Tectonic evolution of the western Superior Province from...

Can. J. Earth Sci. 43: 1085–1117 (2006) doi:10.1139/E06-062 © 2006 NRC Canada

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Tectonic evolution of the western Superior Provincefrom NATMAP and Lithoprobe studies1,2

J.A. Percival, M. Sanborn-Barrie, T. Skulski, G.M. Stott, H. Helmstaedt, and D.J. White

Abstract: Five discrete accretionary events assembled fragments of continental and oceanic crust into a coherent Superiorcraton by 2.60 Ga. They exhibit similar sequences of events at �10 million year intervals: cessation of arc magmatism,early deformation, synorogenic sedimentation, sanukitoid magmatism, bulk shortening, regional metamorphism, latetranspression, orogenic gold localization, emplacement of crust-derived granites, and postorogenic cooling. The NorthernSuperior superterrane recorded 3.7–2.75 Ga events prior to 2.72 Ga collision with the 3.0 Ga North Caribou superterrane.Following 2.98 Ga rifting, the Uchi margin of the North Caribou superterrane evolved in an upper plate setting before2.72–2.70 Ga collision of the <3.4 Ga Winnipeg River terrane, which trapped synorogenic English River turbidites in thecollision zone. The Winnipeg River terrane was reworked in 2.75–2.68 Ga magmatic and tectonic events, including thecentral Superior orogeny (2.71–2.70 Ga) that marks accretion of the juvenile western Wabigoon terrane. In the south, theWawa–Abitibi terrane evolved in a mainly oceanic setting until Shebandowanian collision with the composite Superiorsuperterrane at 2.695 Ga. Synorogenic Quetico turbidites were trapped in the collision zone. The final accretionary eventinvolved addition of the Minnesota River Valley terrane (MRVT) from the south, and deposition and metamorphism ofsynorogenic turbidites of the Pontiac terrane during the �2.68 Ga Minnesotan orogeny. Seismic reflection and refractionimages indicate north-dipping structures, interpreted as a stack of discrete 10–15 km thick terranes. A slab of high-velocitymaterial, possibly representing subcreted oceanic lithosphere, as well as Moho offsets, support a model of progressive ac-cretion through plate-tectonic-like processes.

Résumé : Il y a 2,60 Ga, cinq événements accrétionnaires distincts ont rassemblé des fragments de croûte continentaleet océanique; le craton cohérent du lac Supérieur est le résultat de cette accrétion. Ces fragments montrent des séquencesd’événements semblables, à des intervalles d’environ 10 Ma : la cessation du magmatisme d’arc, une déformation précoce,une sédimentation synorogénique, un magmatisme sanukitoïde, un rétrécissement en vrac, un métamorphisme régional, unetranspression tardive, la localisation de l’or orogénique, la mise en place de granites dérivés de la croûte et un refroi-dissement post-orogénique. Le superterrane du lac Supérieur septentrional a enregistré des événements de 3,7–2,75 Gaavant la collision à 2,72 Ga avec le superterrane de North Caribou de 3,0 Ga. À la suite de la distension il y a 2,98 Ga,la bordure de la sous-province d’Uchi du superterrane de North Caribou a évolué en un environnement de plaque supé-rieure avant la collision, �2,72–2,70 Ga, avec le terrane de Winnipeg River, <3,4 Ga, ce qui a piégé les turbidites sy-norogéniques d’English River dans la zone de collision. Vers 2,75–2,68 Ga, le terrane de Winnipeg River a été remainélors d’événements magmatiques et tectoniques, incluant l’orogène du centre de la Province du lac Supérieur (2,71–2,70 Ga)qui marque l’accrétion du terrane juvénile Wabigoon occidental. Vers le sud, le terrane de Wawa–Abitibi a évolué dansun environnement surtout océanique jusqu’à sa collision au Shebandowanien avec le superterrane composite du lac Su-périeur à 2,695 Ga. Des turbidites synorogéniques de Quetico ont été piégées dans la zone de collision. L’événementaccrétionnaire final a impliqué l’ajout du terrane de Minnesota River Valley provenant du sud ainsi que la déposition etle métamorphisme de turbidites synorogéniques du terrane de Pontiac au cours de l’orogenèse minnesotaine, �2,68 Ga.Les données de sismique réflexion et réfraction indiquent des structures à pendage vers le nord qui sont interprétéescomme un empilement de terranes distincts d’une épaisseur de 10–15 km. Une dalle de matériau à haute vitesse, repré-sentant possiblement une lithosphère océanique accrétée par le dessous, ainsi que des décalages du Moho, supportentun modèle d’accrétions progressives par des processus semblables à celui de la tectonique des plaques.

[Traduit par la Rédaction] Percival at al. 1117

Received 2 June 2005. Accepted 23 May 2006. Published on the NRC Research Press Web site at http://cjes.nrc.ca on4 September 2006.

Paper handled by Associate Editor R. Clowes.

J.A. Percival,3 M. Sanborn-Barrie, T. Skulski, and D.J. White. Geological Survey of Canada, 601 Booth Street, Ottawa, ONK1A 0E8, Canada.G.M. Stott. Ontario Geological Survey, 933 Ramsey Lake Road, Sudbury, ON P3E 6B5, Canada.H. Helmstaedt. Department of Geological Sciences and Geological Engineering, Queen’s University, Kingston, ON K7L 3N6, Canada.

1This article is one of a selection of papers published in this Special Issue on The Western Superior Province Lithoprobe andNATMAP transects.

2Lithoprobe Publication 1459; Geological Survey of Canada Contribution 2005771.3Corresponding author (e-mail: [email protected]).

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Introduction

Processes that led to the formation of the ancient conti-nents are relevant to the understanding of tectonically stableregions and the origin of much mineral wealth. The reach ofearth science has improved in several dimensions over thepast two decades, particularly in deep-earth probing and inresolving events in deep time. These advances have invokeda new level of understanding of Archean tectonic processes,which has in turn led to renewed debate on the role of platetectonics in the production of ancient continental crust (cf.de Wit 1998; Hamilton 1998, 2003).

As the world’s largest Archean craton, the Superior Province(Fig. 1) provides information on both the nature and scale ofancient processes. Its rich mineral wealth has inspired in-tense geological investigation over the past 20 years, includingthree Lithoprobe transects (Percival and West 1994; Luddenand Hynes 2000; White et al. 2003), the western SuperiorNATMAP project (e.g., Percival et al. 2000), OperationTreasure Hunt, Geology of Ontario synthesis (Geology ofOntario 1991), major mapping projects in northern Quebec(e.g., Leclair et al. 1998), studies of the geodynamic settingof greenstone belts (e.g., Kerrich et al. 1999), and many pri-vate exploration initiatives aimed at base-and precious-metaltargets and diamond exploration.

Well-preserved supracrustal sequences of the western Su-perior Province were studied as early as the late 1800s andhave been the basis for many fundamental ideas on Archeangeology (e.g., Lawson 1913). Because of its relatively good ex-posure and detailed knowledge base, this classic greenstone–graniteregion was a natural focus for a Lithoprobe transect; de-signed to test the hypothesis that its linear belt structure re-sulted from the formation and accretion of island arcs andaccretionary prisms, as had been proposed by Langford andMorin in 1976, and since supported by many workers (e.g.,Card 1990; Williams et al. 1992; Stott 1997). Western Supe-rior NATMAP studies provided complementary informationalong Lithoprobe transects and their extensions focusing oncontinental–oceanic crust transitions, many of which hostimportant mineral deposits.

This paper summarizes new discoveries made through coor-dinated Lithoprobe and NATMAP geoscience activities since1997, in the form of a contemporary regional tectonic frame-work built on existing compilations (e.g., Williams et al.1992; Stott 1997; Card and Poulsen 1998; Skulski andVilleneuve 1999), new geophysical images of the crust (e.g.,Kay et al. 1999a, 1999b; White et al. 2003) and mantle (e.g.,Craven et al. 2001; Kendall et al. 2002), recent geologicalcompilation maps (e.g., Bailes et al. 2003; Corkery et al.;4

Percival et al. 2002b; Sanborn-Barrie et al. 2002, 2004; Stoneet al. 2002, 2004; Stott et al. 2002), and subprovince-scalesyntheses (e.g., Percival and Helmstaedt 2004).

Geological and geophysical setting

Lithosphere-scale perspectiveThe western Superior transect area extends from the U.S.

border in the south to the edge of the Superior craton in the

north, and the limit of Phanerozoic cover in the east andwest (Figs. 1, 2). Most work, including seismic reflectionand refraction experiments, was conducted in the southernpart of the area, utilizing the road network. In the north,teleseismic and magnetotelluric data was acquired with por-table instruments, and aircraft-assisted geological–isotopicmapping was targeted to create a north–south profile of vari-able width and detail (Fig. 3).

Several first-order observations on lithosphere structureand properties can be made from the various geophysicalstudies conducted within the western Superior Province. Forreference, seismic reflection images and a coincident seismicrefraction velocity model (from Musacchio et al. 2004) forthe crust and upper mantle, which form the basis of the in-terpretation summarized here (after White et al. 2003 andMusacchio et al. 2004), are shown in Fig. 4 and depictedschematically in Fig. 4b. Crustal thickness values (or depthto Moho, where the Moho is identified by M1–M3 in Fig. 4a)along the north–south transect decrease northward from 45to 38 km with abrupt offsets identified in the seismic reflec-tion image (e.g., at S1 and S2 in Fig. 4a). The crust thinswestward away from the transect to a minimum of 34 km,and eastward reaches a local minimum of ~36 km in theLake Nipigon region which may be related to Mesoproterozoicrifting (Kay et al. 1999a; Musacchio et al. 2004; Calvert etal. 2004). The gross-scale lithospheric architecture in the re-flection images–refraction models is characterized by a pre-dominance of N-dipping features, which have beeninterpreted as the end result of north-dipping subduction,collision, and accretion. Noteworthy features include a high-velocity (Vp = 7.4–7.5 km·s–1) northward-tapering zone (L)of inferred amphibolite or garnet amphibolite composition atthe base of the crust with associated reflectivity that contin-ues into the upper mantle (S2). Within the upper mantle be-neath this, a 15–20 km thick high-velocity layer (H) thatdips shallowly northward from a minimum depth of 48–50 km has an inferred harzburgite peridotitic compositionbased on its associated >6% azimuthal anisotropy and highP-wave velocities (Vp = 8.4–8.8 km·s–1). Both of these lay-ers, based on inferred composition, structural attitude, andtectonic context are interpreted as relic oceanic lithospheresutured at the base of the crust during the final stages oflithospheric assembly. A slightly earlier suture (S1) is ob-served approximately 100 km farther to the north, adjacentto a crustal-scale zone of bivergence (B) approaching themargin of the North Caribou terrane. Comparison with Fig. 3reveals that the interpreted zone of suturing (S1 to S2) isspatially coincident with a lateral transition in properties ofthe mantle that persists to depths exceeding 250 km.

Early tomographic images suggested the presence of litho-sphere beneath parts of the Superior Province between 250 km(Grand 1987) and 350 km thick (Van der Lee and Nolet1997). Analysis of shear-wave splitting indicated prominenteast–west anisotropy in the lithosphere (Silver and Chan1988), which is parallel to surface trends and attributed tomantle deformation during Archean tectonism (Silver 1996).Musacchio et al. (2004) estimated upper mantle velocities inthe 8.3–8.8 km·s–1 range, consistent with depleted harzburgite

4 Corkery, M.T., Skulski, T., Stone, D., Syme, E.C., Bailes, A.H., Cameron, M.T., and Whalen, J.B. Geology and tectono-stratigraphicassemblages, West Sachigo area, Manitoba. Ontario Geological Survey, Preliminary Map P.3463 or Manitoba Geological Survey, Open File2004–3 or Geological Survey of Canada Open File 1522, Scale 1 : 250 000. Manuscript in preparation.

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compositions. Other studies have added resolution to thestructure of the mantle lithosphere. Kendall et al. (2002) dis-tinguished a northern zone of isotropic upper mantle beneaththe North Caribou superterrane in contrast with a southernzone characterized by east–west anisotropy. These domainsare separated by a subvertical high-velocity zone that extends

to �300 km depth (Fig. 3; Sol et al. 2002). To the north, be-neath Paleoproterozoic crust of the Trans-Hudson orogen,the lithosphere is thinner and notably less anisotropic (Kend-all et al. 2002).

Domains are also observed in the electrical conductivitystructure. Craven et al. (2001) reported an essentially isotro-

Fig. 1. (a) Tectonic map of North America (after Hoffman 1990) showing the location of the Superior Province and Western Superior(WS) transect area. Greenland is restored to its position prior to the opening of the Labrador Sea. Grey areas show Paleoproterozoicdomains associated with the amalgamation of Laurentia. (b) Subprovinces and broad structural trends of the Superior Province (modi-fied after Card and Ciesielski 1986; Percival et al. 1992; Leclair et al. 2004). ERT, English River terrane; EwT, Eastern Wabigoonterrane; KU, Kapuskasing uplift; MT, Marmion terrane; NCT, North Caribou superterrane; NSS, Northern Superior superterrane; OSD,Oxford–Stull domain; PS, Pontiac subprovince; QT, Quetico terrane; WAT, Wawa–Abitibi terrane; WRT, Winnipeg River terrane; WwT,Western Wabigoon terrane; P, Paleoproterozoic cover. Note locations of Kirkland Lake (KL), Porpoise Cove (PC), Attawapiskat (A), re-ferred to in text. Box shows location of Fig. 2.

