by Wouter Bleeker

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The Slave craton of the northwestern Canadian Shield is one of the oldest and most distinct building blocks of North American cratonic lithosphere. It hosts Earth’s oldest intact rocks, the Acasta gneisses. These ancient gneisses are embedded in a large Mesoarchean to Hadean basement complex that underlies the west-central parts of the craton. Although itself poorly mineralized, the basement complex is overlain by Neoarchean supracrustal sequences, and is heavily intruded and cannibalized by plutonic suites that range in age from 2720-2670 Ma synvolcanic plutons to 2595-2585 Ma late-orogenic batholithic granites. Supracrustal sequences, collectively known as the Yellowknife Supergroup, are represented by an early cover sequence comprising quartzite and banded iron formation (ca. 2800 Ma), a thick dominantly tholeiitic greenstone sequence (ca. 2700 Ma), younger arc-like sequences (ca. 2690-2660 Ma), extensive turbidite blankets (ca. 2680-2620 Ma), and finally synorogenic conglomerates that were deposited at ca. 2600 Ma or shortly thereafter. The early cover sequence and the overlying tholeiites represent subaerial exposure and then volcanicdominated rifting of the basement. Arc-like sequences formed in part on top of the attenuated basement and in progressively widening, juvenile, back-arc-like basins and contain some of Canada’s largest undeveloped volcanogenic massive sulphide deposits. After 2680 Ma, much of the Slave craton became overlain by the Burwash Basin, one of the largest and best preserved Archean turbidite basins in the world and comparable in size and setting to the Japan Sea. During orogenesis, supracrustal sequences were telescoped, thickened, and multiply folded between ca. 2650 and 2580 Ma, with a peak in crustal anatexis between 2595 and 2585Ma (the “granite bloom”). Numerous orogenic gold deposits formed throughout the Slave craton, either as shear- or vein-hosted deposits in deformed greenstones or within the chemical traps provided by banded iron formations in the turbidites. Proterozoic rift-related magmatic suites and arcs around the margins of the craton host a variety of mineral deposits. Finally, the craton was intruded by several hundred Phanerozoic kimberlite pipes, some of which support Canada’s first diamond mines.

The Archean Slave craton (Figs. 1, 2, 3) (Henderson, 1981; Padgham and Fyson, 1992; King and Helmstaedt, 1997; Bleeker and Davis, 1999a) is a major building block of the Canadian Shield. It is one of approximately 35 Archean cratons preserved around the world (Bleeker, 2003). Amalgamation of the Slave craton with the Rae craton started at ca. 2 Ga, initiating the climactic growth of Laurentia from 2.0 to 1.8 Ga (Hoffman, 1988, 1989), probably within the broader context of the formation of Earth’s first modern supercontinent, Nuna. Much of the Slave craton is old, and within the context of the Laurentian collage it can be considered, for all practical purposes, as a far-travelled, if not exotic, fragment of crust relative to other well known cratons in Laurentia (such as the Superior, Hearne, Rae, and Nain, see Bleeker, 2003, 2004).

As a mere fragment of ancient crust (e.g. Hoffman, 1988; Isachsen and Bowring, 1994; Bleeker, 2003), surrounded by Paleoproterozoic rifted margins, the Slave craton originated from the break-up of a much larger late Archean landmass, possibly the speculative late Archean supercontinent Kenorland (Williams et al., 1991) or, perhaps more likely, a smaller landmass referred to as the supercraton Sclavia (Bleeker, 2003). The late Archean and earliest Proterozoic development and evolution of Slave crust should thus be viewed within the context of the growth and subsequent break-up of this larger Sclavia supercraton, even though the shape and size of this supercraton remain currently unknown. The salient point is that ancient cratons, like the Slave, preserve only parts of the much larger tectonic systems in which they were generated.

In agreement with this conceptual view, latest Archean events¹ are remarkably homogeneous across the Slave craton and may be used, together with pre-2.0 Ga Proterozoic mafic dyke swarms², to help identify neighbouring fragments of Sclavia from among the 35 extant cratons. One such Slave craton-wide event was a voluminous “granite bloom” between ca. 2595 and 2585 Ma (Davis and Bleeker, 1999). This singular event in the craton’s history transferred, irreversibly, a significant fraction of heat-producing elements and lower crustal fluids to the upper crust, thus allowing slow cooling and stiffening of the lower crust and setting the stage for cratonization and long-term preservation (Bleeker, 2002).

Predating this cratonization event, the Slave crust preserves a complex and spatially heterogeneous record of crustal growth spanning more than 1.5 billion years (e.g. Kusky, 1989; Isachsen and Bowring, 1994; Bleeker and Davis, 1999a; 1999b and references therein; Sircombe et al., 2001; Ketchum et al., 2004; see Iizuka et al. (2006) for a discussion on the recent discovery of a 4.2 Ga zircon xenocryst in Acasta gneiss).

The present paper briefly summarizes this crustal growth history and the overall geological evolution of the Slave craton, from the formation of early sialic basement, to the development of the dominant ca. 2.7 Ga greenstone sequences, major orogenesis at ca. 2.6 Ga, and final cratonization at ca. 2.55 Ga, thus providing a framework for discussion of important metallogenic events within the craton. Pertinent information concerning the location, mineralogy, structure, size, grade, age, and stratigraphic and tectonic settings for all significant mineral deposits within or adjacent to the craton is summarized in the Appendix, along with short summaries for some additional metallogenic events. Figure 4 shows the locations of these mineral deposits and occurrences on a geological map of the craton.

An ancient and largely crystalline basement complex underlies much of the central and western parts of the craton (Fig. 5, see also Figs. 2, 3) (Baragar and McGlynn, 1976; Kusky, 1989; Bleeker et al., 1999_a,b; Ketchum and Bleeker, 2001; Ketchum et al., 2004). It is referred to as the Central Slave Basement Complex (Bleeker et al., 2000).

Along the Acasta River, this basement complex consists of polymetamorphic gneisses of tonalitic to gabbroic composition (Fig. 6A) that yield protolith ages up to 4.03 Ga³ (Bowring et al., 1989; Stern and Bleeker, 1998; Bowring and Williams, 1999). The record of inheritance extends back to 4.2 Ga (Iizuka et al., 2006). The Acasta gneisses were essentially a chance discovery (Bowring et al., 1989; M. St-Onge, pers. comm., 2000) and no other rocks of this age have yet been found.Apart from a central core of the craton, with sporadic ages >3.5 Ga (Acasta to Point Lake), the Central Slave Basement Complex is mostly younger with important age modes4, from detrital and protolith U-Pb zircon ages, around 3400 Ma, 3150 Ma, 2950 Ma, and 2826 Ma (Fig. 5B) (e.g. Sircombe et al., 2001; see also Bleeker and Davis (1999b) for a compilation of basement ages).

Interestingly, complementary data from the mantle suggest that at least part of the subcontinental lithospheric mantle below the central part of the craton may be of similar antiquity (e.g. Aulbach et al., 2004). A crude age zonation can be recognized in the basement complex (Ketchum and Bleeker, 2001), although no easily interpretable tectonic pattern has yet emerged. Pre-2.9 Ga supracrustal rocks have been found at the base of some greenstone belts (Bleeker and Davis, 1999b and references therein; Ketchum et al., 2004), but form only a very small component of the craton’s geological inventory. An important example of such an occurrence is a felsic metavolcanic rock, dated at 3118 +11/-8 Ma, at the base of the supracrustal succession on the eastern flank of the Winter Lake belt (Hrabi et al., 1995; Hrabi and Grant, 1999). Elsewhere, enclaves of supracrustal origin are known to occur within pre-2.8 Ga basement but have yet to be systematically dated.

Mineralization

There are few, if any, known mineral occurrences of note within the Central Slave Basement Complex, a statistic that is generally mirrored by other Mesoarchean and older gneis-sic complexes in cratons around the world. To a first degree, this poor endowment is thought to correlate with the generally mid- to lower crustal erosion depths within such crystalline basement complexes and, consequently, the virtual lack of supracrustal rocks.

Indirectly, however, the presence of the ancient basement complex may have influenced or exerted important controls on several classes of younger mineral deposits:

• Seafloor volcanogenic massive sulphide (VMS) mineralization within the Neoarchean bimodal rift volcanic rocks that overlie faulted basement, such as the basal greenstones of the Yellowknife and Courageous Lake greenstone belts.
• Gold deposits hosted by quartz veins and/or shear zones that developed along the flanks of major basement uplifts.
• Late Archean evolved granites and their associated pegmatite swarms, some of which are enriched in rare elements (e.g. Li, Be, Sn, Nb, Ta). Such enrichment typically correlates with a presence of ancient and multiply recycled sialic crust.
• Diamondiferous kimberlite pipes in the Lac de Gras region (Fig. 7), which appear to have ascended through the edge of the Central Slave Basement Complex (e.g. see cross-section of Fig. 3).

