Geology and Metallogeny of the Superior Province

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GEOLOGY AND METALLOGENY OF THE SUPERIOR PROVINCE, CANADA

JOHN A. PERCIVAL

Geological Survey of Canada, 601 Booth St. Ottawa ON K1A 0E8

E-mail: [email protected]

Introduction

The Archean Superior Province represents the miningheartland of Canada, hosting major camps in the Abitibi dis-trict of Ontario-Quebec and Red Lake region of westernOntario. In addition, the Paleoproterozoic Sudbury structure(1.85 Ga), located on the southeastern edge of the Superior Province, is one of the world's largest nickel producingregions. This report summarizes information on the geologyand metallogeny of the Superior Province compiled in recentsyntheses of Ontario (Williams et al., 1992; Fyon et al.,1992) and the Superior Province (Card 1990; Card and Poulsen 1998; Skulski and Villeneuve, 1998), augmented byrecent work in the western (Percival et al., 2005a) and north-eastern Superior Province (Leclair et al., 2004a, 2004b;Labbé and Lacoste, 2004).

Superior Province

The Superior Province forms the core of the NorthAmerican continent (Fig. 1) and is surrounded by provincesof Paleoproterozoic age on the west, north and east, and Mesoproterozoic age (Grenville Province) on the southeast.Tectonic stability has prevailed since ca. 2.5 Ga in large partsof the Superior Province. Proterozoic and younger activity islimited to rifting of the margins, emplacement of numerousmafic dyke swarms (Table 1; Buchan and Ernst 2004), com-

pressional reactivation and large-scale rotation at ca. 1.9 Ga,and failed rifting at ca. 1.1 Ga. With the exception of thenorthwestern Superior margin that was pervasivelydeformed and metamorphosed at ca. 1.85 Ga, the craton hasescaped ductile deformation.

Current views regard the Superior Province as a collagemade up of small continental fragments of Mesoarchean ageand Neoarchean oceanic plates, with a complex history of aggregation between 2.72 and 2.68 Ga (Percival et al.,2004a; 2005a) and subsequent post-orogenic effects.Sedimentary rocks as old as 2.48 Ga unconformably overlie

Superior Province granites, indicating that most erosion had occurred prior to ca. 2.5 Ga.

The southern Superior Province (to latitude 52°N) is amajor source of mineral wealth. Owing to its potential for these and other commodities, the Superior Province attractsmineral exploration in both established and frontier regions.

Geological Setting

The Superior Province forms the Archean core of theCanadian Shield. It has been tectonically stable since ca. 2.6Ga, having occupied a lower plate setting during mostPaleoproterozoic and Mesoproterozoic tectonism that affect-

ed its margins.A first-order feature of the Superior Province is its lin-

ear subprovinces of distinctive lithological and structuralcharacter, accentuated by subparallel boundary faults (e.g.Card and Ciesielski, 1986). Trends are generally E-W in thesouth, WNW in the northwest, and NW in the northeasternSuperior (Fig. 2). Recent work based on isotopic and zirconinheritance studies has provided a means of "seeing through"the latest structural and magmatic events, particularly inregions dominated by granitic rocks, revealing fundamentalage domains across the Superior Province. Several domainsof Mesoarchean age are recognized in spite of pervasive

Neoarchean magmatism, metamorphism and deformation.The oldest continental crust (up to 3.7 Ga) occurs in the

Northern Superior superterrane (NSS; Skulski et al., 2000)and Inukjuak domain of the northeastern Superior province(David et al., 2003). A large region of ca. 3.0 Ga crust, the

North Caribou terrane (Stott and Corfu, 1991), which mayextend into northern Quebec (Stott, 1997), has been inter-

preted as a continental nucleus during assembly of theSuperior Province (cf. Goodwin, 1968; Thurston et al., 1991;Williams et al., 1992; Stott, 1997; Thurston, 2002). Farther south, the Winnipeg River (WR) and Marmion (MM) ter-ranes are relatively small continental fragments dating back to 3.4 and 3.0 Ga, respectively (Beakhouse 1991; Tomlinson

Fig. 1: Tectonic map of North America, showing location of the ArcheanSuperior Province at the core of the Canadian Shield (after Hoffman, 1989).



et al., 2004). In the far south, the Minnesota River Valley ter-rane (MRV) of unknown extent contains remnants of crust asold as ca. 3.6 Ga (Goldich et al., 1984; Bickford et al., 2004).In central Québec, the Opatica belt has Mesoarchean her-itage, as do large parts of the northeastern Superior Province(Leclair et al., 2004a).

Domains of oceanic ancestry, identified by juvenile iso-topic signatures and lack of inherited zircon, separate most

of the continental fragments. These dominantlygreenstone-granite terranes generally have longstrike lengths and record geodynamic environ-ments including oceanic floor, plateaux, island arcand back-arc settings (e.g. Thurston, 1994).Examples include the Oxford-Stull terrane in thenorth, the western Wabigoon in the west, and theWawa-Abitibi subprovince in the southeastern

Superior Province. Rocks formed in these envi-ronments host some of the largest massive sul-

phide deposits of the province, particularly in theAbitibi belt.

Still younger features, the metasedimentary belts (Breaks, 1991; Williams, 1991; Davis, 2002),separate some of the continental and oceanicdomains. Extending across the entire province,these 50-100 km wide belts of metagreywacke,migmatite and derived granite appear to representthick syn-orogenic sequences, deposited,deformed and metamorphosed during collisionalorogeny. Faults developed in association with sev-eral distinct tectonic events host a variety of oro-

genic gold deposits across the province.

Geophysical Setting

Early tomographic images of the Superior Province sug-gested lithosphere thickness of at least 250 km beneath thecraton (Grand 1987), and possibly as much as 350 km (Vander Lee and Nolet, 1997). Analysis of shear-wave splittingindicates prominent east-west anisotropy in the lithosphere(Silver and Chan, 1988; Silver, 1996; Kay et al., 1999a, b),which has been attributed to mantle deformation duringArchean tectonism. Musacchio et al. (2004) estimated upper mantle velocities in the 8.3-8.8 km.s-1 range, consistent with

depleted harzburgite compositions. In the western Superior,Kendall et al. (2002) distinguished a northern zone of isotropic upper mantle beneath the North Caribou terranethat contrasts with a southern zone characterized by east-west anisotropy. These domains, which are separated by asubvertical high-velocity zone, extend to about 300 kmdepth (Sol et al., 2002). To the north beneathPaleoproterozoic crust of the Trans-Hudson orogen, the lith-osphere is thinner and notably more isotropic (Kendall et al.,2002). Domains resembling those defined seismically arealso observed in the electrical conductivity structure. Cravenet al., (2001) reported an essentially isotropic mantle beneaththe North Caribou terrane, in contrast to pronounced east-west anisotropy in the south. Subsequent analysis has mod-

eled a steeply north-dipping, tabular, resistive "slab" separat-ing the two domains, which corresponds approximately tothe subvertical high-velocity zone. Both the seismic and electrical structures have been attributed to formation of theSuperior craton through subduction-accretion processes(Kendall et al., 2002; Craven et al., 2001). Heat flow in theSuperior Province averages 42 ± 8 mW .m-2 (Cheng et al.,2002; Rolandone et al., 2003) and the reduced heat flow(mantle component) is consistent with a thermal lithosphereat least 240 km thick (Jaupart et al., 1998), supporting theconcept of a thick tectosphere.