Fig. 2. Modified tectonic framework for the western Superior Province showing age range of continental domains, distribution of oce-anic domains and metasedimentary belts, Proterozoic cover (P), and location of seismic profiles.

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pic mantle beneath the North Caribou superterrane, in contrastto pronounced east–west anisotropy in the south. Subsequentanalysis has modeled a steeply north dipping, tabular, resis-tive zone separating the two domains (Fig. 3; Craven et al.2004). Both the seismic and electrical structures are consis-tent with slab-like features attributed to formation of the Su-perior craton through subduction–accretion processes (Kendallet al. 2002; Craven et al. 2001).

Heat flow in the western Superior region averages 42 ± 8mW·m–2 (Cheng et al. 2002; Rolandone et al. 2003). Re-duced heat flow (mantle component) is consistent with athermal lithosphere at least 240 km thick (Jaupart et al.1998), supporting the concept of a thick tectosphere.

Direct control on mantle lithosphere composition is con-strained by the petrology of xenoliths recovered from kimberlitepipes. In the western Superior Province, however, the knownpipes are not ideally situated to provide suitable lithospheredata. Pipes in the Attawapiskat area (Fig. 1) sample mantleclose to the Superior margin and indicate mainly lherzoliticcompositions and a cool geotherm (Scully 2000; Scully et al.2004; Armstrong et al. 2004). The compositional range issimilar to that from Kirkland Lake pipes (Fig. 1; Vicker andSchulze 1994; Schulze 1996), where geothermobarometrysuggests geotherms are steeper, corresponding to surface heatflow of about 40 mW·m–2.

Taken together, observations of the Superior Province man-

tle indicate a cool refractory lithosphere typical of Archeancratons (e.g., Jordan 1978). Individual mantle domains cor-respond to recognized surface geological features (terranes)and anisotropies of seismic velocity and electrical conduc-tivity are east trending, coplanar to the dominant penetrativecrustal structures. Despite these advances, the level ofknowledge of the Superior lithosphere is significantly lessthan that for the Kaapvaal and Slave cratons, where diamon-diferous kimberlites provide both abundant samples of themantle and economic incentive to understand its structureand evolution (cf. Jones et al. 2003).

Geological settingThe Superior Province forms the Archean core of the Ca-

nadian Shield (Fig. 1a). It has been tectonically stable since�2.6 Ga, and has subsequently occupied a lower plate settingduring most Paleoproterozoic and Mesoproterozoic tectonismthat affected its margins.

A first-order feature of the Superior Province is its linearsubprovinces of distinctive lithological and structural charac-ter, accentuated by subparallel boundary faults (e.g., Cardand Ciesielski 1986). Trends are generally east–west in thesouth, west-northwest in the northwest, and northwest in thenortheastern Superior (Fig. 1b). Recent work, based on isoto-pic and zircon inheritance studies, has revealed fundamentalage domains across the Superior Province (Fig. 1). Five dis-

Fig. 3. Schematic representation of western Superior Province crust and mantle lithosphere structure as defined by seismic andmagnetotelluric data (modified after Kendall et al. (2002); Sol et al. (2002); White et al. (2003); Craven et al. 2004; Mussachio et al.(2004); and Percival et al. (2004b)). NSS, Northern Superior superterrane. The crustal and upper mantle structure in the outlined areaare taken from relatively high-resolution data sets shown in Fig. 4. To the north, the interpretation is based on broadband teleseismic,magnetotelluric, and reconnaissance geological data sources. Mantle features of note include major changes at the interface betweenthe North Caribou and terranes to the south. The North Caribou superterrane is characterized by relatively high velocity and modestseismic and electrical anisotropy. Across a boundary marked by steeply dipping electrical and seismic anomalies, the lithosphere veloc-ity declines as both electrical and seismic anisotropy increase. Near the northern end of the profile, the Northern Superior – North Car-ibou boundary projects downward towards a south-dipping zone of high resistivity.

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Fig. 4. (a) Seismic reflection data and seismic refraction velocity model (Musacchio et al. 2004) for the main north–south transect (see Fig. 2 for location) shown with no ver-tical exaggeration. The time-migrated seismic reflection data (0–16 s) from White et al. (2003) have been converted to depth using the refraction velocities to obtain the imagesshown. The differences in the geometry of the reflection and refraction profiles are accounted for by ensuring that the two images are spatially coincident at several pointsalong the profiles. It should be noted that the seismic reflection images have inherently higher spatial resolution than the refraction velocity model. Thus, resolution limitations(as presented in Musacchio et al. 2004) should be considered when making detailed comparisons. (b) Tectonic interpretation of the seismic data (after White et al. 2003). Ab-breviations as in Fig. 3 except for S1 and S2 which are dipping lower crust–mantle reflectors interpreted as sutures. Other labels are referred to in the text. The individual re-flection line segments (1ES, 1EN, 1D, and 1A) are labeled.

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tinct Mesoarchean terranes are recognized in spite of pervasiveoverprinting by Neoarchean magmatism, metamorphism, anddeformation. They represent fragments of continental crustthat existed for tens to hundreds of million years prior to thedevelopment of ocean basins. Units defined as terranes havea tectono-stratigraphic history independent to that of neigh-bouring regions prior to amalgamation. Superterranes representtectonic collages of terranes that were assembled before incorpo-ration into the composite Superior superterrane. Superterranesmay show evidence of metamorphic or deformation eventsassociated with their construction. Assemblages are defined asunconformity- or fault-bounded supracrustal sequences linkedby their lithological, age, and geochemical characteristics.

The oldest remnants of continental crust (�3.7 Ga) occurin the Northern Superior superterrane (NSS, Skulski et al.2000). To the south, a large remnant of �3.0 Ga continentalcrust, the North Caribou superterrane (Stott and Corfu 1991;Stott 1997), is thought to be the nucleus around which ter-ranes accreted during assembly of the Superior Province (cf.Goodwin 1968; Thurston et al. 1991; Williams et al. 1992;Stott 1997; Thurston 2002). Farther south, the WinnipegRiver (WR) and Marmion (MM) terranes are relatively smallcontinental fragments dating back to 3.4 and 3.0 Ga, respec-tively (Beakhouse 1991; Tomlinson et al. 2004). In the far

south, the Minnesota River Valley terrane (MRVT), of un-known extent, contains remnants of crust as old as �3.5 Ga(Goldich et al. 1984; Bickford et al. 2006).

Domains of oceanic affinity, identified by juvenile isoto-pic signatures and lack of inherited zircon in volcanic andplutonic rocks, separate most of the continental fragments(Fig. 1). These dominantly greenstone–granite terranes gen-erally have long strike lengths and record environments thatresemble present-day oceanic floor, plateaux, island-arc, andback-arc settings (e.g., Thurston 1994). Examples includeparts of the Oxford–Stull terrane in the north, the westernWabigoon in the west, and the Wawa–Abitibi terrane in thesouthwestern Superior Province (Fig. 1).

Still younger features, the metasedimentary belts (e.g.,English River, Quetico, Pontiac; Breaks 1991; Williams1991; Fralick et al. 2006), separate some of the continentaland oceanic domains. Extending across the entire province,these 50–100 km wide belts of metagreywacke, migmatite,and derived granite appear to represent thick synorogenic se-quences (Davis 1996b, 1998), deposited, deformed, and meta-morphosed during collisional orogeny.

Crustal geophysical perspectiveTrends defined on potential field maps (Fig. 5) correspond

Fig. 5. Potential field derivative maps of the western Superior Province and superimposed geological boundaries. (a) shaded relief totalfield aeromagnetic map (sun angle at 360°); (b) first vertical derivative of Bouguer gravity, highlighting lateral density variations (dataavailable from GSC Geophysical Data Centre, http://gdr.nrcan.gc.ca/index_e.php). Abbreviations as in Fig. 6.

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well with first-order geological features. Metasedimentarybelts generally define gravity and aeromagnetic lows, whereas,greenstone belts form gravity highs. The province-scale cur-vature of continental domains and structural trends is well-reflected in the orientation of aeromagnetic patterns.

Gravity data have been inverted to estimate crustal thick-ness following removal of near-surface effects (e.g., Nitescuet al. 2003). This approach supports seismic observations ofthinner crust beneath the English River terrane where Mohotopography may be as much as 8 km between the “roots” ofnorthern and southern greenstone terranes and the thin Eng-lish River crust.

Seismic reflection profiles (e.g., White et al. 2003; Calvertet al. 2004) generally indicate gently north-dipping struc-tures as far north as the central North Caribou superterrane(Figs. 3, 6). Farther north, dips are to the south denoting achange in vergence, which has been interpreted to be consis-tent with a doubly vergent orogen (White et al. 2003) or acrustal-scale synclinorium (Hynes and Song 2006). Reflec-tors locally extend beneath the Moho where they have beeninterpreted as fossil subduction sutures (White et al. 2003).Owing to the generally low conductivity of western SuperiorProvince crust, little information on the crust has been pro-vided through magnetotelluric surveys.

Tectonic architecture

The tectonic building blocks of the Superior Province aredescribed in terms of their internal makeup and the nature oftheir boundaries. Aspects of the assembly of the tectonic ele-ments, leading to cratonization of the Superior Province, arediscussed in a subsequent section.

Northern Superior superterraneDominated by granitic and gneissic rocks, the poorly ex-

posed Northern Superior superterrane at the northern fringeof the Superior Province (Fig. 2) has been recognized on thebasis of isotopic evidence from Manitoba, Ontario, and Que-bec (Skulski et al. 1999). In the west, the data include �3.5Ga orthogneiss from the Assean Lake block (Böhm et al.2000), >3.5 Ga inherited zircon ages (Skulski et al. 2000;Stone et al. 2004), and detrital zircons with ages up to 3.9Ga (Böhm et al. 2003). Tonalite–trondhjemite–granodiorite(TTG) magmatism occurred at 3.2–3.1 Ga and amphibolite-facies metamorphism at 2.68 and 2.61 Ga (Böhm et al.2003). To the east at Yelling Lake (Fig. 2), magmatism at

2.85–2.81 Ga was followed by metamorphism at 2.74 Ga,indicating a tectonothermal event prior to its assembly withother domains of the Superior Province (Skulski et al. 2000).The Porpoise Cove volcanic–sedimentary sequence of north-ern Quebec (Fig. 1b) contains 3.8 Ga felsic volcanic rocks(David et al. 2003) suggesting that it may belong to theNorthern Superior superterrane.

North Caribou superterraneThe North Caribou superterrane (Fig. 2; Thurston et al.

1991) is the largest Mesoarchean domain of the SuperiorProvince (Stott 1997). It is characterized by widespread evi-dence for crust with �3.0 Ga mantle extraction ages(Stevenson 1995; Stevenson and Patchett 1990; Corfu et al.1998; Hollings et al. 1999: Henry et al. 2000) and displaysevidence for an amalgamation event prior to 2.87 Ga (Stottet al. 1989; Thurston et al. 1991). A characteristic cover se-quence of quartz arenite and mafic to ultramafic volcanicrocks has been interpreted as platformal or rift-type deposits(Thurston and Chivers 1990). Mesoarchean units have beenvariably reworked by subsequent Archean magmatic anddeformational events. The superterrane has wide, transitionalnorthern and southern margins.