The contiguous nature of the basement complex, by at least 2.9 Ga, is indicated by a thin (typically 2-200 m) but widespread ca. 2.9 to 2.8 Ga cover sequence of quartzite, rare rhyolite, and banded iron formation (Fig. 8, see also Fig. 6B) collectively known as the Central Slave Cover Group (see also Covello et al., 1988; Roscoe et al., 1989; Padgham, 1992; Bleeker et al., 1999b, 2000). This sequence, which is locally intruded by ultramafic sills (Fig. 8), marks the onset of the Neoarchean cycle of supracrustal development (Bleeker et al., 1999b), known collectively as the Yellowknife Supergroup (Henderson, 1970).

The supermature and commonly fuchsitic quartzites that are characteristic of this sequence overlie a regional unconformity that marks the emergence and erosional unroofing of the basement complex in what was probably an aggressive, CO₂-rich, Archean atmosphere (e.g. Kasting, 1993). The quartzites, which locally preserve cross-bedding, mark the progressive drowning of this unconformity in a tide-influenced coastal setting. Overlying banded iron formations (BIFs) are indicative of deeper water, more outboard sedimentation in response to continued subsidence.

Abundant detrital chromite in the quartzites (Figs. 8C,E,F) may suggest contemporaneous komatiitic volcanism. Similar fuchsitic quartzite sequences occur in many other cratons worldwide, particularly between ca. 3.1 Ga and 2.8 Ga. After 2.4 Ga, mature quartzites are rarely fuchsitic, reflecting a lesser role for detrital chromite (and komatiites) in the post-Archean world.

Mineralization

The Central Slave Cover Group hosts some of the more prominent BIFs of the Slave craton (Fig. 4), although most are thin (1-10 m) and variable in composition along strike, changing from oxide-iron formation into silicate-rich varieties, or merely ferruginous cherts. Locally, however, folding has thickened highly magnetic Fe-oxide BIFs into substantial thicknesses (e.g. at Amacher Lake, on the eastern flank of the Sleepy Dragon Complex), resulting in some of the highest amplitude total field magnetic anomalies in the Slave craton. Overall, the BIFs appear of low economic value, although some may possibly host epigenetic gold mineralization and may be under-explored for this commodity. However, across the craton, most of the known iron formation- hosted gold mineralization is associated with younger BIFs that occur intercalated within low- to mediumgrade turbidite packages (Fig. 4).

Fuchsitic quartzites that occur stratigraphically below the iron formations are enriched in detrital heavy minerals, including highly stable species such as chromite, zircon, and rutile. Individual, detrital, small black chromite grains are a characteristic feature of these otherwise white to grey quartzites (Bleeker et al., 1999b) and, where recrystallization is strong, allow these clastic rocks to be distinguished from metacherts. During metamorphism and deformation the detrital chromite grains reacted to varying degrees with surrounding minerals to produce bright green fuchsitic mica5 (Figs. 8E,F). In a few localities, chromite was concentrated enough to form seams of “black sand”. These occurrences, although of scientific interest, are too small to be of economic value. If road access were available, some of the green-white quartzite would make attractive building or decorative stone. In Greenland, India, and Australia, similar quartzites are commonly quarried for this purpose. Elsewhere in the world, quartzites similar to those described here contain paleoplacer deposits of gold and/or uranium (e.g. the Witwatersrand quartzites of South Africa; basal Huronian quartzites near Elliot Lake, Canada) (e.g. Roscoe and Minter, 1993). An initial survey of such potential in the Slave craton was carried out by Roscoe (1990, 1992), indeed resulting in anomalous values of gold and uranium.

Ultramafic sills (or flows?) have intruded the cover sequence in several places, and locally contain seams or veins of magmatic chromite (Covello et al., 1988). Economic concentrations have not been found. In one remote locality, on the south shore of Desteffany Lake, one of the authors found sulphide concentrations adjacent to ultramafic rocks within the cover sequence. Overall, however, the volume of komatiitic rocks and the potential for substantial massive Ni-Cu mineralization appears limited, not only at this stratigraphic interval but throughout the Slave craton.

Wherever the thin cover sequence is recognized, it is overlain by a thick and extensive sequence of tholeiitic basalts, with minor komatiite and rhyolite tuff intercalations (Fig. 8). In the Yellowknife greenstone belt, this basalt-dominated volcanic sequence (Fig. 9) is known as the Kam Group (Helmstaedt and Padgham, 1986; Bleeker et al., 1999b). Possible correlative basalt successions are known across the basement domain (Fig. 10), as far east as the Courageous Lake belt, and at least as far north as around the Exmouth antiform in the Acasta area (Bleeker et al., 2000). This basalt sequence typically consists of several hundred metres to several kilometres of pillowed and massive flows, intercalated with thin felsic volcaniclastic horizons, and intruded by numerous dykes and sills of multiple generations (e.g. Henderson and Brown, 1966).

Well dated components of this basalt-dominated sequence yield ages from 2722 to 2697 Ma (Davis and Bleeker, unpublished data, Isachsen and Bowring, 1997). A mafic dyke cutting across the lower part of the sequence north of Yellowknife, dated at 2738 Ma (J. Ketchum, pers. comm., 2004), demonstrates that parts of this basalt sequence are even older. In Yellowknife, the top of the sequence is represented by voluminous basaltic flows and intercalated felsic volcanic rocks of the Yellowknife Bay Formation, dated at ca. 2700 Ma (Fig. 9). Similar 2700 Ma ages have been obtained from the Courageous Lake and Acasta areas and strongly support the overall regional correlation (Bleeker et al., 1999b). Stratigraphy, dense dyke swarms, and isotopic data link the basalt sequence to the basement (Henderson, 1985; Bleeker et al., 1999a,b; Northrup et al., 1999; Cousens, 2000; Bleeker, 2002 and references therein).

If the broad regional correlation of these basalts is valid, the magnitude of volcanism (areal distribution >100,000 km2, typical thickness 1-6 km) approaches large igneous province (LIP) proportions (Coffin and Eldholm, 1994, 2001; Ernst et al., 2005). The widespread basaltic volcanism probably accompanied protracted rifting of the basement complex, possibly assisted by mantle plume activity. The stratigraphic succession in the Yellowknife area is compatible with such a rifting interpretation. At the top of the Kam Group, bimodal volcanic rocks of the Yellowknife Bay Formation become progressively more intercalated with volcaniclastic sediments, before final intrusion by thick tholeiitic sills. One of these, the Kam Point gabbro sill, has a preliminary baddeleyite age of ca. 2697 Ma (Davis and Bleeker, unpublished data; see Fig. 9).

From a mantle perspective, it seems likely that events associated with the voluminous basaltic volcanism across the ancient basement terrain must have involved thinning or at least modification of the lithospheric mantle below the Central Slave Basement Complex. Large-scale melting was probably triggered by adiabatic rise of asthenospheric mantle. Perhaps, then, the ca. 2.7 Ga basaltic volcanism may have contributed to the highly depleted mantle compositions underlying the core of the craton (e.g. Griffin et al., 1999; Grutter et al., 1999; Kopylova and Russell, 2000; Carbno and Canil, 2002).

Mineralization

A volcanically dominated rift environment, characterized by bimodal volcanism and minor aprons of volcaniclastic sedimentary rocks, is a highly favourable environment for seafloor hydrothermal activity and the formation of volcanogenic massive sulphide (VMS) deposits. Indeed, numerous showings of sulphidic horizons occur throughout the basalt-dominated greenstone belts of the central and western Slave craton (Fig. 4).

Of particular interest are intercalated felsic volcanic flows and/or sills, which are direct indicators of proximity to a differentiated magmatic centre, and thus a long-lived subvolcanic heat source. The Bell Lake quartz-porphyritic tonalite sill (Fig. 6C) and the rhyolitic Townsite Formation, dated at 2713 ± 2 Ma and ca. 2709 Ma, respectively (Davis et al., 2004), are examples of such proximal felsic volcanic rocks in the Yellowknife greenstone belt. Hydrothermal alteration and minor sulphide mineralization are known from the Yellowknife Belt (e.g. the Homer Lake showings, see Fig. 4), but to date no deposits of potential economic interest have been found. Similarly, despite at least a first wave of exploration across the other basaltic greenstone belts of the west-central Slave craton, the authors are not aware of any major discoveries.

