tectonic evolution of manitoba-saskatchewan portion of the tho - ansdell 2005

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Tectonic evolution of the Manitoba–Saskatchewansegment of the Paleoproterozoic Trans-HudsonOrogen, Canada1, 2

Kevin M. Ansdell

Abstract: Time slices and schematic cross-sections that attempt to show the spatial and temporal relationship betweengeological entities within the Manitoba–Saskatchewan segment of the Trans-Hudson Orogen and that are consistentwith the available geological, geophysical, geochemical, isotopic, and geochronological data are presented. The Trans-Hudson orogenic belt developed as a result of closure of the Manikewan Ocean, which initially opened at about 2.1 Gaby rifting of a possible Neoarchean supercontinent. The oldest oceanic arc rocks indicate that subduction was ongoingby 1.92 Ga, with the development of a complex Manikewan “Ring of Fire” that lasted for the next 100 Ma.Intraoceanic accretion of arc, ocean-floor, and ocean-island rocks within the Manikewan Ocean at 1.87 Ga formed theFlin Flon – Glennie complex, which then subsequently collided with the accreted terranes along the Hearne cratonmargin at ca. 1.85 Ga. These rocks were then deformed and metamorphosed over the next 75 Ma during collisionswith the Sask craton and the Superior craton, both of which are interpreted to have been drifting generally northwardstowards the Hearne craton. The generation of arc magmas in the orogen ceased at 1.83 Ga, an indication that continentalcollisions were well advanced at that stage. The present arrangement and erosion level of geological entities is relatedto structural reorganization after the peak of regional metamorphism at ca. 1.81. The schematic time slices and sectionsform part of ongoing efforts to better understand the geological evolution of the Paleoproterozoic of Canada.

Résumé : Cet article présente des tranches de temps et des coupes transversales schématiques qui tentent de monter lesrelations spatiales et temporelles entre les entités géologiques à l’intérieur du segment Manitoba–Saskatchewan del’orogène trans-hudsonien et qui concordent avec les données géologiques, géophysiques, géochimiques, isotopiques etgéochronologiques disponibles. Le développement de la ceinture orogénique trans-hudsonienne résulte de la fermeturede l’océan Manikewan qui, à l’origine, s’était ouvert vers 2,1 Ga par la distension d’un possible supercontinent néo-archéen.Les plus anciennes roches d’arc océanique indiquent que la subduction était en cours vers 1,92 Ga avec le développementd’un « anneau de feu » Manikewan complexe qui a duré 100 millions d’années. L’accrétion intra-océanique de rochesd’arc, du fonds océanique et d’îles océaniques à l’intérieur de l’océan Manikewan à 1,87 Ga a formé le complexe deFlin Flon – Glennie, lequel est ensuite entré en collision avec les terranes accrétés le long de la bordure du craton Hearneil y a environ 1,85 Ga. Ces roches ont ensuite été déformées et métamorphosées au cours des 75 millions d’années quiont suivi, durant des collisions avec le craton Sask et le craton Supérieur; tous deux auraient dérivé en direction généralenord vers le craton Hearne. La génération de magmas d’arc dans l’orogène a cessé à 1,83 Ga, une indication que lescollisions continentales étaient bien avancées à ce stade. Les positions et le niveau d’érosion actuels des entités géologiquessont reliés à la réorganisation structurale après la crête de métamorphisme régional vers 1,81 Ga. Les tranches detemps et les sections schématiques forment une partie des efforts en cours pour mieux comprendre l’évolution géologiqueau Paléoprotérozoïque au Canada.

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Introduction

Unravelling the evolution of Earth is the main aim of geo-logical studies and requires integration of lithological andstructural mapping with geochemistry, geophysics, and geo-chronology. The formulation of plate tectonics provided a

mechanism for the generation, movement, and deformationof the Earth’s crust that has implications for the distributionand topographic features of continental land masses overtime. The position of continents affects oceanic circulationpatterns, climate, and the distribution of organisms and sedi-mentary basins, whereas the location of plate boundaries is

Can. J. Earth Sci. 42: 741–759 (2005) doi: 10.1139/E05-035 © 2005 NRC Canada

Received 12 November 2003. Accepted 8 March 2005. Published on the NRC Research Press Web site at http://cjes.nrc.ca on20 June 2005.

Paper handled by Associate Editor R.M. Clowes.

K.M. Ansdell. Department of Geological Sciences, University of Saskatchewan, 114 Science Place, Saskatoon, SK S7N 5E2,Canada (e-mail: [email protected]).

1This article is one of a selection of papers published in this Special Issue on The Trans-Hudson Orogen Transect of Lithoprobe.2Lithoprobe Publication 1403.

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an important control on volcanic activity, mountain building,and the formation of natural resources. The distribution ofcontinental crust and plate boundaries can be observed today,but their location in the past requires paleomagnetic andgeochronological control of specific geological features thatare known or thought to be associated with plate margins orinteriors.