North Caribou, northern margin: Oxford–Stull domainThe Oxford–Stull domain (Fig. 2; Thurston et al. 1991)

represents the largely juvenile, 2.88–2.73 Ga continental north-ern margin of the 3.0 Ga North Caribou superterrane thatwas tectonically imbricated with oceanic crustal fragments(Skulski et al. 2000; Syme et al. 1999; Corkery et al. 2000;Stone et al. 2004). The Oxford–Stull domain tectono-stratigraphy(Corkery et al. 2000 and in preparation) includes 2.84–2.83Ga tholeiitic mafic sequences and calc-alkaline-arc volcanicrocks, with juvenile to locally enriched Nd isotopic compo-sition that are unconformably overlain by <2.82 Ga sedi-ments of the Opischikona assemblage that contain <2.94 Gadetrital zircons (Skulski et al. 2000). Synvolcanic plutons as-sociated with 2.84–2.72 Ga calc-alkaline volcanism are iso-topically juvenile in the Oxford–Stull domain, but have <3 GaNd model ages in the adjacent Munro Lake domain, reflectingthe influence of thicker North Caribou crust (Skulski et al.2000). This package was juxtaposed along D1 faults withsubmarine, depleted tholeiitic basalts of the Seller Lake as-semblage prior to intrusion of 2.78 Ga tonalite (Corkery etal. 2000). Submarine-arc volcanic rocks (2.738 Ga) coveredthe composite basement prior to �2.72 Ga D2 deformation

Fig. 6. Location of features referred to in the text (see also Fig. 1b). Domains: MLD, Munro Lake; ILD, Island Lake; MT,Marmion terrane; MRVT, Minnesota River Valley terrane. Blocks: PB, Pikwitonei; SLB, Split Lake. Subprovinces: BRS, BirdRiver. Greenstone belts: BI, Black Island; FLGB, Favourable Lake; GLGB, Gods Lake (Oxford–Stull domain); ILGB, IslandLake; SJGB, St. Joseph; SLGB, Stull Lake; RLGB, Red Lake; BUGB, Birch–Uchi; MDGB, Meen–Dempster; PLGB, PickleLake; FHMGB, Fort Hope – Miminiska; WLGB, Wallace Lake; GLGB, Garner Lake (Uchi domain); RiLGB, Rice Lake;BLGB, Bee Lake; NCGB, north Caribou; MLGB, Melchett Lake; SeLGB, Separation Lake; SSGB, Savant–Sturgeon; SGB,Shebandowan; WiLGB, Winston Lake; SrGB, Schreiber. Complexes: ALC, Assean Lake; ELMC, English Lake magmaticcomplex; BRPC, Berens River plutonic complex; CL, Cedar Lake gneiss; TL, Tannis Lake gneiss. Plutons: NCP, North Cari-bou; LLB, Lewis Lake batholith; FB, Fletcher Lake batholith. Faults: NKF, North Kenyon; WSWSZ, Wolf Bay – Stull –Wunnumin; SL–LSJF, Sydney Lake – Lake St. Joseph; WF, Wanipigow; SSZ, Seymourville; MRF, Miniss; PLF, Paint Lake; QF,Quetico; SRRLF, Seine River – Rainy Lake; GLTZ, Great Lakes tectonic zone. Lakes: YL, Yelling; PL, Ponask; CWP, centralWabigoon plateau. Towns and reference points: LoW, Lake of the Woods; M, Manitouwadge; MH, McKellar Harbour; RL,Rainy Lake.

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(Lin et al. 2006) that may mark collision of the NorthernSuperior superterrane (Skulski et al. 2000). Unconformablyoverlying the shortened continental margin collage is a 2.722–2.705 Ga successor arc of calc-alkaline to shoshonitic volcanicand associated sedimentary rocks (Oxford Lake assemblage;Brooks et al. 1982; Corkery and Skulski 1998; Corkery et al.2000; Skulski et al. 2000; Stone et al. 2004; Lin et al. 2006).Synorogenic sedimentary rocks of the Cross Lake assem-blage that contain detrital zircons ranging in age from 2.704

to 3.65 Ga (Corkery et al. 1992; Corkery et al. 2000; Lin etal. 2006) lie unconformably on the older rocks. The entirecollage is cut by northwest-trending, dextral shear zones(D3, Osmani and Stott 1988; Lin and Jiang 2001; Lin et al.2006; Parmenter et al. 2006), themselves cut by 2.692 Gagranite (Corkery et al. manuscript in preparation), providinga bracket of <2.704 to >2.692 Ga on D3. Several smallsyntectonic gold deposits and showings occur in associationwith faults such as the Wolf Bay – Stull – Wunnumin shear

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zone (Fig. 6), particularly in the Little Stull Lake area (Jiangand Corkery 1998).

Munro Lake and Island Lake domainsThe Munro Lake and Island Lake domains (Fig. 2) com-

prise plutonic rocks with several small supracrustal belts inthe northern North Caribou superterrane (Stone et al. 2004;Parks et al. 2006). In the Munro Lake domain, quartzite lo-cally interbedded with komatiite overlies 2.883–2.865 Gatonalite (Stone et al. 2004; Corkery et al. manuscript in prep-aration). Tonalite and granodiorite plutons across the Munroand Island Lake domains have U–Pb ages ranging from 2.88to 2.70 Ga and Nd model ages from 3.05 to 2.71 Ga reflect-ing variable recycling of North Caribou age crust (Turek etal. 1986; Stevenson and Turek 1992; Skulski et al. 2000).

To the south, the Island Lake domain includes 2.89, 2.85,and 2.74 Ga volcanic sequences in a series of structural panels(Parks et al. 2003, 2004). Diverse clastic sedimentary se-quences were deposited synvolcanically at <2.84 to >2.744 Gaand post-volcanically at <2.712 Ga (Island Lake assemblage).All of these sedimentary sequences have detrital zircon U–Pbages that range from 2.938 to 2.711 Ga (Corfu and Lin2000), consistent with North Caribou provenance. Penetrativedeformation is slightly younger (�2.695 Ga, Parks et al.2003, 2006), followed by localized strain and shear-zone-hosted gold mineralization (�2.658 Ga, Lin and Corfu 2002).

Central North Caribou domainThe central North Caribou domain, which is dominated by

younger plutonic material (Corfu and Stone 1998a), pre-serves several remnants of �3.0 Ga basement crust. Some ofthe oldest rocks are 3.02 Ga felsic volcanic rocks of theNorth Spirit assemblage (Corfu and Wood 1986), with juve-nile 3.1 and younger Nd model ages (Stevenson 1995). Oneof the largest plutonic remnants occurs in the southwesterncorner where Krogh et al. (1974) first recognized 3.0 Garocks. Here, rocks of the English Lake magmatic complex(Fig. 6) have ultramafic through tonalitic compositions withevidence of mantle derivation (Whalen et al. 2003).

Thin quartzite–komatiite packages are preserved sporadi-cally across the central North Caribou domain (Thurston andChivers 1990; Thurston et al. 1991). They consist of a lower,quartz-rich, coarse clastic unit, locally unconformable onbasement, overlain by carbonate, iron formation, basaltic,and komatiitic volcanic units. In different areas the sequenceshave been interpreted as platformal cover strata (Thurstonand Chivers 1990) or plume-related rift deposits (Hollingsand Kerrich 1999; Hollings 2002; Percival et al. 2002a, 2006).Along the southwestern margin of North Caribou basement,quartz arenite was deposited between �3.0 and 2.93 Ga(Percival et al. 2006; Sasseville et al. 2006). Evidence forplume-related rifting is taken from the presence of komatiite(Hollings and Kerrich 1999; Tomlinson et al. 2001). TheBalmer assemblage (2.99–2.98 Ga) of the Uchi domain ispossibly correlative (Tomlinson et al. 1998; Sanborn-Barrieet al. 2001).

Parts of the North Caribou superterrane have been assem-bled from older fragments (Stott 1997), although the earlyhistory is generally obscured by younger plutonism. Evi-dence of early tectonometamorphism comes from the NorthCaribou greenstone belt, where different volcanic assemblages

are intruded by the 2.87 Ga North Caribou pluton, which isinterpreted to postdate regional deformation and metamor-phism (Stott et al. 1989). Some workers have postulated thatthe iron-formation-hosted Musselwhite lode gold deposit formedduring development of structures associated with 2.87 Gapluton emplacement (Fyon et al. 1992). Alternatively, later(�2.7 Ga) structural reactivation and hydrothermal circula-tion could have utilized preexisting structures.

Neoarchean granitoid rocks of the Berens River plutoniccomplex dominate the central North Caribou domain (Stone1998). The complex comprises tonalitic, dioritic, granodioritic,and granitic plutons that crystallized between 2.745 and 2.708Ga and exhibit calc-alkaline geochemical trends (Stone 1998),as well as younger metaluminous and peraluminous granites(Corfu and Stone 1998a). There is a general trend towardhigher degrees of fractionation with decreasing age (Fig. 7in Corfu and Stone 1998a). A suite of 2.70–2.696 Ga plutonsof sanukitoid affinity (e.g., Shirey and Hanson 1984; Sternand Hanson 1991) shows evidence of extraction from de-pleted mantle, with some crustal contamination (Stevensonet al. 1999). Plutons of arc affinity exhibit Nd isotopic evi-dence of substantial assimilation of �3 Ga evolved crust(Henry et al. 1998). Plutons of the Berens River complexwere emplaced at depths from 18 to 10 km (0.6 to 0.3 GPa),and some of the youngest plutons were emplaced at thehighest crustal levels (Stone 2000). Regional barometric datashow higher paleopressures along the southern margin of thecomplex, indicative of greater exhumation. Together withseveral low-pressure estimates from immediately north ofthe Uchi domain (op. cit.), this observation is consistent withan interpretation based on seismic reflection data that a gentlysouth-dipping normal fault separates the hanging wall RedLake belt from the footwall Berens River complex (Calvertet al. 2004), although no major structure has been recog-nized in this area (Sanborn-Barrie et al. 2004).

A crustal-scale synform beneath the southern Berens Rivercomplex (Fig. 4) was interpreted from seismic reflectiondata (White et al. 2003). Hynes and Song (2006) studiedmetamorphic conditions of supracrustal units in a transectacross the southern limb of the “synform”. Pressure esti-mates increase northward as predicted, but not to the extentexpected for folded subhorizontal layering. Rather, the struc-ture (Fig. 4) appears more akin to that of divergent orogengeometry as described by Snyder et al. (1996) and modeledby Ellis et al. (1998), in which the north-dipping reflectorswould represent south-vergent thrusts.

In general, across the Berens River plutonic complex, plutonswithin and adjacent to greenstone belts cooled quickly frommagmatic temperatures, whereas, those in the interior of thecomplex underwent more prolonged thermal and hydrother-mal activity, as indicated by titanite and apatite U–Pb datesin the 2.66–2.63 Ga range (Corfu and Stone 1998b). Volca-nic equivalents to many of the plutonic suites are recognizedas the Confederation, Graves, and St. Joseph assemblages ofthe Uchi domain.

North Caribou superterrane, southern margin: Uchidomain

The Uchi domain records �300 million years of tectono-stratigraphic evolution along the southern margin of the NorthCaribou superterrane (Figs. 2, 7; Stott and Corfu 1991; Corfu

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and Stott 1993a, 1996; Hollings et al. 2000; Sanborn-Barrieet al. 2001). Chronostratigraphic correlations have been es-tablished within greenstone belts over a strike length of atleast 500 km.

North Caribou basement is exposed near Lake Winnipeg(Fig. 6; Krogh et al. 1974; Whalen et al. 2003) and is inferredfarther east. The 2.99–2.96 Ga rift-related Balmer assem-blage may have equivalents in the Lewis–Storey, Conley,and Overload Bay assemblages (Tomlinson and Sasseville2000; Tomlinson et al. 2001; Sasseville 2002; Sasseville etal. 2006; Percival et al. 2006). Other possible linkages amongisolated Mesoarchean units include the 2.94–2.91 Ga Ballassemblage at Red Lake with plutonic units to the west, andmafic–ultramafic rocks of the Garner Lake belt (Figs. 6, 7;Anderson 2003) that may be correlative with the �2.87 GaWoman assemblage in the Birch–Uchi belt. A deformationevent and unconformity or disconformity separate Meso-archean from Neoarchean strata across the Uchi domain.The 2.745–2.734 Ga Confederation and Graves calc-alkalinevolcanic assemblages are widespread in greenstone belts ofthe eastern Uchi domain and absent in the west; although,plutonic rocks of equivalent age are present in the BerensRiver complex to the north (Corfu and Stone 1998a). Youngerpackages, including the 2.731–2.729 Ga Bidou, 2.723 GaBlack Island, 2.722–2.718 Ga Gem, and 2.718 Ga AndersonLake (Rogers 2001; Rogers and McNicoll, personal communi-cation, 2005) assemblages in the west are temporally correla-

tive with the 2.723–2.713 Ga St. Joseph assemblage in theeast (Fig. 7). Percival et al. (2006) suggested that the BlackIsland assemblage was accreted to the southwestern NorthCaribou margin. Similar juvenile material appears to extendto the west beneath Paleozoic cover based on aeromagneticcharacter and Nd isotopic results from basement drill core(Stevenson et al. 2000).

Coarse clastic sedimentary rocks generally form the youn-gest strata along the southern margin of the North Caribousuperterrane. Where dated, these sequences contain detritalzircons as young as 2.703 Ga and may be facies equivalentsof the marine greywacke turbidites of the English River terraneto the south (e.g., Campbell 1971; Devaney 1999a; Stott1996). The sedimentary rocks have variable age relationshipswith respect to deformation. For example, in the westernarea some <2.704 Ga sedimentary packages carry D1–D5structures (Brommecker 1991; Anderson 2003, 2004), whereas,other <2.705 Ga assemblages were deposited unconformablyon units affected by D1 ± D2 deformation. There are probablyclose temporal and process linkages between sedimentationand deformation.