Although possibly unrelated, it is worth mentioning here the VMS deposits of the High Lake greenstone belt of the north-central Slave craton. These deposits, the High Lake A/B, D, andWest Zone (Figs. 4, 11), along with several other sulphide-rich horizons (Wolfden Resources Inc., 2005), are associated with bimodal volcanic rocks dated at 2705 Ma (Henderson et al., 2000). Thus the setting and age are compatible with a mafic volcanic- dominated rift environment similar to that of the Kam Group. However, basement rocks have not been identified in the immediate area. Metal ratio plots (Fig. 12) discriminate the High Lake deposits from most other VMS deposits in the Slave, suggesting a similarity to Cyprus-type deposits (the “mafic class” of Barrie and Hannington, 1999), which form in mafic volcanism-dominated extensional settings.

Following ca. 2.7 Ga basaltic volcanism and rifting, most areas in the Slave craton show a transition to calc-alkaline volcanism characterized by abundant felsic and intermediate volcanic rocks, tholeiitic to calc-alkaline basaltic rocks, and intercalated volcaniclastic sedimentary rocks (Fig. 10). In nearly all areas, these rocks are stratigraphically overlain by turbiditic greywacke-mudstones (Fig. 10). Ages for the arclike volcanic rocks and their plutonic counterparts typically range from 2690 to 2660 Ma (e.g. Mortensen et al., 1988; Isachsen, 1992; Mortensen et al., 1992a,b; Villeneuve and Henderson, 1998; Northrup et al., 1999; Pehrsson and Villeneuve, 1999; Ketchum et al., 2004).

The arc-like volcanic rocks tend to be geochemically juvenile (e.g. Davis and Hegner, 1992; Cousens et al., 2002). They dominate the eastern part of the craton, where they lack any apparent association with older basement, its cover, and/or the basalt-dominated rift sequence discussed above. These observations have led to models in which the eastern Slave represents an exotic juvenile arc (the “Hackett River arc”) that collided with the basement domain in the west (e.g. Kusky, 1989, 1990). However, similar arc-like rocks, with identical ages, stratigraphically overlie the basement domain and its cover in the west-central parts of the craton (Fig. 10A), where they can be tied to the basement and bimodal rift volcanic rocks by means of unconformities (Bleeker, 2001), crosscutting feeder dykes, and subvolcanic intrusions (see Fig. 13) (Bleeker et al., 1999a).

It thus appears that, if these rocks were generated in an arc-like setting, this arc was constructed marginal to, and on top of, the highly extended continental crust of the Central Slave Basement Complex. This suggests a marginal to continental arc setting. The arc must have been actively extending, evolving into a back-arc basin that was ultimately filled with turbiditic sediments (Figs. 10, 13). The geochemistry of the volcanic rocks and the associated subvolcanic plutons, although isotopically juvenile, typically shows arc-like signatures (negative Nb and Ta anomalies, light rare earth and large-ion lithophile element enrichment, and relative depletions of heavy rare earths) compatible with enriched sources in a supra-subduction zone setting and final equilibration with garnet-bearing residues. Alternative models invoke partial melting of a mafic underplate, perhaps accompanying delamination (e.g. Cousens et al., 2002).

Mineralization

The arc-like volcanic sequences spanning the interval 2687 to 2660 Ma host the Hackett River, Izok Lake, Sunrise, Gondor and Kennedy Lake deposits (Fig. 4).An exception is the High Lake deposits, which are associated with older, ca. 2705 Ma, bimodal volcanic rocks (discussed above).

Nearly all the VMS deposits of this group are associated with proximal felsic volcanic rocks at or near the transition to overlying turbiditic metasedimentary rocks. This transition is characterized by rhyodacite to rhyolite complexes, volcaniclastic sediment aprons, thin sulphidic chert horizons, and, in some localities, banded iron formations. Carbonate rocks (calc-arenites) are associated with some of the felsic complexes (Fig. 6D). This typical stratigraphic evolution, from shallow water or emergent felsic volcanic complexes to deep-water turbidite sedimentation, suggests active extension, tectonic subsidence of the arc environment, and creation of significant accommodation space that was then filled by deep-water sediments (Fig. 13). Such an environment of active faulting, attenuated lithosphere, active volcanism, and high heat flow, has long been recognized as a favourable setting for development of large VMS deposits.

Other characteristic features of this high heat flow regime are the presence of high-level synvolcanic intrusions, for instance in the Hackett River belt (see Fig. 14; the Sandy Hill Pluton, Hanimor Gneiss Dome, and smaller intrusive bodies that lie within the Malley Rapids Anticlinorium). The larger VMS occurrences of the Hackett River belt (A, East Cleaver, Boot Lake, Cleaver Low Grade, Yava, and Musk) exhibit a strong spatial relationship, and hence inferred genetic relationship to these intrusions. It is possible that these synvolcanic intrusions contributed magmatic fluids to the ascending hydrothermal solutions, resulting in the unusually high silver and lead contents for these deposits. Stratiform carbonate rocks are a common feature associated with the Hackett River deposits as well as the BB, Bear, and Turnback Lake (XL) deposits. A likely origin for these carbonate deposits is the venting of Ca-saturated hydrothermal fluids onto the seafloor and rapid mixing with cold, higher pH seawater (i.e. analogues to “white smokers” in modern seafloor settings). If correct, these carbonate deposits represent excellent marker horizons and vectors in the search for new volcanogenic massive sulphide deposits.

Starting at ca. 2680 Ma, a broad turbidite basin — the Burwash Basin — developed across much of the craton and progressively buried the volcanic substrate (e.g. Henderson, 1985; Ferguson et al., 2005). The transition from volcanic and volcaniclastic rocks to deep-water sediments is commonly conformable or disconformable6. In a number of localities, however, the transition is marked by a well developed unconformity or nonconformity7 that cuts down into underlying crystalline rocks and is overlain by shallow-water clastic rocks, including conglomerates with granitoid cobbles (Henderson, 1985; Bleeker et al., 1997; Bleeker, 2001). One such locality occurs along the southwestern flank of the Sleepy Dragon Complex, another at Point Lake, and they demonstrate an autochthonous nature of the sedimentary sequence (Bleeker, 2001). Detrital zircons in the basal clastic rocks indicate that deposition began sometime after 2683 Ma (Bleeker et al., 1997).

A persistence of volcanic intercalations up-section and late mafic sill complexes suggest a volcanically active extensional setting, perhaps best compared with modern back-arcs. The minimum size of this basin was ca. 400 x 800 km (Fig. 15A), making it the largest and possibly best preserved Archean turbidite basin in the world, comparable in size to the Japan Sea. Like the modern Japan Sea, the Burwash Basin was largely ensialic, in agreement with inferences by early workers (e.g. Henderson, 1985).

The Burwash Basin fill consists largely of immature greywackes and mudstones, deposited below wave base, and locally may be up to 10 km thick (Bleeker and Beaumont-Smith, 1995). Intercalated tuff layers have been dated at 2661 Ma (e.g. Bleeker and Villeneuve, 1995). Across the Slave craton, the greywacke turbidites have been given different formational names: the classical Burwash Formation in the Yellowknife Domain (Henderson, 1972); the Contwoyto Formation in central and northern Slave, identical in essentially all aspects to the Burwash Formation further south, except for the presence of intercalated iron formations; the Itchen Formation, a more mud-rich facies in the north-central Slave; and the Beechey Lake Group in the northeastern Slave, which also contains iron formations. Many of the turbidite beds, particularly those of the Burwash and Contwoyto Formation, are sand dominated with only thin silt to mud intervals at the top of the graded beds. In the Yellowknife Domain, thick amalgamated sand beds (2-10 m) are not uncommon (Fig. 6F). Petrography, detrital zircons, and geochemical analysis indicate that the greywacke detritus consists of a mixture of mafic and felsic volcanic rocks and uplifted plutonic infrastructure, with minor input from ancient basement rocks (Henderson, 1972; Jenner et al., 1981; Henderson, 1985; Yamishita and Creaser, 1999; Ferguson et al., 2005). The main axis of the basin and subsequent structural trends appear to have been northeast-southwest (Fig. 15), distinctly across the north-south isotopic boundaries that track the nature of deep basement. This interpretation is based on the following observations:

• Identical Burwash Formation turbidites extend from near the type area of Yellowknife (e.g. Henderson, 1972) to the northeastern Slave (Figs. 10, 15A).
• Banded iron formations in the turbidites are restricted to the northwest half of the craton (Fig. 15A) suggesting a northeast to southwest facies boundary or tectonic trend across the basin (see also Padgham, 1992). This argument is complicated however by the observation that some iron formations in the western part of the craton are hosted by a younger turbidite package, the Damoti Lake sequence (see Fig. 10E).
• Earliest folds in the turbidites, which formed between 2650 and 2630 Ma, have northeasterly trends after qualitative “unfolding” of younger superimposed fold generations (Fig. 15B) (e.g. Bleeker, 1996).
• The earliest plutonic suite that intrudes folded Burwash strata, the ca. 2630 Ma Defeat Suite, appears to form a northeast-southwest-trending magmatic belt across the southeastern half of the craton (Fig. 15C) (Davis and Bleeker, 1999).