John A. Percival

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Fig. 2: Mosaic map of the Superior Province showing major tectonic elements. Datasources: Manitoba: Manitoba Department of Mines and Natural Resources 1965; Ontario:Ontario Geological Survey, 1992; Quebec: Géologie Québec, 2002; Leclair et al., 2004b.



Direct control on mantle lithosphere composition is gen-erally acquired from xenoliths in kimberlite pipes. However,the known pipes are not ideally situated to provide compre-hensive information. Pipes in the Attawapiskat area samplemantle close to the Superior margin, and indicate mainlylherzolitic compositions and a cool geotherm (Scully, 2000).The compositional range is similar to that from Kirkland Lake pipes (Vicker and Schulze, 1994; Schulze 1996), where

geothermobarometry suggests a warmer geotherm corre-sponding to surface heat flow of about 40 mW.m-2. Takentogether, observations of the Superior Province mantle indi-cate a cool refractory lithosphere typical of Archean cratons(Jordan, 1978). Individual mantle domains correspond to

broad geological features (terranes) and anisotropies of seis-mic velocity and electrical conductivity are east-trending,coplanar to the dominant penetrative crustal structures.

Gravity and aeromagnetic trends generally correspond well with first-order crustal geological features.Metasedimentary belts define gravity and aeromagneticlows, whereas greenstone belts form local gravity highs.Gravity data have been inverted to estimate crustal thick-ness, following removal of near-surface effects (e.g. Nitescu

et al., 2003). This approach supports seismic observations of crust thinner by as much as 9 km beneath the English River subprovince. Seismic reflection profiles in both western(White et al., 2003; Calvert et al., 2004) and eastern Superior (Calvert and Ludden, 1999) generally indicate gently north-dipping structures in southern low-grade regions, changingto southward dips in plutonic gneissic terranes farther north.The vergence patterns have been interpreted in terms of adoubly vergent orogen (White et al., 2003) or crustal-scalesynclinorium (Hynes and Song, 2005). Reflectors locallyextend beyond the Moho, where they have been interpreted as subduction scars (Calvert et al., 1995; White et al., 2003).

Western Superior Province

Northern Superior Superterrane

The Northern Superior superterrane (NSS) at the north-ern fringe of the Superior Province is dominated by graniticand gneissic rocks and has been identified based on frag-mentary isotopic evidence from Ontario and Manitoba(Skulski et al., 1999). Supracrustal units in the Assean Lakecomplex of Manitoba include greywacke with detrital zirconages up to 3.9 Ga, iron formation and mafic to intermediatevolcanic rocks (Böhm et al., 2000; 2003). The ancient rockshave been strongly reworked by granitoid magmatism at 3.2-3.1, 2.85-2.81 and 2.74-2.71 Ga, representing evolution in a

continental magmatic arc setting, followed by amphibolite-facies metamorphism at 2.68 and 2.61 Ga that may haveresulted from collisions during tectonic assembly (Skulski etal., 2000; Böhm et al., 2003). The Northern Superior supert-errane is bounded to the south by the North Kenyon fault,which juxtaposes it with the Oxford-Stull terrane. Prominentductile shear zones in the eastern NSS may be Neoarcheanor Paleoproterozoic in age. Extrapolation of aeromagnetictrends eastward beneath Paleozoic cover suggests that unitsof the NSS may host diamondiferous kimberlite such as theVictor Pipe in the Attawapiskat area. Additional exploration

targets have been identified through kimberlite indicator minerals in the Ontario-Manitoba border region. Correlationis uncertain between the NSS and units of similar antiquityin the Inukjuak domain of northern Quebec.

Oxford-Stull Terrane

The Oxford-Stull Lake domain, as defined by Thurston

et al. (1991), contains some of the largest greenstone belts inthe northwestern Superior Province, including the KneeLake-Gods Lake and Stull Lake belts of Manitoba (Syme etal., 1999; Corkery and Skulski, 1998; Corkery et al., 2005).Mainly basaltic rocks of the Hayes River assemblage have

been dated locally in the 2.83 Ga range, and volcanic rocksof the Oxford Lake assemblage fall between 2.729 and 2.719Ga (Corkery et al., 2000). The occurrence of coarse clasticsediments (Corkery et al., 2000) and alkaline volcanic rocks(Brooks et al., 1982) along some faults suggests that theywere active during deposition of the Oxford Lake assem-

blage (Corkery et al., 1992). Mafic intrusions are common insome belts such as at Big Trout Lake, and have some Ni -PGE potential. Plutons of tonalite, granodiorite and granite

underlie large parts of the terrane and yield ages between2.83 and 2.69 Ga. Isotopic data suggest that the Oxford-Stullterrane evolved in an oceanic setting, possibly on the edge of thinned North Caribou crust (Parks et al., 2005), until ca.2.71 Ga, when it was juxtaposed with the NSS along the

North Kenyon fault. It is bounded by the Gods Lake Narrows shear zone in the south.

North Caribou Terrane

The North Caribou terrane (Thurston et al., 1991) is thelargest domain with Mesoarchean ancestry of the Superior Province. Basement consists of ca. 3.0 Ga juvenile plutonicand minor volcanic belts (Stevenson, 1995; Stevenson and

Patchett, 1990; Corfu et al., 1998; Hollings et al., 1999:Henry et al., 2000), upon which were deposited early rift-related (Thurston and Chivers, 1990; 2.98-2.85 Ga) and younger (2.85-2.71 Ga) arc sequences. It was severelyreworked by continental arc magmatism at 2.75-2.70 Ga.The terrane has wide transitional margins in both the northand south.

In the north, the Munro Lake and Island Lake terranes(Thurston et al., 1991) are inferred to have been formed onthinned crust of the North Caribou terrane. These regions aredominated by plutonic rocks with several small supracrustal

belts. In the Ponask Lake area near the Ontario-Manitoba border, detrital zircons from a quartzite-komatiite sequenceindicate a depositional age <2.865 Ga, inferred to date

breakup at the northern North Caribou margin (Stone et al.,2004). The main phase of plutonism was followed by local-ized strain and shear-zone-hosted gold mineralization (ca.2.685 Ga; Lin and Corfu, 2002), particularly in the LittleStull Lake area near the Ontario-Manitoba border (Jiang and Corkery, 1998; Lin and Jiang, 2001; Lin et al., 2005).

The central North Caribou terrane is intruded by wide-spread tonalitic, dioritic, granodioritic and granitic plutonsthat crystallized between 2.745 and 2.697Ga at depths rang-ing from 18 to 10 km (0.6 to 0.3 GPa) (Stone, 1998; Stone,2000; Corfu and Stone, 1998a). Remnants of ca. 3.0 Ga

Superior Metallogeny



tonalite and supracrustal rocks are sporadically preserved through the younger magmatism (Krogh et al., 1974; Corfuand Ayres, 1991; Henry et al., 1998; Whalen et al., 2003).Within the greenstone belts, thin packages of quartz arenite-carbonate-komatiite have variably been interpreted as plat-formal cover strata (Thurston and Chivers, 1990; Thurston etal., 1991) and plume-related rift deposits (Hollings and Kerrich, 1999; Percival et al., 2002a, 2005b). Evidence of

early (>2.87 Ga) deformation is recorded in the NorthCaribou greenstone belt (Stott et al., 1989). The iron forma-tion-hosted Musselwhite lode gold deposit may have formed during development of structures associated with 2.87 Ga

pluton emplacement, or during ca. 2.7 Ga events (Fyon et al.,1992).