The Red Lake greenstone belt is one of Canada’s mostprolific gold producers (Pirie 1982; Corfu and Wallace 1986;Corfu and Andrews 1987; Sanborn-Barrie et al. 2004; Zengand Calvert 2006; J.R. Harris et al. 2006). The �300 millionyear stratigraphic sequence, built on North Caribou base-ment, records rifting (Tomlinson et al. 1998), continental-arc

Fig. 7. Tectono-stratigraphic columns illustrating possible regional correlations among Mesoarchean and Neoarchean units and events inthe Uchi subprovince (modified after Stott and Corfu 1991; Beakhouse et al. 1999; Rogers 2002; Bailes et al. 2003; Hollings andKerrich 2004; Sanborn-Barrie et al. 2004). Assemblage abbreviations for greenstone belts for, Lake Winnipeg – Rice Lake: LS, Lewis–Storey; Bd, Bidou; Bl, Black Island; Ed, Edmunds Lake; Gm, Gem; Ho, Hole River; Sa, San Antonio. Assemblage abbreviations forWallace Lake: Cl, Conley; OB, Overload Bay; Bg, Big Island; Sd, Siderock. Assemblage abbreviations for Garner – Bee Lake: Gn,Garner Lake intrusion; Al, Anderson; Kl, Kangaroo; Assemblage abbreviations for Red Lake: Bm, Balmer; Bl, Ball; Sl, Slate Bay; BC,Bruce Channel; TB, Trout Bay; Cf, Confederation; Gr, Graves; Assemblage abbreviations for Birch–Uchi: Wo, Woman; Sf, Sundown;Sp: Springpole; Assemblage abbreviations for Meen–Dempster: Kg, Kaminiskag; Mn, Meen; Assemblage abbreviations for PickleLake: PC, Pickle Crow; Assemblage abbreviations for Lake St. Joseph: SJ, St. Joseph; Ei, Eagle Island.

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magmatism (Henry et al. 2000; Sanborn-Barrie et al. 2001),intra-arc rifting (Parker 1999), several phases of deforma-tion, and associated sedimentation (Sanborn-Barrie et al. 2004).Lode gold deposits are localized within altered Balmer as-semblage rocks in proximity to the Mesoarchean–Neoarcheanunconformity (Parker 2000; Sanborn-Barrie et al. 2001; Dubéet al. 2004; J.R. Harris et al. 2006). Multiple ages of miner-alization are identified, including main-stage mineralizationassociated with D2 structures prior to 2.712 Ga and late goldremobilization after 2.701 Ga (Corfu and Andrews 1987;Dubé et al. 2004). The timing of metamorphism has notbeen established directly. U–Pb titanite and apatite dates inthe Red Lake belt reveal a rapid thermal decline from theheight of magmatic activity at �2.730 Ga to 2.71–2.70 Ga(Corfu and Stone 1998b). However, a second prograde meta-morphic pulse is indicated by amphibolite-facies metamor-phism of the <2.70 Ga Austin conglomerate (Sanborn-Barrieet al. 2004) and by regionally extensive 40Ar/39Ar dates of�2.66 Ga for hornblende and 2.63 Ga for biotite (Hanes andArchibald 1998).

Across the Uchi domain, the regional-scale extent of de-formation and the fact that strain gradients and associatedmetamorphic imprints transect volcanic assemblages andplutons, indicate that the driving forces for tectonometam-orphism have larger length-scales than individual greenstonebelts or plutons (Stott and Corfu 1991). These observationspoint to complex tectonic processes as the underlying cause.Tectonic models for regional deformation across the south-ern margin of the North Caribou superterrane, attributed tothe Uchian orogeny, are described in a later section.

English River terraneThe English River terrane (Figs. 2, 6) is distinguished

from adjacent regions by supracrustal rocks of metasedimentaryorigin, high metamorphic grade, and a prominent east–weststructural grain (Breaks 1991). The setting of the EnglishRiver has traditionally been considered as a fore-arc basin(Langford and Morin 1976) or accretionary prism (Breaks1991), although, more recently Pan et al. (1999) suggested aback-arc setting. Detrital zircon studies indicate that somesediments were deposited <2.704 Ga after cessation of arcactivity in adjacent volcanic belts (Corfu et al. 1995; Davis1996a, 1996b, 1998) and are broadly syn-collisional, therebyimplying an origin as a synorogenic flysch basin. The smallMelchett Lake greenstone belt (Devaney 1999b) in the east-ern English River terrane comprises a juvenile 2.726 Gacalc-alkaline volcanic sequence (Corfu and Stott 1993a; 1996;Davis et al. 2005), possibly correlative with the St. Josephassemblage to the north.

The main protolith to metasedimentary schist, migmatite,and derived diatexite is turbiditic greywacke, with some ox-ide-facies iron formation. Detrital zircons indicate sourceages between 3.25 and 2.704 Ga (Corfu et al. 1995; Stott etal. 2002). A lower bracket on depositional age is providedby 2.698 Ga plutons (Corfu et al. 1995).

Metamorphic conditions range from middle amphibolitefacies near the margins to interior upper amphibolite (650–750 °C, �0.5 MPa, Pan et al. 1999), and granulite facies(750–850 °C, 0.6–0.7 MPa, Perkins and Chipera 1985; Panet al. 1999). Elevated temperatures may have been attainedthrough addition of mantle-derived magmatic heat (Breaks

1991), an inference supported by the near-isobaric coolingpaths (Hynes 1997). The main tectonothermal event at �2.691Ga was followed by a second thermal pulse at 2.669 Ga(Corfu et al. 1995; Pan et al. 1999), intrusion of �2.65 Gapegmatites (Corfu et al. 1995; Smith et al. 2004), growth ofhydrothermal minerals (Pan et al. 1999), and relatively slowcooling (40Ar/39Ar biotite ages 2.66–2.4 Ga; Hanes andArchibald 2001).

The dominant east–west structural grain of the terrane re-flects upright to north-vergent F2 folds of an S1 foliation(Breaks 1991; Hrabi and Cruden 2001). The early foliationappears to be a composite fabric that includes primary layer-ing and at least one set of early structures (Sanborn-Barrie1988; Hynes 1997, 1998). Most strain (D3–D5 events) coin-cided with or postdated formation of �2.69 Ga migmatiticlayering (Corfu et al. 1995; Hrabi et al. 2000; Hrabi andCruden 2006).

Seismic lines 2a and 2b (Fig. 2) show minimal reflectivityfrom the English River terrane. Gravity (Nitescu et al. 2003),seismic reflection (White et al. 2003), and seismic refraction(Kay et al. 1999b) profiles collectively indicate that the Mohobeneath the combined English River – Winnipeg River terraneis shallower by about 8 km than in adjacent subprovinces. Alate to posttectonic uplift event could account for exposureof high-grade rocks with a protracted cooling history, possi-bly on strike-slip faults with normal components (Stone 1981;Stott 1996; White et al. 2003). Alternatively, the thin crustcould partly be an inherited feature of the thinned continen-tal margin of the Winnipeg River terrane.

The southern margin of the western English River terraneis marked by the Bird River – Separation Lake greenstonebelt (Breaks 1991), which includes mafic metavolcanic rocksand the Cr-bearing Bird River sill. Ages of the largely juve-nile supracrustal units range from �2.78 to 2.73 Ga (Timminset al. 1985).

Winnipeg River terraneThe Winnipeg River terrane is a collective term used to

describe the plutonic domain exposed north and east of thewestern Wabigoon volcanic domain. It consists of two mainelements, (i) the Winnipeg River subprovince of Beakhouse(1991), a >500 km long terrane composed of Neoarcheanplutonic rocks with Mesoarchean to Paleoarchean inheritance;and (ii) a Neoarchean plutonic domain, formerly referred toas the central Wabigoon granitoid complex (Percival et al.2002b, 2004a) or Wabigoon diapiric axis (Edwards and Sutcliffe1980; Thurston and Davis 1985; cf. Schwerdtner 1992) thatcontains scattered remnants of Mesoarchean crust and isoto-pic evidence for recycled 3.4–3.0 Ga material (Tomlinsonand Percival 2000; Tomlinson et al. 2004; Whalen et al.2002, 2004a). With inheritance dating back to �3.4 Ga (Henryet al. 2000; Tomlinson and Dickin 2003), the WinnipegRiver terrane stands apart from the Northern Superior andNorth Caribou superterranes to the north and the Marmiondomain to the south (described later in the text). It also car-ries a long record of magmatic and structural events (Corfu1988; Percival et al. 2004a; Melnyk et al. 2006).

The Mesoarchean history of the Winnipeg River terranehas remained cryptic because of extensive overprinting byNeoarchean magmatism and deformation. Tonalitic rocks arethe oldest units (3.32–3.04 Ga, Krogh et al. 1976; Corfu

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1988; Davis et al. 1988; Melnyk et al. 2006) and some showNd isotopic signatures in excess of 3.4 Ga (Henry et al.2000; Tomlinson and Dickin 2003) and zircon inheritance.Similar isotopic signatures characterize 2.88–2.83 Ga tonali-ties (Beakhouse and McNutt 1991; Beakhouse et al. 1988).Volcanic belts 2.9–3.07 Ga (Davis et al. 1988; Sanborn-Barrie et al. 2002) are also considered part of the WinnipegRiver terrane.

Significant pulses of Neoarchean tonalite–granodioritemagmatism occurred at 2.716–2.705 Ga, followed by em-placement of granites at �2.70–2.69 Ga (Beakhouse 1991;Beakhouse et al. 1988; Cruden et al. 1997, 1998; Corfu1988, 1996). Beakhouse (1991) noted the lack of magmatismin the Winnipeg River terrane north of the western Wabigoonsubprovince between 2.75 and 2.71 Ga, a time of major ac-tivity in the adjacent Wabigoon and Uchi subprovinces. Heconcluded that the Winnipeg River subprovince had remainedtectonically isolated until after �2.71 Ga, when it began tointeract with neighbouring arc terranes.

A complex Neoarchean structural–metamorphic historycharacterizes the Winnipeg River terrane (e.g., Gower andClifford 1981). Rocks as young as 2.72 Ga and older poly-deformed gneisses were folded (D3) between 2.717 and2.712 Ga, prior to syntectonic injection of 2.71–2.707 Gatonalite and granodiorite sheets accompanying D4 deforma-tion (Melnyk et al. 2006). Upright D5 folding took place af-ter 2.705–2.70 Ga, and F6 folds occurred within a dextraltranspressive regime, possibly as late as 2.65 Ga (Melnyk etal. 2006).

The eastern Winnipeg River terrane is a 200 km widetransverse corridor of granitoid rocks separating the volcanic-dominated eastern and western Wabigoon domains (Figs. 2,6). Small greenstone belts with ages >3.075–2.703 Ga (Daviset al. 1988; Tomlinson et al. 2002, 2003) are cut by granitoidunits 3.075–2.680 Ga (Davis et al. 1988; Whalen et al. 2002).Some of the oldest rocks have εNd values of –1 to +1, sug-gesting derivation from even older crustal sources (Tomlinsonet al. 2004). At least five generations of Neoarchean struc-tures (D1–D5) have been recognized in complex tonaliticgneisses (Brown 2002; Percival et al. 2004a), although thedominant tonalite–granodiorite suite (2.723–2.709 Ga, Whalenet al. 2002) has only S3 foliation, F4 folds, and D5 shear zones.

Marmion terraneThe Marmion terrane (Figs. 2, 6), formerly included as

part of the south-central Wabigoon subprovince, is now rec-ognized as consisting of 3.01–2.999 Ga Marmion tonalitebasement (Davis and Jackson 1988; Tomlinson et al. 2004),upon which several greenstone belts formed between 2.99and 2.78 Ga (Stone et al. 2002; Tomlinson et al. 2003). Incontrast with Winnipeg River-type crust with 3.4 Ga ances-try to the north, the Marmion terrane appears to have beenjuvenile at 3.0 Ga. It either accreted to the Winnipeg Riverterrane by �2.92 Ga (Tomlinson et al. 2004) or formed bymagmatic addition of 3.0 Ga juvenile crust at the WinnipegRiver margin. The Marmion terrane experienced little, if any,Neoarchean (i.e., 2.745–2.72 Ga) magmatic activity in con-trast with the Winnipeg River terrane to the north and theWabigoon terranes to the west and east (described in the fol-lowing text).

The eastern Winnipeg River and Marmion terranes are

characterized by steeply dipping structures at surface andsubhorizontal reflectivity at depth. Lithoprobe Line 1, whichcrosses these domains, shows several 10 km scale, gentlynorth-dipping crustal panels, including a lower-crustal, high-velocity (Figs. 3, 4; Musacchio et al. 2004) layer of maficcomposition that terminates as one of two mantle reflectorsalong the line (White et al. 2003). This feature is interpreted,on the basis of its seismic features and gravity expression, tobe made up dominantly of amphibole and inferred to be asubcreted fragment of oceanic crust (White et al. 2003).

Wabigoon subprovinceThe Wabigoon subprovince has long been recognized as a

composite terrane comprising volcanic-dominated domainswith a central axis of variable-age plutonic rocks (Davis andJackson 1988; Percival et al. 2002b). Current understandingis that it comprises distinct western and eastern domains(Fig. 2) separated by rocks of Mesoarchean ancestry(Tomlinson et al. 2002, 2004). The tectonic characteristicsand significance of the western and eastern Wabigoon areoutlined in the following text.