With more and better U-Pb zircon ages, a tentative “volcanic line” of 2661Ma felsic volcanic complexes, coeval with turbidite sedimentation, has begun to emerge (Fig. 15A; Bleeker and Davis, unpublished data). This volcanic line also trends northeast-southwest, from the Hope Bay belt to Yellowknife, and may represent the first recognition of a linear arc system.

Mineralization

The immature greywackes and mudstones of the Burwash Basin contain few primary mineral deposits other than banded iron formations (Fig. 6G). The latter occur intercalated in greywackes scattered across a broad swath in the northern part of the craton (Fig. 15A), from the Goose Lake and George Lake areas to the Point Lake area, and from there to the southwestern Slave. Many of the BIFs are highly magnetic. Although of scientific interest for the understanding of facies boundaries, basin evolution, and geochemistry, they are uneconomic in terms of their ferrous metal content.

Within and adjacent to the Back River Volcanic Complex, Jefferson et al. (1989) described three distinct iron formation horizons: 1) intercalated within the volcanic rocks of the Back River Complex itself; 2) at the upper contact of the Back River volcanics; and 3) an upper horizon well within the turbidites of the Burwash Formation. Lateral facies changes exhibited by the mineralogy of these iron formation units vary from oxide facies (magnetite±hematite and quartz), through silicate facies (chert-grunerite±stilpnomelane± chlorite), to carbonate facies (siderite ±quartz).

The principal type of economic mineralization within Burwash Formation metaturbidites is epigenetic gold mineralization hosted by the intercalated BIFs (Padgham, 1992). The most important example of this deposit type is the Lupin deposit on the southern shores of Contwoyto Lake (Bullis et al., 1991a,b; Geusebroek and Duke, 2004), which has been a significant gold producer from 1982 to 2003, yielding between 3 and 4 million ounces ofAu (Normin, 2005). Other examples, such as George Lake (Fig. 16) and Goose Lake, occur throughout the northern Slave craton and may become economic with elevated gold prices and better access. The general model for these deposits is that the host BIFs served as chemical traps for gold-bearing, low-salinity, mixed H2OCO₂ fluids during metamorphism and deformation (Bullis et al., 1991b; see also Kerswill, 1993; Phillips, 1993). Destabilization of theAu-carrying sulphur complexes, due to interaction with reduced Fe-rich host rocks, led to alteration and gold deposition, either in veins or in fluid-altered and sulphidized zones of the iron formations. The structural timing of these epigenetic deposits is generally syn- to late-kinematic and syn- to late-metamorphic, i.e.,consistent with maximum fluid production deeper in the thickened structural- metamorphic pile. The most likely source for the fluids, and the gold, is metamorphic devolatilization of a voluminous, immature, sediment pile and its volcanic substrate (Phillips, 1993). Sporadic iron formations provided accidental traps to the migrating fluids, with discrete structures locally playing a role in increased focusing of fluid flow. Similar processes also led to gold-bearing quartz veins within metaturbidites (e.g. laminated veins along sheared bedding planes, saddle reefs) (Boyle, 1961, 1986), but without a large-scale focusing mechanism these occurrences and deposits tend to be of small size, although locally of high grade. Examples are the Ptarmigan and Discovery mines in proximity to Yellowknife (Fig. 4) (e.g. Brophy, 1987).

Turbidite sedimentation in the Burwash Basin came to an end sometime before 2650Ma, the age of the oldest recorded granitoid (feldspar porphyry) pluton intruding Burwash strata (Point Lake area, Mueller et al., 2001). Subsequent tectonic events record the closure and folding of the Burwash Basin (D1) prior to 2630 Ma (see F1 fold belt in Fig. 15B). The latter age constraint is provided by early plutons of the Defeat Suite (Davis and Bleeker, 1999), a distinct and possibly subduction-related magmatic suite across the southern (and southeastern) Slave craton (Fig. 15C).

Closure of the highly extended, but largely ensialic back-arc basin allowed considerable shortening and mobility but with a structural style dominated, at least at high structural levels, by fairly systematic, mostly upright, northeastsouthwest trending fold trains (e.g. Bleeker and Beaumont-Smith, 1995). At deeper levels, e.g., along the basement-cover interface, the fold trains must have detached allowing differential shortening of the basement and cover (e.g. Kusky, 1991). The folded Burwash strata do not represent an outboard accretionary prism however (cf. Kusky, 1991), which would require a trench setting rather than the more likely ensialic back-arc setting; hence, there is little evidence for a discrete “Contwoyto terrane” (Kusky, 1989) in the central Slave craton.

Interestingly, the northeastsouthwest structural grain of the F1 fold belt is also recognized in the lithospheric mantle (Grutter et al., 1999). Shallow subduction (either from the southeast or northwest?) may have emplaced distinct mantle slabs (Davis et al., 2003b). These processes terminated with docking of an outboard terrane (e.g. Fig. 13), either in the southeast or the northwest; however, this terrane is not preserved within the present limits of the Slave craton. Finally, crustal thickening led to uplift and erosional exhumation of folded Burwash strata and partial unroofing of Defeat Suite plutons. Detrital zircons of Defeat Suite age are recorded in younger sedimentary packages (e.g. Fig. 10F).

Mineralization

Folding and incipient crustal thickening (D1), and the onset of regional metamorphism, together with Defeat Suite plutonism, must have initiated devolatilization reactions and metamorphic fluid flow. Although these events thus likely kick-started the development of epigenetic gold mineralization, they were followed and overprinted by much more intense metamorphic events ca. 20 to 30 million years later, during D2 deformation.

Arc-generation and subduction processes almost certainly modified the mantle lithosphere below the Slave craton, possibly creating the starting conditions for what is now a thick diamondiferous mantle root (e.g. Davis et al., 2003b). Interestingly, trends of similar mantle domains, based on indicator mineral chemistry, appear to parallel the northeastsouthwest trends of the Burwash Basin and D1 folding (Grutter et al., 1999).

Along the western and northwestern margin of the craton, younger turbidites containing ca. 2630 Ma detrital zircons (Figs. 10E, F, and 15D) (Sircombe and Bleeker, unpublished SHRIMP data, Pehrsson and Villeneuve, 1999; Bennett et al., 2005; Bennett, 2006) record a migration of sedimentation and tectonic activity to the northwest. Deposition was coeval with uplift and erosional unroofing of Defeat plutons and tightly folded Burwash Formation strata. Shortly following their deposition, these younger turbidites were shortened and intruded by ca. 2616 to 2608 Ma tonalite-granodiorite plutons similar to the Concession Suite of the Contwoyto Lake area (Davis et al., 1994; Pehrsson and Villeneuve, 1999).

In the multiply folded, metamorphosed, and intermittently exposed terrain of the western Slave craton, it has proven difficult to distinguish these younger turbiditic greywackes from Burwash Basin turbidites. There is no sharply defined demarcation line that separates the two turbidite packages and recognition of the younger sequence relies heavily on the absence of Defeat Suite-age plutons and the presence of <2640Ma detrital zircons. Preliminary work suggests that the younger turbidite sequence contains abundant intercalated iron formations, mostly of silicate facies, those of the Damoti Lake area representing one of the more significant examples. Many of the iron formations are “lean” (i.e., ironstones of low to moderate Fe content), comprising background turbiditic greywacke variably enriched in metamorphic garnet, other Fe-rich silicates, and/or disseminated Fe sulphides.

A distinct volcaniclastic sediment and greywacke package, associated with felsic volcanic rocks and subvolcanic intrusions, occurs along the tightly folded synclinal core of the High Lake greenstone belt of the northern Slave craton (see Figs. 10D, 15D) (Henderson et al., 2000). This package, dated at approximately 2616 to 2612 Ma (Henderson et al., 2000), is of significance in that it is one of the few examples of a preserved volcano-sedimentary carapace coeval with one of the major plutonic suites, i.e., the Concession Suite.

Mineralization

Types of mineralization within the younger (turbiditic) greywacke packages are similar to those in folded Burwash Basin strata. A principal example of epigenetic gold mineralization is that hosted by silicate facies iron formation in the Damoti Lake area. Similar iron formations occur all along the western margin of the Slave craton, from the Russell Lake area in the south to the Emile River area in the north (see Fig. 15D), and most have been moderately explored for gold and base metals. Scattered gold mineralization also occurs in numerous “lean” iron formations throughout the western Slave (Padgham, 1992), e.g., the Wheeler and Germaine Lake areas (“W” in Fig. 15D). The stratigraphic status of the iron formations and the host turbidites in this particular area remains currently unresolved. Extensive field investigations have shown that an earlier reported 2612 Ma “tuff bed” (e.g. Isachsen and Bowring, 1994) is in fact another crosscutting porphyry dyke. Hence, the turbidites in this area are >2612 Ma and could be of Burwash age.