The Uchi subprovince preserves a ca. 300 m.y. record of tectonostratigraphic evolution along the southern margin of the North Caribou terrane (Stott and Corfu, 1991; Corfu and Stott, 1993a, b; 1996a; Hollings et al., 2000; Sanborn-Barrieet al., 2001; 2004). This region hosts some of the largestmineral deposits of the western Superior region, includingthe Red Lake gold camp. Aeromagnetic trends show thecomplex structural configuration of supracrustal rocks in a

chain of greenstone belts separated by large lobes of pluton-ic material. The stratigraphic record preserved in the RiceLake, Wallace Lake, Red Lake, Confederation Lake, Meen-Dempster, Pickle Lake and Fort Hope greenstone beltsreflects a history of rifting beginning ca. 2.99 Ga (Pirie,1982; Corfu and Wallace, 1986; Corfu and Andrews, 1987;Tomlinson et al., 1998, 2001; Sanborn-Barrie et al., 2004;Percival et al., 2002a; 2005b; Sasseville, 2002; Bailes et al.,2003; Sasseville and Tomlinson, 2005), followed by a pro-tracted history of continental arc magmatism at 2.94-2.91,2.90-2.89, 2.85 and 2.75-2.72 Ga (Stott, 1996; Beakhouse etal., 1999; Henry et al., 2000; Hollings, 2000; Hollings et al.,2000; Rogers et al., 2000; Rogers, 2002; Young, 2003;Sanborn-Barrie et al., 2004). These rocks host both world-class gold deposits (see Andrews et al., 1986; Brommecker,1991; Poulsen et al., 1996; Dubé et al., 2004; Harris et al.,2005) and massive sulphide mineralization (Nunes and Thurston, 1980; Lesher et al., 1986; Rogers, 2002). Severaldeformation episodes are recognized within the greenstone

belts, including pre-2.74, 2.73, 2.72 and 2.70 Ga events thathave produced composite, steep, east-trending fabrics(Parker, 2000; Sanborn-Barrie et al., 2001; 2004; Dubé et al.,2003; 2004; Rogers and McNicoll, 2005; Young et al., 2005;Harris et al., 2005). Multiple ages of gold mineralization areindicated, with the main stage associated with D2 structures

prior to 2.712 Ga and late-stage gold localization after 2.701Ga (Corfu and Andrews, 1987; Dubé et al., 2004). Coarseclastic sedimentary rocks generally represent the youngeststrata along the southern margin of the North Caribou terrane(Devaney, 1999a, b). Where dated, these sequences containdetrital zircons as young as 2.703 Ga, and may be faciesequivalents of the marine greywacke turbidites of theEnglish River subprovince to the south (e.g. Campbell,1971; Stott, 1996).

Over 450 km of strike length, the east-trending SydneyLake - Lake St. Joseph (SL-LSJ) fault separates rocks of the

North Caribou margin to the north from metasedimentaryschists and migmatitic rocks of the English River sub-

province to the south. The steeply dipping, 1-3 km wide,

brittle-ductile fault zone is estimated to have accommodated about 30 km of right-lateral transcurrent displacement and 2.5 km of south-side-up movement (Stone, 1981). TheMiniss River fault, which is cut and offset by the SL-LSJfault, has an age of ca. 2.68 Ga (Bethune et al., 2000; 2005).

English River Subprovince

The English River subprovince (ERS) is thought tomark the suture between the North Caribou and WinnipegRiver terranes. Distinguished from adjacent regions bysupracrustal rocks of metasedimentary origin, the ERS alsodisplays high metamorphic grade, and a prominent east-weststructural grain (Breaks, 1991). Based on the turbiditicnature of its chemically immature greywackes, the setting of the English River has traditionally been considered a fore-arc basin (Langford and Morin, 1976) or accretionary prism(Breaks, 1991). Detrital zircon studies indicate that the sedi-ments were deposited after arc activity in adjacent volcanic

belts (2.705-2.698 Ga; Corfu et al., 1995; Davis, 1996a, b;1998) and close to the time of collisional orogeny, therebyimplying an origin as a syn-orogenic flysch basin. The small

Melchett Lake greenstone belt in the central English River subprovince comprises a juvenile, ca. 2.723 Ga calc-alkalinevolcanic sequence (Corfu and Stott, 1993a; 1996a; Davis etal., 2000), possibly correlative with the Lake St. Josephassemblage to the north. Metamorphic conditions in theEnglish River range from middle amphibolite facies near themargins, to low-pressure granulite facies (750-850°C at 0.6-0.7 MPa; Perkins and Chipera, 1985; Pan et al., 1999), coin-ciding with widespread generation of migmatite and diatex-ite at 2.691 Ga (Corfu et al., 1995). The main tectonothermalevent was followed by a second thermal pulse at 2.669 Ga(op. cit.; Pan et al., 1999), intrusion of ca. 2.65 Ga peg-matites (Corfu et al., 1995), and growth of hydrothermalminerals (Pan et al., 1999).

The dominant east-west structural grain of the subprovince reflects upright to north-vergent F2 folds of an ear-lier foliation, which is defined in many areas by migmatiticlayering (Breaks, 1991; Hrabi and Cruden, 2001; 2005). Theearly foliation appears to be a composite fabric that includes

primary layering and at least one set of early folds and axial planar foliation. It is particularly well expressed in the LacSeul area (Sanborn-Barrie, 1988), where early (F1?), large-scale north-trending, west-vergent folds were delineated (Hynes, 1997).

Gravity (Nitescu et al., 2003), seismic reflection (Whiteet al., 2003) and seismic refraction (Kay et al., 1999b) pro-files collectively indicate that the Moho beneath the com-

bined English River - Winnipeg River subprovince is elevat-

ed by about 8 km over that in adjacent subprovinces. A lateto post-tectonic adjustment of this style could account for metamorphic pressure estimates up to 0.3 GPa higher withinthe English River subprovince, as well as its slow coolinghistory (Corfu, 1996; Hanes and Archibald, 2001). TheEnglish River - Winnipeg River boundary separates domi-nantly metasedimentary rocks of the English River sub-

province from mainly metaplutonic rocks of the WinnipegRiver subprovince to the south. Between the two sub-

provinces lies the metavolcanic Bird River subprovince ineastern Manitoba, with mafic intrusion-hosted Cr deposits,

John A. Percival

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and its narrow eastward extension, the Separation Lake belt(Breaks, 1991). These belts consist dominantly of largely

juvenile mafic rocks with ages of ca. 2.78 to 2.73 Ga(Timmins et al., 1985). Depositional contacts have beeninferred between English River clastic rocks and both vol-canic strata of the Separation Lake belt (Hrabi and Cruden,2005) and gneissic tonalitic basement to the east (Sanborn-Barrie, 1988). The Separation Lake belt appears to be in tec-

tonic and intrusive contact with granitic rocks of theWinnipeg River subprovince to the south (Blackburn and Young, 2000). The boundary zone is a focus for emplace-ment of ca. 2.646 Ga rare metal pegmatites (Larbi et al.,1999; Smith et al., 2004), including the Tanco and Separation Rapids fields (Blackburn and Young, 2000).