The aim of this paper is to integrate available geologicaldata to reconstruct the location of geological entities in-volved in the development of the Manitoba–Saskatchewansegment of the Trans-Hudson Orogen (MS-THO) in centralCanada (Fig. 1). This region is aerially insignificant on aglobal scale, but it forms part of an extensive Paleopro-terozoic orogenic belt that has remnants in North Americaand northern Europe (Fig. 1C), and it preserves evidence forsignificant crustal growth and recycling at that time (Hoffman1988; Zhao et al. 2002). In addition, a voluminous data set,which has resulted from Lithoprobe’s Trans-Hudson OrogenTransect (THOT) and the Geological Survey of Canada-ledNATMAP (National Geoscience Mapping Program) ShieldMargin project (this entire issue; Clowes et al. 1999; Lucaset al. 1999a, 1999b), can be used to construct paleotectonicmaps for specific time periods in the Paleoproterozoic devel-opment of this region. Because the rocks within the orogenhave been metamorphosed and multiply deformed during theTrans-Hudson orogeny, the reconstruction of these time slicesis fraught with difficulty. Nevertheless, they represent a use-ful way of showing geological events that were occurring atthe same time and they force us to examine the evidence forthe spatial relationship among different rock units, some ofwhich remain problematic. Ultimately, these paleotectonicmaps can be linked to other similar maps to produce betterreconstructions of the Paleoproterozoic world.

General geological relationships

The MS-THO represents the southwestern part of theexposed THO in the Canadian Shield. The THO extendsnortheastwards into Nunavut and northern Quebec, as it wrapsaround the Archean Superior craton, and is interpreted toextend southwards into the United States below the Phanero-zoic sedimentary rocks of the Williston basin (Fig. 1B). TheTHO may be cut by the Central Plains orogen (Sims andPeterman 1986) or may extend as far west as northern Arizona(Hill and Bickford 2001). To the west and north of theMS-THO are the Archean rocks of the Hearne and Raecratons, which have been substantially reworked during Paleo-proterozoic orogenesis.

The MS-THO is subdivided into three zones: HearneProvince, Superior Boundary Zone, and the Reindeer Zone(Fig. 1A). The latter consists of volcanic, intrusive, and sedi-mentary rocks that developed in a variety of plate tectonicenvironments ranging from island-arc and back-arc to ocean-island settings and active continental margins. The first twoconsist of sedimentary and volcanic rocks that were laiddown unconformably on the margins of the Archean Hearneand Superior cratons, respectively. All were subsequentlymetamorphosed and deformed during the Trans-Hudsonorogeny. All the rocks of the THO have been deformed andmetamorphosed to varying degrees as a result of collisionsbetween juvenile Paleoproterozoic terranes and the Archean

Superior, Hearne, and Sask cratons. The Sask craton is exposedin two structural windows in the southern part of the MS-THO,underlies much of the Reindeer Zone, and probably formsmuch of the Precambrian basement underneath the Phanerozoiccover to the south.

Time-slice construction

The schematic paleotectonic diagrams shown in this paper(Fig. 2) are constructed by considering the huge volume ofgeological, geochemical, and geochronological data that hasbeen generated in the last 20 years in the MS-THO. Thesedata are used to update orogen-scale paleotectonic recon-structions developed by, for example, Lewry (1981b), Greenet al. (1985a, 1985b), Ansdell et al. (1995), and Clowes etal. (1999). The time period designated for each paleotectonicmap is constrained by the available U–Pb geochronologicalcontrol on geological events (short time periods indicate awell-constrained time period) and on the occurrence of sig-nificant geological events that indicate a change in tectonicscenario. Ansdell et al. (2005) summarize the geochronol-ogical control on geological events in the MS-THO and pro-vide some of the important references for geochronologicaldata. Scale and orientation are not provided, although thereis a voluminous set of paleomagnetic data available for theorogen (Symons and Harris 2005). Constraints provided bypaleomagnetic data for the various diagrams will be describedlater in the text. The paleotectonic diagrams are constructedassuming that the Hearne craton represents a fixed block inspace. Movements are then considered relative to this fixedblock. In addition, it is assumed that plate tectonic processeswere operating in a similar manner as observed at the presentday and that the three Archean cratons (Superior, Hearne,Sask) were located on their own plates prior to the finalstages of continental collision. Finally, geological studies inthe MS-THO are ongoing because there are a number ofunresolved issues, and thus there is an element of personalbias in the construction of the time slices.

Formation of the Manikewan Ocean

The MS-THO preserves the remnants of Paleoproterozoicvolcanic and sedimentary rock units that were deformed andmetamorphosed as Archean cratons collided. Convergent plateboundaries must have been in existence to generate thesecollisions, and by inference there must have been oceaniccrust which was subducted to generate arc magmas (e.g.,Flin Flon domain; Stern et al. 1995a, 1995b). This ocean ba-sin, which closed as the older continents collided, wasnamed the Manikewan Ocean by Stauffer (1984). However,when and how did this ocean basin develop?

