the slave craton geological and metallogenic evolution

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Bleeker, W., and Hall, B., 2007, The Slave Craton: Geology and metallogenic evolution, in Goodfellow, W.D., ed., Mineral Deposits of Canada: A Synthesisof Major Deposit-Types, District Metallogeny, the Evolution of Geological Provinces, and Exploration Methods: Geological Association of Canada, MineralDeposits Division, Special Publication No. 5, p. 849-879.

THE SLAVE CRATON: GEOLOGICAL AND METALLOGENIC EVOLUTION

WOUTER BLEEKER 1 AND BRIAN HALL2

1. Geological Survey of Canada, 601 Booth Street, Ottawa, Ontario K1A 0E82. Department of Earth Sciences, Laurentian University, 935 Ramsey Lake Road, Sudbury, Ontario P3E 2C6

Corresponding author’s email: [email protected]

Abstract

The Slave craton of the northwestern Canadian Shield is one of the oldest and most distinct building blocks of NorthAmerican cratonic lithosphere. It hosts Earth’s oldest intact rocks, the Acasta gneisses. These ancient gneisses areembedded 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 isheavily 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 dom-inantly tholeiitic greenstone sequence (ca. 2700 Ma), younger arc-like sequences (ca. 2690-2660 Ma), extensive tur-

bidite 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 volcanic-dominated rifting of the basement. Arc-like sequences formed in part on top of the attenuated basement and in pro-gressively widening, juvenile, back-arc-like basins and contain some of Canada’s largest undeveloped volcanogenicmassive sulphide deposits. After 2680 Ma, much of the Slave craton became overlain by the Burwash Basin, one of thelargest 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 2580Ma, with a peak in crustal anatexis between 2595 and 2585 Ma (the “granite bloom”). Numerous orogenic gold depositsformed throughout the Slave craton, either as shear- or vein-hosted deposits in deformed greenstones or within thechemical traps provided by banded iron formations in the turbidites. Proterozoic rift-related magmatic suites and arcsaround the margins of the craton host a variety of mineral deposits. Finally, the craton was intruded by several hundredPhanerozoic kimberlite pipes, some of which support Canada’s first diamond mines.

Résumé

Le craton des Esclaves, qui occupe la partie nord-ouest du Bouclier canadien, est l’un des plus vieux et des plusdistinctifs des éléments constitutifs de la lithosphère du craton nord-américain. Il renferme les plus anciennes rochesintactes de la Terre, les gneiss d’Acasta. Ces gneiss anciens sont inclus dans un grand complexe de socle datant duMésoarchéen à l’Hadéen, qui s’étend aux parties centrale et occidentale du craton. Peu minéralisé en soi, le complexede socle est recouvert de séquences supracrustales du Néoarchéen et est intensément traversé et cannibalisé par dessuites plutoniques variant de plutons synvolcaniques âgés de 2720 à 2660 Ma à des batholites de granite tar-diorogéniques datant de 2595 à 2585 Ma. Les séquences supracrustales, collectivement attribuées au Supergroupe deYellowknife, consistent en une séquence de couverture initiale composée de quartzite et de formation de fer rubanée(env. 2800 Ma), en une épaisse séquence de roches vertes à dominante tholéiitique (env. 2700 Ma), en séquences plus

récentes apparentées à celles des arcs (env. 2690 à 2660 Ma), en couvertures étendues de turbidites (env. 2680 à 2620Ma) et, enfin, en conglomérats synorogéniques déposés à 2600 Ma environ, ou peu après. La séquence de couvertureinitiale et les tholéiites sus-jacentes témoignent d'une exposition subaérienne puis d’un rifting du socle, qui s’est accom-

pagné de la mise en place de matériaux à prédominance volcanique. Les séquences apparentées à celles des arcs se sontformées en partie sur un socle aminci et en partie dans des bassins juvéniles apparentés à des bassins d'arrière-arc s’élar-gissant progressivement. Ces séquences renferment certains des plus vastes gîtes de sulfures massifs volcanogéniquesinexploités du Canada. Après 2680 Ma, une bonne partie du craton des Esclaves a été enfouie sous les dépôts du bassinde Burwash, un des plus grands bassins de turbidites archéennes au monde, aux dépôts parmi les mieux conservés, quiest comparable par sa taille et son cadre à l’actuelle mer du Japon. Pendant l'orogenèse, les séquences supracrustales sesont télescopées et épaissies, et ont été plissées à plusieurs reprises entre environ 2650 Ma et 2580 Ma, avec une péri-ode d’intensité maximale d'anatexie crustale de 2595 Ma à 2585 Ma (l’« efflorescence granitique »). De nombreux gîtesd'or orogéniques se sont formés un peu partout dans le craton des Esclaves, soit sous forme de gîtes encaissés dans deszones de cisaillement ou de gîtes filoniens dans des roches vertes déformées, soit sous forme de minéralisations con-tenues dans les pièges chimiques constitués par les formations de fer rubanées au sein des turbidites. Les suites et lesarcs magmatiques liés au rifting, qui sont apparus au Protérozoïque le long des marges du craton renferment toute une

gamme de gîtes minéraux. Enfin, le craton a été percé au Phanérozoïque par plusieurs centaines de cheminées de kimberlite, dont certaines ont permis la mise en exploitation des premières mines de diamants au Canada.

Introduction

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 Archeancratons 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), prob-ably within the broader context of the formation of Earth’sfirst modern supercontinent, Nuna. Much of the Slave cratonis 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 byPaleoproterozoic rifted margins, the Slave craton originatedfrom the break-up of a much larger late Archean landmass,

possibly the speculative late Archean supercontinentKenorland (Williams et al., 1991) or, perhaps more likely, asmaller landmass referred to as the supercraton Sclavia(Bleeker, 2003). The late Archean and earliest Proterozoicdevelopment and evolution of Slave crust should thus beviewed within the context of the growth and subsequent

break-up of this larger Sclavia supercraton, even though theshape and size of this supercraton remain currentlyunknown. The salient point is that ancient cratons, like the

Slave, preserve only parts of the much larger tectonic sys-tems in which they were generated.

In agreement with this conceptual view, latest Archeanevents1 are remarkably homogeneous across the Slave cra-ton and may be used, together with pre-2.0 Ga Proterozoicmafic dyke swarms2, to help identify neighbouring frag-

ments of Sclavia from among the35 extant cratons. One such Slavecraton-wide event was a volumi-nous “granite bloom” between ca.2595 and 2585 Ma (Davis andBleeker, 1999). This singular eventin the craton’s history transferred,irreversibly, a significant fraction

of heat-producing elements andlower crustal fluids to the upper crust, thus allowing slow coolingand stiffening of the lower crustand setting the stage for cratoniza-tion and long-term preservation(Bleeker, 2002).

Predating this cratonizationevent, the Slave crust preserves acomplex and spatially heteroge-neous record of crustal growthspanning more than 1.5 billionyears (e.g. Kusky, 1989; Isachsenand Bowring, 1994; Bleeker andDavis, 1999a; 1999b and referencestherein; Sircombe et al., 2001;Ketchum et al., 2004; see Iizuka etal. (2006) for a discussion on therecent discovery of a 4.2 Ga zirconxenocryst in Acasta gneiss).

The present paper briefly sum-marizes this crustal growth historyand the overall geological evolu-tion of the Slave craton, from theformation 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 aframework for discussion of important metallogenic eventswithin the craton. Pertinent information concerning the loca-tion, mineralogy, structure, size, grade, age, and stratigraphicand tectonic settings for all significant mineral depositswithin or adjacent to the craton is summarized in theAppendix, along with short summaries for some additionalmetallogenic events. Figure 4 shows the locations of thesemineral deposits and occurrences on a geological map of thecraton.

Ancient Basement Complex

An ancient and largely crystalline basement complexunderlies 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., 1999a,b; Ketchum and Bleeker,2001; Ketchum et al., 2004). It is referred to as the CentralSlave Basement Complex (Bleeker et al., 2000).

Along the Acasta River, this basement complex consistsof polymetamorphic gneisses of tonalitic to gabbroic com-

position (Fig. 6A) that yield protolith ages up to 4.03 Ga3

W. Bleeker and B. Hall

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FIGURE 1. Tectonic map of the Precambrian basement of North America, showing the location of the ArcheanSlave craton relative to other first-order crustal elements. Greenland is shown in a pre-drift position (modi-fied after Hoffman, 1988; Ross et al., 1991).

1. In the present context, events younger than ca. 2680 Ma.2. See Bleeker (2004) and Bleeker and Ernst (2006) for general methodology.3. All ages quoted in this paper are U-Pb zircon ages, unless stated otherwise. These were determined either by isotope-dilution thermal ionization mass spec-

trometry (ID-TIMS) or sensitive high-resolution ion microprobe (SHRIMP) methods. Generalized ages, or those with poor precision, are given at the bil-lion year (Ga) level, whereas more precise ages and their errors are given at the million year (Ma) level.



The Slave Craton: Geological and Metallogenic Evolution

8

(Bowring et al., 1989; Stern and Bleeker, 1998; Bowring andWilliams, 1999). The record of inheritance extends back to4.2 Ga (Iizuka et al., 2006). The Acasta gneisses were essen-tially 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 spo-radic ages >3.5 Ga (Acasta to Point Lake), the Central Slave

Basement Complex is mostly younger with important agemodes4, from detrital and protolith U-Pb zircon ages, around3400 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 sug-gest that at least part of the subcontinental lithospheric man-tle below the central part of the craton may be of similar antiquity (e.g. Aulbach et al., 2004). A crude age zonationcan be recognized in the basement complex (Ketchum and

Bleeker, 2001), although no easily interpretable tectonic pat-tern has yet emerged. Pre-2.9 Ga supracrustal rocks have

been found at the base of some greenstone belts (Bleeker andDavis, 1999b and references therein; Ketchum et al., 2004),

but form only a very small component of the craton’s geo-logical inventory. An important example of such an occur-rence 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 knownto occur within pre-2.8 Ga basement but have yet to be sys-tematically dated.

Mineralization

There are few, if any, known mineral occurrences of notewithin the Central Slave Basement Complex, a statistic thatis generally mirrored by other Mesoarchean and older gneis-

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2.73-2.70 Ga tholeiitic greenstone beltswithin the realm of the basement complex

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2.68-2.63 Ga turbidite sequences

2.68-2.58 Ga granitoid rocks, undivided;locally includes pre- 2.8 Ga granitoids and gneisses

Nd isotopic boundary (granitoids)

of Davis and Hegner (1992)

Pb isotopic boundary (syngenetic leads)of Thorpe et al. (1992)

Fold trends of first post-turbiditeregional folds (F1; refolded by F2 folds)

Diamond deposit

Transparent overlay outlining inferred extent of the Central Slave Basement Complex (CSBC)

Basement exposures: areas where pre-2.9 Garocks have been dated by U-Pb methods

Occurrences of cover sequence; arrowsindicate stratigraphic facing directions

Archaean strike-slip f ault zones: Yellowknife River Fault Zone (YRFZ) and Beaulieu River Fault Zone (BRFZ)

Proterozoic platformal cover

orogenic beltsand

Phanerozoic cover

SLAVE CRATON

Acasta GneissComplex

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SLAVE CRATON

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FIGURE 2. Simplified geological map of the Slave craton. Localities mentioned in the text are highlighted. Cross-section line (ENE-WSW) refers to the cra-ton-wide structural section shown in Figure 3.