Starting at ca. 2600 Ma, the entire craton was affected by cross-folding and significant further shortening (D2), probably in response to final collision along a distant active margin of the growing supercraton Sclavia (Bleeker, 2003). Structural trends, such as axial planes of folds and the dominant cleavage resulting from this event, are generally northsouth to northwest-southeast, but may refract strongly around more rigid plutons or other heterogeneities (e.g. Davis and Bleeker, 1999). Moderate overthickening of the crust led to high-T, low-Pmetamorphism (Thompson, 1989), widespread anatexis, the appearance of S-type granites, and a hot and weak lower crust. These processes culminated in ca. 2590 Ma extension and the craton-wide “granite bloom” (Kusky, 1993; Davis and Bleeker, 1999). The intrusion of carbonatites (Villeneuve and Relf, 1998) and involvement of other mantle-derived melts (e.g. sanukitoids, see Bennett, 2006) indicate a role for mantle processes (delamination?). Overall timing relationships are summarized in Figure 17A and B.

While peak temperatures were attained in the lower crust, large basement-cored domes were amplified by buoyancydriven deformation (Fig. 3); lower crustal devolatilization reactions mobilized gold-bearing fluids; and synorogenic clastic basins formed and were immediately infolded into tight synclines (Bleeker, 2002). At least one of these synorogenic clastic basins, the Beaulieu Rapids Formation (Corcoran et al., 1999) may have formed as late as ca. 2580 Ma (Sircombe and Bleeker, unpublished SHRIMP data; see detrital zircon age spectra in Fig. 17B). Late strike-slip faults overprinted and truncated the synclinally infolded clastic basins. Examples of such faults are the Yellowknife River Fault Zone in Yellowknife, and the Beaulieu River Fault Zone along the eastern flank of the Sleepy Dragon Complex (see Fig. 2). Over a period of ca. 70 million years, the lower crust cooled (Bethune et al., 1999), finally coupled with the mantle, and the Slave (within the larger context of Sclavia) became a craton.

Mineralization

Strong penetrative regional deformation, culminating between 2600 and 2590 Ma as constrained by synkinematic granite sheets (Davis and Bleeker, 1999), represents the most obvious deformation throughout much of the Slave craton. It must have driven moderate to significant crustal thickening and led to the main thermal peak of regional metamorphism in most areas. The D2 deformation and associated metamorphism was the main driver for epigenetic gold mineralization throughout the Slave craton. In the Yellowknife greenstone belt, it led to formation of the approximately 15 million ounce Con-Giant Au deposit, along a complex system of mostly reverse shear zones. As is typical for this class of deposits, the Con-Giant system occurs mostly within moderate to strongly deformed basaltic rocks, in proximity to a regional stratigraphic break, the Yellowknife River Fault Zone (Figs. 2, 3). An asymmetric synclinal panel of synorogenic conglomerates (the Jackson Lake Formation) occurs along this fault zone. Similar relationships are observed in several other major Archean gold camps, most notably Timmins, Kirkland Lake, and Kalgoorlie. The critical control common to all these camps is localization of Au mineralization within significant bends of the regional fault zones; these bends were probably dilational during emplacement of the gold-bearing quartz veins, providing favoured pathways for focusing of upward fluid flow.

Although numerous other volcanic-hosted gold vein systems are known from the Slave craton, some of which saw brief production in the past, only one other major camp has emerged in recent years. This camp occurs in the Hope Bay belt, on the Coronation Gulf coast of the Slave craton, and consists of a string of deposits (Boston, Madrid, Doris; see Fig. 18) that are being readied for production. Elsewhere, overall potential for this class of deposits remains excellent, in spite of significant past exploration. Several greenstone belts throughout the Slave craton have very similar struc-tural-stratigraphic characteristics to that of the Yellowknife belt, including a thick, folded turbidite pile adjacent to a basaltic greenstone belt, a regional deformation zone, and a young, tightly infolded, synorogenic conglomerate package. The best examples are the Point Lake andArcadia Bay areas. The Courageous Lake belt is equally interesting although it lacks a synorogenic conglomerate assemblage. This belt hosts the Slave craton’s largest unexploited gold resource (Tundra-Fat deposit at 29.5 million tonnes containing 6.86 g/tonne gold).

Another class of mineral deposits related to final orogenesis is that of rare-element-enriched granitoids (Li, Be, Sn, Nb, Ta), in particular highly evolved anatectic granites and their late-stage pegmatites (e.g. Meintzer and Cerny, 1983; Meintzer et al., 1984; Meintzer and Wise, 1987). Tin (cassiterite) and Li (spodumene) were briefly mined from such pegmatites in the Yellowknife Domain, but several other pegmatite fields are known across the Slave craton (e.g. Tomascak and Cerny, 1992). From a global metallogenic point of view, these occurrences are of interest as they tend to be diagnostic for the presence of ancient, multiply recycled felsic crust.

The youngest granite plutons of the Slave craton are ca. 2590 to 2580 Ma (James and Mortensen, 1992; van Breemen et al., 1992; Davis and Bleeker, 1999; Bennett et al., 2005; Ootes et al., 2005). Only some pegmatitic granites are known to be significantly younger, for instance a pegmatite in the Winter Lake belt dated by U-Pb monazite at 2550 ± 1 Ma (M. Villeneuve, pers. comm., 2006).

Between ca. 2595 and 2585 Ma, an enormous volume of granite was generated throughout much of the Slave craton. This “granite bloom”, driven by moderate tectonic overthickening (D1-D2) and high intrinsic heat production, irreversibly transferred a significant fraction of heat-producing elements and lower crustal fluids (andAu) to the upper crust. In the lower crust, it must have involved large-scale migration of anatectic granitoid magmas, significant horizontal channel flow of partially molten rocks, development of horizontal layering, and flattening of the Moho discontinuity into a stable density configuration. Collectively, these processes allowed the lower crust to cool and stiffen over several tens of millions of years. U-Pb geochronology of lower crustal xenoliths shows that at depth high-grade metamorphic reactions and zircon growth continued to about 2510 Ma (Davis et al., 2003a). Finally, sufficient cooling allowed the crust to mechanically couple with the mantle (Bleeker, 2002). The end product was cratonic crust of high relative strength.

Following cratonization, there is an approximately 300 million year time span during which there are few recorded events within the Slave craton (Fig. 17B). Ca. 2.45 Ga magmatism, known from many other cratons around the world (e.g. Heaman, 1997), so far appears to be absent from the Slave craton.

At 2230Ma, northeast-trending mafic dykes of theMalley swarm intruded the central Slave craton, providing the first evidence for Paleoproterozoic mantle-derived magmatism and attempted rifting events (LeCheminant and van Breemen, 1994). Subsequent to the Malley event, between 2200 and 2000 Ma, Slave crust was affected by as many as ten other mafic dyke swarms (see Ernst and Buchan, 2001, for a global data base, and 2004) and associated extension events before the craton finally broke out of the confines of its ancestral Sclavia supercraton. Details remain sketchy however. The eastern margin of the Slave craton, now involved in and overridden by the Thelon Orogen, almost certainly became established as a passive margin well before the western margin. The latter is referred to as the so-called Coronation margin and later became involved in Wopmay Orogen (e.g. Hoffman, 1980; Bowring and Grotzinger, 1992; Hildebrand and Bowring, 1999). Once liberated out of Sclavia, the Slave continental microplate likely experienced a drift phase as an independent craton, before being progressively incorporated into the growing Laurentian collage and the supercontinent of Nuna (Hoffman, 1988; Bleeker, 2003).

Paleoproterozoic amalgamation processes varied along the margins of the Slave craton. In the east, the Slave acted as a lower plate, being overridden by the west-vergent Thelon Orogen (Fig. 3). In the south, along the shores and islands of Great Slave Lake, deformation was mainly transpressional along an oblique suture that evolved into a major continental transform boundary (Hoffman, 1987; Hanmer et al., 1992). In the west, at least two arc terranes were involved (Hottah terrane and the younger Great Bear Magmatic Zone, Ghandi et al., 2001; Ghandi and van Breemen, 2005), followed by oblique folding and late-stage, dextral, strike-slip deformation along the Wopmay Fault Zone. Development of the Great Bear arc, between about 1880 and 1840 Ma, likely involved subduction of Paleoproterozoic lithosphere below the western Slave craton (e.g. Bostock, 1998; Heaman et al., 2002).