Western Wabigoon domainThe western Wabigoon domain is dominated by mafic vol-

canic rocks with large tonalite–granodiorite plutons (Blackburnet al. 1991). Volcanic rocks range in composition from tholeiiticto calc-alkaline and are interpreted to represent ocean crustand arc environments, respectively (Ayer and Davis 1997;Ayer 1998a; Ayer and Dostal 2000; Wyman et al. 2000).Most of the preserved volcanic rocks were deposited be-tween �2.745 and 2.72 Ga (Corfu and Davis 1992) with rareolder (2.775 Ga), and younger (2.713–2.70 Ga) volcanic–sedimentary sequences. Plutonic rocks range from broadlysynvolcanic batholiths composed of tonalite–diorite–gabbro(�2.735–2.72 Ga, Davis and Edwards 1982; Corfu and Da-vis 1992; Whalen et al. 2004a), to younger granodioritebatholiths and plutons (�2.710 Ga, Davis and Edwards 1986;Sanborn-Barrie 1988; Davis and Smith 1991; Melnyk et al.2006), monzodiorite plutons of sanukitoid affinity (�2.698–2.690 Ga; Stern and Hanson 1991; Ayer 1998b; Stevenson etal. 1999), and plutons and batholiths of monzogranite (2.69–2.66 Ga; Schwerdtner et al. 1979; Sanborn-Barrie 1988;Melnyk et al. 2000). Immature clastic metasedimentary se-quences are preserved in narrow belts within volcanic se-quences. They are commonly younger than the volcanic rocks,as illustrated by local unconformable relationships (Fralick1997) and geochronological constraints, indicating deposi-tion between �2.711 and <2.698 Ga (Davis 1996a, 1996b,1998; Davis et al. 1988; Fralick and Davis 1999; Sanborn-Barrie and Skulski 2006). Virtually all carry ancient (>3 Ga)detrital zircons indicating ancient source regions. At leasttwo phases of deformation affected supracrustal rocks of thewestern Wabigoon subprovince (Blackburn et al. 1991; Ed-wards and Stauffer 1999) with apparent diachroneity in theonset of deformation from pre-2.709 Ga in the Lake of theWoods area (Davis and Smith 1991; Ayer and Davis 1997;Melnyk et al. 2006), to �2.700 Ga in the Sioux Lookout –Savant area in the east (Sanborn-Barrie et al. 1998, 2002;Sanborn-Barrie and Skulski 2006). These events involved atleast local tectonic inversion, through thrust imbrication (Da-

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vis et al. 1988), and formation of nappe-like structures (e.g.,Poulsen et al. 1980).

The Sturgeon–Savant greenstone belt (Fig. 6) hosts the in-terface between the Winnipeg River terrane and its auto-chthonous cover and juvenile rocks of the western Wabigoonsubprovince (Sanborn-Barrie et al. 2002), represented byoceanic plateau and arc volcanic sequences (2.775–2.72 Ga;Davis et al. 1988). Continental rift rocks include a thin 2.93–2.88 Ga mafic–felsic volcanic sequence (Skulski et al. 1998)and the overlying Jutten assemblage comprising a lower sed-imentary unit with 3.4–2.9 Ga detrital zircons, and an uppertholeiitic volcanic sequence with εNd values of +0.5 to +2.0(Davis and Moore 1991; Sanborn-Barrie and Skulski 2006).Juvenile calc-alkaline arc rocks were erupted mainly between2.745 and 2.735 Ga (Davis et al. 1985; Sanborn-Barrie andSkulski 1999; Sanborn-Barrie et al. 2002). Voluminous co-eval (2.735 Ga) tonalitic rocks (Whalen et al. 2004b) mayhave provided the heat source for seawater convection andmassive sulphide mineralization (Galley et al. 2000). Youn-ger (�2.718 Ga; Davis et al. 1988), high Fe, Ti basalt, andminor dacite represent a rifted arc sequence. Associated sed-imentary rocks contain both Neoarchean (2.745–2.730 Ga)and Mesoarchean (3.1–2.8 Ga) detritus based on SHRIMPU–Pb zircon analyses (Skulski et al. 1998). Two youngersedimentary sequences complete the stratigraphic record ofthe Sturgeon–Savant greenstone belt: (i) greywacke – ironformation (2.716–2.711 Ga) of the Warclub assemblage; and(ii) sandstone and arkose (<2.698 Ga) of the synorogenicAment Bay assemblage (Davis et al. 1988). Two sets of duc-tile structures postdate <2.704 Ga rocks: (i) north-trendingupright F1 folds; and (ii) east-trending upright D2 folds andpenetrative foliation. Pre-D1 folds have been inferred locally(Sanborn-Barrie et al. 1998).

Eastern Wabigoon domainThe eastern Wabigoon domain (Fig. 2) is a composite

terrane with greenstone belts and intervening granitoid plutonsthat show variable Mesoarchean (Winnipeg River and Mar-mion) and Neoarchean ancestry. The supracrustal rocks havebeen divided into several assemblages (Stott and Davis 1999;Tomlinson et al. 2000; Stott et al. 2002). In the northwest,the 3.0–2.92 Ga Toronto and Tashota assemblages may rep-resent a continental margin sequence built on the WinnipegRiver terrane. The central part of the belt is dominated byrocks of oceanic affinity including tholeiitic basalts of the2.78–2.769 Ga Onaman assemblage, 2.738 Ga Willet back-arcrocks, and the overlying 2.734–2.722 Ga calc-alkalineMetcalfe–Venus assemblage of continental affinity (Stott etal. 2002). Across the southeastern Wabigoon domain, the2.74–2.734 Ga calc-alkaline Elmhirst–Rickaby assemblageis possibly built on Marmion-age substrate (Tomlinson et al.2004). Unconformably overlying clastic rocks were depos-ited after �2.71 Ga. At least two sets of structures are pres-ent in the eastern Wabigoon domain: east–west-striking D1folds and foliation (<2.709 Ga) and east–west-striking, dextraltranspressive D2 structures and related shear zones most no-table across the Humboldt Bay high strain zone (Stott et al.2002). A 2.694 Ga pluton provides a lower limit on the ageof D2 deformation (Stott and Davis 1999).

Seismic reflection profile 3 (see location on Fig. 2) pro-vides an image of the crustal structure beneath the eastern

Wabigoon subprovince, revealing gently north-dipping reflec-tivity at depth. The structure resembles that of the Marmion andeastern Winnipeg River terranes to the east. Prominentsouth-dipping features at the southern margin of the subprovincecorrespond to the Paint Lake fault. The electrical structure ofthe mantle, imaged with magnetotelluric techniques, showsprominent east–west conductivity anisotropy, interpreted asgraphite films reflecting a tectonic fabric (Craven et al. 2001).

Two general models have been proposed for formation ofthe Wabigoon domains: (i) an ensialic rift setting (Blackburn1980; Blackburn et al. 1991; Cruden et al. 1998; Devaney2000); and (ii) an oceanic setting followed by accretion tothe Winnipeg River terrane (Davis and Smith 1991; Corfu1996; Percival et al. 2004a; Sanborn-Barrie and Skulski 2006;Melnyk et al. 2006). Further discussion of these tectonicmodels follows in the section Central Superior orogeny.

Quetico terraneThe Quetico terrane (Figs. 2, 6) consists dominantly of

greywacke, derived migmatite, and granite. No stratigraphicsequence has been established within the steeply dipping,polydeformed, and variably metamorphosed sedimentarysuccession. Younging directions are dominantly to the north(Percival 1989), yet age constraints indicate older ages ofdeposition for the northern Quetico (<2.698 >2.696 Ga; Da-vis et al. 1990; Davis 1998) relative to the south (<2.692 Ga;Zaleski et al. 1999), consistent with accretionary prismgrowth (Percival and Williams 1989; Valli et al. 2004).

Several plutonic suites cut metasedimentary units includ-ing 2.696 Ga tonalite (Davis 1996a). An early (D1) deforma-tion event pre-dated emplacement of a chain of Alaskan typemafic–ultramafic intrusions in the northern Quetico (e.g.,Pettigrew 2004), which are associated with alkaline plutonsincluding nepheline syenite and carbonatite. These rocks,derived from metasomatized mantle, have ages in the range2.69–2.68 Ga (Lassen 2004) and geochemical affinities withthe Archean sanukitoid suite (cf. Stern et al. 1989; Stevensonet al. 1999; Lassen 2004). Two subsequent deformation events(D2 and D3) were followed by low-pressure, high-temperaturemetamorphism that reached upper amphibolite and localgranulite facies at �2.67–2.65 Ga (Pan et al. 1994, 1998;Valli et al. 2004) in the central region and greenschist faciesat the margins (Percival 1989). Coeval, crust-derived graniticplutons and pegmatites include �2.67 Ga peraluminous gran-ite and �2.65 Ga biotite granite (e.g., Southwick 1991).

Tectonic models for the Quetico terrane have favouredfore-arc settings (e.g., Langford and Morin 1976; Percivaland Williams 1989; Williams et al. 1991; Fralick et al.2006). Depositional ages of �2.698 to 2.690 Ga overlap re-gional deformation and late magmatism in the Wabigoonsubprovince, suggesting a synorogenic origin (Davis 1998).

Wawa–Abitibi terraneMost workers accept a correlation between the Wawa and

Abitibi terranes across the transverse Kapuskasing uplift structure(Fig. 1b; see Percival and West 1994 for a review). Althoughthe Wawa terrane was not the focus of Lithoprobe–NATMAPactivities, we outline its geological history to assess its tec-tonic significance with respect to the western Superior Prov-ince. Within the Wawa terrane, volcanism appears to haveinitiated with the 2.89–2.88 Ga Hawk assemblage (Turek et

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al. 1992). An oceanic setting is indicated for the 2.745 GaWawa and 2.72 Ga Greenwater and Manitouwadge assem-blages (Turek et al. 1992). The latter two formed part of anarc–back-arc system that is characterized by significant mas-sive sulphide deposits (Corfu and Stott 1998) in theShebandowan, Winston Lake, and Manitouwadge greenstonebelts (Sage et al. 1996a, 1996b; Williams et al. 1991). Polatet al. (1999; Polat and Kerrich 2001) reported a variety ofoceanic magma types from the Schreiber belt and interpretedthe belt as a tectonic mélange (Polat et al. 1998; Polat andKerrich 1999).

Relatively late-stage volcanism at �2.695 Ga took placeduring D1 thrusting. Subsequent �2.689 Ga calc-alkalic toalkalic magmatism (Corfu and Stott 1998) and associatedcoarse clastic Timiskaming-type sedimentation (<2.689 Ga)were followed by sanukitoid magmatism (2.685–2.680 Ga)and dextral transpressive D2 deformation (Davis and Lin2003). The youngest rocks are <2.682 Ga conglomerates de-posited near the Quetico subprovince boundary, concurrentwith D2 deformation (Corfu and Stott 1998). The �2.685–2.68Ga tectonic events were termed the Shebandowanian phaseof the Kenoran orogeny (Stott and Corfu 1991).

Minnesota River Valley terraneThe poorly exposed Minnesota River Valley terrane (MRVT,

Figs. 1, 6) contains some of the oldest rocks of the SuperiorProvince, with a history dating back to �3.5 Ga (Bickford etal. 2006). Collision between the MRV and Wawa–Abitibiterranes is inferred to have begun �2.685 Ga, based on thecessation of arc magmatism and onset of penetrative defor-mation in the Wawa–Abitibi terrane, followed by the influxof turbidites into the Pontiac basin in the southeastern Supe-rior Province (2.685–2.682 Ga; Davis 2002). Late granitesof crustal origin have ages in the 2.67–2.65 Ga range (Goldichet al. 1984).

Boundary relationshipsBoundaries between adjacent terranes can be demonstrated

to be long-lived, evolving features consistent with develop-ment within Wilson cycles. For example, the timing of riftsequences, arc magmatism, collisional deformation, overlapsedimentation, and transcurrent faulting at several bound-aries fits the pattern and timescale anticipated for oceanopening and closing, although, these features are rarely allpreserved.

The Northern Superior superterrane is juxtaposed with thejuvenile 2.84–2.71 Ga Oxford–Stull domain along the dextraltranscurrent North Kenyon fault. Orogenic sediments depos-ited <2.71 Ga on Oxford–Stull domain basement record thefirst influx of >3.5 detritus (Corkery et al. 1992; Skulski etal. 2000).

The southern margin of the North Caribou superterranerecords a 300 million year history of rifting, arc magmatism,collision(s), overlap sedimentation, and faulting. The SydneyLake – Lake St. Joseph (SL–LSJ) fault (Fig. 6), traditionallydefined as the boundary with the English River terrane, re-cords only the late dextral movement on this complex zone.The steeply dipping, 1–3 km wide brittle–ductile fault zoneis estimated to have accommodated about 30 km of right-lateral transcurrent displacement and 2.5 km of south-side-upmovement (Stone 1981). Where crossed by seismic line 1a,

the fault is imaged as a steeply north-dipping discontinuitywith normal geometry (White et al. 2003). Bethune et al.(2000, 2006) determined a maximum age of 2.68 Ga for theMiniss River fault, which is cut and offset (�6 km) by theextension of the SL–LSJ fault in the region of the seismicprofile. Hrabi and Cruden (2006) inferred a still younger ageof (<2.646 Ga) for brittle movement on the fault.

At the southern margin of the English River terrane, ductiledeformation and high-grade metamorphism obscure originalcontact relationships with the Winnipeg River terrane. How-ever, depositional contacts have been inferred between EnglishRiver clastic rocks, and both volcanic strata of the SeparationLake greenstone belt (Hrabi et al. 2000; Hrabi and Cruden2006), and gneissic tonalitic basement to the east (Sanborn-Barrie 1988). Coarse siliciclastic rocks <2.701 Ga may uncon-formably overlie the boundary (Hrabi and Cruden 2006),which was the locus of emplacement for �2.646 Ga rare-metal-rich pegmatites (Larbi et al. 1999; Breaks and Tindle2002; Smith et al. 2004), including the Tanco and SeparationRapids fields (Blackburn and Young 2000).