The development of new ocean basins initially requiresthe rifting of older continental blocks, although whether thisoccurs as a result of passive lithospheric stretching or activedoming above a mantle plume remains controversial (e.g.,Storey et al. 1992). Distinguishing between these two sce-narios is often difficult, but there should be differences inrelative timing, volume, and composition of mafic magmaassociated with the rifting event (White 1992). If rifting issuccessful, then the new ocean basins will have sedimentarysequences forming on the margins of the bounding conti-

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Fig. 1. (A) Simplified geological map of the Manitoba–Saskatchewan segment of the THO showing the location and numbers ofLithoprobe seismic reflection lines discussed in the text. The Flin Flon and Glennie domains are termed the Flin Flon – Glennie complex(Ashton 1999). HL, Hanson Lake block; ORSZ, Owl River shear zone. (B) The distribution of Archean and Paleoproterozoic elementswithin the North American craton (modified after Hoffman 1988). The rectangle outlines the location of Fig. 1A. THO, Trans-HudsonOrogen; S, Superior craton; H, Hearne Province; R, Rae Province; GFTZ, Great Falls Tectonic Zone. (C) Present-day global distributionof Paleoproterozoic orogens and Archean cratons (modified after Zhao et al. 2002). Paleoproterozoic orogens: AK, Akitkan; C, Capricorn;CA, Central Australia basement; EB, Eburnian; KK, Kola-Karelian; LP, Limpopo belt; M, Moyar belt; SF, Sveocfennian; TA, Trans-Amazonian; TNC, Trans-North China. The Transantarctic orogen in Antarctica is not shown.

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nents. Thus, any lithological remnants of the initial riftingshould be preserved on the continental margins below passivemargin sedimentary sequences and close to the transitionbetween oceanic and continental crust. Evidence for riftingof the Archean cratons now bordering the THO should thusbe preserved within the Wollaston domain and the SuperiorBoundary Zone (Fig. 1). These are regions in which Archeancontinental crust is unconformably overlain by supracrustalrocks and severely reworked during the Paleoproterozoic.

The oldest supracrustal rocks in the Wollaston domain occurwithin the Courtenay Lake – Cairns Lake fold belt (Money1968) and form part of the Needle Falls Group (Fig. 3; Ray1979; Coombe 1994). This is dominantly a sedimentarysequence consisting of a generally fining-upward clasticsuccession of arkoses, conglomerates, and variably graphiticpelites towards the base, overlain by calcareous clastic units,marble, and iron formation. The youngest unit is a homoge-neous arkose. Overall, these sedimentary rocks are inter-preted as a rift–fill sequence with detritus being derivedpredominantly from Neoarchean sources ranging in age fromabout 2585–2510 Ma (Hamilton et al. 2000; Rayner et al.2005). Igneous rocks are rare, although there are amphibo-lites intercalated with the clastic rocks near the base of theNeedle Falls Group that have a trace-element geochemicalsignature indicative of within-plate continental rift tholeiites(Fossenier et al. 1995; MacNeil et al. 1997). In addition, theCook Lake area also contains quartz–feldspar porphyries,probably derived from crustal melts generated by the influxof mafic magma into the crust, which are only found withinthe lowermost clastic succession. A U–Pb age of 2075 ± 2 Maobtained from one of these porphyries (Ansdell et al. 2000)has been interpreted as the best, albeit minimum, direct esti-mate for the age of rifting along the margin of the Hearnecraton.

The supracrustal rocks of the Superior Boundary Zonethat overlie Archean basement gneisses form part of theOspwagan Group (Fig. 3; Scoates et al. 1977; Bleeker 1990;Zwanzig 1999). The lowermost sequence of clastic sedimen-tary rocks (Manasan Formation) is interpreted to overlieunconformably Archean gneisses and consists of locally de-veloped pebble conglomerates grading up into quartzites andwackes that probably were deposited in a passive margin set-ting. The overlying formations indicate that the basin wasthen starved of clastic detritus. The Thompson Formationconsists of dolomitic and calcareous rocks that are overlainby pelitic sedimentary rocks, cherts, and silicate- and sulphide-facies iron formations (Pipe Formation). There is a return toclastic sedimentation during the deposition of the SettingFormation, which precedes the eruption of the mafic andultramafic volcanic rocks of the Bah Lake assemblage of theOspwagan Group. Associated with this volcanic package aremetagreywackes, which are interpreted as turbidites shed offvolcanic edifices, and mafic and ultramafic intrusive rocks.The ages of detrital zircons indicate that a significant pro-portion of the clastic detritus was derived from the Superiorcraton, although Hamilton and Bleeker (2002) also obtainedan age of ca. 1974 Ma for a detrital zircon from a greywackein the Setting Formation, which they interpreted as beingderived from penecontemporaneous volcanic activity. Thisage provides a minimum age for the deposition of the under-

lying sedimentary rocks of the Ospwagan Group and thus ofrifting along this margin of the Superior craton.

The ages obtained from both the Wollaston domain andthe Superior Boundary Zone are consistent with other esti-mates for the timing of rifting along the margins of theHearne and Superior cratons, respectively (Fig. 2B). Asplerand Chiarenzelli (1998) and Aspler et al. (2001) suggestedthat gabbroic intrusions into the lower part of the HurwitzGroup, which developed in an intracratonic basin on theHearne craton, at ca. 2.1 Ga represent a period of lithosphericextension and mantle melting, which eventually led to rifting.The estimated age of rifting along the margin of the Superiorcraton is considered to be about 2.2 Ga (Wardle et al. 2002);this is consistent with an a U–Pb age of 2025 ± 25 Ma fromdiagenetic apatite cement in the Richmond Gulf Group, onthe east shore of Hudson Bay (Chandler and Parrish 1989),and with the ca. 2.0 Ga age of mafic volcanism and rift sedi-mentary rocks preserved in the Povungnituk Group in theQuebec–Baffin segment of the THO (Parrish 1989; Machadoet al. 1993).

