Evolution of an Archean basement complex and its autochthonous cover, southern Slave Province, Canada

Precambrian Research 135 (2004) 149–176

Evolution of an Archean basement complex and its autochthonous cover, southern Slave Province, Canada John W.F. Ketchuma,∗ , Wouter Bleekerb , Richard A. Sternb,1 a

Jack Satterly Geochronology Laboratory, Royal Ontario Museum, 100 Queen’s Park, Toronto, Ontario, Canada M5S 2C6 b Geological Survey of Canada, 601 Booth Street, Ottawa, Ontario, Canada K1A 0E8 Received 24 July 2003; accepted 25 August 2004

Abstract The Sleepy Dragon Complex (SDC) is one of several antiformal exposures of a 4.0–2.85 Ga basement block underlying the west-central region of the Archean Slave Province, northwestern Canada. This basement block, the Central Slave Basement Complex (CSBC), is overlain by an autochthonous to locally parautochthonous, dominantly quartz-rich clastic sequence, the 2.85–2.80 Ga Central Slave Cover Group. Together these units comprised a fundamental building block during Neoarchean growth and assembly of the Slave craton. The well-exposed Patterson Lake–Morose Lake area represents a type locality for study of the CSBC, Central Slave Cover Group, and overlying rocks of the Yellowknife Supergroup. Field mapping and U–Pb geochronology, employing both thermal ionisation mass spectrometry (TIMS) and sensitive high resolution ion microprobe (SHRIMP), document a prolonged crustal history. Oldest intact basement units consist of foliated to gneissic tonalite dated in two places at 2955 ± 12 Ma and 2944 ± 9 Ma. Indirect evidence for a ca. 3150 Ma basement component is obtained both from xenocrystic zircon in tonalite and detrital zircons from the overlying Central Slave Cover Group, which here consists of the Patterson Lake Formation. A granite boulder from this formation is dated at 2934 ± 3 Ma and may be locally derived. A sheared mafic volcanic unit along the basement-cover contact contains 2942 ± 3 Ma metamorphic titanite which provides a minimum deposition age. This volcanic unit is therefore also a part of the basement complex. Younger events in the SDC mainly reflect tonalite-trondhjemite-granodiorite plutonism during construction of the overlying Cameron River and Beaulieu River greenstone belts of the Yellowknife Supergroup. A unit of K-feldspar megacrystic granodiorite is dated at 2726 ± 3 Ma, and tonalite and granodiorite bodies located toward the centre of the SDC have primary crystallization ages of 2683 + 3/−2 Ma, 2677 ± 2 Ma, and 2672 + 7/−6 Ma. The core of the southern SDC is occupied by a late-tectonic to post-tectonic granite pluton previously dated at 2586 ± 2 Ma. An intravolcanic unconformity along the southern margin of the SDC (developed in part on 2.7 Ga granodiorite) is constrained by existing U–Pb data to have formed sometime between

∗ Corresponding author. Present address: GEMOC ARC National Key Centre, Department of Earth and Planetary Sciences, Macquarie University, New South Wales 2109, Australia. Fax: +612 9850 8943. E-mail address: [email protected] (J.W.F. Ketchum). 1 Present address: Centre for Microscopy and Microanalysis, M010, The University of Western Australia, 35 Stirling Highway, Crawley, Western Australia 6009, Australia.

0301-9268/$ – see front matter © 2004 Elsevier B.V. All rights reserved. doi:10.1016/j.precamres.2004.08.005

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ca. 2685–2660 Ma. An abundance of synvolcanic plutons of similar age and the presence of a >3-km thick mafic volcanic sequence overlying a less dense, magmatically active felsic crust suggest that gravitational instability and magmatic diapirism may have played important roles in doming and uplift of the SDC prior to regional D1 deformation. The younger-toward-centre zonation of plutonic rocks in the SDC is documented in other domal granite–gneiss complexes of Archean age (e.g., Pilbara craton, western Australia) where it is attributed to upward and outward displacement of older crust during repeated magma channeling into a growing mid-crustal dome. In the west-central Slave craton, such a process might account for the regional presence of local unconformities of apparently similar age. The domal basement–plutonic complexes would have been further amplified after 2660 Ma during the development of regional F1 - and F2 -fold belts. © 2004 Elsevier B.V. All rights reserved. Keywords: Slave Province; Archean basement complex; U–Pb geochronology; TIMS; SHRIMP; Magmatic diapirism

1. Introduction Regions of variably deformed, heterogeneous granitoid rocks, here termed granite–gneiss domains, underlie a significant portion of Archean cratons. Models of Archean cratonic development require a thorough understanding of granite–gneiss domains, particularly in regions where their oldest components predate earliest supracrustal rocks. This relationship, which potentially signals the presence of sialic basement to some or all of the supracrustal assemblages, is of fundamental importance in determining the origin of cratons. However, ancient continental basement components may be difficult to identify due to younger magmatic, structural, erosional, or depositional events. This raises the possibility that sialic basement may go unrecognized in some regions or in entire cratons (e.g., Bleeker, 2002). Detailed field, isotopic, and geochemical investigations are required to ‘see through’ younger events to demonstrate or at least strongly infer the presence of basement. It is noteworthy that such studies (e.g., Arndt, 1999; Horstwood et al., 1999; Bleeker et al., 1999a; Van Kranendonk et al., 2001) inexorably yield a net increase in the global volume of ancient crust. Ongoing discoveries of this ‘hidden Archean’ have the potential to influence our understanding of crustal evolution and inferred rates of early continental growth. The Archean Slave Province, northern Canada (Fig. 1), is underlain in central and western regions by a Hadean to Mesoarchean basement whose oldest and best-known components are the 4.0 Ga Acasta gneisses (Bowring et al., 1989; Stern and Bleeker, 1998; Bowring and Williams, 1999). These

rocks and younger gneisses and foliated granitoid plutons form both a stratigraphic and structural basement to widespread 2.73–2.60 Ga volcanic belts and sedimentary assemblages of the Yellowknife Supergroup (Henderson, 1970). The basement rocks are mainly exposed in structural culminations that form granite–gneiss domains. A recent tectonic model for the west-central Slave craton (Bleeker and Ketchum, 1998; Bleeker et al., 1999a,b, 2000) unites all basement rocks at ca. 2.9 Ga in a single block, the Central Slave Basement Complex (CSBC). Pre-2.9 Ga components have been confirmed in many regions by U–Pb dating (e.g., Krogh and Gibbins, 1978; Lambert and van Breemen, 1991; Northrup et al., 1999; Yamashita et al., 2000; Ketchum and Bleeker, 2000), but are more readily identified where they are overlain by a lithologically distinctive, autochthonous to locally parautochthonous supracrustal assemblage. Discontinuous occurrences of this assemblage, which is dominated by quartz-rich siliciclastic rocks, comprise the Central Slave Cover Group (Bleeker and Ketchum, 1998; Bleeker et al., 1999a). The Central Slave Cover Group was deposited between 2.85 Ga and 2.80 Ga (Isachsen and Bowring, 1997; Ketchum and Bleeker, 2000; Sircombe et al., 2001) and is overlain by autochthonous to parautochthonous mafic volcanic rocks of the Yellowknife Supergroup (Bleeker et al., 1999b; Bleeker, 2002). Widespread occurrences of the Central Slave Cover Group mark the regional extent of the CSBC (Fig. 1). Based on geological and isotopic data from the Slave Province and Paleoproterozoic Wopmay orogen (e.g., Hildebrand et al., 1990; Davis and Hegner, 1992; Davis et al., 1996; Yamashita et al., 1999; Ketchum and Bleeker, un-

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Fig. 1. Geological map of Slave craton, Northwest Territories. Various greenstone belts and basement gneiss complexes are indicated, along with occurrences of the autochthonous to parautochthonous, 2.85–2.80 Ga Central Slave Cover Group. PL, BL indicate locations of the Patterson Lake and Brown Lake formations of this group, respectively. TL indicates location of Takijuq Lake. The Pb and Nd isotopic boundaries have been interpreted to broadly indicate the surface and subsurface positions, respectively, of the boundary between the Paleoarchean to Neoarchean western Slave block and the predominantly Neoarchean eastern Slave block (Thorpe et al., 1992; Davis and Hegner, 1992). The inferred surface and subsurface extent of the Central Slave Basement Complex (CSBC) is shown in an inset map. The present study area is located in the southern part of the Sleepy Dragon Complex at Patterson Lake. Modified from Bleeker et al. (1999a).

published data), the minimum subsurface extent of the CSBC is inferred to be much larger (inset map, Fig. 1). The combination of a large basement block and excellent bedrock exposure makes the Slave craton an

ideal location to study an ancient basement complex and its earliest supracrustal cover. In this paper we report U–Pb data from a detailed study of the Sleepy Dragon Complex (SDC; Davidson, 1972; Henderson, 1985) and overlying Patterson Lake Formation of the

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Fig. 2. Regional geological map with location of study area (Fig. 3) indicated by box. U–Pb ages are from previous studies discussed in text. Abbreviations: CRB, Cameron River greenstone belt; BRB, Beaulieu River greenstone belt; RLF, Raquette Lake Formation; DLF, Detour Lake Formation. Modified from Bleeker et al. (1999b).

Central Slave Cover Group (Bleeker et al., 1999a) in the Patterson Lake––Morose Lake area of the southern Slave Province (Figs. 1–3). This is a classic locality for early descriptions of Archean basement lithologies and basement-cover relationships (e.g., Baragar, 1966; Davidson, 1972; Baragar and McGlynn, 1976; Henderson, 1985; Kusky, 1990), although rocks of the Patterson Lake Formation were not always recognized. We outline a detailed chronology of events which should serve as a benchmark for comparisons with basement-cover transitions elsewhere in the westcentral Slave Province. Age data presented below provide several important constraints on the tectonothermal evolution of the CSBC, which represents a protocraton to a much larger continental block that was assembled during the Neoarchean.

2. Regional geology 2.1. Sleepy Dragon Complex (SDC) The SDC forms a prominent granite–gneiss domain that is mantled by the Cameron River and Beaulieu River volcanic belts of the Yellowknife Supergroup (Figs. 2 and 3). Like other granite–gneiss domains of the west-central Slave Province, the SDC consists mainly of three components: (i) >2.9 Ga gneissic and foliated basement rocks of the CSBC; (ii) foliated, mainly synvolcanic granitoid plutons emplaced between ca. 2.73 and 2.66 Ga; and (iii) foliated to massive, ca. 2.64–2.58 Ga granitoid plutons (e.g., van Breemen et al., 1992; Davis and Bleeker, 1999). Some granite–gneiss domains also contain 2.85–2.82 Ga

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Fig. 3. Map of study area showing locations and ages of dated units. Number preceding age corresponds to sample number given in text, Tables 1 and 2, and Figs. 5, 7, and 8. Additional U–Pb ages (indicated by superscript numbers) are from: (1) Bleeker et al. (1999b) and (2) Davis and Bleeker (1999). Modified from Bleeker et al. (1999a).

gneissic to foliated plutonic rocks (Henderson et al., 1987; Northrup et al., 1999; Ketchum and Bleeker, 2001a), but whether these represent depositional basement to the Central Slave Cover Group depends on the precise ages of plutonic and cover group rocks in each area. The 2.73–2.58 Ga granitoid plutons (groups (ii) and (iii) above) are typically the dominant components of the granite–gneiss domains. Within the SDC, basement rocks comprise a heterogeneous assemblage of foliated to migmatitic, upper amphibolite facies tonalite, diorite, granodiorite, granite, and gabbro. These units were deformed and metamorphosed prior to intrusion of younger granitoid bodies and deposition of the flanking volcanic belts (Davidson, 1972; Lambert, 1988; James and Mortensen, 1992; Bleeker et al., 1999a,b). Details of early tectonometamorphic events are poorly known. The oldest basement unit is tonalite gneiss with an age

