zircon grains (ZB, Fig. 14f) and two analyses of subhedral to anhedral monazite (MC and ME, Fig. 14f) define a discordia line with a distinctly younger upper intercept of 1816 + 2/–1 Ma (MSWD = 0·54). A third analysis of subhedral monazite (MD, Fig. 14f) plots slightly above concordia but has an overlapping 207Pb/235U age of 1813 ± 2 Ma. A small zircon prism (ZC, Fig. 14f) is concordant at 1797 ± 2 Ma. Finally, the two youngest analyses of euhedral monazite (MA and MB, Fig. 14f) give a weighted mean 207Pb/235U age of 1791 ± 2 Ma. These results, combined with textural and chemical inhomogeneities of the analysed zircon and monazite populations (Electronic Appendix 1), provide evidence for multiple mineral growth events. The external morphology of the oldest zircon grains (Electronic Appendix 1) is consistent with a magmatic origin at 1834 ± 1 Ma. This result places a minimum age on S1a foliation formation. The slightly younger ca. 1830 Ma age may indicate the time of granulite-facies metamorphism of the monzonite, shortly following its crystallization. The younger 1816 + 2/–1 Ma, 1797 ± 2 Ma, and 1791 ± 2 Ma ages are interpreted to reflect episodic (or prolonged) zircon and monazite growth during and following M2 regional metamorphism as a result of mainly retrograde fluid flow (discussed below).
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DISCUSSION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL BACKGROUNDTECTONIC EVENTS, DEFORMATION...METHODSRESULTSDISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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The integration of mineral growth–deformation fabric relationships, thermobarometric results, and petrological data provided by representative phase diagrams allows the identification and characterization of two contrasting, regionally significant, tectono-metamorphic events (M1a, M2) in the internal zone of the Trans-Hudson Orogen. New U–Pb data document the timing and duration of these two metamorphic events, allow the reinterpretation of previously acquired age data for the study area, and reveal additional complexities in the polymetamorphic evolution of southern Baffin Island. In this section, we summarize the main results and identify the potential causes of metamorphism for each thermal event documented in the study area, and draw parallels with the thermal evolution of the upper-plate domain in the Cenozoic India–Asia collision zone.
M1a metamorphism
The oldest regional metamorphic event identified by this study, M1a, is characterized by granulite-facies assemblages in metapelite and is associated with the development of a pervasive compositional foliation (S1a) in the metamorphic rocks of southern Baffin Island (petrology samples D515, S502, M500, P501, and S501; XRF sample N016). S1a assemblages reflect P–T conditions of up to 810°C and 8·3 kbar (Table 7), and evolved along a steep, compressional P–T path segment (Figs 8 and 9). U–Pb ages from metamorphic monazite in siliciclastic units with pervasive S1a assemblages and fabrics (geochronology samples D513B, D332A, N311, and S502) indicate monazite growth between ca. 1849 and 1835 Ma (Fig. 15). These new data indicate that high-grade M1a metamorphic conditions were maintained for at least 14 Myr, during which time the S1a assemblages and foliation were developed (Fig. 15), and are consistent with previously reported ages for metamorphic monazite from other siliciclastic samples of the Meta Incognita microcontinent (see Table 8; Wodicka & Scott, 1997).
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New dating of anatectic granites (geochronology samples D094B and N056) bracket partial melting at before ca. 1841–1837 Ma (Table 8). These results are consistent with two previously reported 1845 ± 2 Ma age determinations (Scott, 1997
; Wodicka & Scott, 1997) for monzogranitic melt pods and veins in the study area. The results highlight the intimate field and temporal association that can be established between granulite-facies garnet–cordierite (M1a) metamorphism, anatexis, and regional (S1a) foliation development in the study area.
The ca. 1849–1835 Ma age span determined for M1a metamorphism and partial melting partly coincides with the waning stages of ca. 1865–1848 Ma Cumberland batholith magmatism (Jackson et al., 1990; Wodicka & Scott, 1997; Scott & Wodicka, 1998; Scott, 1999; Fig. 15) and largely postdates the timing of accretion of the intra-oceanic Narsajuaq arc to the leading edge of the upper (Churchill) plate (i.e. the study area, Fig. 4c) at ca. 1845 Ma (see above). We therefore propose that the M1a regional metamorphic event and associated partial melting was caused not only by emplacement of the Andean-type Cumberland batholith, which probably played a significant role in the attainment of peak M1a metamorphic conditions in an upper-plate setting (e.g. Wodicka & St-Onge, 1998; St-Onge et al., 2000b), but also by the crustal thickening (D1a) related to the collision and accretion of the Narsajuaq arc to the southern margin of the Meta Incognita microcontinent (and manifest in the study area as S1a) at ca. 1845 Ma.