The configuration of Archean cratons prior to rifting ofthe Hearne and Superior cratons is controversial. Aspler andChiarenzelli (1998) suggested that gabbroic intrusions in theHurwitz Group herald the breakup of a Neoarchean – earliestPaleoproterozoic supercontinent, termed Kenorland. They alsosuggested that the spatial relationship between Archean con-tinental blocks prior to fragmentation of this supercontinentwas similar to that attained after reassembly during the Paleo-proterozoic to form the core of Laurentia. Thus, the presentHearne and Superior cratons were adjacent to each otherprior to rifting, which would mean that the pre-rifting geo-logical history of the two margins should be similar. How-ever, marked differences in the age of geological events alongthe northern margin of the Superior Province and the southernmargin of the Hearne Province (Ansdell et al. 2005) do notsupport a straightforward fit of the two Archean cratons(Bleeker 2002). In addition, the interpretation of paleomagneticsignatures from mafic dykes cutting the Archean cratons andfrom rocks within the Reindeer Zone have led to differingsuggestions as to the maximum separation of the Hearne andSuperior cratons and rotation between them (e.g., Symonsand Harris 2005; Evans 2002). Nevertheless, there are maficdyke systems, such as the ca. 2.45 Ga Matachewan andKaminak dykes (Heaman 1997) that intrude the Superior andHearne cratons, and which may have been derived by meltingof the same mantle plume. Correlation of dykes of this ageand stratigraphy led to the suggestion that the Wyoming andSuperior cratons were adjacent to each other at 2.45 Ga andthus formed part of a larger Archean craton (Fig. 2A; Roscoeand Card 1993; Heaman 1997). It is unclear where theHearne craton was situated with respect to this reconstruc-tion (Fig. 2A), although recently Bleeker (2004) has pro-posed that the Hearne craton may have been adjacent to theHuronian margin of the Superior craton.

The size of the Manikewan Ocean

There is little direct record of the early evolution of theManikewan Ocean, which developed as a result of oceanspreading processes. The oceanic crust that formed was de-

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Fig. 2. Schematic time slices showing the distribution of tectonicelements during the evolution of the THO during thePaleoproterozoic. (A) Possible distribution of Archean Superior,Wyoming, and Karelia cratons (Heaman 1997) prior to develop-ment of the Manikewan Ocean. The relationship between thesecratons and the Hearne–Rae craton is controversial, althoughAspler and Chiarenzelli (1998) suggest that they may have beenlinked as part of a larger supercontinent (Kenorland). In contrast,Bleeker (2004) suggests that the Hearne craton may have beenadjacent to the southern margin of the Superior craton at ca.2.45 Ga. The Sask craton may have originally been part of theWyoming craton. (B) Constraints on the age of rifting along theSuperior and Hearne cratons. Rifting and sea-floor spreadingleads to the formation of the Manikewan Ocean. (C) Possible re-lationship between tectonic elements within the “Pacific-like”Manikewan Ocean. Flower-symbol refers to clastic input intosedimentary basins. (D) Accretion of oceanic terranes to formthe Flin Flon – Glennie complex and a simplified La Ronge –Lynn Lake arc system. (E) Successor arc magmatism, resultingfrom subduction below the Flin Flon – Glennie complex, and ex-tension of the margin of the Superior craton. “??” underlines ourlack of understanding of the geological history of the northeast-ern Reindeer Zone. (F) Collision between the Flin Flon – Glen-nie complex and the Hearne craton margin and development ofthe Wathaman batholith. (G) Formation of the Kisseynew basinon extending Flin Flon – Glennie complex crust. Exhumation ofthe older arc rocks leads to provision of detritus for theKisseynew basin, and smaller fluvial-alluvial basins on the suc-cessor arc. (H) Collapse of the Kisseynew basin as the Sask andSuperior cratons attempt to subduct below the Reindeer Zonerocks. (I) Crustal thickening and peak regional metamorphism inthe THO, including formation of the structural windows into theSask craton. (J) Post-continental collision strike–slip movementalong the Needle Falls shear zone, Tabbernor fault, and Superiorboundary fault leading to southeastwards extrusion of theinternides of the THO.

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Fig. 3. Tectonostratigraphic columns for the supracrustal sequences on the margins of the Hearne craton (A) and the Superior craton (B). See text for details.

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stroyed when subduction zones were generated within oralong the margins of the Manikewan Ocean. The age of theoldest arc-volcanic rocks in the MS-THO is ca. 1.92 Ga;thus, the first ca. 150 Ma history of the ocean is not re-corded by any preserved oceanic rocks. The Wollaston do-main and Superior Boundary Zone do contain supracrustalrocks that formed on the continental margins of this ocean,although these supracrustal sequences are not complete andhave been metamorphosed to at least amphibolite grade andmultiply deformed.