>3.0 Ga (James and Mortensen, 1992; all cited ages are based on U–Pb dating). Other previously dated basement units (Fig. 2) include granodiorite gneiss along the east side of the complex (2936 + 17/−14 Ma; Lambert and van Breemen, 1991), and megacrystic granite gneiss located immediately northeast of the present study area (2819 + 40/−31 Ma; Henderson et al., 1987). Nd isotopic data from gneissic rocks suggest the presence of crust as old as ca. 3.3 Ga (Dud´as et al., 1988). Younger plutons are mainly tonalitic to granitic. Dated bodies include 2683.5 ± 2 Ma foliated granodiorite (Henderson et al., 1987), strongly deformed granite emplaced at 2641 ± 3.5 Ma (James and Mortensen, 1992), and ca. 2585 Ma, massive to weakly foliated, K-feldspar megacrystic granite (James and Mortensen, 1992; Davis and Bleeker, 1999; Fig. 2). Emplacement of these plutons is contemporaneous with known

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times of volcanic and plutonic activity outside the SDC. Metadiabase dykes are common within the SDC and are particularly abundant in the southwestern part of the complex (Baragar, 1966; Henderson, 1985; Lambert, 1988; Lambert et al., 1992; Bleeker et al., 1997). Baragar (1966) first proposed that these dykes represent feeders to overlying mafic volcanic rocks, a conclusion that has been largely confirmed by subsequent field and isotopic studies (e.g., Lambert et al., 1992; Bleeker et al., 1997, 1999b; W. Bleeker, unpublished field data). Bleeker et al. (1997, 1999b) presented field and U–Pb evidence for three temporally distinct mafic dyke swarms in the Patterson Lake area. These dykes are described in more detail below. Younger fabrics and structures that overprint both the SDC and adjacent Yellowknife Supergroup are related mainly to deformation in the vicinity of the basement-cover contact (Kusky, 1990; James and Mortensen, 1992; Bleeker et al., 1997, 1999b) and to development of a regional D1 –D2 -fold interference pattern that is partly responsible for the domal form of the SDC (Bleeker and Beaumont-Smith, 1995; Bleeker, 1996; Fig. 2). 2.2. Supracrustal rocks overlying the Sleepy Dragon Complex 2.2.1. Central slave cover group Sporadic occurrences of deformed, quartz-rich siliciclastic rocks and banded iron formation along the base of western Slave greenstone belts have been documented over several decades (Easton et al., 1982; Helmstaedt and Padgham, 1986; Covello et al., 1988; Kusky, 1989; Padgham, 1992; Thompson et al., 1995). In all places where stratigraphic tops can be determined, these rocks face away from basement, as do overlying volcanic belts (Padgham and Fyson, 1992). Based on work in a few localities, the siliciclastic rocks have been interpreted as remnants of a continental shelf assemblage (Covello et al., 1988; Kusky, 1989; Rice et al., 1990; Padgham, 1992; Isachsen and Bowring, 1994; Pickett and Mueller, 2000). Their lithologic association and characteristics are similar to those of other Archean platformal successions such as those documented in the western Superior Province (Thurston and Chivers, 1990; Percival et al., 2001). A detailed examination of variably tectonized basement-cover con-

tacts across the central and western Slave Province (Bleeker and Ketchum, 1998; Bleeker et al., 1999a, 2000) led to additional discoveries of these rocks and recognition that they form a <200-m thick but regionally extensive cover sequence. This autochthonous to parautochthonous package, the Central Slave Cover Group, may extend as far north as the Arctic Ocean, outlining a Mesoarchean block of considerable extent (Fig. 1). Bleeker et al. (1999a) introduced formation names for regional occurrences of the Central Slave Cover Group, and of these, the Brown Lake, Patterson Lake, and Amacher Lake formations overlie the SDC (Figs. 1 and 2). These formations are marked by up to 100-m thick packages of conglomerate, quartzrich arkosic grit, quartz arenite (±fuchsite, detrital chromite), minor siltstone and pelite, and a capping unit of silicate and/or oxide facies banded iron formation. Thin mafic volcanic units and ultramafic sills or flows are known in several localities, typically near the base of the package (e.g., Lambert and van Staal, 1987; Kusky, 1990; Bleeker et al., 1997; however, see below). Discontinuous exposure of the cover group, both around the SDC and elsewhere, is due to one or more of: (i) structural thinning and excision within high strain zones, (ii) occlusion by plutons, and (iii) removal by erosion at unconformities (Bleeker et al., 1999a,b). Local non-deposition may also be a factor. Contact relationships between the SDC and the Central Slave Cover Group are varied. In many regions the contact is highly tectonized with no clear-cut evidence of a former stratigraphic relationship (e.g., Lambert, 1988; James and Mortensen, 1992). In these areas the high-strain fabrics typically overprint uppermost basement rocks, the Central Slave Cover Group, and lowermost basaltic units of the Yellowknife Supergroup. However, in some lower strain areas, an original unconformable relationship is preserved. This is best demonstrated at Brown Lake (BL, Fig. 1) where quartz cobble conglomerate overlies foliated basement tonalite dated at 2908 ± 2 Ma (Ketchum and Bleeker, 2000). Abundant sillimanite and andalusite in tonalite adjacent to the unconformity suggest the presence of a paleoweathering surface (Bleeker et al., 1999a). Although this unconformity is deformed, the locus of greatest ductile strain occurs in adjacent basement rocks, accounting for local preservation of the unconformity (see also James and Mortensen, 1992).

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Direct dating of the Central Slave Cover Group is limited to U–Pb zircon ages of rare felsic volcanic members found in two occurrences of this group at the base of the Yellowknife greenstone belt (Fig. 1; Ketchum and Bleeker, 2000). At Dwyer Lake, felsic tuff with a primary age of 2853 + 2/−1.5 Ma stratigraphically overlies the basal quartzite, providing a minimum age for its deposition. Eighteen km to the north at Bell Lake, a felsic tuff layer within banded iron formation near the top of the cover group was deposited at 2826 ± 1.5 Ma. In conjunction with a maximum deposition age of 2826 ± 3 Ma for a quartzite layer near the top of the Patterson Lake Formation (Sircombe et al., 2001), the U–Pb data indicate deposition of the cover group over at least 25 Myr. These data may also indicate regionally diachronous deposition (Sircombe et al., 2001). 2.2.2. Cameron River and Beaulieu River greenstone belts The SDC is for the most part directly overlain by rocks of the Cameron River and Beaulieu River greenstone belts (Figs. 2 and 3). These multiply deformed volcanic belts are described in detail by Henderson (1985) and Lambert (1988) and are considered to be broadly correlative. Each is characterized by a lower tholeiitic package dominated by pillowed basalt and mafic dykes and sills, and an upper calc-alkaline package containing flows and breccias of andesite, dacite, and rhyolite. Mafic to felsic tuff and sedimentary rocks form minor components. The belts typically have a highly strained lower contact with the SDC and are in conformable contact with the overlying, turbiditic Burwash Formation. The two-part subdivision of the Cameron and Beaulieu River belts mimics that observed in the Yellowknife greenstone belt to the west (tholeiitic Kam Group and overlying, calc-alkaline Banting Group; Helmstaedt and Padgham, 1986). Although the belts are not correlative in all details (MacLachlan and Helmstaedt, 1995), rhyolite flows at comparably high stratigraphic levels in the Beaulieu River and Yellowknife belts yield identical ages of ca. 2663 Ma (Henderson et al., 1987; Isachsen and Bowring, 1994). There are currently no other precise age data for the Cameron River and Beaulieu River belts, but comparison with the Yellowknife belt (Isachsen and Bowring, 1997) suggests that the lower

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tholeiitic assemblage may contain rocks as old as 2.72 Ga. Lithotectonic models for the Cameron River and Beaulieu River greenstone belts can be broadly grouped under allochthonous and (par)autochthonous headings. Although Kusky (1989, 1990) considered these belts to represent segments of a far-travelled oceanic allochthon, other workers (e.g., Henderson, 1985; Lambert, 1988; Lambert et al., 1992; Bleeker et al., 1999b; Bleeker, 2002) favour eruption onto sialic basement represented by the CSBC, a view now supported by a variety of structural, geochemical, isotopic, and geochronological data (see Bleeker, 2002). Structural and geochronologic data from the SDC basementcover interface (Kusky, 1990; Bleeker et al., 1999b) suggest tectonic transport of these greenstone belts sometime between 2734–2687 Ma, but the maximum amount of displacement is suggested to be on the order of a few tens of km (Bleeker et al., 1999b). Although coeval transport of the underlying Central Slave Cover Group occurred at least locally during this event, the exact position of the major dislocation surface cannot always be identified within the basementcover high-strain zone. Hence in some places the entire supracrustal package has been transported (as at Brown Lake; see above), whereas in others only the greenstone belts are likely to have been displaced. 2.2.3. Raquette Lake and Detour Lake formations (Ross Lake Group) Two volcano-sedimentary formations unconformably overlie the SDC along its southern margin. The Raquette Lake Formation (Fig. 2) consists of a heterogeneous assemblage of conglomerate, quartz-rich arenite, felsic tuff, and calcareous sedimentary rocks. These units comprise a shallow-water, 5-km long package with a maximum thickness of 60 m (Henderson, 1985). However, Bleeker (2001) and Mueller and Corcoran (2001) include overlying rocks (e.g., rhyolite tuff and mudstone) within the Raquette Lake Formation, yielding a total thickness of 150–200 m. Deposition of this formation was contemporaneous with the latter stages of Cameron River volcanic activity (Henderson, 1985), specifically with eruption of a thin rhyolite unit at the base of the upper, calc-alkaline package (Cameron River Rhyolite; Bleeker, 1996). The unconformity at the base of the Raquette Lake Formation also occurs within the Cameron River belt

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between the tholeiitic and calc-alkaline assemblages. This unconformity progressively cuts out tholeiitic basalt toward the southeast until Raquette Lake lithologies directly overlie the SDC (Bleeker, 1996, 2001), which at this location consists of the Ross Lake granodiorite (ca. 2.7 Ga; T. Krogh, unpublished data; W. Davis, personal communication, 2002). Although Mueller and Corcoran (2001) regard the Cameron River basalt and Raquette Lake Formation as lateral facies equivalents, the presence of numerous gabbroic sills and cross-cutting dykes in the former and their absence in the latter (Bleeker, 1996) provides considerable support for the proposed intravolcanic unconformity. The Raquette Lake Formation was deposited between 2683 ± 1 Ma and 2661 ± 2 Ma (Bleeker and Villeneuve, 1995; Bleeker et al., 1997). Detrital zircons from quartzite define two age groups at ca. 2685 Ma and ca. 2935 Ma (Bleeker et al., 1997), identical to the ages of dated units in the underlying SDC. Southeast of the Raquette Lake Formation (Fig. 2), clastic, volcaniclastic, carbonate-rich, and minor felsic volcanic rocks adjacent to the SDC were named the Detour Lake Group by Kusky (1990) and the Detour Lake Formation by Bleeker (2001). Kusky (1990) suggested that these rocks did not resemble the Raquette Lake Formation and represented the oldest supracrustal units in the region. However, based on extensive mapping, Davidson (1972), Henderson (1985), Lambert (1988), and Bleeker (2001) concluded that these rocks are relatively young and were deposited during and after volcanic activity. The unconformity at the base of the Detour Lake Formation (e.g., Kusky, 1990) is continuous with the Raquette Lake unconformity and also occurs on ca. 2.7 Ga granodiorite (Bleeker, 2001). The two formations are considered to be correlative and form the principal components of a new lithostratigraphic unit, the Ross Lake Group (Bleeker, 2001).