Late- to post-D1a crustal magmatism, S1a foliation formation, and M1b metamorphism
U–Pb data from four felsic dykes and plugs (geochronology samples D532, D533, D534, and N312) constrain the age of late- to post-accretion (D1a) crustal magmatism and S1a foliation formation. The oldest dyke (D534), dated at 1844 + 9/–3 Ma (Table 8), was largely coeval with M1a metamorphism, being emplaced during a late stage of S1a foliation development (Fig. 15). A similar, late to post-S1a monzogranite dyke was dated by Scott (1997) at 1840 ± 3 Ma (Table 8). In contrast, a distinctly younger phase of felsic crustal magmatism, dominated by undeformed dykes and plugs (geochronology samples D532, D533, and N312), took place at ca. 1834–1831 Ma. Field observations indicate that these younger intrusive rocks cut the well-developed S1a foliation on southern Baffin Island. Thus, the intrusion ages and structural state of individual dykes and plugs indicate that development of S1a in the southern Baffin Island study area occurred mostly prior to ca. 1844–1840 Ma and must have ended before ca. 1834–1831 Ma. These age determinations indicate that S1a foliation formation was largely coeval with D1a accretion of the Narsajuaq arc at ca. 1845 Ma, and may therefore be related to closure of the Soper River suture (Fig. 15) and consequent Andean-type tectonic thickening along the southern margin of the Meta Incognita microcontinent.
A concordant 1830 ± 2 Ma age from metamorphic monazite in sample S502, as well as previously published monazite and zircon data by Scott & Gauthier (1996), Wodicka & Scott (1997), and Scott & Wodicka (1998) (Table 8) suggest that rocks from southern Baffin Island experienced a thermal perturbation at ca. 1833–1829 Ma. Whether these ages reflect continued mineral growth during the waning stages of a protracted M1a metamorphic event, or whether they represent a distinct thermal event (M1b; Fig. 15) related to intrusion of the ca. 1834–1831 Ma felsic dykes and plugs remains uncertain. The latter hypothesis is favoured by: (1) evidence for two distinct episodes of monazite growth in sample S502 at ca. 1849–1844 Ma and at 1830 ± 2 Ma (Table 8); (2) the presence of ca. 1830 Ma metamorphic zircon in monzonite sample N312; and (3) the post-S1a character of the 1834–1831 Ma magmatic suite. Regardless of the interpretation, the combined data imply that temperatures in the study area remained high for at least 20 Myr, from about 1849 until 1829 Ma.
M2 metamorphism
The second regional metamorphic event, M2, is characterized by the growth of upper amphibolite-facies assemblages at the expense of S1a assemblages, the development of a penetrative, schistose, sub-millimetre-scale S2 foliation, and the growth of sparse, sub-equant S2 garnets that are generally aligned with, and overgrow, the schistose foliation (petrology samples D518, D527, E501, J501, N516, M500, P501, and S501; XRF sample D024). S2 metapelitic assemblages record a clockwise, decompression P–T path segment (Figs 10 and 11) and P–T conditions during M2 are estimated as having involved 710°C and 6·0 kbar from mineral rim compositions (Table 7). New concordant monazite ages for siliciclastic and anatectic granite units (geochronology samples D332A and N056) bracket the retrograde M2 metamorphic event between ca. 1820 and 1813 Ma (Table 8). Effects of M2 are also recorded by new growth and (or) recrystallization of monazite or zircon in geochronology samples N056, D094B, and N312 (Table 8). The ca. 1820–1813 Ma ages overlap with previously reported ca. 1819–1814 Ma metamorphic zircon and monazite ages for the Lake Harbour Group and Blandford Bay assemblage (Wodicka & Scott, 1997; Scott et al., 2002).
The clockwise decompressive P–T path segment documented for M2 is consistent with numerical models for the thermal evolution of tectonically thickened crust (e.g. England & Thompson, 1984), and the ca. 1820–1813 Ma age span for M2 closely corresponds to the early stages of D2 collision in Trans-Hudson Orogen (bracketed between ca. 1820 and 1795 Ma; Wodicka & Scott, 1997; Scott & Wodicka, 1998; Fig. 15). Growth of the S2 assemblages in the study area appears to be closely tied to the intensity of foliation development (Table 1) whereas U–Pb data presented in Electronic Appendix 1, and summarized above, strongly suggest that the D2 event had an influence on the U–Pb systematics of monazite and zircon. In general, greatest recrystallization by S2 mineral assemblages, including new growth and (or) recrystallization of monazite or zircon, occurs in samples that exhibit a strong penetrative foliation and are located in the immediate vicinity of D2 thrusts (e.g. petrology samples D518, D527, E501, J501, N516; geochronology samples N056, D332A, and D094B). Thus, deformation-enhanced fluid circulation during retrograde amphibolite-facies metamorphism (e.g. St-Onge et al., 2000b) directly related to the D2 collision of the upper (Churchill) plate (i.e. including the study area) with the lower-plate Superior craton (Fig. 4e) appears to have been instrumental in promoting partial to total recrystallization and new growth of M2 silicate assemblages, monazite, and zircon in the polymetamorphic rocks of southern Baffin Island.