4. More or less well defined peaks on age histograms, e.g., see Sircombe et al. (2001).




852

FIGURE 3. (A) West-southwest to east-northeast, structural-stratigraphic section across the Slave craton (after Bleeker, 2002, see Fig. 2 for location of pro-file). 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 abovethis level in lighter tones for ease of interpretation. Deep structure in the western part of the profile interpreted from LITHOPROBE’s SNORCLE seismicreflection profile (e.g. Cook et al., 1999). Volcanic units <2690 Ma and overlying turbidites overlap the boundary between isotopically different basementdomains. Field photos: (B) Typical 2.95 Ga foliated tonalites of the Central Slave Basement Complex with transposed 2734 Ma mafic dykes. (C, D, E) Basalquartz pebble conglomerate, fuchsitic quartzite, and banded iron formation of the Central Slave Cover Group that overlies the basement complex (~20 cmtall polished slabs). (F) Variolitic pillow basalts of the Kam Group, Yellowknife (pen for scale). (G) Syn-Kam Group K-feldspar porphyritic granodiorite plu-ton in basement below greenstone belts (pen for scale). (H) Polymict conglomerate, including 10 to 30 cm granitoid cobbles, which occurs locally at the baseof the younger, 2690 to 2660 Ma, volcanic cycle.

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Reflection MohoReflection Moho

Yellowknife River Fault Zone

CameronRiver Belt

BurwashFormation

Yellowknife BeltRussel LakeBelt

BeaulieuRiver Belt

Loop LakeBelt

CourageousLake Belt

Beaulieu River Fault Zone

AntonComplex

CoronationSupergroup

Sleepy DragonComplex Jolly Lake

Complex

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SLAVE PROVINCE LAVE PROVINCEBEAR PROVINCE EAR PROVINCE

Great Bear Magmatic Zone

Wopmay Front

Bear Province (Wopmay Orogen) Slave ProvincePalaeoproterozoic (2.0-1.8 Ga) Hadean to Archaean (4.05-2.58 Ga)

Realm of the Central Slave Basement Complex: central and western parts of craton

Central Slave Basement Complex

Central SlaveCentral Slaveottah TerraneHottah Terrane

Indin-West Bay Fault System

Y1

Y1

Y2

Poorly known basement of Hottah Terrane

Central Slave Basement Complex (4.05-2.86 Ga)Poorly defined basement of the westernmost Slave craton

Proterozoic, brittle strike-slip fault system,sinistral, overprinting western margin of thecraton (Indin-West Bay Fault System)

Proterozoic frontal thrusts

Proterozoic gabbrosills, causing bright reflections (Y1) east of Yellowknife

Great Bear Magmatic Zone volcanic rocks (1.86-1.84 Ga)

Central Slave Cover Group: conglomerate, quartzite, banded iron formation

Overall younging direction of cover sequence

Great Bear Magmatic Zone plutonic rocks (1.86-1.84 Ga)

Kam Group and correlatives: mafic volcanic rocks, minor rhyolites(2.73-2.70 Ga)

Coronation Supergroup (2.0-1.86 Ga):

Snare and Epworth groups

Banting Group and correlatives (2.69-2.66 Ga)

Syn-Banting Group subvolcanic plutons (2.69-2.66 Ga)

Syn-Kam Group subvolcanic plutons (2.73-2.70 Ga)

Hidden and Duckfish lake plutons, Yellowknife area (ca. 2.61 Ga)Defeat Suite plutons: diorite, tonalite, granodiorite (ca. 2.63-2.62 Ma, post-date D1)

Burwash Formation and correlatives: turbiditic greywackes (2.68-2.64 Ga)

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8

sic complexes in cratons around the world. To a first degree,this poor endowment is thought to correlate with the gener-ally mid- to lower crustal erosion depths within such crys-talline basement complexes and, consequently, the virtuallack of supracrustal rocks.

Indirectly, however, the presence of the ancient basementcomplex may have influenced or exerted important controlson several classes of younger mineral deposits:

• Seafloor volcanogenic massive sulphide (VMS) mineral-ization within the Neoarchean bimodal rift volcanic

WB '99

100 km

Lac de Graskimberlites

GeorgeLake

>3.0 Ga lower

crustal xenoliths

Back River Volcanic Complex

Bathurst Fault System

Rifted margin or 2.69 Ga suture?

Lac de GrasStructural Basin Back River

CulminationMalley Rapids Anticlinorium

Hackett River Gneiss Dome

GoulburnSupergroup

ENE ENE

Hackett River Syncline

CHURCHILL PROVINCE (RAE CRATON)

CHURCHILL PROVINCE(RAE CRATON)

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SLAVE PROVINCE LAVE PROVINCE

Kilihigok Basin

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Churchill ProvincePalaeoproterozoic (2.0-1.8 Ga)

Eastern Slave Domain:Palaeoproterozoic with reworked Archaean

Archaean to Palaeoproterozoic

Eastern Slave BasementEastern Slave Basement Rae HinterlandRae Hinterland

Eastern Slave Domain

Poorly defined basement of the eastern Slave craton,isotopically 2.85 Ga≤

Form surface tracesoutlining fold structures Vertical strike-slip faults

Bear Creek GroupChurchill Province (Rae craton)basement, reworked

Peacock Hills and Kuuvik Formations

High-grade cover rocks

Mafic to intermediate volcanic rocks (ca. 2.71-2.70 Ga)

Ellice Formation

Reworked Slave Province metavolcanic rocksand metagreywackes, kyanite bearing close to front

Churchill Province granitoids

Churchill Province granitoids

Late-tectonic conglomerates in asymmetric, fault-bounded panels(2.60-2.58 Ga), e.g. along Yellowknife River Fault Zone and BeaulieuRiver Fault Zone

Crustally derived, anatectic granitoid sheets and batholiths, and presumed source regions (2 .60-2.58 Ga, syn-kinematic withD2 and D3)

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MM J,K J,KI

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FIGURE 3 CONTINUED. (I) Intermediate to felsic volcaniclastic deposits typical of the younger volcanic cycle (hammer for scale). (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 F1fold structures in turbidites of the Yellowknife structural basin. (L) Synorogenic conglomerates, <2600 Ma, unconformably overlying unroofed granitoidrocks, Point Lake (pebbles ~5-10 cm in size). (M) Late-orogenic, ca. 2585 Ma K-feldspar megacrystic granites of the Morose Suite in the core of the domalSleepy Dragon Complex; inset shows 1 to 2 cm large K-feldspar megacrysts.




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Central SlaveBasement Complex


Point LakePoint Lake

ContwoytoLake

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Hope BayBelt


YellowknifeDomain Blatchford Lake

Intrusive Complex


CourageousLake Belt

Back River Volc. Complex

High LakeBelt

ArcadiaBay

Point LakeBelt

Indin LakeBelt

HackettRiver Belt

Lac de Graskimberlite field

SLAVE CRATON Significant mineral

deposits and occurences

Relative Scale of deposits:

0.1 - 1.0 million tonnes

1.0 - 10 million tonnes

10 - 100 million tonnes


~2175-2185 MaSW Slave magmatic

province




8

Eocambrian to Eocene diamond-bearing and allied rocks

Paleoproterozoic alkaline intrusion-related rare element deposits

1 SnapLake2 Ekati(Koala)3 Ekati(Jay)4 Ekati(Fox)5 GahchoKue(Kennady Lake)6 Ekati(Panda)7 Ekati(Sable)8 Diavik (A154-South)9 Diavik(A-418)

10 Ekati(Pigeon)11 Diavik(A-21)12 Ekati(Misery)13 Jericho14 Ekati(Beartooth)15 Ekati(Koala North)16 Ekati(Lynx)17 Diavik (A154-North)

Proterozoic mafic intrusions with Ni,Cu, Cr andreef-stylePGE mineralization

Paleoproterozoic hydrothermalCu-Au (IOCG) deposits

Paleoproterozoic mafic intrusion-relatedNi, Cu, PGEs, V

18 Muskox Intrusion, Mackenzieevent

19 NICO20 Sue-Diane

(21)Booth River Complex,Lac de Grasevent

22 Thor (Lake Zone)23 Thor(TZone)

LateArchean rare element-enriched pegmatites

Archean Au deposits - Vein or shear zone-hosted

Archean Au deposits - Intrusion related

Archean Au deposits - BIF or turbidites-hosted

Archean VMS in arc-like enviornments

Archean VMS in bimodal rift environments

Archean BIFs at top of Central Slave Cover Group

24 Murphy25 F1(Main Dyke)26 MooseNo.227 F1(SouthwestDyke)28 Shorty29 A nn30 VO(Cota)31 Paint32 Jake33 K1(Dyke)34 EchoTor 35 Bil l(36) Peg Tantalum

37 Tundra (Fat)38 Colomac39 Giant Yellowknife40 Madrid41 Co n42 Boston43 Discovery44 K im45 Cass46 Doris47 Tundra (Carbonate)48 Nicolas Lake49 Arcadia50 Supercrest51 MAHE52 Mosher Lake

53 Ptarmigan &Tom54 Saucer Lake55 Salmita56 Thompson-Lundmark57 Bullmoose58 Tundra Mine59 Crestaurum

60 Ulu (Flood)

61 Lupin62 Goose Lake (South)63 Turner Lake64 George Lake (Locale 1)65 RE N66 George Lake (Locale 2)67 Pistol Lake68 Damoti Lake (Horseshoe Zone)69 George Lake (Lone Cow Pond)

70 George Lake (GH)71 George Lake (Slave)72 Butterfly Lake

73 Izok Lake74 Gondar 75 Hackett (A Zone)76 Sunrise77 Hackett (East Cleaver)78 Hackett (Boot Lake)79 Hackett (Cleaver Low Grade)80 Yava81 Bear 82 B B83 Kennedy Lake84 Kennedy Lake (Copper Zone)85 Hackett (Jo)86 Musk87 Turback Lake (XL)88 Susa Lake

89 High Lake (West Zone)90 High Lake (A / B Zone)91 High Lake (D Zone)92 De b93 Hood River (#10)94 Hood River (#41)(95) Homer Lake showings, Yellowknife greenstone belt

(96) Dwyer and Bell Lake(97) Patterson Lake(98) Amacher Lake(99) Brown Lake(100) Winter Lake belt(101) Acasta

211 Approximate extent of Booth River Complexunderneath cover of Kilihigok Basin

Minimum extent of ca. 2175-2185 Ma magmaticprovince in the southwestern Slave craton, comprising:

-Dogrib mafic dyke swarm (ca. 2185-2189 Ma)-Duck Lake sill (2180+/-2 Ma)-Blatchford Lake Intrusive Complex (2176+/-2 Ma)-Big Spruce alkaline complex NW of Yellowknife-Squalus Lake alkaline complex (2180+/-2 Ma)

FIGURE 4. Map of the Slave craton (as Fig. 2) annotated with locations of significant mineral deposits and occurrences. Deposits are grouped bygenetic class and age, and listed on the facing page. See Appendix 1 for more complete information on these deposits and occurrences.