3. Geology of the Patterson Lake–Morose Lake area In the Patterson Lake area, the SDC is broadly divisible into marginal and interior domains dominated by basement rocks and younger granitoid plutons, respectively (Fig. 3). A prominent basement unit of the marginal domain consists of strongly foliated to locally gneissic tonalite with layers of more mafic tonalite and

quartz diorite. The tonalite generally has a porphyroclastic texture due to the presence of numerous small plagioclase phenocrysts. Mylonitic fabrics are variably developed within this unit and become progressively more intense toward the basement-cover contact. Variably foliated, K-feldspar megacrystic granodiorite is also abundant near this contact. The granodiorite consists of at least three rock types that differ mainly in abundance of K-feldspar megacrysts, with internal contacts ranging from sharp to gradational. Limited field observations suggest that the granodiorite intrudes the tonalite. This is borne out by U–Pb data presented below. The interior domain is underlain by foliated biotite granodiorite, biotite-hornblende tonalite, and minor diorite plutons (Bleeker et al., 1997). For clarity, these are depicted in Fig. 3 as a single map unit. Field relationships and U–Pb data indicate that these rocks are younger than basement tonalites but older than the lateto post-tectonic Morose Granite (Henderson, 1985), a large body of K-feldspar megacrystic biotite granite located near the centre of the SDC and dated at 2586 ± 2 Ma (Davis and Bleeker, 1999). This unit is also considered here to be part of the interior domain. The 6-km long Patterson Lake Formation has been described in detail by Bleeker et al. (1997, 1999a). This ∼100-m thick formation is truncated to the southwest by a fault and pinches out to the northeast (Fig. 3). A representative stratigraphic section is shown in Fig. 4. The basal unit is an impure, locally fuchsitic quartzite that in several places hosts deformed lenses of metamorphosed ultramafic rock. Progressively up section are quartz-rich conglomerate with granitoid, chert, vein quartz, and mafic volcanic clasts, impure quartzite, and fuchsite-bearing orthoquartzite. These units are overlain by silicate- and oxide-facies, thinly banded iron formation, locally with graded siliciclastic layers. The siliciclastic layers also occur with siltstone at the top of the section in upward-facing graded beds. A strongly deformed mafic volcanic unit, locally exhibiting stretched pillows, underlies the clastic succession and was provisionally included in the Patterson Lake Formation by Bleeker et al. (1999a). However, age data presented below demonstrate that this volcanic unit is a component of the basement complex. Three temporally distinct mafic dyke swarms intrude the SDC in the study area and form important markers for determining the relative and absolute ages

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swarm 1 dyke yield a U–Pb upper intercept age of 2734 ± 2 Ma which is interpreted as the dyke emplacement age (Bleeker et al., 1999b). This is a minimum age for deposition of the Patterson Lake Formation and a maximum age for high-strain deformation associated with oblique normal movement on the basementcover contact (Bleeker et al., 1999b). A minimum age for this movement is provided by abundant, crosscutting, north- to northeast-trending dykes of swarm 2, one of which is dated at 2687 ± 1 Ma (Bleeker et al., 1999b). Swarm 2 dykes extend into the Cameron River basalts and therefore stitch together basement, autochthonous cover, and overlying mafic volcanic rocks by 2687 Ma. The youngest, north- to northwesttrending mafic dykes (swarm 3) are also abundant within these units. This swarm has not been directly dated but intrudes ca. 2684 Ma granodiorite northeast of the study area (Henderson et al., 1987; Bleeker et al., 1999b). All three dyke swarms are largely absent from the interior domain; the only exception noted during this study was probable swarm 3 dykes that cut a large tonalite enclave in the Morose Granite (see below).

4. Sampling strategy

Fig. 4. Reconstructed stratigraphy for the Patterson Lake Formation, Patterson Lake area, drawn approximately to scale. U–Pb data constrain the basement-cover contact (a strongly deformed unconformity) to the top of a mafic schist unit derived from pillow basalt. Locations of dated samples 1, 8, and 9 indicated. Modified from Bleeker et al. (1999a).

of various events. Cross-cutting relationships between all three swarms are best observed along the wellexposed shores of Patterson and Webb lakes (Fig. 2). The oldest dykes (swarm 1 of Bleeker et al., 1997, 1999b) are transposed within the basement foliation and do not appear to cut an older fabric, although this relationship is somewhat equivocal due to high superimposed strain. These dykes are present in similar abundance in the Patterson Lake Formation but have not been observed in overlying basalts of the Cameron River belt. Bladed, skeletal zircons from a

Field relationships in the study area suggest a protracted, multistage evolution for the SDC (Henderson, 1985; James and Mortensen, 1992; Bleeker et al., 1997, 1999a,b). To determine absolute ages of plutonic, structural, and metamorphic events, seven samples covering the spectrum of mapped lithologies were collected for U–Pb dating in a transect extending from the centre to the western margin of the SDC. Two additional samples were collected from the Patterson Lake Formation to determine the age of detrital zircons and a granite clast, and to constrain maximum deposition age. This follows on earlier detrital zircon dating of this formation by Bleeker et al. (1999a) and Sircombe et al. (2001). Results are presented below in order from oldest to youngest primary age.

5. U–Pb methodology Isotope dilution-thermal ionization mass spectrometry (ID-TIMS) of single and multigrain fractions of zircon and titanite was carried out at both the Memorial

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Fig. 5. U–Pb concordia diagrams of thermal ionization mass spectrometry (TIMS) and sensitive high resolution ion microprobe (SHRIMP) data from the same samples, all from the Sleepy Dragon Complex. Ages given in box represent best estimate of primary crystallization age. All SHRIMP analyses are of zircon whereas TIMS analyses include zircon (Z) and titanite (T). Age uncertainties and ellipses are 2␴. See text for additional details.

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University of Newfoundland (MUN) and the Royal Ontario Museum (ROM). Standard mineral separation, picking, abrasion, dissolution, chemical separation, and mass spectrometric procedures were followed and are described in detail elsewhere (e.g., Ketchum et al., 1997). A notable variation in technique was the lack of ion exchange column chemistry for single-grain zircon fractions analysed at the ROM. This modification appears to yield no adverse effects during mass spectrometry as long as analysed grains are small (<5 ␮g). At MUN, isotope ratios were measured with a FinniganMAT 262 mass spectrometer using four Faraday cups and a secondary electron multiplier-ion counter in multicollection mode. At the ROM, a VG 354 instrument with a single Daly detector was employed in peak jumping mode. Ages were calculated using in-house software and where necessary, used the Pb model isotopic compositions of Stacey and Kramers (1975) to correct for common Pb in excess of laboratory blank. Linear regressions of discordant data follow the method of Davis (1982). All uncertainties given in Table 1 and shown as TIMS error ellipses in Figs. 5, 7 and 8 are given at the 95% confidence level. A follow-up investigation of zircon populations was carried out using the SHRIMP II at the Geological Survey of Canada (GSC). Application of both the ID-TIMS and SHRIMP techniques is advantageous when intragrain and intergrain complexities are apparent in a zircon population, as is the case for several samples in this study. We used the SHRIMP to investigate the ages of potential inherited components and metamorphic overgrowths in three plutonic samples, and to address uncertainty in the interpretation of multigrain ID-TIMS data from a fourth sample. Zircons selected for detailed examination from the plutonic samples were generally larger grains (from both non-magnetic and paramagnetic fractions) with visible evidence of internal complexity. In most instances this complexity suggested the presence of inherited zircon cores, but in only a few cases did these consist of older xenocrystic material. SHRIMP analytical procedures followed those described by Stern (1997), and the standards and U–Pb calibration methods followed Stern and Amelin (2002). Zircons were cast in 2.5 cm diameter epoxy mounts (GSC #IP218, #IP219) along with fragments of the GSC laboratory standard zircon z6266 (206 Pb/238 U age = 559 Ma). The mid-sections of the zircons were exposed using 9, 6, and 1 ␮m diamond compound,

159

and the internal features of the zircons (such as zoning, structures, alteration, etc.) were characterized with backscattered electrons using a Cambridge Instruments scanning electron microscrope. Mount surfaces were evaporatively coated with 10 nm of high purity Au. Analyses were conducted using an O− primary beam, projected onto the zircons at 10 kV as an elliptical spot of ca. 15 ␮m diameter and with a beam current of ca. 4 nA and uniform density. The count rates of 10 isotopes or polyatomic ions of Zr+ , U+ , Th+ , and Pb+ in zircon were sequentially measured (six scans) with a single electron multiplier and a pulse counting system with deadtime of 31 ns. Mass resolution was 5500 (1%) and 206 Pb sensitivity ∼ 9 cps/ppm/nA in zircon. Off-line data processing was accomplished using customized in-house software. The 1␴ external errors of 206 Pb/238 U ratios reported in Table 2 incorporate a ±1.0% error in calibrating the standard zircon (see Stern and Amelin, 2002). No fractionation correction was applied to the Pb-isotope data; common Pb corrections used the measured 204 Pb/206 Pb and compositions modelled after Cumming and Richards (1975). Isoplot v. 2.49 (Ludwig, 2001) was used to generate concordia plots and calculate weighted mean ages. 6. U–Pb data from the Sleepy Dragon Complex 6.1. Sample 1: mafic volcanic schist, Patterson Lake A 10–30-m wide unit of highly deformed mafic schist occurs beneath lowermost siliciclastic rocks of the Patterson Lake Formation (Fig. 4) and is well exposed on a small island at the north end of Patterson Lake. Here, thin alternating bands of mafic and intermediate composition appear to represent highly stretched pillows and interpillow (tuffaceous?) material, respectively, although this interpretation could not be verified. A sample containing both components was collected on the chance that zircon might be present in the intermediate composition material. The sample yielded numerous small, euhedral to subhedral colourless zircon prisms with square crosssections, and colourless, weakly to moderately turbid, subhedral titanite. Because of the small grain size, multigrain zircon fractions were chosen for TIMS analysis. The six analysed fractions have U concentrations

160

Table 1 TIMS U–Pb isotopic data, Patterson–Morose Lakes area Pbc (pg)

206 Pb/

206 Pb/

204 Pb

238 U

20 9 8 6 4 16 301 366

234 170 237 663 290 319 187 136

0.5693 0.5194 0.5328 0.5587 0.5371 0.5505 0.5742 0.5799

0.0066 0.0070 0.0061 0.0045 0.0089 0.0054 0.0023 0.0024

16.426 14.252 14.392 15.704 14.266 15.286 17.010 17.177

2␴

207 Pb/

2␴

207 Pb/ 206 Pb (Ma)

2␴ (Ma)

Discord. (%)

0.190 0.214 0.195 0.113 0.249 0.135 0.076 0.075

2899.9 2817.9 2783.1 2857.5 2764.7 2837.5 2942.4 2942.4

5.5 11.1 7.1 6.6 10.3 8.7 2.9 3.0

−0.2 5.3 0.6 −0.2 −0.3 0.4 0.7 −0.3

235 U

Sample 2 (BNB95-50): tonalite gneiss, Webb Lake (393400 E/6971000 N) Z1 sm lbr pr (4) 0.002 704 0.27 Z2 sm lbr pr (4) 0.002 1444 0.30 Z3 2nd best br pr (1) 0.001 565 0.21 Z4 2nd best br pr (8) 0.003 867 0.43 Z5 sm lbr pr (8) 0.003 1015 0.40 Z6 lbr lrg lpr (5) 0.005 729 0.44

375.4 815.5 331.6 530.5 556.2 398.4

11 18 6 46 25 91

3756 5041 3313 1912 3715 1203

0.4822 0.5081 0.5369 0.5312 0.4791 0.4727

0.0023 0.0023 0.0025 0.0017 0.0010 0.0019

13.059 13.931 15.306 15.122 12.754 12.653

0.062 0.067 0.067 0.049 0.029 0.051

2796.7 2816.7 2880.5 2878.1 2768.5 2777.5

2.5 1.1 3.4 1.4 1.0 1.9

11.2 7.3 4.7 5.6 10.7 12.2

Sample 3 (97JKS-51b): fine-grained tonalite, Patterson Lake (395900 N/6974240 N) Z1 br lrg euh pr (1) 0.009 189 0.53 Z2 lbr euh pr (2) 0.004 605 0.38 Z3 lbr sm euh pr (3) 0.003 1027 0.41 Z4 br sm pr (1) 0.001 780 0.42 Z5 sm eq pr (3) 0.005 566 0.39