Winnipeg River Terrane

The Winnipeg River terrane is a collective term used todescribe the plutonic domain exposed north and east of thewestern Wabigoon subprovince. It consists of two main ele-ments: 1) the Winnipeg River subprovince proper (Beakhouse, 1991), a >500 km long terrane composed of

Mesoarchean metaplutonic rocks variably intruded by Neoarchean plutons; and 2) a largely Neoarchean plutonicdomain, formerly referred to as the central Wabigoon grani-toid complex (Percival et al., 2002b) and Wabigoon diapiricaxis (Edwards and Sutcliffe, 1980; Thurston and Davis,1985), that contains scattered remnants of Mesoarcheancrust (Tomlinson and Percival, 2000; Tomlinson et al., 2004;Whalen et al., 2002; 2004a). With inheritance dating back toca. 3.4 Ga (Henry et al., 2000; Tomlinson and Dickin, 2003),the Winnipeg River terrane stands apart from the NorthernSuperior and North Caribou terranes to the north and theMarmion domain to the south (described below). It also car-ries a distinct record of magmatic and structural events(Percival et al., 2004b; Melnyk et al., 2005), typically char-

acterized by amphibolite to granulite facies metamorphism(Corfu, 1988).

The Mesoarchean history of the Winnipeg River terranehas remained cryptic due to extensive overprinting

Neoarchean magmatism and deformation. Tonalitic rocksare the oldest units recognized, and include both 3.32-3.05Ga gneissic (Corfu, 1988; Davis et al., 1988; Melnyk et al.,2005) and 3.04 Ga foliated varieties (Krogh et al., 1976).Similar isotopic signatures characterize younger (2.88, 2.84and 2.83 Ga) tonalitic rocks, reflecting the antiquity of the

basement (Beakhouse and McNutt, 1991; Beakhouse et al.,1988). Mafic volcanic belts older than ca. 3.0 Ga (Davis etal., 1988) and ca. 2.93-2.88 Ga volcanic rocks in the easternSavant-Sturgeon greenstone belt (Sanborn-Barrie and

Skulski, 1999; Sanborn-Barrie et al., 2002) are also consid-ered part of the Winnipeg River terrane. Significant pulses of

Neoarchean tonalite-granodiorite magmatism occurred at2.715-2.705 Ga followed by emplacement of granites at ca.2.70-2.69 Ga (Beakhouse, 1991; Beakhouse et al., 1988;Cruden et al., 1997; 1998; Corfu, 1988, 1996). A complex

Neoarchean structural-metamorphic history began with dep-osition of metasedimentary rocks after 2.72 Ga (Melnyk etal., 2005). The supracrustal rocks and older gneisses werefolded (D3) between 2.717 and 2.713 Ga, prior to syntecton-

ic injection of 2.713-2.707 Ga tonalite and granodiorite

sheets accompanying D4 horizontal extensional deformation

(Melnyk et al., op. cit.). Upright folding during D5 deforma-

tion took place after 2.705 Ga, and younger upright F6 folds

indicate a period of north-south compression associated withemplacement of 2.695-2.685 Ga granite and granodiorite(op. cit.). Late pegmatites and granites intruded during adextral transpressive (D7) regime (op. cit.).

The eastern Winnipeg River terrane hosts east-trendinggreenstone belts including the Caribou Lake, Obonga,Garden Lake and Heaven Lake belts that have ages >3075 to2703 Ma (Davis et al., 1988; Tomlinson et al., 2002; 2003).Dated granitoid units have ages in the range 3075-2680 Ma(Davis et al., 1988; Whalen et al., 2002) and some of the old-est rocks have Nd values of -1 to + 1, suggesting derivationfrom even older crustal sources (Tomlinson et al., 2004). Atleast five generations of Neoarchean structures (D1-D5) have

been recognized in complex tonalitic gneisses (Schwerdtner,1992; Brown, 2002; Percival et al., 2004a).

The Marmion domain, formerly included as part of thesouth-central Wabigoon subprovince, is now recognized toconsist of 3.01-2.999 Ga tonalite (Davis and Jackson, 1988;

Tomlinson et al., 2004), upon which the Steep Rock, knownfor its economic iron deposits, Finlayson Lake and LumbyLake greenstone belts formed between 2.99 and 2.78 Ga(Stone et al., 2002; Tomlinson et al., 2003).

A period of continental arc magmatism in the WinnipegRiver subprovince (2.72-2.70 Ga; Corfu, 1988; 1996;Whalen et al., 2002; Melnyk et al., 2005) is attributed tonorth- and eastward subduction of oceanic rocks (Sanborn-Barrie and Skulski, 2005) followed by 2.708-2.701 Ga D1

deformation. Post-2.704 Ga regional deformation (D2)

across the Wabigoon outlasted deposition of syncollisionalcoarse clastic sedimentary overlap sequences (i.e. Crowduck and Ament Bay assemblages; Fralick 1997; Ayer and Davis1997; Sanborn-Barrie and Skulski, 2005). Late faults with

both strike-slip and dip-slip motion define the present subprovince boundary (Gower and Clifford, 1981).

Wabigoon Subprovince

The Wabigoon subprovince has long been recognized asa composite terrane comprising volcanic-dominated domains with a central axis of variable-age plutonic rocks(Davis and Jackson, 1988; Percival et al., 2002b; Percivaland Helmstaedt, 2004). It consists of distinct western and eastern segments.

Western Wabigoon Subprovince

The western Wabigoon subprovince is dominated bymafic volcanic rocks with large tonalitic plutons (Blackburnet al., 1991). Volcanic rocks range in composition fromtholeiitic to calc-alkaline, and are interpreted to representocean floor or plateau and arc environments, respectively(Ayer and Davis, 1997; Ayer, 1998a; Ayer and Dostal, 2000;Wyman et al., 2000). Most of the preserved volcanic rockswere deposited between ca. 2.745 and 2.72 Ga (Corfu and Davis, 1992), with rare older rocks, such as the 2.775 Ga(Davis et al., 1988) Fourbay assemblage of oceanic plateauaffinity (Sanborn-Barrie and Skulski, 1999) and minor younger (2.713-2.70 Ga) volcanic-sedimentary sequences.




Plutonic rocks range from broadly synvolcanic batholithscomposed of tonalite-diorite-gabbro (ca. 2.735-2.72 Ga;Davis and Edwards, 1982; Corfu and Davis, 1992; Whalenet al., 2004a), to younger granodiorite batholiths and plutons(ca. 2.710 Ga; Davis and Edwards, 1986; Sanborn-Barrie,1988; Davis and Smith, 1991; Melnyk et al., 2005), monzo-diorite plutons of sanukitoid affinity (ca. 2.698-2.690 Ga;Stern and Hanson, 1991; Ayer, 1998b; Stevenson et al.,

1999), and plutons and batholiths of monzogranite (2690-2660 Ma; Schwerdtner et al., 1979; Sanborn-Barrie, 1988;Melnyk et al., 2000). Immature clastic metasedimentarysequences are preserved in narrow belts within volcanicsequences. They are commonly younger than the volcanicrocks, as illustrated by local unconformable relationships(Fralick, 1997) and geochronological constraints indicatingdeposition between ca. 2.711 and <2.702 Ga (Davis, 1996a;1996b; 1998; Fralick and Davis, 1999). Virtually all carryancient (>3 Ga) detrital zircons indicating old components insource regions. At least two phases of deformation affected supracrustal rocks of the western Wabigoon subprovince(Blackburn et al., 1991), with apparent diachroneity in theonset of deformation from ca. 2.709 Ga in the Lake of the

Woods area (Davis and Smith, 1991; Ayer and Davis, 1997;Melnyk et al., 2005), to ca. 2.700 Ga in the Sioux Lookout -Savant area in the east (Sanborn-Barrie et al., 1998; 2002;Sanborn-Barrie and Skulski, 2005). These events involved atleast local tectonic inversion, through thrust imbrication(Davis et al., 1988) and possible formation of nappe-likestructures (Poulsen et al., 1980).