Post-D2 thermal and fluid activity
Two lines of evidence suggest that the youngest zircon, monazite, and titanite ages of ca. 1797–1785 Ma from the study area (geochronology samples D020, D513A, N312; Table 8) might be related to post-collisional thermal activity or fluid infiltration rather than to cooling. First, the ca. 1797–1785 Ma age range overlaps in time with emplacement of ca. 1795–1784 Ma post-collisional syenogranite pegmatite dykes in the composite upper (Churchill) plate (Scott, 1997; Wodicka & Scott, 1997; Fig. 15). Second, late fluid infiltration would be consistent with petrographic evidence for minor greenschist-facies retrogression, which includes secondary chlorite, muscovite, and epidote spatially associated with fractures and sericitized plagioclase noted in, for example, sample N312.
Upper-plate magmatism and metamorphism in the India–Asia collison zone
In the western upper-plate (Asian) domain of the Cenozoic India–Asia collision zone, the Karakoram batholith includes a series of pre-collisional deformed and metamorphosed 150–95 Ma granitoids (Searle et al., 1989, 1990; Treolar et al., 1989; Searle, 1991; Crawford & Searle, 1992; Fraser et al., 2001; Hildebrand et al., 2001), as well as massive, pre-collisional, mid-Cretaceous, I-type plutonic suites (Fraser et al., 2001). The younger Transhimalayan batholith (Singh & Jain, 2003) is a linear belt of Andean-type plutons that extends the full length of the exposed upper plate and is bracketed between 102 ± 2 Ma and 49·8 ± 0·8 Ma (Scharer et al., 1984; Weinberg & Dunlap, 2000). Taken as a set, these Andean-type plutonic units witness the long-lived nature of the upper-plate margin of Asia, with most magmatic units probably formed above pre-collisional northward-directed subduction zones prior to the India–Asia collision (Searle et al., 1999).
Several episodes of high-grade, regional, Barrovian metamorphism are documented in the Karakoram upper-plate domain (Searle & Tirrul, 1991; Fraser et al., 2001; Hildebrand et al., 2001). U–Pb monazite ages from staurolite-grade schists and sillimanite-grade gneisses pre-dating the India–Asia collision peak at ca. 135–126, 106–102 and 82·9 ± 6·1 to 61·9 ± 4·7 Ma (Fraser et al., 2001; Hildebrand et al., 2001). Metamorphism postdating the collision event 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 continuing today at depth (Searle, 2007).
We surmise that the polymetamorphic evolution documented for the internal zone of the Trans-Hudson Orogen on southern Baffin Island (this study) and that documented for the upper-plate domain of the India–Asia collision zone (Searle, 2007; and references therein) are analogous in type, duration, and extent. In both orogens, early upper-plate thermal events are interpreted as driven by heat advection related to the emplacement of pre-collision, Andean-type plutonic suites, whereas subsequent metamorphic events are related to either the continental collision itself or to post-collisional crustal fluid flow.
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CONCLUSIONS |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL BACKGROUNDTECTONIC EVENTS, DEFORMATION...METHODSRESULTSDISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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The integration of new structural, metamorphic, and geochronological data and observations from the southern Baffin Island study area document that metasedimentary and metaplutonic rocks from the internal zone of the Trans-Hudson Orogen record a complex plate-margin to intraplate polymetamorphic history that includes the following events.
- Three deformation events (D0, D1a, D2), which are the result of the sequential north to south closing of three crustal sutures (the Baffin suture, Soper River suture, and Bergeron suture).
- A granulite-facies regional metamorphic event (M1a) that is characterized by an approximately isothermal, compressive P–T path segment and the development of a dominant (S1a) compositional foliation. M1a is associated with the emplacement of a pre-collision, Andean-type magmatic batholith and crustal thickening related to the accretion of an intra-oceanic arc to the evolving upper plate.
- A retrograde regional metamorphic event (M2), which is characterized by a clockwise, decompression P–T path segment and the development of a penetrative schistose S2 foliation. M2 is associated with crustal deformation and fluid infiltration generated by continent–continent collision.
- At least two localized thermal events (M1b, post-D2), associated with post-accretion and post-collision crustal magmatism.
The protracted (ca. 80 Myr) record of magmatism, deformation, and metamorphism documented for the internal zone of Trans-Hudson Orogen on Baffin Island reveals a Paleoproterozoic tectonothermal evolution, which is similar in character, scope, and scale to that of the upper plate of other, younger orogenic belts, in particular that of the Himalaya–Karakoram Orogen of Asia (e.g. Hildebrand et al., 2001; Fraser et al., 2001). We suggest that the results presented in this study provide tangible evidence for the operation of modern plate-margin to intraplate tectonometamorphic processes back to the early Paleoproterozoic.
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SUPPLEMENTARY DATA |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL BACKGROUNDTECTONIC EVENTS, DEFORMATION...METHODSRESULTSDISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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Supplementary data for this paper are available at Journal of Petrology online.