856

2800280000030002003200

Formationo f t h eMoo nFormation

of the Moon



Crust-forming events fromB leeker & Dav is (1999)Crust-forming events fromBleeker & Davis (1999)

IsuaIsua

Warrawoona andOnverwacht groups

Warrawoona and Onverwacht groups

3400340060036008003800000400020042005004500 26002600

Age (in Ma)

R e

l a t i v

e p r o b a

b i l i t y

II II

IV IV

VV

III III

VIVI

VIIVII

VIIIVIII

Ia,b Ia,b Ic Ic II II IV IV VV II III VIVI VIIVII VIIIVIII IX IX

IX IX

Jack Hi llsdetrital zircons

Jack Hillsdetrital zircons

B

A

Detrital zircons, Central Slave Cover Group-combined data of 5 quartzite samples-296 zircon grains

0 100

km

1 0 5

O

62O

1 0 5

O

6 9O

1 1 7 O

68O

1 1 7

O

61O

YRFZ

Y R F Z

BR FZB R F Z

PbNd

WB 2000

C or onat i on Gu l f

Great Slave Lake


Wo pmay

Orogen

W o p m

a y O

r o g e n

W op

n

maF

Z oe

W o p

n

m a y F a u

l t

Zo

e

Bath

urstFault

B a t h u r s

t F a u l t

Taltson

T a l t s

o n

Magmati

c Zone

M a g m

a t i c

Z o n e

Great

e

Slaev Lak

G r e a t

e

S l a

e vL a

k

Shear Zo

ne S h

e a rZ o n e

The l

on Front

T h

e l o

n F

r o n t

P

aleozoicPla

tform

P a l e o z o

i c

P l a

t f o r m

PointLakePoint Lake

LacdeGras

Lac de Gras

ContwoytoLake

ContwoytoLake

Central SlaveBasementComplex


Pb

Nd

Archean Slave craton, undifferentiated

Nd isotopic boundary (granitoids)of Davis and Hegner (1992)

Geological domains based onon basement characteristics(discussed in text)

Greenstone belts

Pb isotopic boundary (syngenetic leads)of Thorpe et al. (1992)

Transparent overlay outlining inferred extent of the Central Slave Basement Complex (CSBC)

Basement exposures: areas where pre-2.9 Ga rocks have been dated

by U-Pb methods

Occurrences of cover sequence;arrows indicate stratigraphic facing directions

Proterozoic platformal cover

orogenic beltsand

Phanerozoic cover

SLAVE CRATON



AntonComplex

Hope BayBlock

Hackett River Belt

YellowknifeDomain

1

1

2

2

2

2

2

3

3

4

4

5




8

rocks that overlie faulted basement, such as the basalgreenstones of the Yellowknife and Courageous Lakegreenstone belts.

• Gold deposits hosted by quartz veins and/or shear zonesthat developed along the flanks of major basementuplifts.

• Late Archean evolved granites and their associated peg-matite swarms, some of which are enriched in rare ele-ments (e.g. Li, Be, Sn, Nb, Ta). Such enrichment typi-cally correlates with a presence of ancient and multiply

recycled sialic crust.

• Diamondiferous kimberlite pipes in the Lac de Grasregion (Fig. 7), which appear to have ascended throughthe edge of the Central Slave Basement Complex (e.g.see cross-section of Fig. 3).

The Cover Sequence

The contiguous nature of the basement complex, by atleast 2.9 Ga, is indicated by a thin (typically 2-200 m) butwidespread 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 islocally intruded by ultramafic sills (Fig. 8), marks the onsetof the Neoarchean cycle of supracrustal development(Bleeker et al., 1999b), known collectively as theYellowknife Supergroup (Henderson, 1970).

The supermature and commonly fuchsitic quartzites thatare characteristic of this sequence overlie a regional uncon-formity 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). Thequartzites, which locally preserve cross-bedding, mark the

progressive drowning of this unconformity in a tide-influ-enced coastal setting. Overlying banded iron formations(BIFs) are indicative of deeper water, more outboard sedi-

mentation in response to continued subsidence.Abundant detrital chromite in the quartzites (Figs.

8C,E,F) may suggest contemporaneous komatiitic volcan-ism. Similar fuchsitic quartzite sequences occur in manyother cratons worldwide, particularly between ca. 3.1 Ga and2.8 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 mostare thin (1-10 m) and variable in composition along strike,changing from oxide-iron formation into silicate-rich vari-eties, or merely ferruginous cherts. Locally, however, fold-ing has thickened highly magnetic Fe-oxide BIFs into sub-stantial thicknesses (e.g. at Amacher Lake, on the easternflank 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 economicvalue, although some may possibly host epigenetic goldmineralization and may be under-explored for this commod-ity. However, across the craton, most of the known iron for-mation-hosted gold mineralization is associated withyounger BIFs that occur intercalated within low- to medium-grade turbidite packages (Fig. 4).

Fuchsitic quartzites that occur stratigraphically below theiron formations are enriched in detrital heavy minerals,including highly stable species such as chromite, zircon, andrutile. Individual, detrital, small black chromite grains are acharacteristic feature of these otherwise white to greyquartzites (Bleeker et al., 1999b) and, where recrystallizationis strong, allow these clastic rocks to be distinguished frommetacherts. During metamorphism and deformation thedetrital chromite grains reacted to varying degrees with sur-rounding minerals to produce bright green fuchsitic mica5

(Figs. 8E,F). In a few localities, chromite was concentratedenough to form seams of “black sand”. These occurrences,although of scientific interest, are too small to be of eco-nomic value. If road access were available, some of thegreen-white quartzite would make attractive building or dec-orative stone. In Greenland, India, and Australia, similar quartzites are commonly quarried for this purpose.Elsewhere in the world, quartzites similar to those describedhere 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. Roscoeand Minter, 1993). An initial survey of such potential in theSlave craton was carried out by Roscoe (1990, 1992), indeedresulting 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 oneremote locality, on the south shore of Desteffany Lake, one

FIGURE 5. (A) Simplified map showing minimum extent of the Hadean to Mesoarchean basement of the Central Slave Basement Complex (CSBC). Smallshaded spheres (yellow or orange) highlight locations where the diagnostic basement to cover stratigraphy has been observed. Based on basement charac-teristics, five tectono-stratigraphic domains can be recognized across the craton: 1) a westernmost domain, with basement of unknown age; no diagnosticevidence 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”, whichhas subtle characteristics that are reminiscent of the Yellowknife area and may well be underlain by ancient basement; currently, data from this area areinsufficient to resolve this question. (B) Age distribution of the Central Slave Basement Complex, as sampled by 296 concordant to near-concordant detri-tal 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 atca. 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 rec-ognized from U-Pb protolith ages (see Bleeker and Davis, 1999b).

5. Fuchsite is a Cr-bearing white mica in which Cr 3+ replaces Al3+. In the quartzites, it formed as a product of the following generalized reaction: detritalchromite ± K-feldspar + fluid → fuchsitic mica + fluid (e.g. Bleeker et al., 2000). Metamorphic breakdown and hydration of detrital K-feldspar was aided

by shearing.




858




8

of the authors found sulphide concentrations adjacent toultramafic rocks within the cover sequence. Overall, how-ever, the volume of komatiitic rocks and the potential for substantial massive Ni-Cu mineralization appears limited,not only at this stratigraphic interval but throughout theSlave craton.

Ca. 2.73-2.70 Ga Tholeiitic Volcanism

Wherever the thin cover sequence is recognized, it isoverlain 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-dom-inated volcanic sequence (Fig. 9) is known as the KamGroup (Helmstaedt and Padgham, 1986; Bleeker et al.,1999b). Possible correlative basalt successions are knownacross the basement domain (Fig. 10), as far east as theCourageous Lake belt, and at least as far north as around theExmouth antiform in the Acasta area (Bleeker et al., 2000).This basalt sequence typically consists of several hundredmetres 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 sequenceyield ages from 2722 to 2697 Ma (Davis and Bleeker,unpublished data, Isachsen and Bowring, 1997). A maficdyke 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 areeven older. In Yellowknife, the top of the sequence is repre-sented by voluminous basaltic flows and intercalated felsicvolcanic rocks of the Yellowknife Bay Formation, dated atca. 2700 Ma (Fig. 9). Similar 2700 Ma ages have beenobtained from the Courageous Lake and Acasta areas andstrongly support the overall regional correlation (Bleeker etal., 1999b). Stratigraphy, dense dyke swarms, and isotopicdata 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 >100,000km2, typical thickness 1-6 km) approaches large igneous

province (LIP) proportions (Coffin and Eldholm, 1994, 2001;

FIGURE 6 . Field photographs illustrating some key elements of the geology of the Slave craton. (A) Cleaned exposures of the Acasta gneisses at their dis-covery 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-weath-ering 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 somewhatyounger 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 (prox-imal) fragments range is size from 1 to 10 cm. (E) Massive sulphide mineralization of the Sunrise deposit, associated with ca. 2670 Ma felsic volcanicrocks just below the interface with the Burwash Formation turbidites. (F) Thick-bedded sandy turbidites typical of the Burwash Formation in its type areaeast 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 theCoronation 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 pro-ducer, the Ekati Mine of the central Slave cra-ton. Main picture shows the flat barren landsof the central Slave craton, with several pipe-like kimberlite bodies being excavated in cir-cular open pits. Inset (upper left) shows a

close-up of one of the partially excavated pipes. Several millimetre-size gem quality dia-mond octahedra are shown on upper right(photos courtesy of D. Snyder and G. Lockhart).



Ernst et al., 2005). The widespread basaltic volcanism probably accom- panied protracted rifting of the basement complex, possiblyassisted by mantle plume activity.The stratigraphic succession in theYellowknife area is compatible withsuch a rifting interpretation. At the

top of the Kam Group, bimodal vol-canic rocks of the Yellowknife BayFormation become progressivelymore intercalated with volcaniclas-tic sediments, before final intrusion

by thick tholeiitic sills. One of these, the Kam Point gabbro sill,has a preliminary baddeleyite ageof ca. 2697 Ma (Davis and Bleeker,unpublished data; see Fig. 9).

From a mantle perspective, itseems likely that events associatedwith the voluminous basaltic vol-canism across the ancient basementterrain must have involved thinningor at least modification of thelithospheric mantle below theCentral Slave Basement Complex.Large-scale melting was probablytriggered by adiabatic rise of asthenospheric mantle. Perhaps,then, the ca. 2.7 Ga basaltic volcan-ism may have contributed to thehighly depleted mantle composi-tions underlying the core of the cra-ton (e.g. Griffin et al., 1999; Grütter 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 volcaniclasticsedimentary rocks, is a highly favourable environment for seafloor hydrothermal activity and the formation of vol-canogenic massive sulphide (VMS) deposits. Indeed, numer-ous showings of sulphidic horizons occur throughout the

basalt-dominated greenstone belts of the central and westernSlave craton (Fig. 4).

Of particular interest are intercalated felsic volcanic flowsand/or sills, which are direct indicators of proximity to a dif-ferentiated magmatic centre, and thus a long-lived subvol-

canic heat source. The Bell Lake quartz-porphyritic tonalitesill (Fig. 6C) and the rhyolitic Townsite Formation, dated at2713 ± 2 Ma and ca. 2709 Ma, respectively (Davis et al.,2004), are examples of such proximal felsic volcanic rocksin the Yellowknife greenstone belt. Hydrothermal alterationand minor sulphide mineralization are known from theYellowknife Belt (e.g. the Homer Lake showings, see Fig.4), but to date no deposits of potential economic interesthave been found. Similarly, despite at least a first wave of exploration across the other basaltic greenstone belts of thewest-central Slave craton, the authors are not aware of anymajor discoveries.