The only method by which the size of the ManikewanOcean can be determined is paleomagnetism, although thethermal and structural evolution of the rocks during the Trans-Hudson orogeny makes interpretation of paleomagnetic datadifficult. Symons (1991, 1998) has used paleomagnetic dataobtained from granitoid plutons of known age, which areassumed to have acted as rigid bodies that were not tiltedduring deformation, to suggest that the Manikewan Oceanwas about 5500 km wide at ca. 1850 Ma (Symons and Harris2005). This means that the distance between the Superiorand Hearne cratons at 1850 Ma (Fig. 2F), as suggested bythe poles for the Sudbury intrusive complex and the Wathamanbatholith, respectively, was about 5500 km. However, Evans(2002) suggests that there are a number of interpretations ofthe paleomagnetic data sets that allow for varying widths ofthe ocean between the Hearne and Superior cratons, althoughhe emphasizes that there is a lack of data for the period oftime from rifting of the Hearne margin and the developmentof the first subduction-related rocks. Overall, the paleo-magnetic data have been used to imply that the ManikewanOcean may have been Pacific-like in size. Thus, if the inter-pretation of the paleomagnetic data (Symons and Harris 2005)is correct, the distance between the Superior and Hearnecratons (Figs. 2C–2F) was about 5000 km or more. This dis-tance then decreased as the two cratons drifted together onoceanic crust that was being subducted prior to terminal con-tinent–continent collision (Figs. 2G–2J).

Development of the Manikewan Ocean“Ring of Fire”

The oldest volcanic rocks preserved in the MS-THO areca. 1.92 Ga arc rocks that provide a minimum age for theinitiation of subduction within the Manikewan Ocean. Thelocation of arcs with respect to each other during the first 40Ma of subduction history and their continuity during thisperiod (Fig. 2C) is difficult to unravel because of later struc-tural and thermal reworking during the Trans-Hudson orog-eny. However, the geochemical and isotopic compositions ofvolcanic rocks within the Reindeer Zone provide constraintson the location of arcs with respect to craton margins or con-tinental fragments and show that volcanic rocks were alsogenerated in back-arc basins and above mantle plumes.

The oldest volcanic rocks in the Flin Flon belt occur indistinct structural blocks and cannot be correlated across theblock-bounding faults. However, their geochemical charac-teristics and Nd isotopic signatures indicate that the majorityof the volcanic rocks formed in oceanic arcs, back arcs, andocean islands that developed at some distance from signifi-cant blocks of continental crust (Stern et al. 1995a, 1995b;Bailes and Galley 1999). Nevertheless, small areas of arc

crust (e.g., Mystic Lake, Snell basalt, Snow Lake) havelower εNd(t) values than coeval juvenile arc crust (Stern etal. 1992, 1995a) and inherited ca. 2.5 Ga zircons (David etal. 1996; Stern et al. 1999). These data suggest that, withinthe Manikewan Ocean, there were fragments of Neoarcheancontinental crust on which some of the arc rocks may havebeen constructed. Some of this Neoarchean crust is pre-served as fault-bounded blocks in the Northeast Arm shearzone (David and Syme 1994), probably during the intra-oceanic accretionary episodethat generated the Flin Flon –Glennie complex (see later in the text).

Recently, a number of workers have recognized the broadsimilarity between arc rocks in the Flin Flon belt and thoseforming in other domains within the Reindeer Zone at thesame time. Maxeiner et al. (1999) suggest that the tholeiiticand calc-alkaline igneous rocks of the Hanson Lake blockand southeastern Glennie domain have lithological and com-positional similarities to the arc rocks of the western FlinFlon belt. The oldest rocks within the thin arcuate volcanicbelts of the Glennie domain are also dominated by tholeiiticarc and ocean-floor rocks (Delaney 1992). The La Rongedomain, which will be described in more detail later in thetext, is dominated by arc volcanic rocks, although Ashton(1999) emphasized that our understanding of the relationshipbetween the Glennie domain and the southern end of the LaRonge domain (Fig. 1) is still problematic. The arc rockswithin the Reindeer Zone may have all been generated abovesubduction zones that were linked within the ManikewanOcean. Figure 2C provides a possible scenario, although theintense structural reworking of the oldest volcanic assem-blages means that it is impossible to determine the originalgeographical relationship between arcs and intervening ocean-floor and ocean-island rocks. The Manikewan Ocean wasprobably as complex as the present-day southwestern Pacific(Hall 2002) and would have included small fragments ofolder continental crust, back-arc basins, oceanic plateaus,and juvenile arc crust. The Flin Flon belt, for example, hasall these components, and economically important volcano-genic massive sulphide Cu–Zn deposits, which were gener-ated in a rifted arc environment similar to the Lau basin(Syme et al. 1999).

The La Ronge and Lynn Lake volcano-plutonic belts andassociated sedimentary rocks lie along the southern marginof the Hearne craton. The physical links between the LaRonge and Lynn Lake belts and the tectonic elements to thesoutheast (e.g., Glennie and Flin Flon domains) are poorlyunderstood. Figure 1 shows that the La Ronge and LynnLake domains are separated from the Glennie, Flin Flon, andSuperior Boundary Zone by metasedimentary rocks of theKisseynew domain, whereas the nature of the contacts betweenthe southwestern portions of the La Ronge, Rottenstone, andGlennie domains just to the north of the Phanerozoic coverhas not been addressed since reconnaissance mapping byLewry (1981a). In Figures 2C to 2H, the arc terranes alongthe Hearne margin are associated with different subductionzones than the arc terranes preserved in the Reindeer Zone,and thus on different plates. This allows the northward driftof the Sask craton relative to the Hearne craton along atransform or very obliquely convergent plate boundary. Oneof the challenges in paleotectonic reconstructions in the THOis to properly link the rocks along the Hearne margin with

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those in the centre of the orogen. Because of the difficulty inproviding this linkage at the present time, the magmatic andaccretionary evolution of the Hearne craton margin is dis-cussed separately later in the text.