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ACKNOWLEDGEMENTS |
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Samples utilized in this study were collected during a Geological Survey of Canada (GSC) project in the southern Baffin Island region led by the first author. Discussions with Rob Berman (GSC), David Corrigan (GSC), and Dave Scott (GSC) are gratefully acknowledged. We thank the technical staff at the GSC Geochronology and Scanning Electron Microscopy laboratories for their help with U–Pb analytical work and SEM imaging, and Rob Berman for his assistance in producing the two petrological phase diagrams. Helpful and thorough reviews of earlier versions of this paper were provided by Rob Berman, Geoffrey Clarke, David Corrigan, Dave Kelsey, and one anonymous reviewer. The Polar Continental Shelf Project is acknowledged for its generous helicopter and Twin Otter support during the course of the fieldwork. This paper is ESS Contribution 2005041.
Corresponding author. Telephone: (613) 995-4935. Fax: (613) 943-5318. E-mail: mstonge{at}nrcan.gc.ca
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REFERENCES |
|---|
|
TOPABSTRACTINTRODUCTIONGEOLOGICAL BACKGROUNDTECTONIC EVENTS, DEFORMATION...METHODSRESULTSDISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
|---|
Berman RG. (1988) Internally-consistent thermodynamic data for stoichiometric minerals in the system Na2O–K2O–CaO–MgO– FeO–Fe2O3–Al2O3–SiO2–TiO2–H2O–CO2. Journal of Petrology 29:445–522.
Berman RG. (1991) Thermobarometry using multi-equilibrium calculations: a new technique, with petrological applications. Canadian Mineralogist 29:833–855.[Web of Science]
Berman RG and Aranovich LY. (1996) Optimized standard state and mixing properties of minerals: 1. Model calibration for olivine, orthopyroxene, cordierite, garnet, and ilmenite in the system FeO–MgO–CaO–Al2O3–SiO2–TiO2. Contributions to Mineralogy and Petrology 126:1–24.[CrossRef][Web of Science]
Bethune KM and Scammel RJ. (2003) Geology, geochronology, and geochemistry of Archean rocks in the Eqe Bay area, north–central Baffin Island, Canada: constraints on the depositional and tectonic history of the Mary River Group of northeastern Rae Province. Canadian Journal of Earth Sciences 40:1137–1167.
Carr SD, Easton RM, Jamieson RA, Culshaw NG. (2000) Geological transect across the Grenville orogen of Ontario and New York. Canadian Journal of Earth Sciences 37:193–216.
Crawford MB and Searle MP. (1992) Field relationships and geochemistry of pre-collisional (India–Asia) granitoid magmatism in the central Karakoram, northern Pakistan. Tectonophysics 206:171–192.[CrossRef][Web of Science]
Culshaw NG, Jamieson RA, Ketchum JWF, Wodicka N, Corrigan D, Reynolds PH. (1997) Transect across the northwestern Grenville orogen, Georgian Bay, Ontario: Polystage convergence and extension in the lower orogenic crust. Tectonics 16:966–982.[CrossRef][Web of Science]
Davis WJ, McNicoll VJ, Bellerive DR, Santowski K, Scott DJ. (1997) Modified chemical procedures for the extraction and purification of U from titanite, allanite, and rutile in the Geochronology Laboratory, Geological Survey of Canada. Radiogenic Age and Isotopic Studies: Report 10; Geological Survey of Canada, Current Research 1997-F: pp. 33–35.
Davis WJ, Parrish RR, McNicoll V, Bellerive D. (1998) Analytical techniques for the determination of 208Pb/232Th ages of monazite and zircon at the Geochronology Laboratory, Geological Survey of Canada. Radiogenic Age and Isotopic Studies: Report 11; Geological Survey of Canada, Current Research 1998-F: pp. 19–22.
De Capitani C. (1994) Gleichgewichts-Phasendiagramme: Theorie und Software. Berichte der Deutschen Mineralogischen Gesellschaft 72:14.
Dunphy JM and Ludden JN. (1998) Petrological and geochemical characteristics of a Paleoproterozoic magmatic arc (Narsajuaq Terrane, Ungava Orogen, Canada) and comparisons to Superior Province granitoids. Precambrian Research 91:109–142.[CrossRef][Web of Science]
England PC and Thompson AB. (1984) Pressure–temperature–time paths of regional metamorphism. 1. Heat transfer during the evolution of regions of thickened continental crust. Journal of Petrology 25:894–928.
Essene EJ. (1989) The current status of thermobarometry in metamorphic rocks. In Daly JS, Cliff RA, Yardley BWD (Eds.). Evolution of Metamorphic Belts. Geological Society, London, Special Publications 43: pp. 1–44.
Foster G, Gibson HD, Parrish R, Horstwood M, Fraser J, Tindle A. (2002) Textural, chemical and isotopic insights into the nature and behaviour of metamorphic monazite. Chemical Geology 191:183–207.[CrossRef][Web of Science]
Fraser JE, Searle MP, Parrish RR, Noble SR. (2001) Chronology of deformation, metamorphism, and magmatism in the southern Karakoram Mountains. Geological Society of America Bulletin 113:1443–1455.