Although possibly unre-lated, it is worth mentioning here the VMS deposits of theHigh Lake greenstone belt of the north-central Slave craton.These deposits, the High Lake A/B, D, and West Zone (Figs.4, 11), along with several other sulphide-rich horizons(Wolfden Resources Inc., 2005), are associated with bimodalvolcanic rocks dated at 2705 Ma (Henderson et al., 2000).Thus the setting and age are compatible with a mafic vol-canic-dominated rift environment similar to that of the KamGroup. However, basement rocks have not been identified inthe immediate area. Metal ratio plots (Fig. 12) discriminatethe High Lake deposits from most other VMS deposits in theSlave, suggesting a similarity to Cyprus-type deposits (the

“mafic class” of Barrie and Hannington, 1999), which formin mafic volcanism-dominated extensional settings.

Post-2.70 Ga Volcanism

Following ca. 2.7 Ga basaltic volcanism and rifting, mostareas in the Slave craton show a transition to calc-alkalinevolcanism characterized by abundant felsic and intermediatevolcanic rocks, tholeiitic to calc-alkaline basaltic rocks, andintercalated volcaniclastic sedimentary rocks (Fig. 10). Innearly all areas, these rocks are stratigraphically overlain byturbiditic greywacke-mudstones (Fig. 10). Ages for the arc-like volcanic rocks and their plutonic counterparts typically


860

Central SlaveCover Group(100-200 m)


Kam Group(0.3-6 km)

Pillow basalts

Thin BIF horizons

Thin rhyolite tuffs

2.82-2.86 Ga inherited zircons

Komatiite flows

Komatiitic sills

W B

' 9 9

Quartz pebble conglomerateRemnants of older supracrustals

3.0 to 4.0 Ga tonaliticto dioritic gneissesand foliated granitoids

2.9 Ga intrusive rocks,cut by mafic feeder dykes

Banded iron formationThin rhyolite tuff horizons

Fuchsitic quartzite

Detrital chromiteBlack shale

Y e l l o w k n i f e S u p e r g r o u p

Unconformity

2877±3

e.g. 3325±8

2853+2/-1

2908±2

2935±9

B

D

C

E

A

F,G

2722±1

2826±1.5

2738±2

Au

Cu

Au

Fe

Ni

Cr

U

Zn

Mn

2734±2

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 andBowring (1997), Bleeker et al. (1999a,b), and Ketchum and Bleeker (2000). Photos (B) to (E) illustrate char-acteristic lithologies and are keyed to the column. (F) Electron back-scatter image of rounded detritalchromite grain in fuchsitic quartzite of photo (D). (G) As in photo (E), but showing more advanced meta-morphic transformation to fuchsite.




8

“Cemetary tuff” (e.g. Fred Henne Park)Trapper tuff

Thin tuff layers (e.g. “Igloo Inn tuff”)

Thin rhyolite or silicified felsic clastic horizon; complex structure and correlation

A

BC “Townsite flows and sills”: massive (B) and fragmental (C) facies

Kam Point gabbro sills“Kam Point felsite” (Isachsen, 1992)

Townsite Formation: basal sandstone (”tuffite”) horizon (facies A)

--?--

Quartz-feldspar porphyry breccia horizon

Epiclastic sand- and siltstone intercalations

Epiclastic sand- and siltstone intercalations

Epiclastic rhyolitic sand- and siltstone intercalations

Cross-cutting gabbro dykes

Gabbro sills

“Yellorex flows” (e.g., Henderson and Brown, 1966)

“Negus flows”

Variolitic pillowed flows

“Fox flows”

Gabbro sill

“Stock flows”

“Bode Tuff” (dated rhyodacite boulder; Isachsen, 1992)

Thick, epiclastic, sandtone unitsrhyolitic

“Ash-flow tuff” at top of Giant mine section(”Banting Group” of Helmstaedt and Padgham, 1986)

Thin banded iron formation horizon

“Ranney chert”

“Ranney tuff” (multiple horizons)

Epiclastic sand- and siltstones

Thin banded iron formation horizons

Numerous gabbro sills (devoid of zircons)

Komatiite sills (flows?)

Spinifex-bearing komatiite flows in Yellowknife Bay(drill core intersections, probably part of Kam Group)

Komatiitic sills

Quartz pebble conglomerate

3.0 to 4.0 Ga tonaliticto dioritic gneissesand foliated granitoids

2.9 Ga intrusive rocks,

cut by mafic feeder dykes

Banded iron formationThin rhyolite tuff horizons

Fuchsitic quartzite (see Sircombe et al., 2001, for detrital zircon data)

Detrital chromiteBlack shale


Y e

l l o w k n

i f e S

u p e

r g r o u p

Central SlaveCover Group(~0.1-0.2 km)

C h a n F m . ( ~ 5 k m )

C r e s t a u r u m

F m . ( ~ 2 . 5 k m )

Townsite Fm.(~0.5 km)

I

II

III

IV

V

VI

VII

VIII

???

Y e l l o w k n i f e B a y F m . ( ~ 5 - 1 0 k m )

Unconformity

No zircon inheritance documented above this level

No zircons with ages younger than

2750 Ma known below this level

I n c r e a s

i n g v o

l c a n

i c l a

s t i c

s e

d i m e n

t s

Joe Lake tuffs: epiclastic sand- and siltstones, and felsic tuff

Bell Lake quartz-phyric tonalite sill

Crestaurum mine tuff (correlation between tuff and volcaniclastic facies uncertain)

2699±1

2697.3±1.5

2702±1

ca. 2701

ca. 2709

ca. 2705±3

2712±2

2722±1

2713±2

2853+2/-1

2734±2

2908±2 2877±3

WB2006

2935±9

2738±2

2826±1.5

2704±1

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 gab-

bro 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-por-

phyritic 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.




8

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 andVilleneuve, 1999; Ketchum et al.,2004).

The arc-like volcanic rocks tend

to be geochemically juvenile (e.g.Davis and Hegner, 1992; Cousenset al., 2002). They dominate theeastern part of the craton, wherethey lack any apparent associationwith older basement, its cover,and/or the basalt-dominated riftsequence discussed above. Theseobservations have led to models inwhich the eastern Slave representsan exotic juvenile arc (the “HackettRiver arc”) that collided with the

basement domain in the west (e.g.Kusky, 1989, 1990). However,similar arc-like rocks, with identi-cal ages, stratigraphically overliethe basement domain and its cover in the west-central parts of the cra-ton (Fig. 10A), where they can betied to the basement and bimodalrift volcanic rocks by means of unconformities (Bleeker, 2001),crosscutting feeder dykes, and sub-volcanic intrusions (see Fig. 13)(Bleeker et al., 1999a).

It thus appears that, if theserocks were generated in an arc-like

setting, this arc was constructedmarginal to, and on top of, thehighly extended continental crustof the Central Slave BasementComplex. This suggests a marginalto continental arc setting. The arcmust have been actively extending,evolving into a back-arc basin thatwas ultimately filled with turbiditicsediments (Figs. 10, 13). The geo-chemistry of the volcanic rocksand the associated subvolcanic plu-tons, although isotopically juve-nile, typically shows arc-like sig-

natures (negative Nb and Ta anom-alies, light rare earth and large-ionlithophile element enrichment, andrelative depletions of heavy rareearths) compatible with enrichedsources in a supra-subduction zonesetting and final equilibration withgarnet-bearing residues.Alternative models invoke partialmelting of a mafic underplate, per-haps accompanying delamination(e.g. Cousens et al., 2002).

Diorite dykes

Hornblende porphyry

Felsic volcanic rocks

Massive sulphides

Amygdaloidal mafic volcanic flows

HIGH LAKE - WEST ZONE SECTION 7472400N

HLW-03-21

HLW-03-19

HLW-03-24

HLW-03-22

HLW-03-47

100 m

0 m ASL

-100 m

-200 m

200 m

West East

300 m

HLW-03-50

HLW-03-55

HLW-04-68 (failed)

FIGURE 11. Cross-section through one of the High Lake massive sulphide deposits (after Wolfden ResourcesInc., 2005).

Cu Ag(x10)

Pb Zn Au( x 10 )3Zn

Bimodal rift-likeenvironment

Arc-like environment

Bear High Lake (A/B) Hackett (Cleaver L. Grade)High Lake (D)

Hackett (East Cleaver)High Lake (West)

Hackett (Boot Lake)

Hackett (Main)

MuskIzok Lake

Sunrise

Turnback (XL)

Yava

A B

FIGURE 12. Ternary plots of metal ratios for VMS deposits in the Slave craton. The metal ratios define twofields: one for Cu-rich, Pb-poor deposits associated with the older ca. 2.7 Ga basalt-dominated bimodal vol-canism, and another characterized by higher Pb (Ag) values for deposits hosted by the younger ca. 2.69 to2.66 Ga arc-like sequences.



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 isthe 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 associatedwith proximal felsic volcanic rocks at or near the transitionto overlying turbiditic metasedimentary rocks. This transi-tion is characterized by rhyodacite to rhyolite complexes,volcaniclastic sediment aprons, thin sulphidic chert hori-zons, and, in some localities, banded iron formations.Carbonate rocks (calc-arenites) are associated with some of the felsic complexes (Fig. 6D). This typical stratigraphicevolution, from shallow water or emergent felsic volcaniccomplexes to deep-water turbidite sedimentation, suggests

active extension, tectonic subsidence of the arc environment,and creation of significant accommodation space that wasthen filled by deep-water sediments (Fig. 13). Such an envi-ronment of active faulting, attenuated lithosphere, activevolcanism, and high heat flow, has long been recognized asa favourable setting for development of large VMS deposits.

Other characteristic features of this high heat flow regimeare the presence of high-level synvolcanic intrusions, for instance in the Hackett River belt (see Fig. 14; the Sandy HillPluton, Hanimor Gneiss Dome, and smaller intrusive bodiesthat 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 rela-tionship to these intrusions. It is possible that these synvol-canic intrusions contributed magmatic fluids to the ascend-ing hydrothermal solutions, resulting in the unusually highsilver and lead contents for these deposits. Stratiform car-

bonate rocks are a common feature associated with theHackett River deposits as well as the BB, Bear, and Turnback Lake (XL) deposits. A likely origin for these carbonatedeposits is the venting of Ca-saturated hydrothermal fluidsonto the seafloor and rapid mixing with cold, higher pH sea-water (i.e. analogues to “white smokers” in modern seafloor settings). If correct, these carbonate deposits represent excel-lent marker horizons and vectors in the search for new vol-canogenic massive sulphide deposits.

Ca. 2.68-2.66 Ga Sedimentation

Starting at ca. 2680 Ma, a broad turbidite basin — theBurwash Basin — developed across much of the craton and

progressively buried the volcanic substrate (e.g. Henderson,1985; Ferguson et al., 2005). The transition from volcanicand volcaniclastic rocks to deep-water sediments is com-monly conformable or disconformable6. In a number of localities, however, the transition is marked by a well devel-

oped unconformity or nonconformity7 that cuts down intounderlying crystalline rocks and is overlain by shallow-


864

Shallow subduction?