The Manikewan “Ring of Fire”: magmatismand accretion along the Hearne continentalmargin

The geological evolution of the southern margin of theHearne craton has been the focus of recent mapping initia-tives that tie the surface geology with the Lithoprobe seismicimages. These images have convincingly shown that as a resultof the final stages of continental collision the Hearne cratonis now underlain by the dominantly Paleoproterozoic rocksof the internal parts of the orogen, and that the crust containsreflectors that dip towards the Hearne craton. However, forthe development of the time slices (Fig. 2) it is important toreview the age and character of the volcanic and sedimentarypackages that were accreted to the margin of the Hearnecraton, and the timing of this accretion and subsequentcontinent–continent collision. The present-day extent of theArchean Hearne craton is important to determine, particu-larly with respect to interpretation of seismic data.

Historically, the evolution of the Hearne continental marginhas been linked to the development of an island-arc outboardof the margin. Ray and Wanless (1980) and Lewry et al.(1981) suggested that the arc complex developed on oceaniccrust at about 1.9 Ga above a northwest-dipping slab withthe supracrustal rocks of the Rottenstone domain being de-posited in a back-arc environment with respect to this arc.The direction of subduction has also been suggested to havebeen directed away from the Hearne margin prior to thegeneration of the Wathaman batholith (e.g., Bickford et al.1990; Maxeiner et al. 2001). The lack of arc volcanic rocksin the Wollaston domain was used as evidence that the Hearnecraton and the overlying Paleoproterozoic sedimentary rockwere on the lower plate.

New interpretations of the geological evolution of the rocksaccreted to the Hearne craton margin have been derived frommapping and associated geochemistry and geochronology ina superb transect across the northern margin of the ReindeerZone preserved on Reindeer Lake (e.g., Maxeiner et al. 2001,2005; Corrigan et al. 1999a, 1999b). Units on Reindeer Lakecan be correlated along structural strike with rock units inthe La Ronge, Rottenstone, Lynn Lake, and Southern Indiandomains (Fig. 1A). The ca. 1.88 Ga Reed Lake volcanic as-semblage, which is temporally equivalent to arc volcanicrocks in the Central Metavolcanic belt of the La Ronge do-main (e.g., Van Schmus et al. 1987), consists of an arc thatwas constructed on the older Lawrence Point volcanic as-semblage (Fig. 4). Maxeiner et al. (2005) suggested that theLawrence Point assemblage represents a supra-subductionzone ophiolite, which may have originally been formed in aback-arc environment. These rocks are bounded to the southby an imbricate zone containing ca. 1.87 Ga turbiditic sedi-ments (Duck Lake assemblage, Fig. 2D), and ca. 1.9 Gaocean-floor basalts and metasedimentary rocks (LevesqueBay supracrustal assemblage), which together are consideredto represent sediments deposited in a fore-arc setting withrespect to the La Ronge – Lynn Lake arc system and obducted

fragments of back-arc (?) ocean-floor rocks (Figs. 2C, 2D).The Clements Island arc-volcanic rocks (ca. 1.9 Ga, Corriganet al. 2001), at the north end of Reindeer Lake, are separatedin time and space from the Reed Lake volcanic assemblage.The back-arc ocean-floor rocks preserved in the Laxdal–Doucet islands area (Wright 2001) may be equivalent to theocean-floor rocks preserved in the Lawrence Point assem-blage.

There is enough evidence provided by internally deformedxenoliths and crosscutting relationships to show that the earliestductile deformation preserved in the Paleoproterozoic rocksalong the Hearne margin predated the intrusion of theWathaman batholith (Maxeiner et al. 2001; Coolican 2001;Beaumont-Smith and Bohm 2002). In the Lynn Lake belt,Zwanzig et al. (1999) and Beaumont-Smith and Bohm (2002)suggest that the earliest deformation was related to accretionof arc and ocean-floor rocks at about 1.87 Ga (Fig. 4C). Theage of the volcanic rocks that were accreted, either to theHearne margin or to form an intraoceanic collage, vary from1910 1

15−+ Ma in the Lynn Lake belt to 1878 ± 3 Ma in the

Rusty Lake belt (e.g., Baldwin et al. 1987). In addition, thegeochemical signatures of the volcanic rocks in the La Ronge,Lynn Lake, Rusty Lake, and Clements Island belts indicatearc and ocean-floor environments (Watters and Pearce 1987;Syme 1985, 1990; Zwanzig et al. 1999; Maxeiner et al.2005). It is of interest that the volcanic rocks of the northernLynn Lake belt have Nd isotopic compositions indicative ofcontamination by Neoarchean crust, whereas those in thesouthern Lynn Lake belt have more juvenile Nd isotopecompositions (Beaumont-Smith and Bohm 2002). Maficintrusions in the Lynn Lake belt also yield εNd(t) valuesranging from –0.6 to +4.0 (Hegner et al. 1989), which suggeststhat Lynn Lake crust may be complicated and include frag-ments of older, possibly Archean crust. In contrast, the vari-ation in Nd isotopic compositions of volcanic and plutonicrocks in the La Ronge domain is more restrictive (+2.0 to+4.7; Chauvel et al. 1987; Thom et al. 1990, Hegner et al.1989; Kyser and Stauffer 1995), which lends support to itsorigin as an oceanic island-arc system (Fig. 2C). Overall, thevariety of tectonic environments in which volcanic rocksformed along the Hearne margin and the range in age is asdiverse as seen in the Flin Flon belt, and thus the geometryof plate boundaries along the margin was probably verycomplicated. In addition, the variations observed in the Ndisotopic composition of volcanic rocks in the Lynn Lake andLa Ronge belts suggest that there may have been fragmentsof older crust outboard of the Hearne craton in the ManikewanOcean with which the magmas may have interacted.