Harrison TM, Ryerson FJ, Le Fort P, An Yin P, Lovera OM, Catlos EJ. (1997) A late Miocene–Pliocene origin for the Central Himalayan inverted metamorphism. Earth and Planetary Science Letters 146:E1–E7.[CrossRef][Web of Science]
Hildebrand PR, Noble SR, Searle MP, Waters DJ, Parrish RR. (2001) Old origin for an active mountain range: geology and geochronology of the eastern Hindu Kush, Pakistan. Geological Society of America Bulletin 113:625–639.
Hoffman PF. (1988) United Plates of America, the birth of a craton: Early Proterozoic assembly and growth of Laurentia. Annual Review of Earth and Planetary Sciences 16:543–603.[CrossRef][Web of Science]
Holland TJB and Powell R. (1998) An internally consistent thermodynamic data set for phases of petrological interest. Journal of Metamorphic Geology 16:309–343.[CrossRef][Web of Science]
Huerta AD, Royden LH, Hodges KV. (1998) The thermal structure of collisional orogens as a response to accretion, erosion, and radiogenic heating. Journal of Geophysical Research 103:15287–15302.[CrossRef][Web of Science]
Huerta AD, Royden LH, Hodges KV. (1999) The effects of accretion, erosion and radiogenic heat on the metamorphic evolution of collisional orogens. Journal of Metamorphic Geology 17:349–366.[CrossRef][Web of Science]
Indares A, Dunning G, Cox R. (2000) Tectono thermal evolution of deep crust in a Mesoproterozoic continental collision setting: the Manicouagan example. Canadian Journal of Earth Sciences 37:325–340.
Jackson GD. (1969) Reconnaissance of north–central Baffin Island. Geological Survey of Canada Paper 69-1:171–176.
Jackson GD. (2000) Geology of the Clyde–Cockburn Land Map Area, North–Central Baffin Island, Nunavut. Geological Survey of Canada Memoir 440: pp. 303.
Jackson GD. (2002) Geology, Isurtuq River–Nedlukseak Fiord, Nunavut. Geological Survey of Canada Open File Map 4259: scale 1:250 000.
Jackson GD, Hunt PA, Loveridge WD, Parrish RR. (1990) Reconnaissance geochronology of Baffin Island, N. W. T. Geological Survey of Canada Paper 89-2:123–148.
Jamieson RA, Culshaw NG, Corrigan D. (1995) North-west propagation of the Grenville orogen: Grenvillian structure and metamorphism near Key Harbour, Georgian Bay, Ontario, Canada. Journal of Metamorphic Geology 13:185–207.[Web of Science]
Jamieson RA, Beaumont C, Hamilton J, Fullsack P. (1996) Tectonic assembly of inverted metamorphic sequences. Geology 24:839–842.
Jamieson RA, Beaumont C, Fullsack P, Lee B. (1998) Barrovian regional metamorphism: where's the heat? In Treloar PJ and O’Brien P (Eds.). What Drives Metamorphism and Metamorphic reactions? Geological Society, London, Special Publications 138: pp. 23–45.
Jamieson RA, Beaumont C, Ngyen MH, Lee B. (2002) Interaction of metamorphism, deformation and exhumation in large convergent orogens. Journal of Metamorphic Geology 20:9–24.[CrossRef][Web of Science]
Johnson TE, Brown M, Solar GS. (2003) Low-pressure subsolidus and suprasolidus phase equilibria in the MnNCKFMASH system: constraints on conditions of regional metamorphism in western Maine, northern Appalachians. American Mineralogist 88:624–638.
Kretz R. (1983) Symbols for rock-forming minerals. American Mineralogist 68:277–279.[Abstract]
Krogh TE. (1982) Improved accuracy of U–Pb ages by creation of more concordant systems using an air abrasion technique. Geochimica et Cosmochimica Acta 46:637–649.[CrossRef][Web of Science]
Krogh TE and Davis GL. (1973) The effect of regional metamorphism on U–Pb systems in zircon and comparison with Rb–Sr systems in the same whole rock. Carnegie Institution of Washington Yearbook 72:601–610.
Lewry JF and Collerson KD. (1990) The Trans-Hudson Orogen; extent, subdivisions, and problems. In Lewry JF and Stauffer MR (Eds.). The Early Proterozoic Trans-Hudson Orogen of North America. Geological Association of Canada, Special Papers 37: pp. 1–14.
Lucas SB. (1989) Structural evolution of the Cape Smith Thrust Belt and the role of out-of-sequence faulting in the thickening of mountain belts. Tectonics 8:655–676.[Web of Science]
Ludwig KR. (2003) Isoplot/Ex Version 3.00: a Geochronological Toolkit for Microsoft Excel. Berkeley Geochronology Center Special Publications 4: pp. 70.