The Sturgeon - Savant greenstone belt, which hosted theSturgeon Lake massive sulphide camp (e.g. Franklin et al.,1975; Galley et al., 2000), consists of several tectonostrati-graphic packages, including the previously described Juttenassemblage (Sanborn-Barrie et al., 2002), the ca. 2775 MaFourbay assemblage, and 2745-2735 Ma sequence, theHandy Lake and South Sturgeon assemblages (Davis et al.,1985; Sanborn-Barrie and Skulski, 1999; Sanborn-Barrie etal., 2002). The 2735 Ma Lewis Lake batholith (Whalen et al.,2004b) may have provided the heat source for seawater con-vection and massive sulphide mineralization (Galley et al.,2000). Younger (ca. 2718 Ma; Davis et al., 1988), high Fe, Ti

basalt and minor dacite of the central Sturgeon assemblagerepresent a rifted arc sequence. Associated sedimentaryrocks contain both arc (2745-2730 Ma) and continental (3.1-2.8 Ga) detritus (Skulski et al., 1998a). Two younger sedi-mentary sequences complete the stratigraphic record: 1)greywacke - iron formation (ca. 2.705 Ga) of the Warclubassemblage; and 2) sandstone and arkose (<2.698 Ga) of thesyn-orogenic Ament Bay assemblage (Sanborn-Barrie et al.,2002). Two sets of ductile structures postdate <2.704 Garocks: 1) north-trending upright F1 folds; and 2) east-trend-

ing upright D2 folds and penetrative foliation. Pre-D1 folds

have been inferred locally (Sanborn-Barrie et al., 1998).

Eastern Wabigoon Subprovince

The eastern Wabigoon subprovince is a composite ter-rane with greenstone belts and intervening granitoid plutonsthat show variable Mesoarchean (Winnipeg River and Marmion) and oceanic affinity (Stott and Davis, 1999;Tomlinson et al., 2000; Stott et al., 2002). In the northwest

part of the belt the 3.0-2.92 Ga Toronto and Tashota assemblages may represent a continental margin sequence built onthe Winnipeg River terrane. Calc-alkaline rocks of the 2.74Ga Marshall assemblage have small massive sulphidedeposits (Stott et al., 2002). The central part of the belt isdominated by rocks of oceanic affinity including tholeiitic

juvenile pillowed basalt of the 2.78-2.738 Ga Onaman and Willet assemblages and the overlying calc-alkaline 2.725-

2.715 Ga Metcalfe-Venus assemblage (Stott et al., 2002).Parts of these assemblages contain widespread hydrothermalalteration and host small massive sulphide deposits (op. cit.).Across the southern part of the eastern Wabigoon domain,the 2.78-2.74 Ga calc-alkaline Elmhirst-Rickaby assemblageis possibly built on Marmion-age substrate (Tomlinson et al.,2004). Unconformably overlying clastic rocks (Albert-Gledhill and Conglomerate assemblages) were deposited after ca. 2.71 Ga. At least two deformation events affected the eastern Wabigoon domain: east-striking D1 structures

(<2.706 Ga) and east-striking, dextral transpressive D2 shear

zones (Stott et al., 2002). The 2.694 Ga Deeds Lake pluton provides a lower limit on the age of D2 deformation (Stott

and Davis, 1999).The Quetico - western Wabigoon boundary is welldefined as the Seine River - Rainy Lake fault, in an areaknown for numerous gold showings and small deposits(Poulsen, 1985). East of Lake Nipigon the boundary with theeastern Wabigoon is an imbricate zone with an early historyof structural telescoping (Devaney and Williams, 1989;Tomlinson et al., 1996) and the significant, structurally host-ed Long Lac vein gold deposits (Lafrance et al., 2004). TheWabigoon-Quetico interface is also marked sporadically by<2.692 Ga coarse clastic rocks of the Seine assemblage(Fralick and Davis, 1999) that were deposited in transten-sional basins (Blackburn et al., 1991) or delta fan environ-ments (e.g. Fralick et al., 2005).

Quetico Subprovince

The Quetico subprovince consists dominantly of greywacke, derived migmatite and granite. No stratigraphicsequence has been established within the steeply dipping,

polydeformed and variably metamorphosed sedimentarysuccession; however, younging directions are dominantly tothe north (Percival, 1989). Depositional age constraints indi-cate slightly older ages for the northern Quetico(<2.698>2.696 Ga; Davis et al., 1990) than for the south(<2.692 Ga; Zaleski et al., 1999).

Several plutonic suites cut metasedimentary units,including early (2.696 Ga) tonalite (Davis, 1996b). An early

(D1) deformation event predated emplacement of a chain of Alaskan type mafic-ultramafic intrusions in the northernQuetico, which have some Ni-PGE potential (e.g. Pettigrew,2004). They are associated with alkaline plutons includingnepheline syenite and carbonatite with ages in the range2.69-2.68 Ga (Lassen, 2004) and geochemical affinities withthe Archean sanukitoid suite (cf. Stern et al., 1989;Stevenson et al., 1999; Lassen, 2004). Two subsequentdeformation events (D2, D3) were followed by low-pressure,

high-temperature metamorphism that reached upper amphibolite and local granulite facies at ca. 2.67-2.65 Ga (Pan etal., 1994; 1998) in the central region and greenschist facies

John A. Percival

6



at the margins (Percival, 1989). Coeval, crust-derived granitic plutons and pegmatites, including ca. 2.67 Ga pera-luminous granite and ca. 2.65 Ga biotite granite (e.g.Southwick, 1991) have sporadic rare-element mineralization(Breaks et al., 2003).

Tectonic models for the Quetico subprovince havefavoured forearc settings (e.g. Langford and Morin, 1976;Percival and Williams, 1989; Williams, 1991; Fralick et al.,

2005). Depositional ages of ca. 2.698 to 2.690 Ga overlapthose of late arc magmatism in the Wabigoon. The domi-nantly sanukitoid plutons of this age may have been trig-gered by slab breakoff (cf. Sajona et al., 2000).

The southern Quetico boundary separates metasedimen-tary rocks from the dominantly volcano-plutonic Wawa-Abitibi subprovince to the south. Stratigraphic linkages

between subprovinces are evident in some areas (Fralick etal., 2005; Zaleski et al., 1999), although at ca. 2.685 Ga dex-tral transpressive shear zones are common (Corfu and Stott,1996b).

Wawa Subprovince

Most workers accept a correlation between the Wawaand Abitibi subprovinces across the transverse Kapuskasinguplift. Within the Wawa subprovince, volcanism appears tohave begun with the 2.89-2.88 Ga Hawk assemblage. The2.775 Ga Hemlo-Black River, 2.745 Ga Wawa and 2.72 GaGreenwater and Manitouwadge assemblages indicate anoceanic setting, the latter contains significant massive sul-

phide mineralization (Corfu and Stott, 1986; Sage et al.,1996a, b; Williams et al., 1991). Polat et al. (1999) reported a variety of oceanic magma types from the Schreiber belt,and interpreted the belt as a tectonic mélange (Polat et al.,1998; Polat and Kerrich, 1999; 2001).