Post-dating the assembly of Laurentia and Nuna, the Slave craton, particularly along its margins, became partially buried beneath intracontinental Proterozoic basins (e.g. Coppermine Homocline and Elu Basin, see Kerans et al., 1981; Hoffman and Hall, 1993). At ca. 1269 to 1267 Ma, the craton was partly uplifted and intruded by the giant Mackenzie dyke swarm, radiating from a plume centre west of Victoria Island (LeCheminant and Heaman, 1989, 1991; Baragar et al., 1996). Coeval flood basalts were erupted in basins now preserved as the Coppermine Homocline. This is the last major event affecting the core of the craton, although some younger mafic magmatic events did affect its edges (e.g. the 723 Ma Franklin event and the ca. 780 Ma Hottah sheets, Heaman et al., 1992; Harlan et al., 2003). Since that time, Slave crust has been “bobbing” gently up and down, with interior seas expanding and receding across the craton. Ordovician and Cretaceous-Tertiary sedimentary rocks and fossils are known as wall-rock fragments in some of the central Slave kimberlites (Nassichuk and McIntyre, 1995; Stasiuk and Nassichuk, 1995; Cookenboo et al., 1998; Stasiuk et al., 1999).

Despite the relative stability at the surface, melting events were triggered in the subcontinental mantle lithosphere and underlying asthenosphere, leaving their traces as clusters of kimberlites across the craton (Fig. 19). From the several hundreds of kimberlites now known across the craton, the fol-lowing ages have been recorded: Cambrian, Siluro-Ordovician, Permian, Jurassic, Cretaceous, and finally Eocene (e.g. Heaman et al., 2003, 2004). The exact triggering events are still unclear but may relate to 1) mantle plumes or convective upwellings disturbing the stress and thermal state of the lithosphere (e.g. Sleep, 2003), or 2) rapidly changing far-field stress patterns within the North American plate due to events along its periphery, such as break-up and opening of the Iapetus Ocean, or changing subduction patterns along the Cordilleran margin (e.g. Snyder and Lockhart, 2005).

Mineralization

From the first attempted rifting events at ca. 2230 Ma (Malley dyke swarm) to final break-up of the supercraton Sclavia at 2.1 to 2.0 Ga, at least ten mafic and alkaline magmatic suites were emplaced within Slave crust, as dyke swarms and intrusive complexes. Several of these events are of interest for their known or potential mineral occurrences.

A principal example is the ca. 2175 to 2185 Ma mafic and alkaline province within the southwestern part of the Slave craton (Fig. 4), comprising the Big Spruce (Cavell and Baadsgaard, 1986), Squalus (Stubley, 1996), and Blatchford (Davidson, 1978; Bowring et al., 1984; Sinclair et al., 1994) alkaline complexes, and the essentially contemporaneous Duck Lake mafic sill (2181 ± 2 Ma, Bleeker and Kamo, 2003) and Dogrib dyke swarm near Yellowknife (ca. 2180- 2190 Ma, LeCheminant et al., 1997). All these events are proximal to either the southern or western margin of the Slave craton and are initial precursor events prior to final break-up⁸.

The Blatchford Lake intrusive complex, on the north shore of Great Slave Lake, consists of a number of nested intrusions ranging from gabbro to granite and syenite.Acentral syenite body, the Thor Lake syenite, hosts rare metal mineralization (Be, Y, rare earths, Nb, Ta, Zr, and Ga) in hydrothermally altered breccias and veins (Fig. 20) (Trueman et al., 1988; see also Sinclair et al., 1994; Taylor and Pollard, 1996).Whereas the main granitic phases, and by inference the related syenite phases, have been dated at ca. 2176 Ma (Bowring et al., 1984; Sinclair et al., 1994), the age of the mineralization remains controversial and has been postulated to be significantly younger (ca. 2094 Ma, Bowring et al., 1984; and discussion in Sinclair et al., 1994). It seems counterintuitive, however, to dissociate in time (by ca. 80 million years) the high-temperature hydrothermal mineralization from the alkaline intrusions with which they are closely associated. In general terms, the Thor Lake deposits can be divided into a series of nested zones with a keel-shaped cross-section (Fig. 20): 1) a “Wall Zone”, composed chiefly of quartz and feldspar, containing fluorite, sulphides, magnetite, carbonate, columbite, phenacite and several unidentified rare earth (REE)-bearing species; 2) a “Lower Intermediate Zone” that occurs predominately in the keels of the T-Zone deposits (Fig. 20), consisting of quartz, plagioclase, amphibole, biotite, muscovite, lepidolite, chlorite, carbonate, magnetite, with important accessory minerals consisting of fluorite, zircon, sulphides, phenacite, columbite and several REE minerals; 3) an “Upper Intermediate Zone” characterized by quartz, muscovite, albite, acmite and phenacite, with chlorite, fluorite, carbonates, magnetite, sulphides, and REE minerals as accessories; and 4) a vuggy “Quartz Zone” occupying the core of this complex (Trueman et al., 1988).

Another important example of rifting/extension related mafic magmatism is the large Booth River complex (Roscoe, 1985; Roscoe et al., 1987; Hulbert, 2005b) of the north-central Slave craton, outcropping from underneath the western extent of the Kilihigok Basin strata (see Fig. 4 for location). This complex has been dated at 2025 ± 2 Ma (Roscoe et al., 1987; Davis et al., 2004) and appears a magmatic focal point along the more extensive and slightly southward-fanning Lac de Gras dyke swarm, which is of similar age. The large, differentiated Booth River complex has attracted attention as a potential host for platinum group element (PGE) mineralization.

The much younger Muskox Intrusion (1267 Ma), on the northwestern margin of the craton and related to the large Mackenzie event (LeCheminant and Heaman, 1989), is another funnel-shaped layered intrusion that has attracted attention for reef-style PGE mineralization (Hulbert, 2005a).

The emplacement of diamondiferous kimberlite pipes represents the final metallogenic event for the Slave craton, with kimberlite ages spanning several discrete pulses throughout the Phanerozoic (Heaman et al., 2004).At Lac de Gras, in the central part of the craton, kimberlite pipes of Eocene age (ca. 55-50 Ma) form a dense cluster and now support two highly profitable diamond mines, Ekati and Diavik. Several other diamond mines are in various stages of development (e.g. Jericho, Kennady Lake). In just over a decade, diamonds have become the most profitable commodity within this ancient craton.

The Slave craton is a relatively small Archean craton with a geological knowledge base that is relatively mature. However, the following major questions remain:

1. What is the nature of the Hope Bay block in the northeast part of the craton? Does cryptic ancient basement reappear in this part of the craton? Is it perhaps a rifted fragment of the Central Slave Basement Complex?
2. What is the tectonic significance of pre-2687 Ma volcanic rocks in the eastern Slave? Do they form remnants of a ca. 2.7 Ga, exotic, intra-oceanic juvenile arc that collided with the extended Central Slave Basement Complex between 2697 and 2687 Ma? If so, where is the suture? Or do these volcanic rocks represent the oldest fill of narrow back-arc-like troughs formed by progressive rifting of the Central Slave Basement Complex in an overall arc setting?
3. What are the detailed trend and extent of particular magmatic (e.g. Defeat and Concession suites) and sedimentary belts (e.g. the post-Defeat turbidite basin along the western margin of the Slave)?
4. How did events inferred from the crustal geological record contribute to, or interfere with, formation of the subcontinental mantle lithosphere below the Slave craton?
5. How far does Slave mantle lithosphere extend below the Rae craton to the east?
6. What is the detailed break-up history for each of the Proterozoic margins of the Slave craton? In other words, how was the Slave crustal fragment liberated out of its ancestral supercraton Sclavia?
7. What is the detailed depositional record associated with rifting, thermal subsidence, and finally collision, along each of the margins of the Slave craton?
8. And, does Slave lithosphere extend northward all the way to the Innuitian front in the highArctic (see Fig. 1)?

To many of these first-order questions, we currently have only rudimentary answers or mere guesses. More sophisticated answers will require more complete and more refined data sets. In particular, a greatly expanded geochronological database, both in quantity and precision, in conjunction with targeted fieldwork across the craton and its marginal belts, should quickly advance the state of knowledge.

In terms of mineral potential, much of the craton and all major greenstone belts have seen at least a first wave of exploration for major commodities. These investigations quickly led to the discovery of a number of large VMS deposits (e.g. Izok Lake), which await road access for economic production. Gold potential remains high, particularly in more remote greenstone belts that may not have seen adequate drill testing. In this respect, the Point Lake greenstone belt, and its extension further north, appears attractive as it has all the major attributes of a world-class gold camp.

We thank colleagues at the Geological Survey of Canada, particularly Bill Davis, John Kerswill, Dave Snyder, Sally Pehrsson, and Ken Buchan, for sharing their insights into the geology and metallogeny of the Slave craton. Kim Nguyen assisted with preparation of some of the figures. Thoughtful reviews by John Ketchum, Tim Kusky, and Jan Peter helped improve the manuscript.