114.5 337.1 587.6 499.0 320.9

7 8 14 6 6

7350 9199 6869 4484 14426

0.5153 0.4902 0.5000 0.5591 0.4975

0.0024 0.0015 0.0023 0.0021 0.0022

14.368 13.379 13.796 16.250 13.668

0.067 0.043 0.064 0.057 0.061

2844.2 2809.4 2827.1 2911.8 2820.0

1.6 1.5 1.5 2.7 1.6

7.1 10.3 9.2 2.1 9.3

Sample 4 (BNB96-50a): K-feldspar megacrystic granite, Patterson Lake (395708 E/6971628 N) Z1 lyel eq pr (1) 0.009 2469 0.29 1142.7 Z2 lbr sm lpr (4) 0.004 830 0.37 403.8 Z3 lbr sm lpr (4) 0.005 903 0.23 387.9 Z4 yel pr (1) 0.003 1120 0.34 472.8 Z5 yel pr (1) 0.001 1716 0.36 909.4 Z6 lbr sm lpr (4) 0.002 518 0.37 277.9 Z7 lbr lpr (1) 0.001 1150 0.12 684.8 Z8 lyel sm pr (2) 0.003 1743 0.65 807.8 Z9 sm clr lpr (8) 0.005 1940 0.24 904.0 T1 lrg turb br frag (18) 0.109 32 0.43 18.8 T2 lbr subh frag (62) 0.111 31 0.23 16.4

132 64 40 28 19 12 10 18 80 582 183

4508 1438 2869 2895 2772 2656 4129 7100 3336 214 596

0.4271 0.4361 0.3993 0.3825 0.4775 0.4823 0.5542 0.3838 0.4350 0.5268 0.4957

0.0023 0.0018 0.0014 0.0024 0.0024 0.0027 0.0036 0.0012 0.0019 0.0022 0.0019

9.367 10.434 9.068 8.058 11.711 12.030 16.289 7.913 9.935 13.666 11.887

0.051 0.046 0.034 0.048 0.059 0.065 0.104 0.026 0.046 0.061 0.048

2445.7 2592.1 2504.7 2377.2 2633.1 2661.1 2929.8 2340.6 2514.1 2725.9 2595.6

1.7 2.3 1.8 4.0 2.2 3.7 2.9 1.6 1.2 2.5 1.9

7.4 11.9 15.9 14.2 5.4 5.6 3.7 12.3 8.8 −0.1 0.0

Sample 5 (BNB95-65): tonalite enclave in Morose granite, Morose Lake (400625 E/6966000 N) Z1 lbr 2:1 pr (1) 0.001 503 0.53 295.2 Z2 lbr 2:1 pr (1) 0.001 275 0.40 155.2 Z3 clr 2:1 pr (1) 0.001 174 0.50 101.2 Z4 clr lpr (2) 0.002 276 0.44 156.7 Z5 lbr tab pr (4) 0.003 316 0.40 179.0 T1 lbr euh frag (36) 0.087 17 0.10 8.7 T2 lbr euh frag (40) 0.080 13 0.07 6.7

11 5 9 4 6 91 80

1462 1847 614 4531 4973 519 422

0.5115 0.5060 0.5091 0.5021 0.5077 0.4882 0.4882

0.0034 0.0037 0.0029 0.0022 0.0025 0.0019 0.0017

12.927 12.641 12.878 12.671 12.804 11.551 11.543

0.070 0.081 0.075 0.055 0.063 0.045 0.040

2683.0 2663.8 2684.5 2680.6 2679.5 2573.4 2572.2

6.8 6.8 4.7 2.3 2.5 2.7 3.1

0.1 1.1 1.5 2.6 1.5 0.5 0.5

J.W.F. Ketchum et al. / Precambrian Research 135 (2004) 149–176

Weight U Th/U Pb* (mg) (ppm) (ppm) Sample 1 (97JKS-53): sheared mafic volcanic, Patterson Lake (NAD 27: 395700 E/6974350 N) Z1 sm high clar eq pr (12) 0.008 15 0.05 9.2 Z2 sm clr euh pry (15) 0.008 5 0.04 3.0 Z3 sm clr euh pr (14) 0.007 7 0.10 4.1 Z4 sm high clar pr (28) 0.016 7 0.03 3.9 Z5 sm high clar pr (20) 0.011 3 0.03 1.7 Z6 sm clr euh pr (48) 0.027 5 0.09 3.0 T1 clr lrg subh pr (54) 0.205 7 0.77 5.0 T2 lrg clr subh pr (44) 0.180 7 0.79 4.9 Fraction and description

8 7 8 12 19 19 28 344 585 547

18278 7554 11403 16110 965 1845 633 656 2171 2066

0.4619 0.4856 0.4805 0.5082 0.5040 0.5081 0.5146 0.5113 0.5123 0.5115

0.0022 0.0022 0.0026 0.0020 0.0026 0.0015 0.0025 0.0011 0.0012 0.0010

11.107 12.145 11.728 12.779 12.657 12.777 12.945 12.822 12.838 12.807

0.054 0.057 0.063 0.053 0.064 0.042 0.062 0.032 0.032 0.027

2600.2 2665.5 2625.3 2674.5 2672.4 2674.8 2675.3 2670.0 2669.1 2667.3

1.9 1.7 2.2 1.6 4.2 2.0 4.4 1.3 1.0 1.0

7.0 5.2 4.4 1.2 1.9 1.2 0.0 0.4 0.1 0.2

Sample 7 (BNB96-52a): foliated granodiorite, Patterson Lake (395324 E/6970857 N) Z1 br lpr frag (1) 0.002 265 0.39 Z2 lbr sm lpr (11) 0.004 553 0.34 Z3 yel 2:1 pr (1) 0.002 4222 0.35 Z4 lyel pr (1) 0.001 3932 0.25 Z5 lyel pr tip (1) 0.001 819 0.51 Z6 lbr sm lpr (17) 0.008 570 0.37 Z7 lbr sm lpr (17) 0.008 702 0.37 Z8 lyel pr (1) 0.001 2433 0.29 Z9 lbr sm lpr (9) 0.003 441 0.45 Z10 br pr (1) 0.003 589 0.38 Z11 sm lbr lpr (8) 0.007 658 0.47 T1 lbr subh frag (40) 0.062 45 0.15 T2 lbr subh frag (33) 0.065 46 0.14

14 26 55 55 22 69 56 30 9 15 30 113 124

1371 2625 4352 1804 1062 1877 2894 2182 4514 3607 4492 791 760

0.5696 0.4822 0.4482 0.4011 0.4503 0.4493 0.4584 0.4273 0.4957 0.4996 0.4593 0.5011 0.4973

0.0028 0.0015 0.0023 0.0015 0.0030 0.0015 0.0016 0.0016 0.0021 0.0013 0.0015 0.0021 0.0015

17.380 11.952 10.320 8.510 10.533 10.830 11.099 9.384 12.290 12.449 10.994 12.127 12.042

0.083 0.041 0.055 0.034 0.070 0.039 0.041 0.037 0.052 0.035 0.035 0.050 0.037

2990.2 2650.8 2527.7 2389.3 2554.4 2604.6 2611.9 2448.2 2651.2 2659.7 2592.8 2611.1 2612.1

3.8 2.0 1.6 2.2 3.6 1.9 1.7 2.1 2.4 1.3 1.7 1.9 1.9

3.5 5.2 6.6 10.6 7.4 9.8 8.3 7.5 2.6 2.2 7.2 −0.3 0.5

173.4 294.3 2085.6 1691.2 424.3 286.0 358.5 1125.6 247.7 327.4 343.7 23.7 23.8

Sample 8 (99JKS-223): epiclastic sedimentary unit, Patterson Lake Formation (396183 E/6974712 N) Z1 br tab pr (1) 0.002 211 0.56 157.9 Z2 dk br pr (1) 0.001 414 0.06 253.0 Z3 lbr pr (1) 0.001 177 0.82 141.1 Z4 lbr pr (1) 0.001 130 0.70 100.6 Z5 lbr tab pr (1) 0.001 149 0.66 114.1 Z6 lbr pr (1) 0.001 186 0.59 125.4 Z7 clr 2:1 pr (1) 0.001 115 0.65 88.1

0.3 0.4 0.5 0.3 0.6 2.1 0.3

52030 38475 14384 18208 9677 3219 15287

0.6248 0.5770 0.6307 0.6276 0.6266 0.5699 0.6271

0.0041 0.0014 0.0042 0.0023 0.0028 0.0015 0.0017

21.022 17.110 21.222 21.138 21.097 16.895 21.107

0.139 0.046 0.143 0.079 0.098 0.051 0.062

3146.5 2944.0 3146.6 3148.0 3147.5 2943.8 3147.0

1.4 1.7 2.0 2.0 1.5 1.8 1.6

0.7 0.3 −0.2 0.3 0.5 1.5 0.4

Sample 9 (00JKS-324): granite boulder in chert breccia, Patterson Lake Formation (395244 E/6974231 N) Z1 lbr cr lpr (1) 0.002 91 0.57 60.1 Z2 sm lbr lpr frag (2) 0.001 177 0.58 110.0 Z3 sm lbr lpr frag (2) 0.001 253 0.41 146.1 Z4 lbr lpr frag (1) 0.001 271 0.42 153.7 Z5 clr lpr (1) 0.001 64 0.61 43.5 Z6 br pr frag (1) 0.001 144 0.38 90.1

7.2 0.4 1.3 1.2 0.4 0.4

929 14638 6677 7115 4697 12257

0.5596 0.5298 0.5116 0.5028 0.5724 0.5549

0.0022 0.0012 0.0012 0.0023 0.0023 0.0016

16.270 14.839 14.143 13.874 16.822 16.076

0.084 0.037 0.037 0.066 0.071 0.048

2912.1 2851.5 2830.3 2827.4 2929.5 2906.4

4.4 1.8 1.9 1.8 2.1 1.8

2.0 4.8 7.2 8.7 0.5 2.6

J.W.F. Ketchum et al. / Precambrian Research 135 (2004) 149–176

Sample 6 (BNB96-33): foliated tonalite, southeast of Patterson Lake (400300 E/6971000 N) Z1 br lpr (1) 0.007 741 0.54 396.0 Z2 br lpr (1) 0.004 420 0.55 236.7 Z3 br pr (1) 0.006 495 0.45 269.2 Z4 lbr cr pr (1) 0.011 545 0.78 335.8 Z5 lbr 2:1 pr (1) 0.002 278 0.62 164.5 Z6 lbr tab cr frag (1) 0.005 221 0.46 127.1 Z7 lbr 2:1 pr (1) 0.002 267 0.44 154.5 T1 clr lrg subh frag (49) 0.218 32 0.67 19.1 T2 lrg dk br frag (32) 0.233 169 1.48 119.6 T3 lrg dk br frag (32) 0.196 179 1.40 124.5

All data are from air-abraded zircon and titanite (Krogh, 1982). Samples 8 and 9 analysed at the Royal Ontario Museum. All other samples analysed at the Memorial University of Newfoundland. Abbreviations: Z, zircon; T, titanite; euh, euhedral; subh, subhedral; eq, equant; tab, tabular; sm, small; lrg, large; pr, prism; lpr, long prism (>3:1); frag, fragment; clr, colourless; dk, dark; br, brown; lbr, light brown; yel, yellow; lyel, light yellow; cr, cracked; turb, turbid; 2:1, 3:1, etc., length:breadth ratio. Number in brackets indicates number of grains analysed. Th/U calculated from radiogenic 208 Pb/206 Pb ratio and 207 Pb/206 Pb age. Pb* = total radiogenic lead. Pbc = total common lead. 206 Pb/204 Pb ratio is uncorrected. All other ratios corrected for fractionation, blank, spike, and initial common Pb. Discord. = percent discordance for the given 207 Pb/206 Pb age. Decay constants used are from Jaffey et al. (1971).