Rifted Mesoarcheanbasement

CSBC Hackett River Burwash Basin(back-arc)

Arc to rifted arc

Large-scale rejuvenationof lower crust

Active margin at Burwash time

Possible accretionary prism

west-central Slave

easternSlave

terrane X

Hope Bay block? ESE WNW

Collision (D1)

Colliding terrane

Pb

“Tonalite factory”

Refertilized mantle wedge

Thick tholeiitic basaltsJuvenile volcanicsQuartzite-BIF cover sequence

Crystalline basement withsyn-volcanic intrusions

Arc pluton “basement”

Rifted Mesoarchean basementand its autochthonous cover

Legend for basementdomains of Slave craton:

Juvenile arc-like volcanic rocksand their plutonic infrastructure;

oldest dated flows 2705 Ma

Unknown terrane responsible for D1 collision: closure and foldingof the Burwash Basin

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 asupra-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 compo-nents 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 andmid-crust. Collision of an unknown terrane (X), between 2650 and 2630 Ma, led to closure and F1 folding of the Burwash Basin.




8

water clastic rocks, including con-glomerates with granitoid cobbles(Henderson, 1985; Bleeker et al.,1997; Bleeker, 2001). One suchlocality occurs along the south-western flank of the Sleepy DragonComplex, another at Point Lake,and they demonstrate an autochtho-

nous nature of the sedimentarysequence (Bleeker, 2001). Detritalzircons in the basal clastic rocksindicate that deposition begansometime after 2683 Ma (Bleeker et al., 1997).

A persistence of volcanic inter-calations up-section and late maficsill complexes suggest a volcani-cally active extensional setting,

perhaps best compared with mod-ern back-arcs. The minimum sizeof this basin was ca. 400 x 800 km(Fig. 15A), making it the largestand possibly best preservedArchean turbidite basin in theworld, comparable in size to theJapan Sea. Like the modern JapanSea, the Burwash Basin waslargely ensialic, in agreement withinferences by early workers (e.g.Henderson, 1985).

The Burwash Basin fill consistslargely of immature greywackesand mudstones, deposited belowwave base, and locally may beup to 10 km thick (Bleeker

and Beaumont-Smith, 1995).Intercalated tuff layers have beendated at 2661 Ma (e.g. Bleeker andVilleneuve, 1995). Across theSlave craton, the greywacke tur-

bidites have been given differentformational names: the classicalBurwash Formation in theYellowknife Domain (Henderson,1972); the Contwoyto Formation incentral and northern Slave, identi-cal in essentially all aspects to theBurwash Formation further south,except for the presence of interca-

lated iron formations; the ItchenFormation, a more mud-rich faciesin the north-central Slave; and theBeechey Lake Group in the northeastern Slave, which alsocontains iron formations. Many of the turbidite beds, partic-ularly those of the Burwash and Contwoyto Formation, aresand dominated with only thin silt to mud intervals at the topof the graded beds. In the Yellowknife Domain, thick amal-

gamated sand beds (2-10 m) are not uncommon (Fig. 6F).Petrography, detrital zircons, and geochemical analysis indi-cate that the greywacke detritus consists of a mixture of mafic and felsic volcanic rocks and uplifted plutonic infra-structure, with minor input from ancient basement rocks

Goulburn Supergroup

Regan Intrusive SuiteHackett River Group

Beechey Lake Group

Mara River Complex

Siltstone, quartzite, argillite, conglomerate

Undifferentiated plutonic rocksBimodal arc-like volcanic rocks

Carbonate facies with dacitic clasts

Synvolcanic plutons

Turbiditic metasedimentary rocks, undifferentiated

VMS deposit (>100,000 tonnes)

VMS occurrence (no tonnage reported)

Undifferentiated migmatitic and plutonic rocks

Mixed metavolcanic rocks and subvolcanic intrusives

Paleoproterozoic

Archean

0 10 20 km

FingerLake

Finger Lake

JoJoA Zone A Zone

BootLakeBoot Lake

EastCleaver

East Cleaver

YavaYava

MuskMusk

N

Hanimor GneissComplex

Sandy Hill pluton

65°56’

108°10’

Malley Rapids Anticlinorium

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. Alsoshown is the carbonate unit that bears a close spatial relationship to the VMS deposits.

6. A disconformable contact is conformable but with a considerable interval of nondeposition, i.e.,a significant time hiatus or diastem.7. A nonconformity: an unconformity in which a younger sedimentary sequence overlies nonstratified crystalline rocks, in this case either older synvolcanic

plutons or ancient crystalline basement rocks (e.g. Point Lake).




866

0 100

km

0 100

km

0 100

km

0 100

km

1 0 5

O

62O

1 0 5

O

6 9O

1 1 7 O 68

O

1 1 7

O

61O

1 0 5

O

62O

1 0 5

O

6 9O

1 1 7

O 68O

1 1 7

O

61O

1 0 5

O

62O

1 0 5

O

6 9O

1 1 7 O 68

O

1 1 7

O

61O

1 0 5

O

62O

1 0 5

O

6 9O

1 1 7

O 68O

1 1 7

O

61O

YRFZ

Y R F

Z

BR FZB R F Z

YRFZ

Y R F Z

BR FZB R F Z

YRFZ

Y R F

Z

B RFZB R F Z

YRFZ

Y R F Z

B RFZB R F Z


Great Slave Lake


Wopmay

Orogen

W o p m

a y O r o

g e n

W op

n

ma

Z oe

W o p

n

m a y F a u

l t

Z o

e

Bath

urstFault

B a t h

u r s t F a u l t

Talson

T a l t s

o n

Magmati

c Zone

M a g m

a t i c

Z o n e

Great

e

Slaev Lak

G r e a t

e

S l a

e vL a k

Shear Zo

ne S h

e a rZ o n e

Thel

onFro

nt

T h

e l o

n F

r o n

t

P

aleozoicPla

tform

P a l e o z o

i c P

l a t f o r m

PointLakePoint Lake

LacdeGras

Lac de Gras

ContwoytoLake

ContwoytoLake

CSBCCSBC


Great Slave Lake


Wopmay

Orogen

W o p m

a y O r o

g e n

W op

n

ma

Z oe

W o p

n

m

a y F a u

l t

Z o

e

Bath

urs

tFault


Talson

T a l t s

o n

Magmati

c Zone

M a g m

a t i c

Z o n e

Great

e

Slaev Lak

G r e a t

e

S l a

e vL a k

Shear Zo

ne S h

e a rZ o n e

Thel

on Fro

nt

T h

e l o

n F

r o n

t

P

aleozoicPla

tform

P a l e o z o

i c P

l a t f o r m

PointLakePoint Lake

LacdeGras

Lac de Gras

ContwoytoLake

ContwoytoLake

CSBCCSBC


Great Slave Lake


Wopmay

Orogen

W o p m

a y O r o

g e n

W op

n

ma

Z oe

W o p

n

m a y F a u

l t

Z o

e

Bath

urstFault

B a t h

u r s t F a u l t

Talson

T a l t s

o n

Magmati

c Zone

M a g m

a t i c

Z o n e

Great

e

Slaev Lak

G r e a t

e

S l a

e vL a k

Shear Zo

ne S h

e a rZ o n e

Thel

onFro

nt

T h

e l o

n F

r o n

t

P

aleozoicPla

tform

P a l e o z o

i c P

l a t f o r m

PointLakePoint Lake

LacdeGras

Lac de Gras

ContwoytoLake

ContwoytoLake

CSBCCSBC


Great Slave Lake


Wopmay

Orogen

W o p m

a y O r o

g e n

W op

n

ma

Z oe

W o p

n

m

a y F a u

l t

Z o

e

Bath

urs

tFault


Talson

T a l t s

o n

Magmati

c Zone

M a g m

a t i c

Z o n e

Great

e

Slaev Lak

G r e a t

e

S l a

e vL a k

Shear Zo

ne S h

e a rZ o n e

Thel

on Fro

nt

T h

e l o

n F

r o n

t

P

aleozoicPla

tform

P a l e o z o

i c P

l a t f o r m

PointLakePoint Lake

LacdeGras

Lac de Gras

ContwoytoLake

ContwoytoLake

CSBCCSBC D

EE

RR

WW

HLHL

A

C

B

D

F1 foldbelt

F 1f o l d

b e l t

Defeat Su

ite

D e f e a t

S u i t e

Burwash

Basin

B u r w a s h

B a s i n

Burwash

Basin

B u r w

a s h

B a s i n

BIFBIF

BIFBIFNoBIFNo BIF

NoBIFNo BIF

BIFBIF

BIFBIF

Ca. 2680-2660 Ma

Ca. 2635-2625 Ma

Ca. 2650-2630 Ma

Ca. 2625-2605 Ma

Minimum extent of Burwash Basin

Defeat Suitemagmatism

Early fold trendsin Burwash Basin

Post-Burwash (volcano-)sedimentary basins




8

(Henderson, 1972; Jenner et al., 1981; Henderson, 1985;Yamishita and Creaser, 1999; Ferguson et al., 2005). Themain axis of the basin and subsequent structural trendsappear to have been northeast-southwest (Fig. 15), distinctlyacross the north-south isotopic boundaries that track thenature of deep basement. This interpretation is based on thefollowing observations:

• Identical Burwash Formation turbidites extend from near the type area of Yellowknife (e.g. Henderson, 1972) tothe northeastern Slave (Figs. 10, 15A).

• Banded iron formations in the turbidites are restricted tothe northwest half of the craton (Fig. 15A) suggesting anortheast to southwest facies boundary or tectonic trendacross the basin (see also Padgham, 1992). This argu-ment is complicated however by the observation thatsome iron formations in the western part of the craton arehosted by a younger turbidite package, the Damoti Lakesequence (see Fig. 10E).

• Earliest folds in the turbidites, which formed between2650 and 2630 Ma, have northeasterly trends after qual-itative “unfolding” of younger superimposed fold gener-

ations (Fig. 15B) (e.g. Bleeker, 1996).• The earliest plutonic suite that intrudes folded Burwash

strata, the ca. 2630 Ma Defeat Suite, appears to form anortheast-southwest-trending magmatic belt across thesoutheastern half of the craton (Fig. 15C) (Davis andBleeker, 1999).

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

Mineralization

The immature greywackes and mudstones of the BurwashBasin contain few primary mineral deposits other than

banded iron formations (Fig. 6G). The latter occur interca-lated in greywackes scattered across a broad swath in thenorthern part of the craton (Fig. 15A), from the Goose Lakeand George Lake areas to the Point Lake area, and from thereto the southwestern Slave. Many of the BIFs are highly mag-netic. Although of scientific interest for the understanding of facies boundaries, basin evolution, and geochemistry, theyare 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 theBack River Complex itself; 2) at the upper contact of theBack River volcanics; and 3) an upper horizon well within

the turbidites of the Burwash Formation. Lateral facieschanges exhibited by the mineralogy of these iron formationunits vary from oxide facies (magnetite±hematite andquartz), through silicate facies (chert-grunerite±stilpnome-lane±chlorite), to carbonate facies (siderite ±quartz).

The principal type of economic mineralization withinBurwash Formation metaturbidites is epigenetic gold miner-alization hosted by the intercalated BIFs (Padgham, 1992).