Machado N, David J, Scott DJ, Lamothe D, Philippe S, Gariépy C. (1993) U–Pb geochronology of the western Cape Smith Belt, Canada: new insights on the age of initial rifting and arc magmatism. Precambrian Research 63:211–223.[CrossRef][Web of Science]
Marmo BA, Clarke GL, Powell R. (2002) Fractionation of bulk rock composition due to porphyroblast growth: effects on eclogite facies mineral equilibria, Pam Peninsula, New Caledonia. Journal of Metamorphic Geology 20:151–165.[CrossRef][Web of Science]
Moecher DP. (1999) The distribution, style and intensity of Alleghanian metamorphism in south–central New England: petrologic evidence from the Pelham and Willimantic Domes. Journal of Geology 107:449–471.[CrossRef][Web of Science]
Parrish RR. (1989) U–Pb geochronology of the Cape Smith belt and Sugluk block. Geoscience Canada 16:126–130.[Web of Science]
Parrish RR, Roddick JC, Loveridge WD, Sullivan RW. (1987) Uranium–lead analytical techniques at the Geochronology Laboratory, Geological Survey of Canada. Radiogenic Age and Isotopic Studies: Report 1; Geological Survey of Canada Paper 87-2: pp. 3–7.
Pouchou JL and Pichoir F. (1984) A new model for quantitative X-ray microanalysis. Recherche Aerospatiale 3:167–192.
Roddick JC. (1987) Generalized numerical error analysis with applications to geochronology and thermodynamics. Geochimica et Cosmochimica Acta 51:2129–2135.[CrossRef][Web of Science]
Rowley DB, Currie BS, Kidd WSF, Zhu B. (2004) India–Asia collision and Himalaya–Tibet orogenesis sensu stricto. Geological Society of America, Abstracts with Programs 36:408.
Scharer U, Xu RH, Allègre CJ. (1984) U–Pb geochronology of Gandase Transhimalyan plutonism in the Lahasa–Xigaze region, Tibet. Earth and Planetary Sciences Letters 69:311–320.
Scott DJ. (1997) Geology, U–Pb, and Pb–Pb geochronology of the Lake Harbour area, southern Baffin Island: implications for the Paleoproterozoic tectonic evolution of north-eastern Laurentia. Canadian Journal of Earth Sciences 34:140–155.
Scott DJ. (1998) U–Pb ages of Archean crust in the southeast arm of the Rae Province, southeastern Ungava Bay, Quebec. Geological Survey of Canada Paper 1998-F:41–45.
Scott DJ. (1999) U–Pb geochronology of the eastern Hall Peninsula, southern Baffin Island, Canada: a northern link between the Archean of West Greenland and the Paleoproterozoic Torngat orogen of northern Labrador. Precambrian Research 93:5–26.[CrossRef][Web of Science]
Scott DJ and Gauthier G. (1996) Comparison of TIMS (U–Pb) and laser ablation microprobe ICP-MS (Pb) techniques for age determination of detrital zircons from Paleoproterozoic metasedimentary rocks from northeastern Laurentia, Canada, with tectonic implications. Chemical Geology 131:127–142.[CrossRef][Web of Science]
Scott DJ and St-Onge MR. (1995) Constraints on Pb closure temperature in titanite based on rocks from the Ungava orogen, Canada: implications for U–Pb geochronology and P–T–t path determinations. Geology 23:1123–1126.
Scott DJ and Wodicka N. (1998) A second report on the U–Pb geochronology of southern Baffin Island. Geological Survey of Canada Paper 1998-F:47–57.
Scott DJ, St-Onge MR, Wodicka N, Hanmer S. (1997) Geology of the Markham Bay–Crooks Inlet area, southern Baffin Island, Northwest Territories. Geological Survey of Canada Paper 1997-C:157–166.
Scott DJ, Stern RA, St-Onge MR, McMullen SM. (2002) U–Pb geochronology of detrital zircons in metasedimentary rocks from southern Baffin Island: implications for the Paleoproterozoic tectonic evolution of Northeastern Laurentia. Canadian Journal of Earth Sciences 39:611–623.
Scott DJ, St-Onge MR, Corrigan D. (2003) Geology of the Archean Rae Craton and Mary River Group and the Paleoproterozoic Piling Group, Central Baffin Island, Nunavut. Geological Survey of Canada Paper 2003-C26: pp. 12.
Searle MP. (1991) Geology and Tectonics of the Karakoram Mountains(John Wiley, Chichester) pp. 358.
Searle MP. (2007) Diagnostic features and processes in the construction and evolution of Oman, Zagros, Himalayan, Karakoram and Tibetan-type orogenic belts. Geological Society of America, Special Papers (in press).
Searle MP and St-Onge MR. (2004) The Trans-Hudson Orogen of North America and the Himalaya–Karakoram–Tibetan Orogen of Asia, Part 2: Structural and thermal evolution of the upper plate. Geological Society of America, Abstracts with Programs 36:408.