The southern margin of the Winnipeg River terrane is acomplex boundary that evolved over a 250 million year pe-riod. The earliest manifestation is Mesoarchean bimodal vol-canism (2.93–2.88 Ga;) followed by rifting (>2.75 Ga) andNeoarchean continental arc magmatism (2.74–2.70 Ga)(Sanborn-Barrie and Skulski 1999; Whalen et al. 2002).Wabigoon oceanic domains were accreted to the margin be-tween 2.71 and 2.70 Ga with subsequent collisional orogen-esis (2.70–2.69 Ga) (Sanborn-Barrie and Skulski 2006;Percival et al. 2004a). Finally, the boundary became the lo-cus of ductile–brittle transpressive faulting (2.685–2.65 Ga).

The boundary between the composite Winnipeg River –Marmion – western Wabigoon terrane and Quetico terrane isgenerally defined by late dextral faults such as the SeineRiver – Rainy Lake fault (SR–RL) (Fig. 6). Early nappe-likestructures in the Rainy Lake area suggest early structuraltelescoping in the boundary zone (Poulsen et al. 1980). Thelate dextral Quetico fault cuts the SR–RL fault and forms theWabigoon–Quetico boundary farther east (Mackasey et al.1974). East of Lake Nipigon the boundary is a transitionzone with an early history of structural imbrication (Devaneyand Williams 1989; Tomlinson et al. 1996). The Wabigoon–Quetico interface is marked sporadically by <2.692 Ga coarseclastic rocks of the Seine assemblage (Fralick and Davis1999) that were deposited in transtensional basins (Blackburnet al. 1991) or delta fan environments (e.g., Fralick et al.2006).

An irregular boundary separates the Quetico from theWawa–Abitibi terrane to the south. Dextral transpressive shearzones active at �2.685 Ga define the boundary in several ar-eas (Corfu and Stott 1998); however, stratigraphic linkagesare evident in <2.696 Ga sedimentary overlap sequences inthe McKellar Harbour area (Fig. 6; Fralick et al. 2006) andin the Manitouwadge area to the east (Zaleski et al. 1999).

The Great Lakes tectonic zone (Fig. 6) is the unexposedboundary between the Minnesota River Valley terrane andWawa–Abitibi terrane, identified from aeromagnetic images(Sims and Day 1993). It is inferred to dip northward basedon the presence of isotopic inheritance in plutons of the Ver-milion district of the southern Wawa–Abitibi subprovince(Sims et al. 1997). White et al. (2003) postulated that the

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high-velocity lower crust at the southern end of line 1 (Fig. 4)is the remains of the ocean basin that separated the Minne-sota River Valley and Abitibi terranes.

History of tectonic assembly

In this section we review the timing and significance oftectonic events recorded throughout the western SuperiorProvince (Fig. 8). Five Neoarchean events in which tectonismcan be related to terrane juxtaposition are considered. Theevents may include some or all of the following processes:terrane collision, flysch and (or) molasse deposition, pene-trative regional fabric development, regional metamorphism,granitic magmatism, and strike-slip faulting. These eventshave been earlier defined as “phases” of the Kenoran orog-

eny (Stott 1997) but are here designated as spatially andtemporally discrete accretionary orogenies following the ap-proach of Stott and Corfu (1988) and based on comparisonsto younger orogens (see Discussion). The term “Kenoranorogeny” can be used in the sense of the 2.72–2.68 Ga tec-tonic dynasty (C.F. Gower, personal communication, 2004)during which assembly of the Superior Province took place.

Neoarchean tectonism that led to assembly of the SuperiorProvince began earliest in the north and continued progres-sively southward over a ca. 40 million year period. The pro-gressive assembly (Stott and Corfu 1991) of the compositeSuperior superterrane is illustrated in a time–space correla-tion diagram (Fig. 8), which summarizes the age range ofMesoarchean terranes, intervening tracts of Neoarchean vol-canic rocks, and the timing of assembly events. The spatial

Fig. 8. Time–space correlation diagram illustrating timing of assembly of the western Superior Province from continental and oceanicfragments. Note successive amalgamation of terranes between 2720 and 2680 Ma expressed through linking deformation events. BL–RL–BI, Bee Lake – Rice Lake – Black Island; v, volcanic rocks; p, plutonic rocks.

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Fig. 9. Schematic evolutionary model for accretionary growth of the western Superior Province. NSS, northern Superior superterrane;NCS, North Caribou superterrane; WwT, western Wabigoon terrane; WRT, southwestern Winnipeg River terrane; WAT, Wawa–Atibiterrane; MRVT, Minnesota River Valley terrane. (a) At 2.75 Ga, independent microcontinental fragments separated by tracts of oceaniccrust of unknown dimension. (b) By 2.72 Ga, the NSS had started its collision with the northern margin of the NCS to initiate thecomposite Superior superterrane. Continental magmatism continued within the central and southern NCS. A D1 deformationevent, �2.735 Ga at the southern NCS margin, may have led to initiation of the 2.732–2.718 Ga Black Island – Rice Lake arc–back-arcterrane to the south. The WwT begins to impinge on the WRT margin. (c) Around 2.7 Ga, ongoing convergence between the NCS andWRT plates results in orogenic deposition of English River sediments, their subsequent burial, and metamorphism. Convergence to thesouth results in WwT–WRT collision and continued magmatism related to independent subduction zones to the south. (d) Between2.70 and 2.69 Ga, the WAT docks with the composite Superior superterrane, accompanied by deposition of synorogenic Quetico flyschin the intervening trench, its burial, and metamorphism. Arc magmatism continues in the oceanic WAT and postorogenic graniticmagmatism is widespread across the composite Superior superterrane to the north. (e) At �2.68 Ga, the MRVT docks with the compositeSuperior superterrane, leading to deposition and burial of the Pontiac metasedimentary belt. The remnants of the final oceanic slab aresubcreted beneath the composite Superior superterrane.

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dimension is provided in a series of tectonic reconstructions(Figs. 9a–9e).

Northern Superior orogenyThis �2.72–2.71 Ga event united the Northern Superior

superterrane and North Caribou superterrane, trapping thepreviously imbricated Oxford–Stull domain of continentalmargin and oceanic affinity (Skulski et al. 2000). Subductionpolarity is inferred to have been southward, based on south-over-north shear-zone movement (Lin et al. 2006) and thepresence of arc magmatic activity between 2.775 and 2.733Ga across the Island Lake and Oxford–Stull domains (Parkset al. 2006; Corkery et al. 2000; Skulski et al. 2000). Thepattern is mirrored by steep slabs of high resistivity in themantle to 150 km depths, reflecting anisotropy of the mantleconsistent with lateral growth by crustal accretion (Fig. 3;Craven et al. 2004). The inferred suture zone is in the vicin-ity of the North Kenyon fault, which was later reactivated asa broad transcurrent structural zone. Detrital zircon agespectra in Oxford–Stull sedimentary rocks reflect tectonicisolation from northern Superior basement until after �2.72Ga D2 deformation that affected both the Oxford–Stull domainand northern Superior superterrane. Docking of the northernSuperior superterrane is recorded by the appearance of >3.5Ga detrital zircons in <2.711 Ga synorogenic sedimentaryrocks (Corkery et al. 2000). The tectonic event is alsomarked by eruption of �2.71 Ga shoshonitic volcanic rocks,which are preserved in strike-slip basins. Regional shorten-

ing was accommodated through D1 and D2 folds and folia-tion (Corkery et al. 2000), as well as development of promi-nent northwest-striking dextral shear zones (Lin et al. 2006).This collision represents initial formation of a composite Su-perior superterrane (CSS) (Percival et al. 2004b).

Uchian orogenyArc magmatism (2.748–2.708 Ga) across the Berens and

Uchi domains of the North Caribou superterrane was theprecursor to the Uchian orogeny wherein northward subductionled to collision between the �3 Ga North Caribou super-terrane and �3.4 Ga Winnipeg River terrane (Stott and Corfu1991; Corfu et al. 1995; Stott 1997). The suture zone be-tween the continental blocks appears to be mainly obscuredby sedimentary rocks of the English River terrane, whichwere deposited after 2.713–2.704 Ga and overridden as thecollision progressed. Structural elements of the suture maybe exposed in the Garner Lake – Lake Winnipeg corridor,where D1 transcurrent shear zones separate juvenile volcanicrocks from those of North Caribou affinity (Poulsen et al.1996; Anderson 2003; Percival et al. 2006).

The broad geometry of the Uchian orogen can be deducedfrom chronological and seismic constraints. Surface struc-tures are consistently steep and north dipping. Rocks formingthe south-central margin of the North Caribou superterrane(i.e., Red Lake – Birch – Uchi region) have undergone pene-trative deformation at �2.718–2.712 Ga (Andrews et al. 1986;Sanborn-Barrie et al. 2001; Dubé et al. 2004) followed by

Fig. 9 (concluded).

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emplacement of posttectonic plutons and initial cooling by�2.70 Ga (Corfu and Stone 1998b). Following this, rocks ofthe English River and Winnipeg River terranes underwentrapid burial and heating. Inferred southward overthrusting ofthe North Caribou superterrane onto the English River terraneis supported by gently north-dipping reflectivity on seismicreflection profiles (Fig. 4; White et al. 2003). Elsewherealong the southern margin of the North Caribou superterrane,arc magmatism continued until at least 2.71 Ga, followed bypenetrative deformation between 2.714–2.702 Ga in the east(Corfu and Stott 1993b) and after 2.704 Ga, in the west(Rogers and McNicoll, personal communication, 2005;Percival et al. 2006). The complex, diachronous history ofdeformation along this margin has yielded several modelswithin the context of microcontinent–continent collision.

(i) Deformation at �2.718 Ga marks subduction of an oce-anic terrane beneath the Uchi margin, followed bycrustal thickening and emplacement of �2.704 Gaposttectonic plutons (Percival et al. 2006). Subsequentcollision of the Winnipeg River terrane led to depositionof <2.704 Ga synorogenic sediments including the prox-imal San Antonio, Kangaroo, and Austin conglomeratesin the north, English River turbidites, and distal forelandsequences in the Wabigoon terrane to the south, fol-lowed by burial and metamorphism as the Uchi marginover-rode the Winnipeg River lower plate (Corfu et al.1995).

(ii) Deformation at �2.718 Ga marks the initial collision ofthe Winnipeg River and North Caribou superterranes,and consumption of the Winnipeg River lower platecontinued until �2.69 Ga. Magmatic and structuraldiachroneity could reflect an irregular North Cariboumargin. Syn-collisional deposition of English Riverturbidites occurred in a peripheral foreland basin(Sanborn-Barrie et al. 2004).

(iii) Development of the English River terrane occurred in aback-arc basin, with deformation by subsequent basininversion (Pan et al. 1999).

(iv) English River turbidites were deposited in a fore-arc toperipheral foreland basin setting, followed by extension(�2.701 Ga) and renewed compression (Hrabi andCruden 2006).

Several authors have remarked on the rapidity of burialand heating of English River sediments (Corfu et al. 1995;Hynes 1997; Pan et al. 1999). Rocks deposited after 2.704Ga were metamorphosed at �800 °C within 15 million years,less than the timescale for thermal relaxation following tec-tonic burial (e.g., England and Thompson 1986). This maybe partly due to magmatic heat input in the form of mantle-derived sanukitoid plutons (Corfu et al. 1995; Nitescu et al.2006). However, most plutonic rocks of the English Riverterrane are crustally derived granitic rocks, more likely theproducts of high crustal temperatures than the cause (Breaks1991). Additional factors may be relevant to the rapid attain-ment of high temperatures:

(i) Slab breakoff (cf. Sajona et al. 2000) could have led tosanukitoid magmatism and asthenospheric rise to subcrustallevels, providing a driving force for upward heatadvection;

(ii) Preconditioning of the North Caribou and Winnipeg

River terranes by arc magmatism prior to collision couldhave established hot upper and lower plates, leading tosimultaneous crustal thickening, burial, and heating;

(iii) Extension �2.70 Ga (D3 of Hrabi and Cruden 2006)could have led to elevated geotherms.

Resolution of these alternatives hinges on further work inthe English River terrane and fringing metasedimentary se-quences to better understand the timing and provenance ofdepositional events and subsequent tectonothermal history.For example, present data indicate a range of depositionalages (<2.718 to <2.701 Ga, Corfu et al. 1995; Stott et al.2002) and provenance (North Caribou and Winnipeg Riverterranes) in different locations within the English River terrane.