161

162

J.W.F. Ketchum et al. / Precambrian Research 135 (2004) 149–176

of <15 ppm and Th/U ratios of <0.1 (Table 1), suggesting but not demonstrating a metamorphic origin. A range of mainly concordant ages between 2900 and 2765 Ma was obtained (Fig. 5a) which is surprising given the apparent uniformity of the zircon population. Possible reasons for the age variation include variable Pb loss, the presence in some or all grains of inherited cores, or mixing of grains of different ages, perhaps reflecting both metamorphic and primary zircon populations or the presence of detrital grains. As a unique interpretation of these data could not be made, additional zircon grains were analysed by SHRIMP. The results (Fig. 5b) confirm the presence of at least two age populations. Six low-U grains yield concordant data with a mean 207 Pb/206 Pb age of 2727 ± 18 Ma, whereas one distinctly higher U (111 ppm) grain is dated at 2956 ± 10 Ma. A backscattered electron image of this older grain shows faint growth zoning with no indication of an inherited core (Fig. 6a). Based on the SHRIMP results, the range of TIMS ages for multigrain zircon fractions most likely reflects variable laboratory mixing of ca. 2727 Ma and ca. 2956 Ma grains. This is consistent with data for

TIMS fraction Z1, which has both the highest U content and the oldest age (Table 1), suggesting that older, higher-U zircons are an important component of this fraction. Two multigrain fractions of colourless subhedral titanite analysed by TIMS are concordant and nearconcordant and provide identical 207 Pb/206 Pb ages of 2942 ± 3 Ma. The morphological similarlity of titanite grains and the duplicated age provide strong evidence for a single age population. As the mafic volcanic unit is penetratively deformed, metamorphosed to amphibolite facies, and contains no relict primary minerals, this titanite is undoubtedly of metamorphic origin. The U–Pb data therefore constrain the mafic volcanic unit to be at least as old as 2942 ± 3 Ma. 6.2. Sample 2: tonalite gneiss, Webb Lake A sample of grey tonalite gneiss, representative of typical gneissic basement in the study area, was collected near the margin of the SDC at Webb Lake (Fig. 3). In this area the tonalite is protomylonitic and is cut by all of the three mafic dyke swarms of Bleeker et al. (1999b).

Fig. 6. Back-scattered electron images and location of SHRIMP spot analyses for zircons from samples 1–4 (images a–d, respectively). Italicized number next to grain is linked to designation in Table 2. Two ages of zircon growth are apparent in images b and c. Age uncertainties are 2␴. See text for additional details.

J.W.F. Ketchum et al. / Precambrian Research 135 (2004) 149–176

163

Fig. 8. Concordia diagrams of TIMS U–Pb age data for clastic units of the Patterson Lake Formation. Boxed age is best estimate of primary crystallization age. Z, zircon. Age uncertainties and ellipses are 2␴.

Fig. 7. Concordia diagrams of TIMS U–Pb age data for younger plutonic units of the Sleepy Dragon Complex. Ages given in box represent best estimate of primary crystallization age. Z, zircon; T, titanite. Age uncertainties and ellipses are 2␴.

The zircon population consists of small brown prismatic grains. Although six fractions of this material yield a roughly linear array of discordant analyses (Fig. 5c), only four are collinear within analysis uncertainties. A discordia line through these fractions has an upper intercept age of 2936 + 7/−6 Ma (24% probability of fit). The remaining two zircon analyses plot to the left of the discordia line and are characterized by high U concentrations (>1000 ppm) relative to those defining the discordia line (560–870 ppm). This suggests that the higher U zircons suffered initial Pb loss at an earlier time, consistent with more rapid accumulation of ␣-decay induced crystal lattice damage expected in

164

Table 2 SHRIMP U–Pb isotopic data, Patterson–Morose Lakes area Spot name

U (ppm)

Th (ppm)

Th/U

Pb (ppm)

204 Pb

204 Pb/

(ppb)

206 Pb

1␴

f206

208 Pb/

1␴

206 Pb

207 Pb/

1␴

235 U

206 Pb/

1␴

R

238 U

207 Pb/

1␴

206 Pb

Apparent ages 206 Pb/ 238 U

1␴ (Ma)

(Ma)

207 Pb/ 206 Pb

1␴ (Ma)

Conc. (%)

(Ma)

0.000175 0.000112 0.000081 0.000010 0.000086 0.000037 0.000009 0.000035

0.000188 0.000051 0.000247 0.000010 0.000148 0.000401 0.000007 0.000080

0.0030 0.0019 0.0014 0.0002 0.0015 0.0006 0.0002 0.0006

0.0024 0.0015 0.0044 0.0113 0.0006 0.0052 0.0003 0.0035

0.0071 0.0019 0.0092 0.0039 0.0056 0.0149 0.0003 0.0031

13.431 13.529 14.173 13.948 13.563 12.687 17.803 13.449

0.442 0.249 0.404 0.616 0.338 0.564 0.207 0.293

0.5125 0.5213 0.5430 0.5362 0.5243 0.4891 0.5959 0.5207

0.0108 0.0082 0.0085 0.0202 0.0081 0.0108 0.0064 0.0077

0.7274 0.9085 0.6418 0.9057 0.7064 0.5957 0.9632 0.7622

0.1901 0.1882 0.1893 0.1887 0.1876 0.1881 0.2167 0.1873

0.0043 0.0015 0.0042 0.0036 0.0033 0.0068 0.0007 0.0027

2667 2705 2796 2767 2717 2567 3013 2702

46 35 35 85 34 47 26 33

2743 2727 2736 2731 2721 2726 2956 2719

38 13 37 31 30 61 5 24

97 99 102 101 100 94 102 99

Sample 2 (BNB95-50): tonalite gneiss, Webb Lake 6913-3.1 969 151.1 0.16 671 6913-3.2 597 106.0 0.18 376 6913-7.1 313 84.3 0.28 201 6913-7.2 884 188.3 0.22 567 6913-10.1 247 146.1 0.61 167 6913-12.1 191 93.5 0.50 127 6913-14.2 143 86.1 0.62 96 6913-3.4 1061 161.8 0.16 774 6913-3.5 685 26.7 0.04 436

5 3 58 4 5 2 7 4 6

0.000010 0.000010 0.000373 0.000010 0.000039 0.000020 0.000105 0.000006 0.000018

0.000010 0.000010 0.000053 0.000010 0.000018 0.000021 0.000038 0.000007 0.000008

0.0002 0.0002 0.0065 0.0002 0.0007 0.0004 0.0018 0.0001 0.0003

0.0438 0.0477 0.0925 0.0607 0.1714 0.1383 0.1692 0.0420 0.0098

0.0006 0.0008 0.0027 0.0006 0.0012 0.0038 0.0059 0.0008 0.0003

21.080 17.155 16.890 17.381 16.712 16.972 16.778 21.981 19.790

0.321 0.193 0.267 0.202 0.235 0.246 0.250 0.267 0.239

0.6246 0.5797 0.5709 0.5834 0.5681 0.5666 0.5649 0.6608 0.5908

0.0079 0.0061 0.0067 0.0060 0.0070 0.0068 0.0067 0.0069 0.0066

0.8859 0.9655 0.8173 0.9326 0.9222 0.8893 0.8608 0.9084 0.9614

0.2448 0.2146 0.2146 0.2161 0.2134 0.2173 0.2154 0.2412 0.2430

0.0017 0.0006 0.0020 0.0009 0.0012 0.0015 0.0017 0.0012 0.0008

3128 2948 2912 2962 2900 2894 2887 3270 2993

31 25 28 25 29 28 28 27 27

3151 2941 2940 2952 2931 2960 2947 3128 3139

11 5 15 7 9 11 12 8 5

99 100 99 100 99 98 98 105 95

Sample 3 (97JKS-51b): fine-grained tonalite, Patterson Lake 6923-2.1 253 69.4 0.28 163 3 6923-2.2 1046 56.9 0.06 550 50 6923-6.1 367 173.2 0.49 244 3 6923-6.2 951 64.8 0.07 528 1 6923-3.1 361 182.3 0.52 244 1 6923-3.2 1054 63.0 0.06 637 2 6923-5.1 1013 231.2 0.24 614 10 6923-5.2 817 69.9 0.09 503 22 6923-1.1 390 64.2 0.17 240 23 6923-2.3 949 55.9 0.06 548 18

0.000027 0.000112 0.000016 0.000002 0.000007 0.000004 0.000022 0.000054 0.000121 0.000041

0.000011 0.000022 0.000027 0.000003 0.000006 0.000003 0.000015 0.000013 0.000045 0.000012

0.0005 0.0019 0.0003 0.0000 0.0001 0.0001 0.0004 0.0009 0.0021 0.0007

0.0786 0.0166 0.1321 0.0182 0.1417 0.0169 0.0652 0.0294 0.0607 0.0175

0.0021 0.0009 0.0015 0.0002 0.0010 0.0003 0.0009 0.0008 0.0018 0.0008

17.165 13.884 17.115 14.600 17.275 16.760 15.806 16.697 16.445 15.577

0.230 0.160 0.202 0.208 0.239 0.202 0.198 0.199 0.248 0.216

0.5772 0.5026 0.5739 0.5307 0.5772 0.5717 0.5536 0.5778 0.5617 0.5482

0.0068 0.0055 0.0061 0.0063 0.0074 0.0062 0.0062 0.0064 0.0078 0.0060

0.9309 0.9763 0.9415 0.8934 0.9645 0.9381 0.9432 0.9645 0.9604 0.8499

0.2157 0.2004 0.2163 0.1995 0.2171 0.2126 0.2071 0.2096 0.2123 0.2061

0.0011 0.0005 0.0009 0.0013 0.0008 0.0009 0.0009 0.0007 0.0009 0.0015

2937 2625 2924 2744 2937 2915 2840 2940 2874 2818

28 24 25 27 30 25 26 26 32 25

2949 2829 2953 2822 2959 2926 2883 2902 2923 2875

8 4 7 11 6 7 7 5 7 12

100 93 99 97 99 100 99 101 98 98

Sample 4 (BNB96-50a): K-feldspar megacrystic granite, Patterson Lake 6926-2.1 2393 578.1 0.25 1344 563 0.000522 6926-2.2 6327 1035.9 0.17 3532 1616 0.000559 6926-3.1 635 388.3 0.63 439 3 0.000011 6926-3.2 650 96.5 0.15 401 6 0.000018 6926-4.1 146 105.6 0.75 102 3 0.000047 6926-5.1 973 57.8 0.06 484 114 0.000282 6926-4.2 1825 407.0 0.23 1068 127 0.000149 6926-6.1 493 111.5 0.23 274 2 0.000010 6926-6.2 722 55.3 0.08 392 8 0.000024

0.000008 0.000012 0.000011 0.000010 0.000020 0.000039 0.000007 0.000010 0.000009

0.0090 0.0097 0.0002 0.0003 0.0008 0.0049 0.0026 0.0002 0.0004

0.0606 0.0491 0.1672 0.0428 0.2076 0.0201 0.0623 0.0639 0.0220

0.0011 0.0006 0.0013 0.0005 0.0021 0.0015 0.0005 0.0014 0.0009

13.767 13.165 17.179 16.699 17.021 11.859 14.526 13.078 13.199

0.354 0.227 0.220 0.249 0.329 0.150 0.157 0.161 0.149

0.5213 0.5276 0.5825 0.5710 0.5705 0.4832 0.5413 0.5170 0.5238

0.0098 0.0087 0.0070 0.0072 0.0083 0.0053 0.0055 0.0056 0.0055

0.8032 0.9809 0.9738 0.8973 0.8219 0.9169 0.9628 0.9216 0.9645

0.1916 0.1810 0.2139 0.2121 0.2164 0.1780 0.1946 0.1835 0.1828

0.0030 0.0006 0.0006 0.0014 0.0024 0.0009 0.0006 0.0009 0.0006

2705 2732 2959 2912 2910 2541 2789 2686 2715

41 37 29 30 34 23 23 24 23

2755 2662 2935 2922 2954 2634 2782 2684 2678

26 6 5 11 18 8 5 8 5

98 103 101 100 99 97 100 100 101

Uncertainties reported at one sigma and are calculated by numerical propagation of all known sources of error (Stern, 1997). f206 refers to mole fraction of total 206 Pb that is due to common Pb; data have been common Pb corrected according to procedures outlined in Stern (1997). Conc. = 100 × (206 Pb/238 U age)/(207 Pb/206 Pb age).