Searle MP and Tirrul R. (1991) Structural and thermal evolution of the Karakoram crust. Journal of the Geological Society, London 148:65–82.
Searle MP, Rex AJ, Tirrul R, Rex DC, Barnicoat A. (1989) Metamorphic, magmatic and tectonic evolution of the central Karakoram in the Biafo–Baltoro–Hushe regions of north Pakistan. In Melinconico LL Jr. and Lillie RG (Eds.). Tectonics of the Western Himalayas. Geological Society of America, Special Papers 232: pp. 47–74.
Searle MP, Parrish RR, Tirrul R, Rex DC. (1990) Age of crystallization and cooling of the K2 gneiss in the Baltoro Karakoram. Journal of the Geological Society, London 147:603–606.
Searle MP, Asif Khan M, Fraser JE, Gough SJ. (1999) The tectonic evolution of the Kohistan–Karakoram collision belt along the Karakoram Highway transect, north Pakistan. Tectonics 18:929–949.[CrossRef][Web of Science]
Searle MP, Simpson RL, Law RD, Parrish RR, Waters DJ. (2003) The structural geometry, metamorphic and magmatic evolution of the Everest massif, High Himalaya of Nepal–South Tibet. Journal of the Geological Society, London 160:345–366.
Singh S and Jain AK. (2003) In Singh S (Ed.). Himalayan granitoids. Granitoids of the Himalayan Collisional Belt Virtual Explorer, Electronic Edition, v. 11, Paper 01.
Spear FS. (1993) Metamorphic Phase Equilibria and Pressure–Temperature–Time Paths(Mineralogical Society of America, Washington, DC).
Steiger RH and Jäger E. (1977) Subcommission on geochronology: convention on the use of decay constants in geo- and cosmochronology. Earth and Planetary Science Letters 36:359–362.[CrossRef][Web of Science]
Stern RA and Sanborn N. (1998) Monazite U–Pb and Th–Pb geochronology by high-resolution secondary ion mass spectrometry. Geological Survey of Canada Paper 1998-F:1–18.
St-Onge MR. (1984) Geothermometry and geobarometry in pelitic rocks of north–central Wopmay Orogen (early Proterozoic), Northwest Territories, Canada. Geological Society of America Bulletin 95:196–208.
St-Onge MR and Ijewliw OJ. (1996) Mineral corona formation during high-P retrogression of granulitic rocks, Ungava Orogen, Canada. Journal of Petrology 37:553–582.
St-Onge MR and Searle MP. (2004) The Trans-Hudson Orogen of North America and the Himalayan Orogen of Asia, Part 1: Structural and thermal evolution of the lower plate. Geological Society of America, Abstracts with Programs 36:408.
St-Onge MR, Lucas SB, Parrish RR. (1992) Terrane accretion in the internal zone of the Ungava orogen, northern Quebec: Part 1. Tectonostratigraphic assemblages and their tectonic implications. Canadian Journal of Earth Sciences 29:746–764.
St-Onge MR, Hanmer S, Scott DJ. (1996) Geology of the Meta Incognita Peninsula, south Baffin Island: tectonostratigraphic units and regional correlations. Geological Survey of Canada Paper 1996-C:63–72.
St-Onge MR, Lucas SB, Scott DJ, Wodicka N. (1999a) Upper and lower plate juxtaposition, deformation and metamorphism during crustal convergence, Trans-Hudson Orogen (Quebec–Baffin segment), Canada. Precambrian Research 93:27–49.[CrossRef][Web of Science]
St-Onge MR, Scott DJ, Wodicka N. (1999b) Geological maps and descriptive notes and legend, parts of Meta Incognita Peninsula, Nunavut. Geological Survey of Canada Maps 1979A to 1985A: scale 1:100 000.
St-Onge MR, Scott DJ, Lucas SB. (2000a) Early partitioning of Quebec: microcontinent formation in the Paleoproterozoic. Geology 28:323–326.
St-Onge MR, Wodicka N, Lucas SB. (2000b) Granulite- and amphibolite-facies metamorphism in a convergent-plate-margin setting: synthesis of the Quebec–Baffin segment of the Trans-Hudson Orogen. Canadian Mineralogist 38:379–398.[CrossRef][Web of Science]
St-Onge MR, Scott DJ, Wodicka N. (2001) Terrane boundaries within Trans-Hudson Orogen (Quebec–Baffin segment), Canada: changing structural and metamorphic character from foreland to hinterland. Precambrian Research 107:75–91.[CrossRef][Web of Science]
St-Onge MR, Scott DJ, Wodicka N. (2002) Review of crustal architecture and evolution in the Ungava Peninsula–Baffin Island area: connection to the Lithoprobe ECSOOT transect. Canadian Journal of Earth Sciences 39:589–610.
St-Onge MR, Wodicka N, Searle M. (2005a) The Trans-Hudson Orogen of North America and the Himalayan–Karakoram–Tibetan Orogen of Asia: structural and thermal evolution of the lower and upper plates. Geological Association of Canada, Program with Abstracts 30:188.