Relatively late-stage volcanism at ca. 2.695 Ga took place during D1 thrusting. Subsequent calc-alkalic to alkalic

magmatism (ca. 2.689 Ga; Corfu and Stott, 1998b) and asso-ciated coarse clastic sedimentation (Timiskaming type;<2.689 Ga) was followed by emplacement of sanukitoid plu-tons (2.65-2.68 Ga) and dextral transpressive D2 deforma-

tion.Mineralization occurs in two main regions: the

Michipicoten-Mishubishu belt in the Wawa area, and theShebandowan-Schreiber belt to the west. The Michipicoten-Mishubishu belt consists mainly of Fe and Au deposits withsome Ni and vein Cu deposits. Iron deposits are in oxide-,sulphide- and carbonate-facies iron formations within 2.74-2.73 Ga volcanic sequences. Gold deposits in this regionoccur in veins associated with shear zones within plutonicrocks of variable composition and age. The Shebandowan-Schreiber mineral belt hosts important gold, iron, volcanic-hosted massive sulphide (e.g. Manitouwadge; Zaleski et al.,1999; Peterson and Zaleski, 1999) and intrusion-hosted Nideposits. The most significant is the Hemlo gold camp, alarge disseminated deposit (Muir, 2003) in a stronglydeformed, ca. 2.693-2.685 Ga volcano-sedimentarysequence of probable Timiskaming affinity (Davis and Lin,2003). Gold was deposited during D2 sinistral wrench defor-

mation between 2.680 and 2.677 Ga (op. cit.), likely fromfluids derived from granitoid rocks.

The Great Lakes tectonic zone is the unexposed

boundary between the Minnesota River Valley terrane and Wawa subprovince, identified from aeromagnetic images(Sims and Day, 1993). It is inferred to dip northward based on the presence of isotopic inheritance in plutons of theVermilion district of the southern Wawa-Abitibi subprovince(Sims et al., 1997).

Kapuskasing Uplift The Kapuskasing uplift represents a 500-km long fault-

bounded structure that divides the Superior Province intoeastern and western halves (Percival and West, 1994).Amphibolite- to granulite-facies tonalite gneiss, paragneiss,mafic gneiss and anorthosite (2.765-2.66 Ga) represent mid to lower crustal levels of the Abitibi-Wawa and Quetico

belts, exposed through east-directed thrusting and sinistraltranscurrent motion during the Paleoproterozoic (Percivaland McGrath, 1986; Percival and Peterman, 1994). Krogh(1993) noted the similarity between 2.66-2.62 Ga ages of high-grade metamorphism of Kapuskasing rocks and thetiming of lode gold deposition in the Abitibi belt (e.g. Zwenget al., 1993; Davis et al., 1994), and proposed a genetic link-

age. Brittle faults are cut by 1.89 and 1.1 Ga alkalic complexes, which have Nb, REE and phosphate showings and prospects (Sage, 1991). Recent work suggests that theKapuskasing structure accommodated a 20° counterclock-wise rotation of the western Superior with respect to the east(Halls, 2004; Halls and Davis, 2004).

Eastern Superior Province

Abitibi Subprovince

The Abitibi subprovince hosts some of the richest min-eral deposits of the Superior Province, including the giant

Kidd Creek massive sulphide deposit (Hannington et al.,1999) and the large gold camps of Ontario and Quebec(Robert and Poulsen, 1997). Views of the tectono-strati-graphic evolution of the Abitibi subprovince have changed markedly from the allochthonous terrane concept introduced in the early 1990s (cf. Jackson and Fyon, 1991; Jackson etal., 1994; Desrochers et al., 1993), to a more traditionalautochthonous stratigraphic framework supported bydetailed geochronology (e.g. Heather, 1998; Ayer et al.,2002; Mueller and Mortensen, 2002). Stratigraphic com-

plexities are explained in terms of evolution of oceanic geo-dynamic settings from plateau, to arc and rift environments(e.g. Thurston, 1994; Bédard and Ludden, 1997; Kerrich etal., 1999; Wyman et al., 1999; 2002).

The Abitibi subprovince has been subdivided into threedomains with overlapping tectonostratigraphic histories. Inthe northern Abitibi, volcanic assemblages are mainly 2.735-2.72 Ga (Ludden et al., 1986; Chown et al., 1992; Legault etal., 2002) and associated with layered intrusions, whereasvolcanic rocks of 2.71-2.695 Ga and their associated miner-al deposits are restricted to the southern Abitibi (Dimroth etal., 1984; Daigneault et al., 2002). In addition, the southernAbitibi has relatively young sedimentary-volcanic depositsincluding ca. 2.69 Ga greywackes of the Porcupine Group(Bleeker and Parrish, 1996) and 2.677-2.673 Ga conglomer-




atic and alkaline volcanic rocks of the Timiskaming Group(Davis 2002). The central zone is dominated by plutonicrocks (Chown et al., 2002).

The northern Abitibi belt is best known for theMatagami-Chibougamau mineral belt, characterized mainly

by Cu-Zn massive sulphide deposits, Cu-Zn vein depositsand some lode gold deposits. Volcanic strata hosting the sul-

phide mineralization range in age from ca. 2.73 to 2.715 Ga.

Gold occurs in veins within shear zones and iron formations,or as disseminated mineralization associated with felsicintrusions (Card and Poulsen, 1998). In the central zone,mineralization is restricted to the Normétal belt, and includesmassive sulphides in 2.728 Ga volcanic rocks (Mortensen,1993), as well as vein gold deposits. To the south, the south-ern Abitibi belt hosts the Timmins-Val d'Or mineral belt,known for its numerous gold deposits, major Cu-Zn massivesulphide deposits (Hannington et al., 1999; Wyman et al.,1999; 2002), komatiite and intrusion-related Ni deposits,some pegmatite-hosted deposits and minor porphyry typemineralization (Card and Poulsen, 1998; Jébrak and Doucet,2002). The richly gold-mineralized Cadillac-Larder Lake

break, which forms part of the southern boundary of the belt,

is considered to represent a south-verging thrust carrying theAbitibi rocks over the Pontiac (Dimroth et al., 1984; Fengand Kerrich, 1991; 1992; Calvert et al., 1995; Calvert and Ludden, 1999; Ludden and Hynes, 2000; Davis, 2002).

Pontiac Subprovince

Metasedimentary schist and paragneiss derived fromturbiditic greywacke and minor conglomerate dominate thenorthern Pontiac subprovince. Detrital zircons indicate depo-sitional ages <2.685 Ga (Mortensen and Card, 1993; Davis,2002), and tonalite, granodiorite and granite plutons range inage from 2.682 to 2.66 Ga. The southern part of the beltincludes the Baby-Belleterre greenstone belt, including vol-

canic rocks as young as 2.682 Ga (Mortensen and Card,1993). The Pontiac belt has been interpreted as a south verg-ing fold-thrust belt that was over-ridden by the southernAbitibi belt (Benn et al., 1994; Calvert and Ludden, 1999;Davis, 2002).

The Benny-Belleterre mineral belt contains the quartz-vein hosted Belleterre gold deposit, as well as syn-volcanic

Ni-Cu sulphides (with some PGEs) in gabbroic sills and small plugs of dunite-gabbro.

Opatica Belt

This predominantly metaplutonic belt bounding theAbitibi belt to the north consists of units ranging from 2.82

Ga tonalite, through 2.77-2.70 tonalite-granodiorite, to 2.68Ga granite and pegmatite (Benn et al., 1992; Sawyer and Benn, 1993; Davis et al., 1994). Polyphase deformationincludes early, west-verging shear zones (<2.72 Ga), over-

printed by 2.69-2.68 Ga south-vergent structures (Sawyer and Benn, 1993). The Frotet-Evans-Troilus greenstone belthas showings including Cu-Au veins, volcanogenic massivesulphide, iron formation and intrusion-hosted Ni-Cu(Gosselin 1996).