Sedimentary rocks preserved along the margin can yieldinformation on the accretionary history. Detrital zircons ortheir Nd isotopic composition represent proxies of the sourceof clastic material in the sedimentary rocks, and thus canrecord the juxtaposition of distinct tectonic units. For example,the Archean rocks of the Hearne craton are unconformablyoverlain by sedimentary rocks of the Wollaston domain, allof which were severely deformed and metamorphosed duringthe Trans-Hudson orogeny (Lewry and Sibbald 1980). Thebasal sedimentary units (Needle Falls Group) have beendescribed above and record the rifting of the Hearne margin.Overlying these rocks is the Wollaston Group (Fig. 3), whichrepresents the thickest preserved part of the section. Tran et

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Fig. 4. Schematic cross-sections showing the timing of magmatism, sedimentation, and accretion along the margin of the Hearne craton culminating in underthrusting by theSask craton. Slab delamination in D and G may have assisted in the generation of the voluminous Wathaman batholith at ca. 1860 Ma and the small volume potassic andalkalic intrusive rocks in the northern MS-THO at ca. 1830 Ma. Based on Tran (2001), Maxeiner et al. (2001), and Hollings and Ansdell (2002). See text for more details.

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al. (2003) and Tran (2001) obtained Nd isotope data fromwhole rocks and sensitive high-resolution ion microprobe(SHRIMP) U–Pb geochronological data from detrital zirconsfrom selected locations within the Wollaston Group stratig-raphy. Both sets of data show that the clastic material inthese sedimentary rocks was not derived solely from theArchean rocks of the adjacent Hearne craton. Yeo and Savage(1999) suggested that the Paleoproterozoic components ofthe sediment could have been provided by rivers drainingthe Taltson orogen to the west and flowing over the inter-vening Archean Rae and Hearne cratons. However, theSHRIMP U–Pb data (Tran 2001) convincingly supports der-ivation of Paleoproterozoic detritus from rocks similar inage (1.92–1.88 Ga) to the terranes being accreted to the mar-gin of the Hearne, and not to the older Paleoproterozoicrocks of the Taltson orogen. Tran (2001) interpreted theWollaston Group sediments to have been deposited initiallyin a back-arc basin behind a continental magmatic arc, partsof which may be preserved within the supracrustals of theRottenstone domain (Figs. 4B, 4C). There is a possibilitythat the Clements Island arc volcanic rocks may have beenpart of this arc. Sedimentation in the Wollaston domainceased prior to the intrusion of the Wathaman batholith, acontinental margin batholith (Lewry et al. 1981), which wasemplaced between ca. 1865 Ma and 1850 Ma (Meyer et al.1992) and has typically been related to the accretion of theLa Ronge arc to the margin of the Hearne craton. The sedi-mentary rocks were then incorporated into a westward-vergingfold and thrust belt (Tran 2001). The overall development ofthe Wollaston Group in a back-arc – foreland settingsupports the general model suggested by Ray and Wanless(1980) and Lewry and Collerson (1990).

Other sedimentary sequences related to the accretionaryhistory of the Hearne margin include the turbiditic sedimentsof the Milton Island assemblage (Corrigan et al. 1998). Theserocks contain detrital zircons derived from the Hearne cratonand adjacent emerging volcano-plutonic terranes (Fig. 2E;Ansdell et al. 1999b), and were thus deposited in a sedimentarybasin between the two. The clastic rocks of the Park Islandassemblage, which overlie the Milton Island assemblage, areinterpreted to have been deposited in a foreland basin withlocal derivation of material from Paleoproterozoic crust thatwas thickened during collisions along the Hearne margin(Corrigan et al. 1998).

Corrigan et al. (1998, 1999a) indicated that the Wathamanbatholith and smaller intrusive bodies of the same age andcomposition intrude Archean metamorphic rocks of the PeterLake domain to the north, and older Paleoproterozoic rocksof the Rottenstone and La Ronge domains to the south. Thissupports the interpretation that the Wathaman batholith stitchesthe La Ronge arc to the margin of the Hearne craton (Fig. 4D).The interpretation of paleomagnetic data by Symons (1991)and Symons and Harris (2005) conflicts with this geologicalobservation, as they suggest that there was still a significantocean basin between the La Ronge arc and the Hearne marginat the time of Wathaman batholith formation. The processleading to magma production is less clear, but may have re-sulted from breakoff of a subducted slab associated with anearlier accretionary event which allowed hot asthenosphereto initiate significant partial melting of hydrated mantle overan extensive area. The volume of magma produced also was

dependant on the degree of obliquity of subduction with agreater volume being generated on the broadly east–west-oriented margin of the Hearne craton as a result of nearperpendicular plate movement as opposed to highly obliquerelative movement further to the southwest.