The most important example of this deposit type is the Lupindeposit on the southern shores of Contwoyto Lake (Bullis etal., 1991a,b; Geusebroek and Duke, 2004), which has been asignificant gold producer from 1982 to 2003, yielding

between 3 and 4 million ounces of Au (Normin, 2005). Other examples, such as George Lake (Fig. 16) and Goose Lake,occur throughout the northern Slave craton and may becomeeconomic with elevated gold prices and better access. Thegeneral model for these deposits is that the host BIFs servedas chemical traps for gold-bearing, low-salinity, mixed H2O-CO2 fluids during metamorphism and deformation (Bullis etal., 1991b; see also Kerswill, 1993; Phillips, 1993).Destabilization of the Au-carrying sulphur complexes, due tointeraction with reduced Fe-rich host rocks, led to alterationand gold deposition, either in veins or in fluid-altered andsulphidized zones of the iron formations. The structural tim-ing of these epigenetic deposits is generally syn- to late-kine-matic and syn- to late-metamorphic, i.e.,consistent withmaximum fluid production deeper in the thickened struc-tural-metamorphic pile. The most likely source for the fluids,and the gold, is metamorphic devolatilization of a volumi-nous, immature, sediment pile and its volcanic substrate(Phillips, 1993). Sporadic iron formations provided acciden-tal traps to the migrating fluids, with discrete structureslocally playing a role in increased focusing of fluid flow.Similar processes also led to gold-bearing quartz veinswithin metaturbidites (e.g. laminated veins along sheared

bedding planes, saddle reefs) (Boyle, 1961, 1986), but with-out a large-scale focusing mechanism these occurrences anddeposits tend to be of small size, although locally of highgrade. Examples are the Ptarmigan and Discovery mines in

proximity to Yellowknife (Fig. 4) (e.g. Brophy, 1987).

Ca. 2.65-2.63 Ga Closure of the Burwash Basin

Turbidite sedimentation in the Burwash Basin came to anend sometime before 2650 Ma, the age of the oldest recordedgranitoid (feldspar porphyry) pluton intruding Burwashstrata (Point Lake area, Mueller et al., 2001). Subsequenttectonic events record the closure and folding of theBurwash Basin (D1) prior to 2630 Ma (see F1 fold belt inFig. 15B). The latter age constraint is provided by early plu-

tons of the Defeat Suite (Davis and Bleeker, 1999), a distinctand possibly subduction-related magmatic suite across thesouthern (and southeastern) Slave craton (Fig. 15C).

FIGURE 15. Thematic maps illustrating key stratigraphic and structural aspects of the Slave craton through time. (A) Minimum extent of the ca. 2680 to2660 Ma Burwash Basin, based on continuity and geochronological similarity of turbiditic greywackes across large parts of the craton. The dash-dot lineseparates areas with intercalated iron formations (northwest) from those lacking iron formations (southeast). Yellow spheres highlight localities with pre-cisely 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) DefeatSuite 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 toWheeler Lake (W). Its full extent is not known.



Closure of the highly extended, but largely ensialic back-arc basinallowed considerable shorteningand mobility but with a structuralstyle dominated, at least at highstructural levels, by fairly system-atic, mostly upright, northeast-southwest trending fold trains (e.g.

Bleeker and Beaumont-Smith,1995). At deeper levels, e.g., alongthe basement-cover interface, thefold trains must have detachedallowing differential shortening of the basement and cover (e.g.Kusky, 1991). The folded Burwashstrata do not represent an outboardaccretionary prism however (cf.Kusky, 1991), which wouldrequire a trench setting rather thanthe more likely ensialic back-arcsetting; hence, there is little evi-dence for a discrete “Contwoyto

terrane” (Kusky, 1989) in the cen-tral Slave craton.

Interestingly, the northeast-southwest structural grain of theF1 fold belt is also recognized inthe lithospheric mantle (Grütter etal., 1999). Shallow subduction(either from the southeast or north-west?) may have emplaced distinctmantle slabs (Davis et al., 2003b).These processes terminated withdocking of an outboard terrane(e.g. Fig. 13), either in the south-east or the northwest; however,this terrane is not preserved withinthe present limits of the Slave cra-ton. Finally, crustal thickening ledto uplift and erosional exhumationof folded Burwash strata and par-tial unroofing of Defeat Suite plu-tons. Detrital zircons of DefeatSuite 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 initiateddevolatilization reactions and metamorphic fluid flow.Although these events thus likely kick-started the develop-ment of epigenetic gold mineralization, they were followedand overprinted by much more intense metamorphic eventsca. 20 to 30 million years later, during D2 deformation.

Arc-generation and subduction processes almost certainlymodified the mantle lithosphere below the Slave craton, pos-sibly 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 northeast-southwest trends of the Burwash Basin and D1 folding(Grütter et al., 1999).

Post-2.63 Ga Turbidites

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, unpublishedSHRIMP data, Pehrsson and Villeneuve, 1999; Bennett etal., 2005; Bennett, 2006) record a migration of sedimenta-tion and tectonic activity to the northwest. Deposition wascoeval with uplift and erosional unroofing of Defeat plutons


868

112 25’

107 30’

65 56’

66 55’

o

o

o

o

0 500 1000 m

L o n g

L a k e

Camp

F o r e s t

Z o n e

F o r d

L a k e

B o b L a k e

G e o r g e

L a k e

?

L o n e

C o w

P o n d

?

?

?

?

70 30

5

S l a v e

L

McCoy Lake L

o c a l e

3

L o c a l e

4

L o c a l e

2

L o c a l e

1

L o c a l

e

1 - N o r t

h

N o r t h

E x t e n s

i o n

p,c

d

O c c u r r e n c e

L a k e

LookoutHill

F o l d

N o s e

?

L a m b L .

G H

Z o n e

Paleoproterozoic

Neoarchean

Goulburn Supergroup

Yellowknife Supergroup, Beechey Lake Group

Arkose, quartzite, siltstone,dolomite, basalt, rhyolite

Mackenzie diabase dykes

Zones of elevated strain

Feldspar-porphyritic rhyolite (2683 Ma)

Turbiditic rocks

Sandstone to granulestone

Silicate facies iron formation

N

K o m

a t i

o

k D

i

n

e

e

f

t

o r m

a

o n

Z

M c C o y

e f r m

a

D

o

t i o n Z o n e

FIGURE 16. Geology of the George Lake area, northeastern Slave craton, where significant epigenetic goldmineralization is hosted by tightly folded and metamorphosed iron formation in Burwash-age metaturbidites(after Jefferson et al., 1992).




8

and tightly folded Burwash Formation strata. Shortly fol-lowing their deposition, these younger turbidites were short-ened and intruded by ca. 2616 to 2608 Ma tonalite-granodi-orite plutons similar to the Concession Suite of theContwoyto Lake area (Davis et al., 1994; Pehrsson andVilleneuve, 1999).

In the multiply folded, metamorphosed, and intermittentlyexposed terrain of the western Slave craton, it has proven dif-ficult to distinguish these younger turbiditic greywackes fromBurwash Basin turbidites. There is no sharply defined demar-cation line that separates the two turbidite packages andrecognition of the younger sequence relies heavily on theabsence of Defeat Suite-age plutons and the presence of <2640 Ma detrital zircons. Preliminary work suggests that theyounger turbidite sequence contains abundant intercalatediron formations, mostly of silicate facies, those of the DamotiLake 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 tur-

biditic greywacke variably enriched in metamorphic garnet,other Fe-rich silicates, and/or disseminated Fe sulphides.

A distinct volcaniclastic sediment and greywacke pack-age, associated with felsic volcanic rocks and subvolcanicintrusions, 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 examplesof a preserved volcano-sedimentary carapace coeval withone of the major plutonic suites, i.e., the Concession Suite.

MineralizationTypes of mineralization within the younger (turbiditic)

greywacke packages are similar to those in folded BurwashBasin strata. A principal example of epigenetic gold miner-alization is that hosted by silicate facies iron formation in theDamoti Lake area. Similar iron formations occur all alongthe western margin of the Slave craton, from the RussellLake 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 alsooccurs in numerous “lean” iron formations throughout the

27302740 2720 2710 2700

2700

2690 2680 2670 2660 2650

2650

2640 2630 2620

Defeat SuiteDefeat Suite

ConcessionSuite

ConcessionSuite

VOLCANISMVOLCANISM

2680-2650MA SEDIMENTATION2680-2650 MA SEDIMENTATION

EARLYPLUTONISMEARLY PLUTONISM

POST-DEFEAT SEDIMENTATIONPOST-DEFEAT SEDIMENTATION

PAN-SLAVE GRANITOID PLUTONISMPAN-SLAVE GRANITOID PLUTONISM

METAMORPHISMMETAMORPHISM

FABRICS & DEFORMATIONFABRICS & DEFORMATION

Time (in Ma)Time (in Ma)

Davis & Bleeker, 1999

e.g. Pehrsson & Villeneuve, 1999(e.g. Strachan tonalite, Indin Lake belt)

cut by Consession Suitedykes and stocks

Kam Point gabbro sills(Davis et al., 2004)

Pehrsson & Villeneuve,1999


Davis & Bleeker, 1999van Breement et al., 1992

No precise timing constraints, but post-Defeat and pre-S cleavage development 2 Suite

Villeneuve & Relf, 1998

in van Breemen et al., 1992

WB 05

James and Mortensen, 1992 Dudás et al., 1990

Dudás et al., 1990



No detailed timing

James & Mortensen, 1992

Anton & Suseplutons?

Anton & Suse plutons?

W o o

l B a y

E a r l y

D e

f e a

t

S u

i t e p

l u t o n s

U W

a t t a

L k

S u s e

L k

g r a n

i t e

W D

e f e

a t

p l u t o n

H i d d e n

L k

D u c

k f i s h

L k

S p a r r o w

L k

L e

i t h L

k

c a r b o n a

t i t e

M o r o s e

L k

A g a s s

i z L

k

HiddenplutonHidden

pluton

Morose GraniteMorose Granite

2610 2600

2600

2590 2580 2570 2560 2550

D2D2 EXTENSIONEXTENSION ARTIAL

EXHUMATIONPARTIAL

EXHUMATION

SHALLOW CRUSTAL LEVELSSHALLOW CRUSTAL LEVELS

(Qualitative t-T curves)(Qualitative t-T curves)

DISCRETEAUREOLESDISCRETE AUREOLES

M1M1 M2AM2A

M2BM2BM2CM2C

D1D1

S sla tycleavage

1S slaty cleavage

1

Young turbidite deposits, e.g. Damoti LakeYoung turbidite deposits, e.g. Damoti Lake

S crenulationcleavage

2S crenulationcleavage

2

Subhorizontal layering

in midto lower crust

Subhorizontal layering

in mid to lower crust S crenulationcleavage

3S crenulationcleavage

3

TemperatureTemperature

TimeTime

ARC RIFTING ARC RIFTING IFTINGRIFTING PRE-2687 MA DEFORMATIONPRE-2687 MA DEFORMATION D3D3

M2M2

Banting GroupBanting Group

Kam Group

ProsperousSuite

ProsperousSuite

Subvolcanic TTG plutons

Note: not included is the K-feldspar megacrystic Stagg PlutonicSuite, westof thestudy area, dated at 2588±7 Ma (seevanBreemenet al.,1992).