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Sample 1 (97JKS-53): sheared mafic volcanic, Patterson Lake 6925-7.1 11 0.1 0.01 6 1 6925-12.1 38 0.2 0.01 20 2 6925-15.1 7 0.1 0.01 4 0 6925-17.1 2 0.0 0.02 1 0 6925-22.1 21 0.2 0.01 11 1 6925-22.2 6 0.0 0.00 3 0 6925-11.1 111 0.1 0.00 69 1 6925-13.1 17 0.1 0.01 9 0

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these grains. This relationship between U concentration and Pb loss appears to be common in zircons analysed for this study. SHRIMP analyses of core and rim areas of five weakly to strongly growth-zoned zircons confirm the TIMS age and also reveal a xenocrystic core in one grain. Six analyses provide a mean 207 Pb/206 Pb age of 2944 ± 9 Ma, which, because of the concordance of these data (Fig. 5d), is taken as the best age estimate for the tonalite protolith. One zircon has a weakly growthzoned, ca. 2940 Ma outer portion and a more strongly zoned, distinctly older core (Fig. 6b). Three analyses of the latter provide a weighted mean 207 Pb/206 Pb age of 3138 ± 21 Ma; the only concordant analysis is 3151 ± 23 Ma. An inherited component with an age of ca. 3140–3150 Ma is therefore demonstrated by these data. 6.3. Sample 3: fine-grained tonalite layer, Patterson Lake On an island at the north end of Patterson Lake, strongly deformed, medium-grained tonalite gneiss similar to sample 2 contains numerous <0.2 m wide layers of fine-grained tonalite to quartz diorite. These layers parallel the host rock foliation and were clearly deformed at the same time as the medium-grained tonalite. Their probable intrusive origin could not be confirmed. A fine-grained tonalite layer was sampled for dating. Numerous small brown prismatic zircons were recovered. Five TIMS analyses of this material are moderately to strongly discordant but provide an upper intercept age of 2933 + 8/−7 Ma (Fig. 5e; 49% probability of fit). SHRIMP data for five cracked and variably growth-zoned grains from this sample are not as easily interpreted. Analyses are progressively less concordant with decreasing age (Fig. 5f), suggesting either variable Pb loss from zircons of uniform age or the presence of grains or grain components of different age. Younger ages are obtained mainly from distinctly growth-zoned rims with Th/U ratios <0.1 (e.g., Fig. 6c), which may indicate a metamorphic overgrowth origin. However, uniform rim ages were not obtained, and the influence of both younger overgrowths and variable Pb loss may account for the observed age variation.

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Relatively low U, unzoned central regions of three zircon grains provide overlapping concordant analyses with a weighted mean 207 Pb/206 Pb age of 2955 ± 12 Ma. These are the oldest analyses obtained and there is no clear indication from back-scattered electron images (e.g., Fig. 6c) that they date an inherited component. We therefore interpret 2955 ± 12 Ma as the primary crystallization age of the fine-grained tonalite layer. This age is identical within uncertainty to the 2944 ± 9 Ma age of the tonalitic host, but is older than the TIMS-determined age of 2933 + 8/−7 Ma. Given the discordance of the TIMS data and the strong dependence of the upper intercept age on a single fraction (Z4; Fig. 5e), it is possible that the upper intercept has been biased toward a younger age due to complex and non-uniform Pb loss behaviour. 6.4. Sample 4: K-feldspar megacrystic granodiorite, Patterson Lake Much of Patterson Lake is underlain by a distinctive unit of variably foliated K-feldspar megacrystic granodiorite that becomes progressively more deformed toward the basement-cover contact (Bleeker et al., 1997). This unit does not contain swarm 1 mafic dykes but is cut by swarms 2 and 3, providing an age bracket of 2734–2687 Ma for its emplacement (see above and Bleeker et al., 1999b). The granodiorite was sampled at a low-strain locality on a large island in Patterson Lake. A variety of zircon types including yellow prismatic grains with and without visible cores, large brown prisms with cores, and smaller brown and colourless needles characterize the sample. Composite grains were avoided during grain selection due to the possibility that cores represent inherited zircon. Titanite in the sample consists mainly of light brown anhedral grains and grain fragments, and subordinate large, dark brown turbid fragments. U–Pb data for single and multigrain zircon fractions indicate a multistage Pb loss history, particularly for colourless and yellow grains with high U concentrations (∼1100–2500 ppm, Table 1). In contrast, small, light brown zircon needles have much lower U contents (∼500–900 ppm), and three discordant fractions of this material are collinear with a concordant fraction of turbid, dark brown titanite with an age of 2726 ± 3 Ma (Fig. 5g). The upper intercept age of a

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discordia line calculated from the three zircon fractions alone is 2722 + 9/−8 Ma but with a low probability of fit. With inclusion of the titanite fraction, the upper intercept age becomes 2724 ± 6 Ma (uncertainty calculated using the error expansion routine of Davis, 1982). It would appear from the collinearity of these analyses and two additional lines of evidence that the dark brown titanite analysis provides the best estimate for the igneous crystallization age of the granodiorite: (i) a fraction of light brown subhedral titanite is concordant at 2596 ± 2 Ma and probably dates a time of metamorphic titanite growth, which implies an alternate origin for dark brown grains; and (ii) the 2726 ± 3 Ma titanite age falls within the previously determined 2734–2687 Ma age bracket for the sampled unit. Based on this evidence, we suggest that the megacrystic granodiorite body was emplaced at 2726 ± 3 Ma. Five zircons analysed by SHRIMP confirm the high U content of some grains but not the postulated emplacement age. Three analyses of two grains provide a weighted mean 207 Pb/206 Pb age of 2934 ± 21 Ma (Fig. 5h), consistent with a single-grain TIMS age of ca. 2930 Ma and interpreted to date inherited basement zircons. Most other analysed grains are characterized by high U and high common Pb and yield inconsistent ages, perhaps due to variable Pb loss. However, two analyses of a lower U zircon provide a weighted mean 207 Pb/206 Pb age of 2680 ± 8 Ma. Based on evidence for ca. 2680 Ma plutonism in the SDC (see below), this grain might reflect secondary metamorphic growth, although its internal characteristics (Fig. 6d) do not strongly favour this interpretation. The postulated 2726 ± 3 Ma age of the megacrystic granodiorite appears to constrain displacement along the basement-cover contact to the period 2726–2687 Ma as this unit is overprinted by obliquenormal shear fabrics. This is a slightly narrower age bracket for displacement on this contact than the 2734–2687 Ma interval established by Bleeker et al. (1999b). 6.5. Sample 5: foliated tonalite enclave in Morose Granite, Morose Lake The ca. 2586 Ma, syn- to post-tectonic Morose Granite pluton contains screens of older, foliated leucocratic tonalite (Davidson, 1972) that are cut by mafic dykes likely belonging to swarm 3 (Bleeker et al.,

1999b). A tonalite sample from a large screen exposed on Morose Lake yielded numerous colourless and light brown zircon prisms, some which exhibit coreovergrowth relationships. Five zircon fractions lacking visible cores were analysed, and of these, four have similar 207 Pb/206 Pb ages and define a Pb loss line with upper and lower intercept ages of 2683 + 18/−5 Ma and 258 + 987/−717 Ma, respectively. The large age uncertainties reflect the tight clustering of analyses (Fig. 7a), with the result that the upper intercept uncertainty is far greater than uncertainties on the individual, nearconcordant analyses. This geometric effect can be overcome by constraining the lower intercept age. Using 258 ± 100 Ma as this constraint, which factors in evidence for non-zero age Pb loss (the more discordant analyses generally have younger 207 Pb/206 Pb ages), a more realistic upper intercept age of 2683 + 3/−2 Ma (34% probability of fit) is obtained, which we interpret as the best estimate for primary crystallization of the leucocratic tonalite. One zircon fraction has a younger 207 Pb/206 Pb age of 2664 ± 7 Ma; the significance of this result is unknown. Two fractions of light brown titanite fragments derived from euhedral grains are concordant at 2573 ± 3 Ma. This age has been shown by Bethune et al. (1999) to record cooling of the Morose pluton through the 600 ± 50 ◦ C titanite closure temperature. Bethune et al. (1999) also report 40 Ar/39 Ar ages of hornblende (2536 ± 14 Ma) and biotite (2425 ± 14 Ma) from sample 5; the temperaturetime path derived from these results and data from a nearby sample constrain the late Archean cooling history of the region. The 2683 Ma age of the leucotonalite is identical to that for a foliated granodiorite pluton at Sleepy Dragon Lake (Henderson et al., 1987; Fig. 2) which also contains swarm 3 dykes (Bleeker et al., 1997, 1999b). 6.6. Sample 6: foliated melanocratic tonalite between Patterson and Morose Lakes As described earlier, foliated tonalite and granodiorite units of the interior domain are not cut by the three mafic dykes swarms observed in basement rocks to the west. This absence suggests that intrusion of tonalite and granodiorite postdated 2683 Ma, the age of the youngest granitoid bodies hosting swarm 3 dykes. A sample of melanocratic tonalite collected between Patterson and Morose lakes represents a distinctive

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compositional endmember of the plutonic suite. The sample yielded abundant brown prismatic zircons, and five of seven single-grain analyses provide an upper intercept age of 2677 ± 2 Ma (89% probability of fit; Fig. 7b). Two other single grain analyses are characterized by high U contents and plot to the left of the regression line, most likely due to the influence of an earlier Pb loss event. We interpret 2677 ± 2 Ma as the primary crystallization age of the tonalite. Titanite in the sample consists of colourless to dark brown fragments. Analysis of three multigrain fractions (one colourless, two dark brown) suggests no significant age differences and a mean age of 2669 ± 2 Ma is obtained. This age likely records either the time of post-magmatic cooling of igneous titanite or growth of metamorphic titanite. 6.7. Sample 7: biotite granodiorite, Patterson Lake Foliated biotite granodiorite at the south end of Patterson Lake is also a unit of the interior domain. The granodiorite lacks cross-cutting mafic dykes but locally hosts trains of angular mafic xenoliths that may be derived from one or more of the basement dyke swarms. A sample from this locality contains both yellow (>800 ppm U) and brown (<700 ppm U) prismatic zircons, many of which contain visible cores. Analyses of small, relatively low U, light brown zircon needles without visible cores appear to provide the most useful age data as four variably discordant fractions are collinear with an upper intercept age of 2672 + 7/−6 Ma (Fig. 7c; error expansion calculation after Davis, 1982). A fifth fraction lying slightly to the right of this discordia line was not included in the regression. Most of the remaining zircon analyses consist of single, high U yellow prisms or tip overgrowths. No meaningful age information was obtained from this material. A single brown prism with a 207 Pb/206 Pb age of 2990 Ma (3.5% discordant) demonstrates the presence of an inherited basement component. Consistent with our focus on lower U grains in other samples, we consider the 2672 + 7/−6 Ma upper intercept age to date igneous crystallization of the granodiorite. Two fractions of light brown subhedral titanite fragments are concordant at 2612 ± 2 Ma, which potentially dates an episode of metamorphic titanite growth.