Opinaca Belt

Metagreywacke, derived migmatite and granite charac-terize the Opinaca belt. Polydeformed schists occur at the

belt margins, whereas the interior portions are metamorphosed to amphibolite and granulite facies. These rocks arecut by the 2.67 Ga Broadback River granite (Davis et al.,1994). Mineralization is scant in the Opinaca belt, which

based on its similarity to the Quetico and English River subprovinces, could have some potential for rare metal peg-matites.

Ashuanipi Complex

The ca. 300 x 300 km Ashuanipi complex consists of high-grade metamorphic and plutonic rocks (Percival et al.,1992). Early sedimentary and volcanic rocks were intruded

by a ca. 2.725 Ga tonalite-diorite suite (Percival et al., 2003),then deformed and metamorphosed to granulite facies and subsequently intruded by orthopyroxene-bearing diatexite(Percival, 1991), granite, granodiorite and syenite plutons(2.65-2.62 Ga) and by A-type granites (2.57 Ga)(Leclair etal., 2004a).

Mineralization is restricted to small, high-grade gold showings in iron formation and skarns (Lapointe and Chown, 1993; Moritz and Chevé, 1992; Chevé and Brouillette, 1995). The occurrences are unusual in their set-ting in a granulite-facies terrane.

La Grande Subprovince

The La Grande belt consists of distinct sectors with vari-able tectonic setting. In the west, Mesoarchean basement(3.33-2.79 Ga) is unconformably overlain by the clasticApple formation, which hosts uranium-gold occurrences(Roscoe and Donaldson, 1988), and 2.75-2.73 Ga volcanicstrata (Goutier and Dion, 2004). Older strata are present inthe Guyer-LG4 sector, including komatiites and related ca.2.82 Ga sills, which contain Cu and massive sulphide miner-alization. Juvenile volcanic rocks (2.75-2.70 Ga) of theEastmain sector are characterized by porphyry and other magmatic mineralization.

Bienville Subprovince

Plutonic rocks of the Bienville subprovince intrude thenorthern margin of the La Grande belt and mark a transitionto dominantly granitic rocks to the north. The 200-km widedomain is underlain by foliated, gneissic and massive gran-odiorite to granite with crystallization ages in the 2.73-2.68Ga range (Skulski et al., 1998b; Percival et al., 2001).

Xenocrystic zircons and Nd model ages indicate sources asold as 3.3 Ga.

Northeastern Superior Province (Minto Block)

This remote, far-north region of Quebec was under-explored prior to the early 1990s. Following the work of Percival et al. (1992, 1994, 2001), large parts were mapped for the first time at 1:250 000 (see synthesis in Leclair et al.,2004a). Metallogenic studies were carried out in parallelwith regional mapping, and the major mineralization typesidentified (Labbé and Lacoste, 2004).

John A. Percival

8



Domains of the Minto block were defined on the basisof lithology, structural and aeromagnetic trends and intensi-ty (Percival et al., 1992; 2001; Stern et al., 1994) and refined

by Leclair et al. (2004a) in light of regional map and isotopicconstraints. The Inukjuak domain in the west (Domain I of Leclair et al., 2004a) hosts the ancient (ca. 3.825 Ga)

Nuvvuagittuq supracrustal belt (David et al., 2003) and ca.3.65 Ga tonalitic gneisses but is dominated by younger plu-

tonic rocks (2.84-2.69 Ga). To the east, the Tikkerutuk domain (II) consists of pyroxene-bearing plutonic rocks,3.02-2.71 Ga, that merge into the Bienville domain (III) tothe south, which is characterized by 2.71-2.69 Ga graniteand granodiorite. Farther east, Lake Minto domain (IV) con-tains metasedimentary and derived migmatitic rocks, inaddition to 2.76-2.70 Ga pyroxene-bearing plutonic rocks(Percival and Mortensen, 2002). The Goudalie domain (V)forms a central spine consisting of relatively large volcano-sedimentary belts (2.88-2.71 Ga) with sparse 3.0 Ga tonaliteand abundant pyroxene-bearing plutons (2.73-2.68 Ga). Inthe east-central Minto, the Utsalik domain (VI) consists of highly magnetic (Pilkington and Percival, 1999), 2.74-2.69Ga pyroxene-bearing plutonic rocks with model ages in the

3.0 Ga range (Leclair et al., 2004a). To the northeast,Douglas Harbour domain (VII) consists of two largecharnockitic massifs (2.74-2.73 Ga) within an older (2.88-2.75 Ga) tonalitic complex (Leclair et al., 2004a; Bédard,2003; Bédard et al., 2003). Percival and Skulski (2000)attributed the ca. 2.70 Ga deformation and high-grade meta-morphism in the western Minto block to collisional process-es whereas Bédard et al., (2003) inferred magmatic and diapiric controls.

The northeastern Superior hosts a variety of mineralshowings and small deposits that have been classified intosyngenetic and epigenetic types (Labbé and Lacoste, 2004).Syngenetic mineralization, occurring mostly withinsupracrustal belts, includes Algoma-type iron formation,volcanogenic massive sulphide, magmatic Ni-Cu, maficintrusion-hosted Fe-Ti-V, and porphyry U-Th and Mo (op.cit.). Supracrustal rocks also host most epigenetic types,including disseminated polymetallic sulphide occurrences(Cu-Zn-Ag-Au), gold in iron formation, vein quartz, and shear zones, sulphide Cu-Ni-Ag-Au in veins, rare earth ele-ments in carbonated rocks, and minor U in veins.

Discussion

Mineralization events are plotted with respect to strati-graphic age in Figure 3. It is evident that volcanogenic mas-sive sulphide mineralization is tightly bracketed between ca.

2.74 and 2.70 Ga, whereas gold deposits have a wider rangeof ages, between 2.71 and 2.61 Ga. The older VMS depositsoccur in the Uchi and Wabigoon belts, and the younger onesin the Abitibi-Wawa subprovince in the south, reflecting thenorth-to-south pattern of progressive accretion of oceanicand microcontinental fragments (Percival et al., 2004a).Examination of the setting of these VMS deposits has gener-ally shown rifted arc or back-arc environments (e.g.Hannington et al., 1999). Causes for the Neoarchean massivesulphide "bonanza" are not immediately evident. Barley etal. (1998) proposed a Neoarchean global event (cf. Condie,

1998) in which plumes produced enriched oceanic crust that,when recycled in subduction zones, generated metal-richderivative magmas. Similar circumstances of rifted oceanicarc settings surround the Paleoproterozoic VMS "bonanza"(Franklin et al., 1998). Formation of both the Neoarchean(2.73-2.70 Ga) and Paleoproterozoic (1.88-1.86 Ga) VMSdeposits preceded collisional tectonism by 10-20 m.y., and itis plausible that impending collisions forced changes in plate

stresses (subduction zone shallowing or roll-back), inducingextension within active arc systems that permitted access of deep-seated, metal-rich magmas.

Gold deposits generally formed during or within tens of millions of years after major tectonism (Robert and Poulsen,1997; Dubé et al., 2004). The timing of mineralization may

be polyphase, and coincides with late metamorphic and hydrothermal effects observed throughout the Superior Province (e.g. Krogh 1993; Davis et al., 1994; Moser et al.,1996; Percival and Skulski, 2000). Percival and Pysklywec(2004) suggested that the late hydrothermal overprint in theSuperior Province is related to inversion of mantle litho-sphere cells, whose main effect was stabilization of the cra-tonic lithosphere ("cratonization").