St-Onge MR, Scott DJ, Corrigan D, Wodicka N. (2005b) Geological maps and descriptive notes and legend, central Baffin Island area, Nunavut. Geological Survey of Canada Maps 2077A to 2082A: scale 1:100 000.
Stüwe K. (1997) Effective bulk composition changes due to cooling: a model predicting complexities in retrograde reaction textures. Contributions to Mineralogy and Petrology 129:43–52.[CrossRef][Web of Science]
Thériault RJ, St-Onge MR, Scott DJ. (2001) Nd isotopic and geochemical signature of the Paleoproterozoic Trans-Hudson Orogen, southern Baffin Island, Canada: implications for the evolution of eastern Laurentia. Precambrian Research 108:113–138.[CrossRef][Web of Science]
Tinkham DK, Zuuaga CA, Stowell HH. (2003) Metapelite phase equilibria modelling in MnNCKFMASH: the effect of variable Al2O3 and MgO/(MgO + FeO) on mineral stability. American Mineralogist 88:1174.
Treloar PJ, Rex DC, Guise PG, Coward MP, Searle MP, Windley BF, Petterson MG, Jan MQ, Luff IW. (1989) K–Ar and Ar–Ar geochronology of the Himalayan collision in NW Pakistan: constraints on the timing of suturing, deformation, metamorphism and uplift. Tectonics 8:881–909.[Web of Science]
Vannay J-C and Hodges KV. (1996) Tectonometamorphic evolution of the Himalayan metamorphic core between the Annapurna and Dhaulagiri, central Nepal. Journal of Metamorphic Geology 14:635–656.[CrossRef][Web of Science]
Walker JD, Martin MW, Bowring SA, Searle MP, Waters DJ, Hodges KV. (1999) Metamorphism, melting and extension: age constraints from the High Himalayan Slab of Southeast Zanskar and Northwest Lahaul. Journal of Geology 107:473–495.[CrossRef][Web of Science]
Weinberg RF and Dunlap WJ. (2000) Growth and deformation of the Ladakh Batholith, northwest Himalayas: implication for timing of continental collision and origin of calc-alkaline batholith. Journal of Geology 108:303–320.[CrossRef][Web of Science][Medline]
White RW, Powell R, Holland TJB. (2001) Calculation of partial melting equilibria in the system Na2O–CaO–K2O–FeO–MgO–Al2O3–SiO2–H2O (NCKFMASH). Journal of Metamorphic Geology 19:139–153.[CrossRef][Web of Science]
Wodicka N and Scott DJ. (1997) A preliminary report on the U–Pb geochronology of the Meta Incognita Peninsula, southern Baffin Island, Northwest Territories. Geological Survey of Canada Paper 1997-C:167–178.
Wodicka N and St-Onge MR. (1998) Implications of U–Pb geochronology for the thermal evolution of supracrustal rocks from the Baffin segment of the Trans-Hudson Orogen. Geological Association of Canada, Program with Abstracts 23:A200.
Wodicka N, Ketchum JWF, Jamieson RA. (2000) Grenvillian metamorphism of monocyclic rocks, Georgian Bay, Ontario, Canada: implications for convergence history. Canadian Mineralogist 38:471–510.[CrossRef][Web of Science]
Wodicka N, Madore L, Larbi Y, Vicker P. (2002a) Géochronologie U–Pb de filons-couches mafiques de la Ceinture de Cape Smith et de la Fosse du Labrador. Séminaire d’information sur la recherche géologique, Programme et résumés 2002. Ministère des Ressources naturelles, Québec V 2002-1048.
Wodicka N, St-Onge MR, Scott DJ, Corrigan D. (2002b) Preliminary report on the U–Pb geochronology of the northern margin of the Trans-Hudson Orogen, central Baffin Island, Nunavut. Geological Survey of Canada Paper 2002-F7: pp. 12.
York D. (1969) Least squares fitting of a line with correlated errors. Earth and Planetary Science Letters 5:320–324.[CrossRef][Web of Science]
Zhao G, Cawood PA, Wilde SA, Min Sun SA. (2002) Review of global 2·1–1·8 Ga orogens: implications for a pre-Rodinia supercontinent. Earth-Science Reviews 59:125–162.
Zhu XK, O’Nions RK, Belshaw NS, Gibb AJ. (1997) Significance of in situ SIMS chronometry of zoned monazite from the Lewisian granulites, northwest Scotland. Chemical Geology 135:35–53.[CrossRef][Web of Science]
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  M. R. St-Onge, J. A. M. Van Gool, A. A. Garde, and D. J. Scott Correlation of Archaean and Palaeoproterozoic units between northeastern Canada and western Greenland: constraining the pre-collisional upper plate accretionary history of the Trans-Hudson orogen Geological Society, London, Special Publications, January 1, 2009; 318(1): 193 - 235. [Abstract] [Full Text] [PDF]
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