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Thorpe, R.I., Cumming, G.L., and Mortensen, J.K., 1992, A significant Pb isotope boundary in the Slave Province and its probable relation to ancient basement in the western Slave Province, in Richardson, D.G., and Irving, M., comps., Project Summaries, Canada-Northwest Territories Mineral Development Agreement 1987-91: Geological Survey of Canada, Open File 2484, p. 179-184.

Tomascak, P., and Cerny, P., 1992, Granitic pegmatites in the Aylmer Lake basin, in Richardson, D.G., and Irving, M., comps., Project Summaries, Canada-Northwest Territories Mineral Development Subsidiary Agreement 1987-1991: Geological Survey of Canada, Open File 2484, p. 51-54.

Trueman, D.L., Pedersen, J.C., de St. Jorre, L., and Smith, D.G.W., 1988, The Thor Lake rare-metal deposits, Northwest Territories, in Taylor, R.P., and Strong, D.F., eds., Recent Advances in the Geology of Granite-Related Mineral Deposits: Canadian Institute of Mining and Metallurgy, Special Volume 39, p. 280-290.

van Breemen, O., Davis, W.J., and King, J.E., 1992, Temporal distribution of granitoid rocks in the Archean Slave Province, northwest Canadian Shield: Canadian Journal of Earth Sciences, v. 29, p. 2186-2199.

Villeneuve, M.E., and Henderson, J.B., 1998, U-Pb geochronology of Wijinnedi Lake area, Slave Province, District of Mackenzie, Northwest Territories, in Radiogenic age and isotopic studies, Report 11: Geological Survey of Canada, Current Research 1998-F, p.  99-106.

Villeneuve,M.E., and Relf, C., 1998, Tectonic setting of 2.6 Ga carbonatites in the Slave Province, northwest Canada: Journal of Petrology, v. 39, p. 1975-1986.

Williams, H., Hoffman, P.F., Lewry, J.F., Monger, J.W.H., and Rivers, T., 1991, Anatomy of North America: Thematic geologic portrayals of the continent: Tectonophysics, v. 187, p. 117-134.

Wolfden Resources Inc., 2005, Company website, High Lake property map and surrounding area, Figure no. 2, http://www.wolfdenresources.com.

Yamishita, K., and Creaser, R.A., 1999, Geochemical and Nd isotopic constraints for the origin of late Archean turbidites from the Yellowknife area, Northwest Territories, Canada: Geochimica et Cosmochimica Acta, v. 63, p. 2579-2598.

[Click on an image thumbnail to view a larger image, notice]

 