Paleoproterozoic and Mesoproterozoic Mafic Dykes andRelated Rocks

Numerous Proterozoic dykes swarms transect theSuperior Province, particularly in the south (Table 1; Buchanand Ernst, 2004). The Matachewan swarm is volumetricallysignificant, making up almost 5% of bedrock area in north-eastern Ontario. Some Ni-Cu mineralization is present inassociated mafic sills within lower Huronian strata (e.g.James et al., 2002). Significant mineralization is associated with the Paleoproterozoic Nipissing magmatic event, in formof vein silver deposits in sills of the Cobalt plate (Marshall

and Watkinson 2000). Mineralization related to theKeweenawan event includes Cu, Ag in basalts and Ni-PGEshowings in Logan sills in the Thunder Bay area. Alkalineintrusions of similar age host Ni-PGE mineralization, and carbonatites host numerous Nb-REE prospects (Sage, 1991).

Sudbury Structure

Situated on the southern edge of the Superior Province,the 1.85 Ga Sudbury Intrusive Complex represents one of the most richly mineralized bodies of the Canadian Shield.Current thinking regards the gabbro-noritic intrusion to have

been generated in response to meteorite impact (Ames,

1999; Rousell et al., 2002; 2003; Therriault et al., 2002).Large Ni, PGE deposits, distributed around the perimeter of the intrusion, are localized primarily in the sublayer at the

base of the differentiated sill (Therriault et al., 2002; Naldrett, 2003).

Knowledge gaps in Superior Province

Most of the Superior Province has been mapped atscales of 1:250,000 or better. This regional information dates




from the 1960s in parts of northern Ontario and Manitoba,and is very recent in northern Quebec. Mapping in regionswith high mineral potential has been updated periodically.Remaining knowledge gaps are primarily the result of cover

by Phanerozoic strata, glacial deposits and water. In northernOntario, recent interpretations using aeromagnetic data and drill core information have extended Superior Province units

beneath the James Bay lowlands (Stott and Berdusco, 2000).The James Bay and southern Hudson Bay region remains

poorly understood as a result of cover and lack of aeromag-netic coverage. Improving our understanding of this area hasseveral purposes: 1) resolving tectonic questions concerningthe "big bend" in Superior Province structural trends, fromeast-west in the west and south, to north-south in the north-east (Fig. 2); 2) assessing the distribution of ancient rocks of the Northern Superior superterrane, with applications for diamond exploration; and 3) tracing northwest-trending,gold-mineralized faults and shear zones between northwest-ern Ontario and north-central Quebec.

References

Ames, D.E., 1999, Geology and regional hydrothermal alteration of thecrater-fill, Onaping Formation; association with Zn-Pb-Cu mineralization,Sudbury Structure, Canada. Unpublished Ph.D. thesis: Carleton University,Ottawa, ON, Canada, 638p.

Andrews, A.J., Hugon, H., Durocher, W.E., Corfu, F., and Lavigne, M.J.,1986, The anatomy of a gold-bearing greenstone belt; Red Lake, north-western Ontario, Canada. in Macdonald, A.J., ed., Gold '86; AnInternational Symposium on the Geology of Gold Deposits; ProceedingsVolume, p.3-22.

Ayer, J.A., 1998a, Petrogenesis and tectonic evolution of the Lake of theWoods greenstone belt, western Wabigoon Subprovince, Ontario, Canada.Unpubl. Ph.D. thesis: University of Ottawa, Ottawa ON, 213 p.

Ayer, J.A., 1998b, The mafic minerals of the Falcon Island ultrapotassic pluton, Lake of the Woods, Ontario; progressive reduction during fraction-ation: Canadian Mineralogist 36, p.49-66.

Ayer, J.A., Davis, D.W., 1997, Neoarchean evolution of differing con-vergent margin assemblages in the Wabigoon Subprovince: geochemicaland geochronological evidence from the Lake of the Woods greenstone belt,Superior Province, northwestern Ontario: Precambrian Research 81, p.155-178.

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Fig. 3: Time-stratigraphic correlation chart showing major tectonic elements of the Superior Province and ages of important mineral deposit types. Age datafrom Skulski and Villeneuve (1999) and more recent sources (see text for details).



Ayer, J.A., and Dostal, J., 2000, Nd and Pb isotopes from the Lake of theWoods greenstone belt, northwestern Ontario; implications for mantle evo-lution and the formation of crust in the southern Superior Province:Canadian Journal of Earth Sciences 37, p.1677-1689.

Ayer, J., Amelin, Y., Corfu, F., Kamo, S., Ketchum, J., Kwok, K. and Trowell, N., 2002, Evolution of the southern Abitibi greenstone belt based on U-Pb geochronology: autochthonous volcanic construction followed by

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Bailes, A., Percival, J.A., Corkery, M.T., McNicoll, V.J., Tomlinson,K.Y., Rogers, N., Sasseville, C., Whalen, J.B., and Stone, D., 2003, Geologyand tectonostratigraphic assemblages, West Uchi area, Geological Survey of Canada Open File 1522, scale 1:250 000.

Barley, M.E., Krapez, B., Groves, D.I., and Kerrich, R., 1998, The lateArchaean bonanza; metallogenic and environmental consequences of theinteraction between mantle plumes, lithospheric tectonics and global cyclic-ity; in Percival, J.A., and Ludden, J.N., eds., Earth's evolution through thePrecambrian: Precambrian Research 91, p.65-90.

Beakhouse, G.P., 1991, Winnipeg River subprovince; in Thurston, P.C.,Williams, H.R., Sutcliffe, R.H., and Stott, G.M., eds., Geology of Ontario:Ontario Geological Survey Special Volume 4, Part 1, p.279-301.

Beakhouse, G.P., McNutt, R.H., 1991, Contrasting types of late Archean plutonic rocks in northwestern Ontario; implications for crustal evolution inthe Superior Province: Precambrian Research 49, p.141-165.

Beakhouse, G.P., McNutt, R.H., Krogh, T.E., 1988, Comparative Rb-Sr

and U-Pb geochronology of late to post tectonic plutons in the WinnipegRiver belt, northwestern Ontario, Canada: Chemical Geology 72, p.283-291.

Beakhouse, G.P., Heaman, L.M., and Creaser, R.A., 1999, Geochemicaland U-Pb zircon geochronological constraints on the development of a LateArchean greenstone belt at Birch Lake, Superior Province, Canada:Precambrian Research 97, p. 77-97.

Bédard, J.H., 2003, Evidence for regional-scale, pluton-driven, high-grade metamorphism in the Archaean Minto Block, northern Superior Province, Canada: Journal of Geology 111, p.183-205.

Bédard, J.H., Brouillette, P., Madore, L., and Berclaz, A., 2003, Archaeancratonization and deformation in the northern Superior Province, Canada;an evaluation of plate tectonic versus vertical tectonic models: PrecambrianResearch 127, p.61-87.

Bédard, L.P., and Ludden, J.N., 1997, Nd-isotope evolution of Archean plutonic rocks in the Opatica, Abitibi and Pontiac subprovinces, Quebec,Canada: Canadian Journal of Earth Sciences 34, p.286-299.

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John A. Percival

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Geology and Metallogeny of the Superior Province

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