Polymict conglomerates and sandstones, e.g. Jackson LakeFmPolymict conglomerates and sandstones, e.g. Jackson Lake Fm

detrital zircon

Bleeker & Villeneuve, 1995

see Ketchum et al., 2004

Burwash FormationBurwash Formation

DetritusDetritus

DEEPCRUSTAL LEVELSDEEP CRUSTAL LEVELS

GRANULITESGRANULITES

M1M1

M2M2

Note: Suse Lake granite agelong judged to be anomalousand suspect. New data indicateit to be 2690±4 Ma (Davis and Bleeker,upublisheddata).

Zircon age, 2σ errors

Baddeleyite age, 2σ errors

U-PbAge Data:

Monazite age, errors2σ

Titanite age, errors2σ

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). (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 = quartzitesof the Central Slave Cover Group; 2 = turbiditic greywackes of Burwash Formation (s.s., 2a) and correlatives in the eastern Slave (Back River, 2b); theDamoti Lake turbidites (3); and synorogenic conglomerates (4). Known mafic dyke swarms are shown by narrow black bars (or grey, if poorly dated).

A



western Slave (Padgham, 1992), e.g., the Wheeler andGermaine Lake areas (“W” in Fig. 15D). The stratigraphicstatus of the iron formations and the host turbidites in this

particular area remains currently unresolved. Extensive fieldinvestigations have shown that an earlier reported 2612 Ma“tuff bed” (e.g. Isachsen and Bowring, 1994) is in factanother crosscutting porphyry dyke. Hence, the turbidites inthis area are >2612 Ma and could be of Burwash age.

2.60-2.58 Ga, Final Orogenesis

Starting at ca. 2600 Ma, the entire craton was affected bycross-folding and significant further shortening (D2), proba-

bly in response to final collision along a distant active mar-gin of the growing supercraton Sclavia (Bleeker, 2003).Structural trends, such as axial planes of folds and the dom-inant cleavage resulting from this event, are generally north-south to northwest-southeast, but may refract stronglyaround more rigid plutons or other heterogeneities (e.g.Davis and Bleeker, 1999). Moderate overthickening of thecrust led to high-T, low-P metamorphism (Thompson, 1989),

widespread anatexis, the appearance of S-type granites, anda hot and weak lower crust. These processes culminated inca. 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 buoyancy-driven deformation (Fig. 3); lower crustal devolatilizationreactions mobilized gold-bearing fluids; and synorogenicclastic basins formed and were immediately infolded intotight synclines (Bleeker, 2002). At least one of these synoro-genic clastic basins, the Beaulieu Rapids Formation(Corcoran et al., 1999) may have formed as late as ca. 2580Ma (Sircombe and Bleeker, unpublished SHRIMP data; seedetrital zircon age spectra in Fig. 17B). Late strike-slip faultsoverprinted and truncated the synclinally infolded clastic


870

Volcanism

Sedimentation

Plutonism

Detrital ZirconRecord

Tectonic Setting

Deformation

BIF

BIF

Plume uplift &volcanism?

Rifting &bimodal volc.

Pume impacts & incipient extension RiftingRifting Arc

volcanism

BurwashBasin

Damoti Lk assemblage

Defeat Suite

Arc?

Arc?

Kam 1 Kam 2 MalleySleepy Dragon 1 Sl. Dragon 2 MacKay Indin(NW)Duck Lake

Sl. Dragon 3 Dogrib

Syn-Kamplutons

Concession Suite

Yellowknife Supergroup

Rifting????

Blatchford LakeComplex

S-type granitesCarbonatites

“Granitebloom”

earliest granites

Late pegmatites

transitional

Syn-orogenic conglomerates

Burwash Fm.

High Lake assemblage

High Lake assemblage

Back-arcsedimentation

Arc magmatism &supra-subduction setting

Back-arccollapse

Collision

Collision(s)Extension

Izok

VMS? VMS Au? Au?Orogenic

AuDetrital chromite

Au, U?Rare metal

mineraliztion

Rare metalmineraliztion

M2

M1Cryptic

metamorphismCryptic metamorphism

in basement

Crypticmetamorphism

Widespreadregional High-Tmetamorphism

Cratonization

Uplift & exhumation Arccollision?

D2

Banting Group & correlativesOldest volcanic assemblages in eastern Slave

Kam Group &correlatives

Central Slave Cover Groupbimodal (ultramafic?) volc.

2900 2800 2700 2600 2500 2300 2200

Central Slave Cover Group

Fuchsiticquartzites

???

? ?? ? ?? ? ??

D3???

Metamorphism

Mineralization

BIFBIFBIF

BIFBIF

BIF

Big Spruce Complex

& Squalus Lake Complex

S U P E R C R A

T O N

S C L A V

I A

Syn-Bantingplutons

C E N T R A L S L A V E B A S E M E N T C O M P L E X

1

4

3

2

1

2a

2b

34

C R A T O N I Z A T I

O N

C R A T O N I Z A

T I O N

BIFBIF

D1

Early aureolesaround high level

TTG plutons

FIGURE 17 CONTINUED. (B) Extended time chart for entire Slave craton (excluding pre-2900 Ma history of the Central Slave Basement Complex). Note breaksin 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 DamotiLake turbidites (3); and synorogenic conglomerates (4). Known mafic dyke swarms are shown by narrow black bars (or grey, if poorly dated).

B




8

basins. Examples of such faults are the Yellowknife River Fault Zone in Yellowknife, and the Beaulieu River FaultZone 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 themantle, 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 synkinematicgranite sheets (Davis and Bleeker, 1999), represents the mostobvious deformation throughout much of the Slave craton. Itmust have driven moderate to significant crustal thickeningand led to the main thermal peak of regional metamorphismin most areas. The D2 deformation and associated metamor-

phism was the main driver for epigenetic gold mineralizationthroughout the Slave craton. In the Yellowknife greenstone

belt, it led to formation of the approximately 15 millionounce 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 moder-

ate to strongly deformed basaltic rocks, in proximity to aregional stratigraphic break, the Yellowknife River FaultZone (Figs. 2, 3). An asymmetric synclinal panel of synoro-genic conglomerates (the Jackson Lake Formation) occursalong this fault zone. Similar relationships are observed inseveral other major Archean gold camps, most notablyTimmins, Kirkland Lake, and Kalgoorlie. The critical con-trol common to all these camps is localization of Au miner-alization 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 sys-tems are known from the Slave craton, some of which saw

brief production in the past, only one other major camp hasemerged in recent years. This camp occurs in the Hope Bay

belt, on the Coronation Gulf coast of the Slave craton, andconsists of a string of deposits (Boston, Madrid, Doris; seeFig. 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-

RiftingRifting RiftingBack-ar c

rifting

Calderian orogeny(Slave-Hottah collision)

Oblique folding& strike-slip onwester n margin

Thelon orogeny(Slave-Rae collision)

Oblique docking of Slave microplate anddevelopment of theGreat Slave Lakeshear zone Continental

ar c

HearneLac de Gras Ghost Lake Hood River Mackenzie “E-W dykes”Indin(NW)k Lake

Bathurst Inlet Mar a River sills (& Morel?) “305 dykes”??

Spruce Complexualus Lake Complex

Passive margin sequence overlyingeastern rifted margin??

Coronation Supergroup

Kilihigok Basin

Rift?

Rifting????

Shelf

Tu ff s

Tu ff Foredeep

Foredeep Foreland basin

Great Bear Magmatic Ar c

Great Bear Magmatic Ar c

latchford LakeComplex

alon

Cu-AumineralizationPGEs? PGEs

Widespread low-grade metamorphicoverprint throughout much of Slave cratonresetting of low-T K-Ar systems

2000 1900 1800 1700 1300 1200 1100

Booth River

Complex

Hepbur n IntrusiveSuite

MuskoxIntrusion

SUPERCONTINENT

(LAURENTIA)N

UN

A Thermal transientin lower crust

and lithospher e

Vo lcanism

Sedimentation

Plutonism

Detrital Z irconRecor d

Te ctoni c S etting

Deformation

Metamorphism

Mineralization

FIGURE 17B CONTINUED.



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 ayoung, tightly infolded, synorogenic conglomerate package.The best examples are the Point Lake and Arcadia Bay areas.The Courageous Lake belt is equally interesting although itlacks a synorogenic conglomerate assemblage. This belthosts the Slave craton’s largest unexploited gold resource

(Tundra-Fat deposit at 29.5 million tonnes containing 6.86g/tonne gold).

Another class of mineral deposits related to final orogen-esis is that of rare-element-enriched granitoids (Li, Be, Sn,

Nb, Ta), in particular highly evolved anatectic granites andtheir late-stage pegmatites (e.g. Meintzer and Cerny, 1983;Meintzer et al., 1984; Meintzer and Wise, 1987). Tin (cassi-terite) 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 tendto be diagnostic for the presence of ancient, multiply recy-cled felsic crust.

Cratonization and Beyond

The youngest granite plutons of the Slave craton are ca.2590 to 2580 Ma (James and Mortensen, 1992; van Breemenet al., 1992; Davis and Bleeker, 1999; Bennett et al., 2005;Ootes et al., 2005). Only some pegmatitic granites areknown to be significantly younger, for instance a pegmatitein 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 over-thickening (D1-D2) and high intrinsic heat production, irre-versibly transferred a significant fraction of heat-producingelements and lower crustal fluids (and Au) to the upper crust.In the lower crust, it must have involved large-scale migra-tion of anatectic granitoid magmas, significant horizontalchannel flow of partially molten rocks, development of hor-izontal layering, and flattening of the Moho discontinuityinto 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 meta-morphic reactions and zircon growth continued to about2510 Ma (Davis et al., 2003a). Finally, sufficient coolingallowed the crust to mechanically couple with the mantle(Bleeker, 2002). The end product was cratonic crust of highrelative strength.

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

At 2230 Ma, northeast-trending mafic dykes of the Malley

swarm intruded the central Slave craton, providing the firstevidence for Paleoproterozoic mantle-derived magmatismand attempted rifting events (LeCheminant and vanBreemen, 1994). Subsequent to the Malley event, between2200 and 2000 Ma, Slave crust was affected by as many asten other mafic dyke swarms (see Ernst and Buchan, 2001,for a global data base, and 2004) and associated extensionevents before the craton finally broke out of the confines of its ancestral Sclavia supercraton. Details remain sketchyhowever. The eastern margin of the Slave craton, nowinvolved in and overridden by the Thelon Orogen, almost


872

0 5 10 km106°30’

68°10’

(Modified from Anonby et al., 1998; Hebel, 1999)

Boston

Madrid

Doris

Amphibolite faciesmafic volcanic rocks

Mafic volcanic rocksand gabbro sills

Felsic volcanic rocks

Granite to granodioriteintrusive rocks

Undifferentiated gra-nitoids and migmatite,

partly gneissic

Yellowknife Supergroup

Hope

Bay

N

2716±3 Mainheritance 2734 Ma?

Minor turbidites along centre of the belt are <2675 Ma and containa substantial proportion of olddetrital zircon grains (2.8-3.3 Ga)

Syn-orogenic conglo-merates, <2647 Ma

2677+3/-1 Ma

Shear zone

Quartz-carbonate-altered shear zone

Gold showing

Gold deposit (>10 tonnes)

6

FIGURE 18. Simplified geology of the Hope Bay greenstone belt of thenortheastern-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 containsinherited 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-sedi-mentary cover (a rifted fragment thereof?), rather than a continuation of the

juvenile greenstone belts of the eastern Slave craton.