	Figure 1: Tectonic map of the Precambrian basement of North America, showing the location of the Archean Slave craton relative to other first-order crustal elements. Greenland is shown in a pre-drift position (modified after Hoffman, 1988; Ross et al., 1991).
	Figure 2: Simplified geological map of the Slave craton. Localities mentioned in the text are highlighted. Cross-section line (ENE-WSW) refers to the craton-wide structural section shown in Figure 3.
	Figure 3: A) West-southwest to east-northeast, structural-stratigraphic section across the Slave craton (after Bleeker, 2002, see Fig. 2 for location of profile). Letters (B,C…M) refer to illustrations below section. No vertical exaggeration with present erosion level at ca. 0 km depth; some units are shown above this level in lighter tones for ease of interpretation. Deep structure in the western part of the profile interpreted from LITHOPROBE's SNORCLE seismic reflection profile (e.g. Cook et al., 1999). Volcanic units <2690 Ma and overlying turbidites overlap the boundary between isotopically different basement domains. Field photos: B) Typical 2.95 Ga foliated tonalites of the Central Slave Basement Complex with transposed 2734 Ma mafic dykes. C, D, E) Basal quartz pebble conglomerate, fuchsitic quartzite, and banded iron formation of the Central Slave Cover Group that overlies the basement complex. F) Variolitic pillow basalts of the Kam Group, Yellowknife. G) Syn-Kam Group K-feldspar porphyritic granodiorite pluton in basement below greenstone belts. H) Polymict conglomerate, including 10 to 30 cm granitoid cobbles, which occurs locally at the base of the younger, 2690 to 2660 Ma, volcanic cycle.
	Figure 3 continued: I) Intermediate to felsic volcaniclastic deposits typical of the younger volcanic cycle. J) Well preserved sub-biotite-grade turbidites in the core of the Yellowknife structural basin, showing graded bedding and load casts. K) Aerial photo of large-scale, upright F1 fold structures in turbidites of the Yellowknife structural basin. L) Synorogenic conglomerates, <2600 Ma, unconformably overlying unroofed granitoid rocks, Point Lake. M) Late-orogenic, ca. 2585 Ma K-feldspar megacrystic granites of the Morose Suite in the core of the domal Sleepy Dragon Complex; inset shows 1 to 2 cm large K-feldspar megacrysts.
	Figure 4a: Map of the Slave craton (as Fig. 2) annotated with locations of significant mineral deposits and occurrences. Deposits are grouped by genetic class and age, and listed on the facing page. See Appendix 1 for more complete information on these deposits and occurrences.
	Figure 4b: Map of the Slave craton (as Fig. 2) annotated with locations of significant mineral deposits and occurrences. Deposits are grouped by genetic class and age, and listed on the facing page. See Appendix 1 for more complete information on these deposits and occurrences.
	Figure 5a: ) Simplified map showing minimum extent of the Hadean to Mesoarchean basement of the Central Slave Basement Complex (CSBC). Small shaded spheres (yellow or orange) highlight locations where the diagnostic basement to cover stratigraphy has been observed. Based on basement characteristics, five tectono-stratigraphic domains can be recognized across the craton: 1) a westernmost domain, with basement of unknown age; no diagnostic evidence for ancient basement; 2) the main extent of the Central Slave Basement Complex; 3) the buried eastern edge of the ancient basement complex; 4) the eastern Slave "Hackett River domain", which lacks stratigraphic and isotopic evidence for ancient basement; and 5) the "Hope Bay block", which has subtle characteristics that are reminiscent of the Yellowknife area and may well be underlain by ancient basement; currently, data from this area are insufficient to resolve this question.
	Figure 5b: Age distribution of the Central Slave Basement Complex, as sampled by 296 concordant to near-concordant detrital zircon grains from five quartzite samples across the basement complex (orange spheres in Fig.5A: Yellowknife, Cameron River belt, northern Beaulieu belt, Point Lake, and quartzite overlying Acasta basement, see Sircombe et al., 2001). Although biased by detrital sampling and variable preservation of basement zircons, this age spectrum allows a rapid assessment of major age components of the basement complex. Note significant age peaks starting at ca. 3400 Ma. The depositional age of these quartzites is ca. 2850 to 2800 Ma. Roman numerals refer to major crust-forming events in the basement recognized from U-Pb protolith ages (see Bleeker and Davis, 1999b).
	Figure 6: Field photographs illustrating some key elements of the geology of the Slave craton. A) Cleaned exposures of the Acasta gneisses at their discovery site. Ancient tonalites (4.03 Ga) occur on left side of the picture, and are intruded by highly deformed younger granite sheets and mafic dykes. B) Basal quartzites of the Central Slave Cover Group overlying basement of the Central Slave Basement Complex. Low terrain to the right is low-weathering basement gneiss; dark ridge in background consists of ca. 2.7 Ga basalts overlying the quartzites. C) Syn-Kam Group quartz-porphyritic tonalite (ca. 2713 Ma), intruding into the northern part of the Yellowknife greenstone belt (geologist Val Jackson for scale). The large sill-like body is cut by somewhat younger mafic dykes that likely fed the upper part of the greenstone belt stratigraphy. Inset shows close-up of altered quartz-porphyritic tonalite. D) Quartz-porphyritic rhyolite breccia with carbonate matrix, typical for the uppermost part of 2690 to 2660 Ma felsic volcanic edifices. Highly angular (proximal) fragments range is size from 1 to 10 cm. E) Massive sulphide mineralization of the Sunrise deposit, associated with ca. 2670 Ma felsic volcanic rocks just below the interface with the Burwash Formation turbidites. F) Thick-bedded sandy turbidites typical of the Burwash Formation in its type area east of Yellowknife. The oblique areal photo (approximately 500 m across, looking towards northeast) shows a typical mushroom interference pattern of an F1 syncline refolded by north-northwest trending F2 folds. G) Silicate-facies iron formation interlayered with turbiditic greywackes, George Lake, northeastern Slave. This banded iron formation hosts significant epigenetic gold mineralization (pickets ca. 50 cm high). H) Passive margin strata of the Coronation Supergroup (Epworth Group) overlying the western margin of the rifted Slave craton, structurally at the base of Wopmay Orogen externides (east facing cliff is approximately 100 m high). I) Dense Proterozoic mafic dyke swarms cutting extended Slave crust and its cover (individual dykes of approximately 20-40 m across).
	Figure 7: North America's first diamond producer, the Ekati Mine of the central Slave craton. Main picture shows the flat barren lands of the central Slave craton, with several pipe-like kimberlite bodies being excavated in circular open pits. Inset (upper left) shows a close-up of one of the partially excavated pipes. Several mm-size gem quality diamond octahedra are shown on upper right (photos courtesy of D. Snyder and G. Lockhart).
	Figure 8: A) Generalized stratigraphic column of the Central Slave Cover Group, the autochthonous cover of the Central Slave Basement Complex (Bleeker et al., 1999a,b). Relevant age data are from Isachsen and Bowring (1997), Bleeker et al. (1999a,b), and Ketchum and Bleeker (2000). Photos B to E illustrate characteristic lithologies and are keyed to the column. F)Electron back-scatter image of rounded detrital chromite grain in fuchsitic quartzite of photo D. G) As in photo E, but showing more advanced metamorphic transformation to fuchsite.
	Figure 9: Stratigraphy and geochronology of the Kam Group, Yellowknife greenstone belt. All major units except for the basal Chan Formation have now been dated, showing a monotonic, younging upwards age progression through the volcanic pile. Sources of age data: Isachsen et al. (1991), Isachsen (1992), Isachsen and Bowring (1997), Bleeker et al. (1999b), Sircombe et al. (2001), Ketchum and Bleeker, unpublished data, Davis et al. (2004). A crosscutting gabbro dyke in the Chan Formation, dated recently at 2738 Ma (J. Ketchum, pers. comm., 2004), shows that much of the Chan Formation must be >2738 Ma. The youngest event recognized to date is the large intrusive gabbro sills at Kam Point, with a preliminary baddeleyite age of 2697 Ma. A large quartz-porphyritic tonalite sill (see Fig. 6C) has been dated at 2713 Ma and acted as a heat engine for extensive seafloor hydrothermal activity and associated alteration.
	Figure 10: Generalized stratigraphic columns for different parts of the Slave craton. A) Typical stratigraphy for much of the central part of the craton, with Yellowknife Supergroup rocks overlying the Central Slave Basement Complex. B) Enlargement of the basement and its autochthonous cover. C) Typical stratigraphy for much of the eastern Slave (e.g. Hacket River, Back River), lacking the association with ancient basement. D) Typical stratigraphy for the north-central Slave, e.g. the High Lake belt. An interesting feature is the occurrence of a ca. 2612 to 2616 Ma volcano-sedimentary package distinct from nearby Burwash-type metagreywackes. This "High Lake assemblage" is tentatively interpreted as the volcano-sedimentary carapace of Concession Suite magmatism. E) Generalized column for the westernmost Slave craton. Much of the stratigraphy is similar to that overlying the CSBC (as in Fig. 10A), but is capped by a younger turbidite sequence with abundant silicate-facies iron formations. F) Detrital zircon data for typical Burwash Formation turbidites (lower panel), and the younger greywacke sequence of Damoti Lake (upper panel). Note that the latter sample contains a robust younger age peak, indicating a significant contribution from erosional unroofing of Defeat Suite-age plutons. Data obtained by SHRIMP (Sircombe and Bleeker, unpublished). Shaded intervals highlight typical age ranges for Kam Group (K), Banting Group (B), and Defeat Suite (D) magmatism. Arrows indicate depositional ages. Zircon symbol highlights a single, concordant, and precise zircon age determination obtained by TIMS (Pehrsson and Villeneuve, 1999).
	Figure 11: . Cross-section through one of the High Lake massive sulphide deposits (after Wolfden Resources Inc., 2005).
	Figure 12: . Ternary plots of metal ratios for VMS deposits in the Slave craton. The metal ratios define two fields: one for Cu-rich, Pb-poor deposits associated with the older ca. 2.7 Ga basalt-dominated bimodal volcanism, and another characterized by higher Pb (Ag) values for deposits hosted by the younger ca. 2.69 to 2.66 Ga arc-like sequences.
	Figure 13: Tectonic model for the general setting of the Slave craton between ca. 2690 and 2660 Ma. The entire Slave appears to have been situated in a supra-subduction zone setting, with abundant and widespread calc-alkaline volcanism and plutonism. Arc-like assemblages (e.g. Hackett River, HR) were built across rifted basement and evolved into a large back-arc basin filled with turbidites: the Burwash Basin. Stratigraphically lowest, pre-2687 Ma components of the eastern arc-like domain could be exotic; alternatively, these rocks could represent juvenile volcanics in narrow back-arc rifts. The Hope Bay block may represent a rifted fragment of the Central Slave Basement Complex (CSBC). Voluminous tonalite intrusions rejuvenated much of the lower and mid-crust. Collision of an unknown terrane (X), between 2650 and 2630 Ma, led to closure and F1 folding of the Burwash Basin.
	Figure 14: Geological map of the Hackett River area (modified from Frith and Fryer, 1985), showing the locations of major VMS deposits as well as smaller occurrences in relation to synvolcanic plutons. Also shown is the carbonate unit that bears a close spatial relationship to the VMS deposits.
	Figure 15: Thematic maps illustrating key stratigraphic and structural aspects of the Slave craton through time. A) Minimum extent of the ca. 2680-2660 Ma Burwash Basin, based on continuity and geochronological similarity of turbiditic greywackes across large parts of the craton. The dash-dot line separates areas with intercalated iron formations (northwest) from those lacking iron formations (southeast). Yellow spheres highlight localities with precisely dated 2661 Ma volcanism closely associated with turbidite sedimentation. The linear trend may reflect a 2661 Ma magmatic line in a general arc-like setting. B) General trends of the F1 fold belt in the Burwash Formation and correlatives, trending northeast-southwest across the craton. C) Defeat Suite plutons dated between 2635 and 2625 Ma. This apparent trend of arc-like plutons parallels the F1 fold belt. D) Areas in the Slave craton with younger volcano-sedimentary packages that post-date deposition and folding of the Burwash Formation: the ca. 2620 Ma Damoti Lake (D) assemblage and the ca. 2612 to 2616 Ma High Lake (HL) assemblage. The Damoti Lake assemblage extends at least from Emile River (E) to Russell Lake (R), and possibly to Wheeler Lake (W). Its full extent is not known.
	Figure 16: Geology of the George Lake area, northeastern Slave craton, where significant epigenetic gold mineralization is hosted by tightly folded and metamorphosed iron formation in Burwash-age metaturbidites (after Jefferson et al., 1992).
	Figure 17: Time charts of stratigraphic and structural events in the Slave craton. A) Detailed chart for key events in the Yellowknife Domain (updated from Davis and Bleeker, 1999).
	Figure 17 continued: B) Extended time chart for entire Slave craton (excluding pre-2900 Ma history of the Central Slave Basement Complex). Note breaks in horizontal scale (time) at two places. Typical detrital age spectra (1, 2a, b, 3, and 4) are shown for some of the main stratigraphic units: 1 = quartzites of the Central Slave Cover Group; 2 = turbiditic greywackes of Burwash Formation (s.s., 2a) and correlatives in the eastern Slave (Back River, 2b); the Damoti Lake turbidites (3); and synorogenic conglomerates (4). Known mafic dyke swarms are shown by narrow black bars (or grey, if poorly dated).
	Figure 17B continued:
	Figure 18: Simplified geology of the Hope Bay greenstone belt of the northeastern-most Slave craton, showing localities of gold mineralization (modified after Anonby et al., 1998; Hebel, 1999). The presence of a thick tholeiitic greenstone package with intercalated rhyolite, dated at 2716 3 Ma, is reminiscent of the Yellowknife belt. The dated rhyolite contains inherited zircons that are 2734 Ma or older (Hebel, 1999), whereas minor turbidites along the centre of the belt contain a substantial fraction of Mesoarchean zircon grains. Collectively, these characteristics suggest an affinity with the Central Slave Basement Complex and its volcano-sedimentary cover (a rifted fragment thereof?), rather than a continuation of the juvenile greenstone belts of the eastern Slave craton.
	Figure 19: Kimberlite pipes and kimberlite fields of the Slave craton (Heaman et al., 2004; Lockhart et al., 2004). Nearly all kimberlites occur within the realm of the Central Slave Basement Complex or its buried eastern edge extending to the east of Lac de Gras (see Fig. 3). The Victoria Island Field, of Permian age (~256-286 Ma; off the map, towards the north), may also be associated with such basement as the Hope Bay block has many characteristics similar to that of the Yellowknife area. BRFZ = Beaulieu River Fault Zone, YRFZ = Yellowknife River Fault Zone.
	Figure 20: . Geology of the Thor Lake rare element deposit (after Trueman et al., 1988). It occurs in the central syenite pluton of the Blatchford Lake Complex, an alkaline rift-related intrusive complex on the north shore of Great Slave Lake and a major component of the ca. 2180 Ma "SW Slave magmatic province" (see text).

See Appendix 1