Central Superior orogenyElements of the central Superior Province were assembled

into a superterrane just prior to incorporation into the com-posite Superior superterrane. The Winnipeg River terranemay have begun to break up as early as �2.93 Ga, leading toformation of �2.775–2.72 Ga western Wabigoon oceanic crust.Ocean width has not been constrained and opinions vary sig-nificantly. Blackburn et al. (1991) inferred an in situ rift,whereas Sanborn-Barrie and Skulski (1999) regarded thewestern Wabigoon as an oceanic terrane. Interpretation ofthe deformation history is strongly influenced by the startingposition of the western Wabigoon with respect to the Winni-peg River margin. In the view of Blackburn et al. (1991), de-formation relates to inversion of the volcanic basin in theform of marginal thrust faults (e.g., Devaney 2000). Alterna-tively, the strain history could relate to a collision betweenthe western Wabigoon and Winnipeg River terranes, an in-terpretation favoured by Sanborn-Barrie and Skulski (1999,2006) and Percival et al. (2004a), although uncertainty re-mains as to the age and nature of collision. In the model ofSanborn-Barrie and Skulski (2006), the western Wabigoonforms the lower plate, subducting north and eastward beneaththe Winnipeg River terrane and giving rise to 2.715–2.70 Gatonalite and associated intermediate pyroclastic rocks. Theturbiditic Warclub assemblage represents an interveningfore-arc succession deposited between 2.711 and <2.703 Gaon the Winnipeg River margin just prior to collision, and re-gional deformation results from over-riding of the westernWabigoon by the Winnipeg River terrane. In contrast, south-west-dipping subduction was inferred by Davis and Smith(1991), Melnyk et al. (2006), and Percival et al. (2004a), andin these models collision occurred prior to 2.710 Ga formingductile fabrics in the Winnipeg River lower plate and openfolds in the overriding western Wabigoon plate (cf. Edwardsand Stauffer 1999). In these models, subsequent events thataffected the composite Winnipeg River – western Wabigoonsuperterrane include 2.71 Ga tonalitic magmatism attributedto continued northward subduction of the Wawa–Abitibiplate from the south, and deposition of the Warclub sedi-ments, the distal equivalents of English River flysch, andtheir deformation during the Uchian orogeny.

Shebandowanian orogenyThis event brought the Abitibi–Wawa terrane into juxtapo-

sition with the composite Superior superterrane at �2.695 Ga(Corfu and Stott 1986, 1998; Stott 1997). Subduction polar-

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ity is inferred to have been to the north, based on the cessa-tion of arc-related magmatic activity at approximately 2.695Ga in the composite Wabigoon – Winnipeg River terrane andemplacement of 2.695–2.685 Ga sanukitoid plutons (e.g.,Whalen et al. 2004a), which may signal slab breakoff (cf.Sajona et al. 2000). The suture is inferred to be beneath theQuetico terrane.

Influx of clastic sediments into the northern Quetico be-gan at 2.698–2.696 Ga (Davis et al. 1990; Davis 1998;Fralick et al. 2006) and continued to <2.69 Ga (Zaleski et al.1999; Fralick et al. 2006). The ages of sedimentation sup-port an accretionary wedge to foreland basin transition (cf.Percival and Williams 1989; Williams et al. 1992; Valli et al.2004). Two deformation events are recognized within thenorthern Wawa–Abitibi terrane. A �2.695 Ga D1 event is as-sociated with calc-alkaline magmatism and probably relatesto intra-arc deformation (Corfu and Stott 1998). Docking be-tween the Wawa–Abitibi and Quetico terranes is indicatedby transgressive sedimentary assemblages (Fralick et al. 2006)and common transpressive deformation (D2), which affectedboth subprovinces between 2.685 and 2.680 Ga (Corfu andStott 1998). Alaskan-type ultramafic (Pettigrew 2004) andalkaline magmatism (�2.68 Ga) in the Quetico may relate tobreakoff of the leading edge of the Wawa–Abitibi slab.

Seismic profiles across the boundary zone indicate gentlynorthward-dipping reflectivity (White et al. 2003). The Queticosubprovince is poorly imaged but does not appear to projectnorthward into the southern Wabigoon (Fig. 4), an inferencein accord with the lack of peraluminous granitic plutons inthe southern Wabigoon – Winnipeg River superterrane. Theseobservations are consistent with accretion rather than subductionof Quetico sedimentary material.

Minnesotan orogenyThis �2.68 Ga event is responsible for collision between

the ancient Minnesota River Valley terrane and the compositeSuperior superterrane. An additional terrane, the <2.682 Ga(Mortensen and Card 1993; Davis 2002) Pontiac metase-dimentary belt, intervenes between the two in the easternSuperior Province. There, the polarity of subduction has beeninferred to be northward, based on north-dipping seismic re-flectivity (Calvert and Ludden 1999) and the presence ofperaluminous granite in the southern Abitibi (Feng and Kerrich1991, 1992; Chown et al. 2002). Similar northward polarityis indicated in the west by reflection geometry (Fig. 4; Whiteet al. 2003) and by the isotopic signature of old crust be-neath the southern Wawa–Abitibi terrane (Sims et al. 1997).The unexposed Great Lakes tectonic zone (Fig. 6) is theprobable suture (Sims and Day 1993).

Deformation at �2.68 Ga within the Wawa–Abitibi, Pon-tiac, and Minnesota River Valley terranes is attributed to theMinnesotan orogeny. Previous tectonic models had viewedthe Minnesota River Valley terrane as a rigid block, analo-gous to the jaw in “vice” models (Ellis et al. 1998). How-ever, in light of reflection images, the terrane appears to bethe lowest structural level in a thrust stack. It projects down-ward into the high-velocity lower crustal layer identified onrefraction (Musacchio et al. 2004) and reflection (White etal. 2003) profiles and interpreted as a slab of oceanic crust(Fig. 4). Tectonic subcretion of this material at �2.68 Gamay account for the rapid cooling and uplift of the Winnipeg

River terrane between the eastern and western Wabigoonterranes at that time (Percival et al. 2004a).

Post-cratonization events

Late orogenic to postorogenic effectsWhere comprehensive geochronological control is avail-

able, most deformation, metamorphic, and crustal meltingevents in individual terranes of the Superior Province followa predictable pattern within a plate collisional tectonic frame-work. The following sequence of events generally occurswithin a �20 million year period: cessation of arc magmatism;deposition of flysch-like sedimentary rocks; intrusion ofsanukitoid plutons; deformation and burial; metamorphism;and emplacement of crust-derived granitic plutons. Late-stagemagmatic and thermal events that outlast these discrete orogenicpulses are widely recorded across the Superior Province. Forinstance, small-volume posttectonic granites and pegmatitesof crustal derivation were emplaced between 2.66 and 2.64Ga, up to 70 million years after defined collisions (e.g.,Corkery et al. 1992; Corfu et al. 1995; Smith et al. 2004). Atapproximately the same time, metamorphic and hydrother-mal activity, recorded in zircon, monazite, and titanite growth,occurred in deeply eroded crustal terranes (Corfu 1988; Krogh1993; Corfu et al. 1995), as well as in some lower-grade en-vironments (e.g., Davis et al. 1994). Several gold depositshave evidence for hydrothermal activity of comparable age(e.g., Jemielita et al. 1990; Zweng et al. 1993; Krogh 1993).Extensional deformation and metamorphism between 2.66and 2.45 Ga in the deep crust exposed in the Kapuskasinguplift (Moser et al. 1996) reflect high temperatures at thistime. Widespread resetting of the Rb–Sr isotopic system(e.g., Beakhouse et al. 1988) reflects a fluid-related distur-bance. Similarly, Ar–Ar ages for hornblende and biotite aregenerally significantly younger than U–Pb ages (e.g., Hanesand Archibald 2001), indicating open system behaviour forsome minerals.

Several theories have been advanced to explain these late-tectonic to posttectonic effects, including continued subductionand tectonic underplating (Krogh 1993), magmatic under-plating (Zweng et al. 1993), and repeated delamination events(Moser et al. 1996). The absence of observed mafic mag-matic rocks of this age, coupled with lower crustal seismicvelocities in the 7 km·s–1 range, does not appear to supportwidespread mafic underplating. Although, it could be main-tained that such rocks were subsequently delaminated, Archeancratons generally have buoyant lithospheric keels (e.g.,Poudjom Djomani et al. 2001) that may have been presentsince the time of crust formation (e.g., Griffin et al. 2003).Recently, Percival and Pysklywec (2004) suggested that theobservations can be reconciled through a process of lithosphericinversion, which would have transported �1300 °C lowerlithospheric mantle into juxtaposition with the lower crust,causing widespread metamorphism, melting, and fluid release.

Paleoproterozoic eventsThe Superior Province is transected by at least 20 swarms

of diabase dykes, some of which are related to incipient Pro-terozoic breakup of the craton (Buchan and Ernst 2004). Theoldest dykes (2.502 Ga, Buchan et al. 1998) attest to cratonstability by that time. However, work by Halls and Davis

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(2004) and Halls (2004) on the paleomagnetic record of Pro-terozoic dyke swarms shows that the western half of the Su-perior Province has rotated counterclockwise by about 20°relative to the eastern half, across the Kapuskasing zone.The precise age of the deformation is unknown, but Paleo-proterozoic reactivation has been documented in several lo-cations (e.g., Peterman and Day 1989; Kamineni et al. 1990;Percival and Peterman 1994; M.J. Harris et al. 2006).

Significant internal deformation of the Superior craton isrecorded in the Kapuskasing uplift of the central SuperiorProvince. This �1.9 Ga event involved about 25 km of east–west shortening and strike-slip displacement that resulted inbrittle uplift of lower-crustal material expressed as promi-nent geophysical anomalies including crust >50 km thick(Percival and West 1994). The intracratonic deformation isthought to be a far-field effect of Paleoproterozoic collisionat the margin of the Superior Province.

A second major incursion into stable Archean lithosphereof the Superior Province occurred in the midcontinent rift(1.11–1.09 Ga, Davis and Green 1997). This arcuate zonerepresents a �20 km wide rift filled with basalt and clasticsedimentary rocks. Associated gabbroic sills, lamprophyredykes, and alkalic rock–carbonatite complexes intrude Supe-rior Province crust to the north and east of Lake Superior.

Discussion

Several first-order characteristics of the western SuperiorProvince may be cited collectively in support of the exis-tence of a plate tectonic regime during the Neoarchean. Thesefeatures include(i) the presence of discrete continental and oceanic domains,

a consequence of continental breakup and seafloor for-mation in the modern tectonic framework;

(ii) the existence of juvenile mantle-derived magmas includ-ing calc-alkaline basalt and sanukitoid-suite rocks, bear-ing evidence of LREE, LILE enrichment. These traceelement and isotopic characteristics are difficult to ex-plain through processes other than mantle metasomatismby fluids or melts in suprasubduction zone environments;

(iii) a history of five, southward propagating, temporally dis-crete orogenic events over the 40 million year period be-tween 2.72 and 2.68 Ga (Fig. 8) — a characteristic oflateral accretion; (iv) orogenic belts with length scales>1000 km, comparable to some modern plate margin di-mensions;

(v) calc-alkaline granitoid batholiths with dimensions andcompositions comparable to those of modern continen-tal magmatic arcs such as the South Patagonian batholith;

(vi) long strike-slip faults, indicating lateral movement (e.g.,Sleep 1992); and (vii) gently dipping crustal panels andMoho offsets, teleseismic and magnetotelluric patternscharacteristic of modern accretionary orogens.

Collectively, these observations provide compelling evi-dence that the Superior Province evolved at �2.75–2.68 Gathrough plate tectonic processes akin to those active today.Significant differences in tectonic style are examined in thefollowing text.

Structural StyleSteeply dipping foliation and steeply plunging folds domi-

nate most domains of the western Superior Province. In manyregions, these fabrics trend east–west or northwest and arethe youngest penetrative structures in a polyphase chronology.Their nature has generally been attributed to transpressivestrain late in the shortening history (e.g., Williams et al.1992; Stott 1997; Parmenter et al. 2006). Evidence forearlier thrusting has been obtained through detailed geo-chronological studies of greenstone belts (e.g., Davis et al.1988; Corfu and Ayres 1991; Corfu and Stott 1993b) wherethe structures responsible for stratigraphic repetition are gen-erally not obvious. Accordingly, this style of deformationcould be more common than currently recognized. For ex-ample, early recumbent folds have been noted in the south-ern Wabigoon (Poulsen et al. 1980) and Quetico (Sawyer1983) subprovinces. Other early (D1) structures appear tohave formed in upright orientations. In the Red Lake andSturgeon Lake belts, F1 folds have steep plunges (Sanborn-Barrie et al. 1998, 2001; Sanborn-Barrie and Skulski 1999,2006). Similarly the D1 shear zone at Lake Winnipeg formedas a steep transcurrent structure (Percival et al. 2006).

The early (�2.73 Ga) north-northwest-trending structuresin the Red Lake and Confederation Lake belts are anoma-lous in light of their belt-scale extent and implication ofeast–west shortening. Similarly, the first folds in <2.704 Gasedimentary rocks of the English River terrane are also ori-ented north-northwest (Hynes 1997). This recurring patterncould reflect structures defining the margins of early nappesor parautochthonous sheets, overprinted by subsequent fold-ing and shortening (cf. Stott and Corfu 1991).