8

certainly became established as a passive margin well before thewestern margin. The latter isreferred to as the so-calledCoronation margin and later

became involved in WopmayOrogen (e.g. Hoffman, 1980;Bowring and Grotzinger, 1992;

Hildebrand and Bowring, 1999).Once liberated out of Sclavia, theSlave continental microplate likelyexperienced a drift phase as anindependent craton, before being

progressively incorporated into thegrowing Laurentian collage and thesupercontinent of Nuna (Hoffman,1988; Bleeker, 2003).

Paleoproterozoic amalgamation processes varied along the marginsof the Slave craton. In the east, theSlave acted as a lower plate, beingoverridden by the west-vergentThelon Orogen (Fig. 3). In thesouth, along the shores and islandsof Great Slave Lake, deformationwas mainly transpressional alongan oblique suture that evolved intoa major continental transform

boundary (Hoffman, 1987; Hanmer et al., 1992). In the west, at leasttwo arc terranes were involved(Hottah terrane and the younger Great Bear Magmatic Zone,Ghandi et al., 2001; Ghandi andvan Breemen, 2005), followed byoblique folding and late-stage, dex-tral, strike-slip deformation alongthe Wopmay Fault Zone.Development of the Great Bear arc,

between about 1880 and 1840 Ma,likely involved subduction of Paleoproterozoic lithosphere belowthe western Slave craton (e.g.Bostock, 1998; Heaman et al.,2002).

Post-dating the assembly of Laurentia and Nuna, the Slave cra-ton, particularly along its margins,

became partially buried beneath

intracontinental Proterozoic basins (e.g. CoppermineHomocline and Elu Basin, see Kerans et al., 1981; Hoffmanand Hall, 1993). At ca. 1269 to 1267 Ma, the craton was

partly uplifted and intruded by the giant Mackenzie dykeswarm, 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 pre-served as the Coppermine Homocline. This is the last major event affecting the core of the craton, although someyounger mafic magmatic events did affect its edges (e.g. the723 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, withinterior seas expanding and receding across the craton.Ordovician and Cretaceous-Tertiary sedimentary rocks andfossils are known as wall-rock fragments in some of the cen-tral 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 eventswere triggered in the subcontinental mantle lithosphere andunderlying asthenosphere, leaving their traces as clusters of kimberlites across the craton (Fig. 19). From the several hun-dreds of kimberlites now known across the craton, the fol-

0 100

km

1 0 5

O

62O

1 0 5

O

6 9O

1 1 7 O 68

O

1 1 7

O

61O

YRFZ

Y R F Z

BRFZB R F Z

PbNd


Great Slave Lake


Wopmay

Orogen

W o p m

a y O r o

g e n

W op

n

maFZ o

e

W o p

n

m a y F a u

l t

Zo

e

BathurstFault


Taltson

T a l t s

o n

Magmati

c Zone

M a g m a t

i cZ o n e

Great

e

Slaev Lak

G r e a t

e

S l a

e vL a

k

Shear Zon

e S h

e a r Z o n

e

The l

onFro

nt

T h

e l o

n

F r o n t

P

aleozoicPla

tform

P a l e o z o

i c

P l a t f o r m

Central SlaveBasementComplex


SLAVE CRATON

Kimberlite Fields

Approximate locationof kimberlite pipes

Approximate outline of kimberlite field

Lacd e G r as

Lac de Gras

PointLakePoint Lake

ContwoytoLake

ContwoytoLake

Western Slave~430-460 Ma

Lac de Gras~75-45 Ma

Jericho Field~170-175 Ma

Southern Slave~520-545 Ma

Coronation Gulf ~613 Ma (Anuri)

Anuri

Jericho

Ekati

Diavik

Snap

Kennady

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 buriedeastern 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.



lowing ages have been recorded:Cambrian, Siluro-Ordovician,Permian, Jurassic, Cretaceous, andfinally Eocene (e.g. Heaman et al.,2003, 2004). The exact triggeringevents are still unclear but mayrelate to 1) mantle plumes or con-vective upwellings disturbing the

stress and thermal state of the litho-sphere (e.g. Sleep, 2003), or 2) rap-idly changing far-field stress pat-terns within the North American

plate due to events along its periph-ery, such as break-up and openingof the Iapetus Ocean, or changingsubduction patterns along theCordilleran margin (e.g. Snyder and Lockhart, 2005).

Mineralization

From the first attempted riftingevents at ca. 2230 Ma (Malley dyke

swarm) to final break-up of thesupercraton Sclavia at 2.1 to 2.0Ga, at least ten mafic and alkalinemagmatic suites were emplacedwithin Slave crust, as dyke swarmsand 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 andalkaline province within the southwestern part of the Slavecraton (Fig. 4), comprising the Big Spruce (Cavell andBaadsgaard, 1986), Squalus (Stubley, 1996), and Blatchford(Davidson, 1978; Bowring et al., 1984; Sinclair et al., 1994)

alkaline complexes, and the essentially contemporaneousDuck 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 theSlave craton and are initial precursor events prior to final

break-up8.

The Blatchford Lake intrusive complex, on the northshore of Great Slave Lake, consists of a number of nestedintrusions ranging from gabbro to granite and syenite. A cen-tral syenite body, the Thor Lake syenite, hosts rare metalmineralization (Be, Y, rare earths, Nb, Ta, Zr, and Ga) inhydrothermally 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 byinference the related syenite phases, have been dated at ca.2176 Ma (Bowring et al., 1984; Sinclair et al., 1994), the ageof 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 (byca. 80 million years) the high-temperature hydrothermalmineralization from the alkaline intrusions with which theyare closely associated. In general terms, the Thor Lake

deposits can be divided into a series of nested zones with akeel-shaped cross-section (Fig. 20): 1) a “Wall Zone”, com-

posed chiefly of quartz and feldspar, containing fluorite, sulphides, magnetite, carbonate, columbite, phenacite and sev-eral unidentified rare earth (REE)-bearing species; 2) a“Lower Intermediate Zone” that occurs predominately in thekeels of the T-Zone deposits (Fig. 20), consisting of quartz,

plagioclase, amphibole, biotite, muscovite, lepidolite, chlo-rite, carbonate, magnetite, with important accessory mineralsconsisting 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, carbon-ates, magnetite, sulphides, and REE minerals as accessories;and 4) a vuggy “Quartz Zone” occupying the core of thiscomplex (Trueman et al., 1988).

Another important example of rifting/extension relatedmafic magmatism is the large Booth River complex (Roscoe,1985; Roscoe et al., 1987; Hulbert, 2005b) of the north-centralSlave craton, outcropping from underneath the western extent

of the Kilihigok Basin strata (see Fig. 4 for location). Thiscomplex has been dated at 2025 ± 2 Ma (Roscoe et al., 1987;Davis et al., 2004) and appears a magmatic focal point alongthe more extensive and slightly southward-fanning Lac deGras dyke swarm, which is of similar age. The large, differen-tiated Booth River complex has attracted attention as a poten-tial host for platinum group element (PGE) mineralization.

The much younger Muskox Intrusion (1267 Ma), on thenorthwestern margin of the craton and related to the largeMackenzie event (LeCheminant and Heaman, 1989), is


874

N

A

A

A’

A’

(From Trueman et al., 1988)

Diabase dyke

Thor Lake Syenite

Grace Lake Syenite (2176 Ma)

Wall Zonemicrocline-albite-quartz,with Nb

Lower Intermediate Zonequartz-magnetite-chlorite-fluorite, with Be, Y, REE, Nb

Upper Intermediate Zonequartz-polylithonite-albite,with Be, Y, REE, Nb

Quartz Zone

quartz-carbonate-sulphides-fluorite, with REE

Thor Lake Mineralization

Legend:

Blatchford Lake Complex

T-Zone

T-Zone

Section

See Fig. B

1 km

150 m50 m

Lu Ke

Zone

A B

c

FIGURE 20. Geology of the Thor Lake rare element deposit (after Trueman et al., 1988). It occurs in the cen-tral syenite pluton of the Blatchford Lake Complex, an alkaline rift-related intrusive complex on the northshore of Great Slave Lake and a major component of the ca. 2180 Ma “Southwest Slave magmatic province”(see text).

8. Based on these approximately coeval magmatic events, we tentatively define a “southwest Slave magmatic province”, as shown in Figure 4.




8

another funnel-shaped layered intrusion that has attractedattention for reef-style PGE mineralization (Hulbert, 2005a).

The emplacement of diamondiferous kimberlite pipes rep-resents the final metallogenic event for the Slave craton,with kimberlite ages spanning several discrete pulsesthroughout the Phanerozoic (Heaman et al., 2004). At Lac deGras, 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 andDiavik. Several other diamond mines are in various stages of development (e.g. Jericho, Kennady Lake). In just over adecade, diamonds have become the most profitable com-modity within this ancient craton.

Summary

The Slave craton is a relatively small Archean craton witha 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 north-east part of the craton? Does cryptic ancient basementreappear 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 vol-

canic rocks in the eastern Slave? Do they form remnantsof a ca. 2.7 Ga, exotic, intra-oceanic juvenile arc thatcollided with the extended Central Slave BasementComplex between 2697 and 2687 Ma? If so, where isthe suture? Or do these volcanic rocks represent the old-est fill of narrow back-arc-like troughs formed by pro-gressive rifting of the Central Slave Basement Complexin an overall arc setting?

3. What are the detailed trend and extent of particular mag-matic (e.g. Defeat and Concession suites) and sedimen-tary belts (e.g. the post-Defeat turbidite basin along thewestern margin of the Slave)?

4. How did events inferred from the crustal geologicalrecord contribute to, or interfere with, formation of thesubcontinental mantle lithosphere below the Slave cra-ton?

5. How far does Slave mantle lithosphere extend below theRae craton to the east?

6. What is the detailed break-up history for each of theProterozoic margins of the Slave craton? In other words,how was the Slave crustal fragment liberated out of itsancestral supercraton Sclavia?

7. What is the detailed depositional record associated withrifting, thermal subsidence, and finally collision, alongeach of the margins of the Slave craton?

8. And, does Slave lithosphere extend northward all theway to the Innuitian front in the high Arctic (see Fig. 1)?

To many of these first-order questions, we currently haveonly rudimentary answers or mere guesses. More sophisti-cated answers will require more complete and more refineddata sets. In particular, a greatly expanded geochronologicaldatabase, both in quantity and precision, in conjunction withtargeted 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 allmajor greenstone belts have seen at least a first wave of

exploration for major commodities. These investigationsquickly led to the discovery of a number of large VMSdeposits (e.g. Izok Lake), which await road access for eco-nomic production. Gold potential remains high, particularlyin more remote greenstone belts that may not have seen ade-quate drill testing. In this respect, the Point Lake greenstone

belt, and its extension further north, appears attractive as ithas all the major attributes of a world-class gold camp.

Acknowledgements

We thank colleagues at the Geological Survey of Canada, particularly Bill Davis, John Kerswill, Dave Snyder, SallyPehrsson, and Ken Buchan, for sharing their insights into thegeology and metallogeny of the Slave craton. Kim Nguyenassisted with preparation of some of the figures. Thoughtfulreviews by John Ketchum, Tim Kusky, and Jan Peter helpedimprove the manuscript.

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