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7. U–Pb data from the Patterson Lake Formation 7.1. Sample 8: epiclastic sedimentary layer, Patterson Lake At Patterson Lake, the top of the Patterson Lake Formation is marked by a massive to thinly bedded mudstone-siltstone unit containing epiclastic layers dominated by medium- to coarse-grained quartz and feldspar crystals. A 25-cm wide epiclastic layer was sampled for detrital zircon dating. This layer occurs at a similar stratigraphic level to, but about two km southwest of a graded siliciclastic bed for which Bleeker et al. (1999a) reported detrital zircon ages of 2943 Ma and 3147–3160 Ma. Sample 8 yielded mainly brown prismatic zircons varying both in size and aspect ratio, and less abundant light pink and colourless grains. Mechanical pitting suggestive of detrital transport was rarely observed. Single-grain analyses indicate that these zircons comprise two distinct age populations with weighted average 207 Pb/206 Pb ages of 2944 ± 1 Ma (two analyses) and 3147 ± 1 Ma (five analyses; Fig. 8a). The maximum deposition age of 2944 Ma is identical to that for the siliciclastic bed dated by Bleeker et al. (1999a). A 3147 Ma source, which was earlier suggested from the data of Bleeker et al. (1999a), is precisely dated here. 7.2. Sample 9: granite boulder in chert breccia, Patterson Lake Chert breccia exposed near the north shore of Patterson Lake is part of a lithologically varied conglomerate unit that occurs near the top of the Patterson Lake Formation. This matrix-supported breccia is dominated by angular chert clasts but locally contains well-rounded granite boulders and cobbles. The breccia is moderately to strongly deformed with pronounced clast elongation observed in near-vertical exposures. A weakly foliated granite boulder was removed from the breccia for dating. Cracked, colourless to light brown, equant to long prismatic grains dominate the zircon population. Six analyses of long prismatic grains and grain fragments are weakly to strongly discordant, with four collinear, lowest U analyses (with one exception) providing an upper intercept age of 2934 ± 3 Ma (Fig. 8b; 92% prob-

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ability of fit). The remaining two analyses plot slightly to the left of this line, suggesting a more complex Pb loss history. For all zircon fractions, greater discordance is directly correlated with higher U content (Table 1). We interpret 2934 ± 3 Ma as the primary crystallization age of the granite body from which the boulder was derived.

8. Discussion 8.1. Chronology of basement-cover events and regional comparisons U–Pb ages presented here, along with data from earlier studies, provide a comprehensive picture of the plutonic and tectonic development of a domal basement complex and its autochthonous cover in the southern Slave Province. This evolution is depicted schematically in Fig. 9 and is described in detail below. Age data linked with field observations highlight a protracted history that both predates and accompanies development of the overlying volcano-sedimentary belts of the Yellowknife Supergroup. As for much of the CSBC (Ketchum and Bleeker, 2001a,b), oldest components of the SDC in the Patterson Lake area are recognized on the basis of isotopic data. Indications of a ca. 3150 Ma component are obtained both from inherited zircons in basement tonalite (sample 2) and detrital zircons from the autochthonous Patterson Lake Formation (sample 8; see also Bleeker et al., 1999a; Sircombe et al., 2001). The 3147 ± 1 Ma detrital zircon population in sample 8 may precisely date this cryptic component. Nd depleted mantle model ages as old as 3.3 Ga suggest the presence of even older crust in the SDC (Dud´as et al., 1988). Based on detailed field mapping and preliminary U–Pb dating, pre-3.0 Ga gneissic basement rocks are likely to be exposed in the northern SDC (James and Mortensen, 1992). Based on our own observations of these rocks and the results of earlier studies (e.g., Davidson, 1972; James and Mortensen, 1992), these early components of the SDC were transformed to amphibolite-facies gneisses prior to widespread ca. 2950 Ma tonalitic magmatism. A significant finding of this study is that a thin mafic volcanic unit along the basement-cover contact is older than 2942 ± 3 Ma. This unit must therefore represent a supracrustal component of the basement com-

Fig. 9. Age–event column depicting the Mesoarchean and Neoarchean geological development of the Sleepy Dragon Complex and Patterson Lake Formation in the study area. Superscript numbers indicate age sources: (1) Dud´as et al. (1988; Nd data); (2) Bleeker et al. (1999a); (3) Bleeker et al. (1999b); (4) Davis and Bleeker (1999); (5) Sircombe et al. (2001); (6) this study. Additional events in the vicinity of the study area (not shown) include ca. 2820 Ma granitoid plutonism at Sleepy Dragon Lake (Henderson et al., 1987), ca. 2700 Ma granitoid plutonism along the southern margin of the SDC (Bleeker, 2002), and deposition of the Raquette Lake Formation (part of the larger Ross Lake Group) on an uplifted and eroded SDC substrate after 2683 Ma but before 2661 Ma (Bleeker and Villeneuve, 1995; Bleeker et al., 1997). Deposition of the Ross Lake Group is potentially related to diapiric uplift of the SDC in response to volumetrically significant 2683–2672 Ma plutonism.

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plex. Supracrustal rocks appear to be uncommon in known or suspected basement complexes of the westcentral Slave Province, although there are few data to directly address this question. Sparse remnants of pre-Yellowknife Supergroup mafic volcanic rocks are known elsewhere in the western Slave Province (e.g., 3142 Ma mafic-felsic assemblage near Takijuq Lake (TL, Fig. 1); Villeneuve et al., 1993), but no regional correlation of these remnants is currently possible. The actual age of the metavolcanic unit at Patterson Lake is unknown, but it must be older than a highly deformed, ca. 2954 Ma tonalitic layer that intrudes it (W. Davis, personal communication, 2002). It is possible that the 2956 ± 10 Ma SHRIMP age of a single zircon from this unit (Fig. 6a) dates a coeval, more felsic volcanic component, but this interpretation cannot be confirmed with the available data. A major ca. 2990–2910 Ma magmatic event, characterized by tonalite-trondhjemite-granodiorite (TTG) plutonism across the southern half of the CSBC (Bleeker and Davis, 1999; Ketchum and Bleeker, 2001b), is well represented in the southern SDC. Two tonalite samples (2 and 3) yield ca. 2950 Ma primary ages, and two younger plutonic units (samples 4 and 7) contain inherited zircons of comparable age. In addition, units of the Patterson Lake Formation contain ca. 2950–2920 Ma detrital zircons, including a precisely dated 2944 Ma component (sample 8) (Bleeker et al., 1999a; Sircombe et al., 2001), and a granite boulder (sample 9) from this formation is dated at 2934 ± 3 Ma. These results collectively suggest that 2950–2930 Ma TTG plutons once formed a significant portion of Sleepy Dragon crust, and that they contributed detritus during ca. 2.85–2.80 Ga deposition of the Central Slave Cover Group. With the exception of the northern CSBC, this TTG event can be documented regionally (Bleeker and Davis, 1999; Ketchum and Bleeker, 2001a,b; Sircombe et al., 2001). Widespread, voluminous plutonism at 2990– 2910 Ma was followed by stabilization, uplift, development of a regional unconformity, and deposition of the Central Slave Cover Group. The upper part of the Patterson Lake Formation was deposited after 2826 Ma (Sircombe et al., 2001) but predates 2734 Ma emplacement of swarm 1 mafic dykes (Bleeker et al., 1999a). These constraints are similar to those outlined by Sircombe et al. (2001) for Central Slave Cover Group deposition across the western Slave Province,

169

with the exception of an occurrence at Dwyer Lake, about 25 km north of Yellowknife (Fig. 1). There, quartz arenite is overlain by (and is therefore older than) a 2853 Ma felsic volcanic unit (Ketchum and Bleeker, 2000). As pointed out by Sircombe et al. (2001), the data suggest either diachronous deposition of the Central Slave Cover Group or the presence of more than one quartzite-dominated autochthonous package. Sircombe et al. (2001) infer a local provenance for the detritus, an interpretation that is consistent with data from this study. It is not clear from field evidence whether 2950–2930 Ma basement units in the study area were deformed prior to deposition of the Patterson Lake Formation. In contrast, at Brown Lake in the northern SDC (BL, Fig. 1), plutonic and high-grade tectonometamorphic activity occurred between 2840–2820 Ma (Ketchum and Bleeker, 2001a). Ca. 2820 Ma granitoid plutonism immediately northeast of the study area at Sleepy Dragon Lake (Fig. 2; Henderson et al., 1987) is considered here to be genetically linked to this tectonothermal event, but no contemporaneous high-grade metamorphic event has been documented in the southern SDC. This difference may indicate that a deeper level of the basement complex is exposed in the Brown Lake area. Intrusion of swarm 1 mafic dykes into units of the SDC and Patterson Lake Formation at 2734 ± 2 Ma was likely contemporaneous with an early stage of Yellowknife Supergroup mafic volcanism (Bleeker et al., 1999b). Although the Cameron River basalts have not been dated, a spatial linkage of basement and autochthonous cover with these rocks is indicated both by the abundance of mafic dykes in all units and Nd isotope data indicating variable crustal contamination of Cameron River dykes and mafic flows (Lambert et al., 1992). Isotopic data from the Point Lake (Northrup et al., 1999) and Yellowknife greenstone belts (Cousens, 2000) also indicate crustal contamination of mafic volcanic rocks. Largely autochthonous development of greenstone belts across the western Slave Province is also suggested by contemporaneous (rift-related?) plutonism within adjacent granite–gneiss domains (Davis et al., 2003). The K-feldspar megacrystic granodiorite pluton dated here at 2726 ± 3 Ma (sample 4) represents the oldest-known example of this activity. Following 2726 Ma granodiorite plutonism, the basement-cover interface in the Patterson Lake area

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became a locus of strong deformation, with development of mylonitic fabrics. Oblique, northwest-side down displacement (present-day coordinates) is indicated (Kusky, 1990; Bleeker et al., 1999b). This event began at or after 2726 Ma as megacrystic granodiorite is strongly deformed within the high-strain zone. Bleeker et al. (1999b) suggest that this zone is part of a regional d´ecollement that transported pre-2687 Ma supracrustal rocks southwestward across the CSBC. As discussed previously, displacement is thought to be no more than several tens of kilometres. If our age and field constraints from the Patterson Lake area are valid regionally, then this d´ecollement was not active until ≤2726 Ma. A minimum age for this event is provided by cross-cutting 2687 Ma mafic dykes of swarm 2. Following this tectonic activity, magmatic events within the southern SDC mainly reflect synvolcanic plutonism during continued development of volcanic stratigraphy. Tonalite and granodiorite bodies in the interior domain of the SDC were emplaced at 2683 + 3/−2 Ma, 2677 ± 2 Ma, and 2672 + 7/−6 Ma (samples 5, 6, and 7, respectively). These ages are similar to dated volcanic and plutonic units across the western Slave Province (e.g., Henderson et al., 1987; Isachsen et al., 1991; Villeneuve, 1993; Villeneuve et al., 1997; Yamashita et al., 2000) and reflect a major magmatic event that significantly modified the CSBC. In the study area, swarm 3 mafic dykes of Bleeker et al. (1999b) are found in 2683 Ma tonalite gneiss (sample 5) but are not observed in well-exposed, ca. 2677–2672 Ma units to the northwest (Fig. 3). The northwest-trending swarm 3 dykes are therefore indirectly dated here at ca. 2680 Ma. Subsequent regional events include post-ca. 2670 Ma deposition of the turbiditic Burwash Formation, D1 deformation between 2660–2630 Ma and D2 deformation at ca. 2600 Ma (Davis and Bleeker, 1999; Davis et al., 2003). The regional deformation events generated F1 –F2 -fold interference patterns that characterize much of the presentday architecture of the southern Slave Province. 8.2. Synplutonic doming of the Sleepy Dragon Complex? The relative contribution of vertical versus horizontal tectonic processes in the development of Archean cratons has long been debated. Although a dominance

of horizontal transport is consistent with plate tectonic models of crustal growth and modification, the architecture of several well-preserved Archean cratons (e.g., Dharwar, Zimbabwe, Pilbara) suggests that diapiric processes were locally or even regionally dominant (e.g., Macgregor, 1951; Hickman, 1984; Choukroune et al., 1997). The history of the SDC outlined above bears some resemblance to classic granite–gneiss domains attributed to magmatic diapirism. To explore the possible role of diapirism during SDC development, below we compare this basement complex to the well-studied granite–gneiss domains of the eastern Pilbara granite–greenstone terrain of northwestern Australia. This comparison, along with evidence for synmagmatic unroofing and unconformity development, suggests that magmatic diapirism may have played an important role in the late plutonic development of the southern SDC. The eastern Pilbara granite–greenstone terrain is characterized by domal granite–gneiss complexes and adjacent, steeply dipping, synclinal greenstone belts. Although aspects of the structural evolution of this region remain controversial, recent work (e.g., Collins et al., 1998; Van Kranendonk et al., 2001, 2002, 2004; Sandiford et al., 2004) has shown that the granitoid domes are sites of long-lived magmatic and nonmagmatic diapirism, and that stratigraphic and thermal conditions at the time of Mesoarchean diapiric activity were favourable for partial convective overturn of the crust. The principal conditions leading to regional diapirism were the presence of a 12–18-km thick greenstone sequence overlying a less dense quartzofeldspathic crust, and the burial of radiogenic heatproducing elements contained within this quartzofeldspathic crust during greenstone accumulation. These conditions are likely to have resulted in significant midand lower-crustal heating, anatexis, and ductile flow which, due to the inverted crustal density profile, may have triggered development of the dome-and-keel regional structure (Sandiford et al., 2004). Within the granitoid complexes, an older (typically tonalitic)-to-younger (typically granitic) zonation toward complex centres is attributed to the displacement of older rocks upward and outward during intrusion of younger plutons (Van Kranendonk et al., 2001, 2002). Younger, syn- and post-volcanic granitoid bodies are typically the dominant components of the complexes relative to older gneissic units. Structural features, in-