Erosion levelsEvidence from field relationships and seismic images sug-

gest that north–south shortening, accommodated by crustalstacking, folding, penetrative flattening, and localized shear-ing accounts for regional orogenic activity. However, erosionlevels are modest but variable throughout most the westernSuperior Province, ranging from minimum levels of �8 kmin some greenstone belts to more than 20 km in EnglishRiver granulites. From present Moho depths on the order of�40 km, maximum crustal thicknesses in the range 45–60 km can be inferred; values that are significantly less thanthose expected to result from continental doubling throughHimalayan-style orogenesis. Extensional faults showing ap-preciable offset are rare. Several possible explanations arise,(i) the panels that were stacked were thin, such that the ag-

gregate crustal thickness never exceeded 60 km, as inthe Appalachians and Canadian Cordillera. Possibilitiesinclude crustal flakes or immature island-arc crust thatmay have been <20 km thick;

(ii) the continental lithosphere was mechanically weak atthe time of collision as a result of magmatic precondi-tioning and was therefore unable to support much topo-graphic load. A steady state may have prevailed in whichshortening produced orogen-parallel extensional flow ratherthan significant topographic expression. Some and perhapsmost of the flow may have been accommodated in theductile deep crust as was inferred beneath the centralWawa–Abitibi terrane (Moser et al. 1996);

(iii) erosion rates were higher than at present owing to harshatmospheric conditions that promoted rapid exhumation.

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Repeated orogenyEvidence reviewed here supports a model whereby the Su-

perior Province was assembled over a �40 million year pe-riod through five separate accretionary orogenic events. Aconsequence of the short polyorogenic history is widespreadstructural–metamorphic overprints in different parts of theSuperior Province. Based on regional patterns of gaps be-tween zircon and titanite U–Pb ages (Corfu 1988; Corfu andStone 1998b), it is apparent that terranes metamorphosedduring early orogenies did not return to thermal equilibriumstates before acting as the foreland or hinterland in the sub-sequent orogeny. These areas were, therefore, susceptible toreworking and so recorded complex polyphase deformation.Because some of the deformation episodes were coaxial,identifying multiple generations of structures has provedchallenging in some areas.

Tectonic analoguesNo single analogue encompasses the complex set of tec-

tonic interactions that developed during the five orogeniesrecognized within the western Superior Province. However,many elements of the history are comparable to that of thePhilippine Sea region, where interaction among the large Pa-cific, Eurasian, and Indian–Australian plates from the Creta-ceous to the present day has driven rotation, strike-slipfaulting and intermittent subduction, arc formation, back-arcspreading, and thrusting within smaller plates (see Fig. 10 ofHall 2002; cf. Card 1990; Jackson and Cruden 1995; de Wit1998). Similar analogies can be drawn with the Paleozoicevolution of the Laurentian margin in the Appalachians (e.g.,van Staal et al. 1998).

The Uchian orogeny united continental rocks of the com-posite Superior superterrane (Northern Superior and NorthCaribou superterranes) and the Winnipeg River terrane, lead-ing to the first (English River) of three periods of orogenicsedimentation (Davis 1996a). The Winnipeg River may cor-respond to the Bird’s Head microcontinent, thought to be afragment of the Australian margin (Hall and Wilson 2000)that lies among active arcs in the southern Philippine Sea. Inthis analogy, the backstop, corresponding to the compositeSuperior superterrane, is the Eurasian plate at the edge of thewestern Pacific (Hall 2002).

An analogue for Wabigoon arc – Winnipeg River terraneassembly as recorded by the central Superior orogeny maybe the ongoing collision between the Banda arc and Indian–Australian plate (e.g., Snyder et al. 1996; Hall and Wilson2000). The highly arcuate form of the Banda arc matchesthat of the western Wabigoon margin, as do isotopic patternsreflecting increasing continental input to arc magmatism(Vroon et al. 1993; Davis et al. 2000). Following the centralSuperior orogeny, the Wabigoon – Winnipeg River super-terrane became the upper plate with respect to the Wawa–Abitibi plate and underwent renewed arc magmatism (Percivalet al. 2004a).

The locus of the �2.69 Ga Shebandowanian orogeny tothe south exhibits many of the same characteristics as the�2.72–2.70 Ga Uchian orogeny, including collisional geom-etry and a trapped sedimentary prism that was buried andheated rapidly. A significant difference is the juvenile natureof the arc on the lower plate. The setting for this arc–arccollisional orogeny resembles that of the Hidaka collision

zone of Hokkaido, where the Kuril fore-arc collided with theNortheast Japan arc across the Horobetsugawa accretionarycomplex in the Cretaceous through Miocene. Common fea-tures include steep bedding and foliation, late strike-slip strain,and metamorphism to granulite facies of parts of the clasticprism (Ueda et al. 2001). Similar crustal velocity profiles,deep crustal layering, and Moho characteristics are also evi-dent (Iwasaki et al. 2002).

The Minnesotan orogeny corresponds to the �2.68 Ga ac-cretion of the Minnesota River Valley terrane to the compos-ite Superior superterrane, and as the final recorded event,could represent the terminal collision. This would be analo-gous to collision of the Indian–Australian plate with the Eur-asian plate. Late strain, in the form of transcurrent faultingwithin the Wawa–Abitibi subprovince, may be a far-field ef-fect of the Minnesotan orogeny.

Outstanding questions and research opportunities

Events of unidentified significanceTectonometamorphic events of limited areal extent or sin-

gle-parameter observations without supporting context areindicated in both the Mesoarchean and Neoarchean record.For example, evidence for �2.94 Ga tectonism from thesouthwestern North Caribou superterrane comes from datedshear zones (Percival et al. 2006). Tectonic inversion ofsupracrustal rocks occurred prior to 2.92 Ga in the same re-gion (Sasseville et al. 2006). Within the central North Cari-bou superterrane, the 2.87 Ga North Caribou pluton cutsdeformed 2.98 and 2.93 Ga assemblages (Stott et al. 1989;Thurston et al. 1991). Within the Winnipeg River terrane,2.92 Ga metamorphism may be a contact metamorphic effectof 2.92 Ga tonalite plutonism (Melnyk et al. 2006) or mayhave been induced by a regional deformation event prior tothis time. Samples of �3.0 Ga Marmion tonalite also yieldtitanite with an estimated age of �2.81 Ga (Davis and Jack-son 1988), the significance of which is unknown in terms ofcooling or deformational events.

Neoarchean tectonometamorphic events not directly tiedto orogenic events described previously in the text includemetamorphic zircon overgrowths with ages in the 2.75 Garange in tonalite from the Northern Superior superterrane(Skulski et al. 2000). Overturning of strata in the Red Lakeand Pickle Lake belts, prior to deposition of continental arcvolcanic rocks of the �2.748–2.735 Ga Confederation as-semblage (Sanborn-Barrie et al. 2004; Young et al. 2006),support the hypothesis of an early deformation event (Stottand Corfu 1991), the extent and significance of which is notyet understood. Similarly, �2.735 Ga belt-scale D1 deforma-tion recorded at Red Lake (Sanborn-Barrie et al. 2001) mayrelate to precollisional intra-arc adjustments or to plate reor-ganization (Sanborn-Barrie et al. 2004; Percival et al. 2006).

Regional questionsThe dominantly linear architecture of the western Supe-

rior Province lends itself to accretionary orogen interpreta-tions. However, the “terrane” structure is less obvious east ofHudson and James bays, and broadscale correlations are im-perfect. For example, the English River and Quetico beltswest of James Bay are discrete tectonic entities separated bythe western and eastern Wabigoon terranes, whereas, east ofJames Bay they appear to merge into a single Opinaca belt

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(Stott and Berdusco 2000). Better understanding of isotopicdomains, correlative units, and timing of sedimentation is re-quired before the tectonic significance of this observationcan be considered in a regional context. A corollary questionconcerns the diminishment of the proportion of supracrustalbelts in the eastern Superior compared with the west. It iscurrently unknown whether this observation relates todeeper erosion levels in the east or to different tectonic envi-ronments.

The eastern extension of the Great Lakes tectonic zone,proposed as the suture between the Minnesota River Valleyand Abitibi–Wawa terranes, is unknown. Ancient crust hasnot been reported from the southeastern Superior Province,either autochthonous or in the Grenville Province.

The broad concave shape defined by the tectonic grain ofthe northeastern Superior Province (Fig. 1b; cf. Percival etal. 1994, 2001) should be reevaluated in light of the discov-ery of the northern Superior superterrane (see Leclair et al.2004). For instance, a region of ancient crust positioned inthe central part of the arc may be floored by very old, stablelithospheric mantle and be of interest to the diamond explo-ration community.

Broader considerationsConsiderable progress has been made in understanding the

correlation between seismic reflectors and geological struc-tures as a result of coordinated surface and subsurface stud-ies. Some uncertainty remains in bridging the gap betweensteep surface structures and the appearance of subhorizontalseismic structures at �3 km depth. Explanations range froma universally listric nature of structures to strain-partitioneddomains and different structural generations at different crustallevels.

As in the exploration industry, targeted drilling may be auseful way to test integrated geological–geophysical inter-pretations. From the accumulated Lithoprobe data sets, thecommunity could formulate a plan to enhance interpretativepower through selective drilling of representative seismicfeatures located at shallow depth.

Conclusions

Knowledge of the Superior Province has increased overthe past two decades as a result of new mapping and associ-ated acquisition of modern information. The tectonic frame-work for the western Superior Province as an accretionaryorogen (Goodwin 1968; Langford and Morin 1976; Card1990; Williams et al. 1992; Stott 1997) has been refinedthrough the collaborative NATMAP and Lithoprobe programs,augmented by independent research activity. New geophysi-cal images of the lithosphere, coupled with structural, geo-chemical, and geochronological information, provide the basisfor a four-dimensional interpretation of the western SuperiorProvince over its 1.3 billion year geological evolution.

Three-dimensional seismic images of the crust illustratecontinuous north-dipping reflectors beneath the first-ordersubprovince structure. The deep structure has been inter-preted as a stack of discrete, �10–15 km-thick terranes. Theseinclude both microcontinental and oceanic terrane types thatwere amalgamated into the composite Superior superterrane

between 2.72 and 2.68 Ga and variably reworked bymagmatism and metamorphism.

At the surface, several microcontinental blocks with inde-pendent geological histories outlined by U–Pb geochronol-ogy and tracer isotopic studies are separated by terranesdominated by juvenile volcanic rocks. The northern Superiorsuperterrane contains some of the oldest (>3.8 Ga) rocks ofthe Superior Province and may constitute an ancient nucleus.It is bound by late transcurrent shear zones from the NorthCaribou superterrane to the south. This juvenile, �3.0 Gamicrocontinent appears to have undergone extension, as re-corded by widespread komatiitic sequences, during �2.98Ga and later rifting events and underwent reworking in acontinental magmatic arc setting along both its northern andsouthern margins between �2.75 and 2.70 Ga. Similarly, ex-tensive continental magmatism affected the Winnipeg Rivermicrocontinental terrane to the south, which bears �3.4 Gaancestry. The North Caribou, Winnipeg River, and westernWabigoon terranes were assembled and English Riverturbidites deposited in the period between 2.72 and 2.70 Ga.Two orogenies, the Uchian and central Superior, were proba-bly responsible for polyphase deformation and greenschist togranulite facies metamorphism in what may have been a rap-idly evolving tectonic system. Accretion along the southernmargin of the composite Superior superterrane led to additionof the Quetico sedimentary prism and juvenile Wawa–Abitibiterrane at �2.69 Ga, followed by terminal collision (�2.68 Ga)with the Minnesota River Valley microcontinental block with3.5 Ga ancestry. Processes associated with final 2.68–2.60Ga “cratonization” include transcurrent faulting, deep-crustalmetamorphism and deformation, generation and emplace-ment of crust-derived granites, circulation of hydrothermalfluids, and formation of some lode gold deposits.

With increasing precision on ages of deformation and as-sociated tectonic events, the concept of a single NeoarcheanKenoran orogeny to explain the evolution of the SuperiorProvince has been replaced by definition of several tempo-rally and spatially discrete orogenies. Recognition of theNorthern Superior (2.71 Ga), Uchian (�2.72–2.70 Ga),Central Superior (2.71–2.70 Ga), Shebandowanian (2.69 Ga),and Minnesotan (2.68 Ga) orogenies and their orderly northto south progression, coupled with crustal-scale images ofnorth-dipping structures, lends support to the hypothesis ofaccretionary growth of the Superior Province driven by pro-cesses akin to modern plate tectonics. Challenges remain inextending the western Superior framework to other parts ofthe Superior Province.

Acknowledgments

Our understanding of the tectonic evolution of the westernSuperior Province is founded on the input and cooperationof many individuals and organizations. The Western Supe-rior NATMAP project involved the Geological Surveys ofCanada, Ontario, and Manitoba, and the Lithoprobe projectoperated with the support of these organizations as well asmajor funding from the Natural Sciences and EngineeringResearch Council (NSERC). It is a pleasure to acknowledgethe financial and intellectual contributions of participatinguniversities, including Alberta, Lakehead, Laurentian, Mani-toba, McGill, Ottawa, Quebec at Montreal, Queen’s, Sas-

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katchewan, Toronto, Waterloo, and Windsor, and the JackSatterly Geochronology Laboratory, including scientists andstudents too numerous to name. Special thanks are due toDon Davis and Maarten de Wit for their thorough and con-structive reviews of the submitted manuscript, and to RonClowes for substantive editorial guidance.

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