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cluding ring faults and a complicated internal geometry that contrasts with a simple external outline of the domes in plan view, are similar to those documented in salt diapirs (e.g., Jackson et al., 1990). Adjacent greenstone belts face away from the granite–gneiss complexes and are steeply dipping to locally overturned. These belts are strongly deformed and contain radially dispersed, down-dip stretching lineations, suggesting a diapiric flow regime. All these characteristics are attributed to repeated magma channeling and ductile flow into progressively amplifying, mid-crustal highs that eventually produced the present-day domal granitoid complexes (Collins et al., 1998; Wellman, 2000; Van Kranendonk et al., 2001, 2002). The tight synclinal geometry and steep dip of the flanking greenstone belts reflect the accompanying downward flow of dense, basalt-dominated crust. Diapirism was driven by a combination of radiogenic heat production (which was 2–3 times greater during the Mesoarchean than today; e.g., Sandiford et al., 2004), the presence of the thick greenstone load on less dense quartzofeldspathic crust, thermal blanketing of the crust by this greenstone cover, and possibly also heat from mantle plumes. Although regional compressive tectonic activity likely characterized some stage(s) of crustal deformation (e.g., Zegers et al., 2001; Blewett, 2002), the present-day crustal geometry suggests that this activity was subordinate to and/or outlasted by diapirism. Several temporal, geometric, and plutonic features of the southern SDC resemble those documented in the eastern Pilbara craton: (i) Neoarchean plutonism was spatially focused, episodic, and prolonged, indicating the existence of a long-lived magmatic centre; (ii) plutonic activity yielded an older-to-younger magmatic sequence toward the centre of the complex (Figs. 2 and 3); (iii) syn- and post-volcanic plutons dominate over basement units (Fig. 2), (iv) the parautochthonous greenstone cover dips steeply away from the basement complex and is locally overturned (Lambert, 1988), and; (v) tight, generally steeply plunging folds characterize parts of the complex interior (particularly northeast of the study area; Davidson, 1972), in contrast to a relatively simple outline of the SDC in plan view (Fig. 2). Although these features are comparable to those observed within and adjacent to the Pilbara granitoid domes, we note that post-2660 Ma regional deformation has modified the SDC and surrounding rocks to the extent that diapir-related structures and tec-

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tonic fabrics, if present, are likely to have been extensively modified, thus precluding a classic structural analysis of diapir tectonics. For instance, the current high-amplitude domal form of the southern SDC is largely attributed to the interference of large-scale F1 and F2 -folds (Fig. 2; Bleeker, 1996, 2002). Steeply plunging lineations within the greenstone belts wrapping the SDC (e.g., Kusky, 1990) mainly reflect D1 and D2 events, and locally, earlier displacement on the basement-cover contact (Bleeker et al., 1999b). In addition, the internal structural development of the SDC has not been documented in detail. The prominent northeast-trending fold within the SDC first mapped by Davidson (1972) could potentially represent an F1 structure rather than one related to diapirism; additional study of this and other internal structures is required. Because of younger deformation, it is difficult to determine whether structural fabrics consistent with diapir tectonics once existed, or still exist in modified form, within and adjacent to the SDC. Therefore, a hypothesis of magmatic diapirism within the SDC must be based mainly on non-structural evidence. Previous work has shown that significant uplift and unroofing of the southern SDC occurred sometime between ca. 2700–2660 Ma, prior to regional D1 deformation. As discussed above, the Raquette Lake–Detour Lake unconformity (Fig. 2) cuts down through and truncates the Cameron River basalt package along the southwestern boundary of the SDC. This unconformity has likely removed a minimum 3 km thickness of basalt over a 15 km strike length (see Bleeker, 1996, 2001) and has also excised an unknown amount of ca. 2700 Ma basement granodiorite. This unconformity was therefore responsible for exposing an ancient basement culmination that is located very close to the present-day structural culmination, suggesting that this basement high represents a long-lived feature. Deposition of clastic and felsic volcanic rocks of the Raquette Lake Formation occurred between 2683 and 2661 Ma, with field and U–Pb zircon data pointing to the SDC as a detrital source (Bleeker and Villeneuve, 1995; Bleeker et al., 1997). Deposition of this formation was therefore broadly contemporaneous with the emplacement of volumetrically significant, 2683–2672 Ma tonalite and granodiorite plutons within the SDC, and with intrusion of swarm 3 mafic dykes at ca. 2680 Ma. Using sedimentological evidence, Mueller and Corcoran (2001) argued for basement uplift on normal faults and

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the development of significant topography during deposition of the Raquette Lake Formation. There are no documented tectonic events of appropriate age that could readily account for this uplift, and supracrustal packages temporally equivalent to the Raquette Lake and Detour Lake formations are mainly limited to scattered regional occurrences (Bleeker, 2001). This suggests a local cause of uplift and unconformity development. The data from the southern SDC collectively support the notion of a thermally and tectonically active regime at ca. 2680–2670 Ma, with spatially and temporally linked plutonism, basement uplift, erosion, and sedimentation. We suggest that all this activity reflects plutonic doming and partial unroofing of the SDC. These events followed displacement on the basement-cover d´ecollement, which occurred before 2687 Ma (Bleeker et al., 1999b) and may have generated conditions favourable for magmatic activity and possibly also partial convective overturn of the crust (i.e., sinking of the Cameron River and Beaulieu River basalts into a less dense felsic substrate). In this regard, extensional weakening and/or rupturing of the upper crust represents an effective trigger for the initiation of diapiric flow (e.g., Sandiford et al., 2004). Given that the basement-cover d´ecollement may have formed in response to regional extension (Bleeker et al., 1999b), links between tectonic and diapiric processes in this instance remain a possibility. The magmatically generated SDC dome would naturally have been amplified during subsequent F1 folding, resulting in a prominent, high-amplitude F1 anticline within the complex (Fig. 2; Bleeker, 1996, 2002). This fold is likely responsible for tilting and preservation of the Raquette Lake–Detour Lake unconformity. Its doubly plunging geometry (Bleeker, 1996) potentially reflects a basement-cover contact that had already been steepened during earlier plutonic doming. The proposed diapiric uplift of the SDC was part of an evolving process that may have initiated with intrusion of ca. 2726 Ma megacrystic granodiorite, a hypothesis consistent with the inward-younging structure of the SDC. Progressive amplification of this basement high may have generated the north-striking faults that die out in both directions from the basement-cover contact (Henderson, 1985; Fig. 3), and the northeast versus

northwest orientation of swarms 2 and 3 mafic dykes in the study area, signalling either a change in the ambient stress field or rigid block rotation (Bleeker et al., 1997). We suggest, however, that magmatic diapirism was particularly pronounced at ca. 2680–2670 Ma because it resulted in erosional unroofing of the SDC. The 2669 ± 2 Ma titanite age from sample 6 potentially records uplift-related cooling of the SDC following this event.

9. Conclusions U–Pb geochronological data outline a prolonged geological evolution for the southern Sleepy Dragon Complex, the initial stages of which are interpreted to document crustal growth and modification within the protocratonic Central Slave Basement Complex. Sleepy Dragon plutonic units dated here at ca. 2.95–2.93 Ga are representative of a major 3.0–2.9 Ga plutonic episode that affected much of the CSBC. A thin, >2.94 Ga mafic volcanic unit is also a part of the basement complex. Evidence for a cryptic, ca. 3.15 Ga basement component is obtained both from inherited and detrital zircons and highlights the antiquity and multicomponent character of the SDC. After 2.9 Ga, the CBSC was eroded and then overlain regionally by a thin, dominantly clastic sequence, the 2.85–2.80 Ga Central Slave Cover Group. In agreement with earlier studies, our ages of detrital zircons and a granite clast from the Patterson Lake Formation match those for underlying Sleepy Dragon basement units, suggesting a predominance of local detritus. Although transient stability of the CBSC is inferred from deposition of the Central Slave Cover Group, evidence of coeval plutonic and high-grade metamorphic activity suggests that the protocraton remained tectonically active during this interval, perhaps reflecting an overall extensional regime. Following a 75 Myr interval that is poorly understood, volcanism became pronounced, forming the principal mafic and (mainly younger) felsic volcanic elements of the Yellowknife Supergroup. In common with many Archean terrains worldwide, voluminous synvolcanic plutonism (mainly TTG suites) significantly modified the underlying crustal section, displacing and occluding older basement units. Within the SDC, emplacement of synvolcanic

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plutons and three mafic dyke swarms occurred between ca. 2.73–2.67 Ga, with preservation of basement units mainly along the margins of the complex. Elsewhere where this preservation pattern is observed (e.g., Pilbara craton), it has been interpreted to indicate plutonic doming. This feature, along with an unconformity of appropriate age along the southern margin of the SDC, suggests that the domal form of the SDC was initiated during pluton emplacement within the complex. The presence of a >3 km-thick mafic volcanic cover overlying a less dense, magmatically active middle to lower crust may have allowed magmatic diapirism and possibly also partial convective overturn of the crust to occur locally. Diapiric uplift and exhumation of Sleepy Dragon crust mainly at ca. 2.68–2.67 Ga is suggested here to be the primary reason for development of the Raquette Lake–Detour Lake unconformity. Subsequent regional folding resulted in a high-amplitude cross-sectional profile for the complex (Fig. 3a of Bleeker, 2002). It is not known how applicable this model is to other domal basement complexes of the Slave Province, but if local unconformities temporally equivalent to that beneath the Raquette Lake and Detour Lake formations occur elsewhere (e.g., Bleeker, 2001), the diapiric uplift model provides one means of explaining their similar ages. Acknowledgements This study was funded by LITHOPROBE, the Natural Sciences and Engineering Research Council of Canada (through a grant to Greg Dunning, Memorial University), and the Geological Survey of Canada. We thank all our field assistants and the staff of the geochronology labs at the Memorial University of Newfoundland, Royal Ontario Museum, and Geological Survey of Canada for their generous support during the course of this project. Use of the Department of Indian and Northern Affairs’ float-equipped aircraft at various times was greatly appreciated. Manuscript preparation was kindly supported by the ARC National Key Centre for Geochemical Evolution and Metallogeny of Continents (www.es.mq.edu.au/GEMOC/). Reviews by Bill Davis, Fernando Corfu, and Nick Culshaw helped to improve the manuscript. This is LITHOPROBE contribution no. 1359 and Geological Survey of Canada contribution no. 2003164.

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