Historically, the eastern edge of the Hearne craton inSaskatchewan has been correlated with the eastern marginof the Wollaston domain and the Peter Lake domain, whichis marked by the presence of late tectonic shear zones, suchas the Needle Falls shear zone (Stauffer and Lewry 1993).The rocks to the east and south of the Wathaman batholithwere all considered to be Paleoproterozoic in age, until thediscovery of ca. 2.5 Ga augen gneiss in the core of a struc-tural dome in the southern Rottenstone domain (Fig. 1A;Bickford et al. 2001). The best interpretation of this inlier isthat it is a structural window into the Hearne basement belowthe veneer of supracrustal rocks that were originally depositedon the Hearne craton. Thus, the Wathaman batholith is anintrusion that was emplaced into complex crust consisting ofArchean and Neoarchean rocks of the Hearne craton, andPaleoproterozoic rocks that had already been accreted to themargin (Fig. 2F). The variation in initial εNd(t) values fromthe Wathaman batholith (–0.3 to –8.0; Chauvel et al. 1987;Kyser and Stauffer 1992; MacHattie et al. 2001) emphasizethe complexity of both the magma source and the crust intowhich it was emplaced. At this point, the southern margin ofArchean Hearne crust in Manitoba is difficult to determine,but may be significantly further south than the northern edgeof the Wathaman batholith. Nd isotope mapping may be veryuseful to define better the buried extent of older Archeancrust. On Reindeer Lake, for example, the Nd isotopic com-position of Wathaman-age plutons indicates significantly lessinvolvement of older crust from north to south (MacHattie etal. 2001).

The Flin Flon – Glennie complex

The Flin Flon – Glennie complex (Ashton 1999) repre-sents a collage of volcanic, intrusive, and sedimentary rocksthat were amalgamated to form a crustal entity that had asignificant effect on the structural development of the internidesof the MS-THO (Fig. 1A). The original subdivisions withinthe Reindeer Zone included the volcano-plutonic Flin Flon,Hanson Lake, and Glennie domains, and the sedimentaryrock-dominated Kisseynew domain, and were based onvariations in lithology and structural style (Macdonald andBroughton 1980). These two-dimensional subdivisions wereused to imply that each of the domains had a distinct geologicalhistory, even though their boundaries were poorly delineatedand that they were juxtaposed during the latest stages ofcontinental collision. Lewry et al. (1990) made an attempt toredefine domains by using structural data to constrain thethree-dimensional relationships within the internides of theMS-THO and provided the first suggestion that the Flin Flon,Hanson Lake, and Glennie domains may have been part ofthe same structural entity. They also inferred that the internidesof the MS-THO were allochthonous with respect to thestructurally underlying rocks of the Sask craton, which aredescribed later in the text, and this relationship has beensupported by seismic reflection lines (Lucas et al. 1993).

The Flin Flon belt contains assemblages of rocks that

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developed in distinct tectonic settings in a largely oceanicenvironment as part of a complex Manikewan Ring of Fireand that are juxtaposed along faults. U–Pb zircon ages fromgranitoid intrusions, which cut the earliest shear zone fabricsin the Mystic Lake, Meridian–West Arm, and Elbow Lakeshear zones, indicate that these volcanic assemblages werejuxtaposed at about 1.87 Ga (Figs. 2C, 2D; Lucas et al.1996; Ansdell and Ryan 1997). The chemical compositionof granitoid intrusions and volcanic rocks that range in agefrom 1.86–1.83 Ga indicate that subduction continued beneaththis accretionary complex (Figs. 2E–2G; Whalen et al. 1999;Ansdell et al. 1999a; Ansdell and Connors 1995). In addition,the intraoceanic accretionary complex emerged above sealevel, as highlighted by the development of a paleoweatheringprofile (Holland et al. 1989; Pan and Stauffer 2000), and thedeposition of fluvial–alluvial sedimentary rocks at ca. 1.84–1.83 Ga that contained clastic material derived locally fromerosion of successor arc rocks that must have been rapidlyexhumed (Figs. 2G, 2H; Stauffer 1990; Ansdell 1993; Ansdellet al. 1992, 1999a).

Even though the timing of accretion between distinct fault-bounded assemblages is only known in the Flin Flon belt,the geological evolution of the Glennie domain (Delaney1992) is remarkably similar. In particular, the age of volcanicand plutonic rocks and the timing of fluvial–alluvial sedi-mentation in the Glennie domain (McNicoll et al. 1992) aresimilar to the Flin Flon belt. This lends credence to the sug-gestion that volcanic terranes may be continuous betweenthe Flin Flon and Glennie domains, and that both domainswere thickened and emergent at the same time.

The Sask craton

A major achievement of the THOT has been the identifi-cation of seismically distinct crust that can be traced to sur-face in two areas in the Glennie and Flin Flon domains(Lucas et al. 1993). Within these areas (Hunter Bay, Istwatikan,and Nistowiak gneiss domes in the Glennie domain and thePelican window in the Flin Flon domain) are exposures ofArchean rocks (Fig. 1A) that are separated from the structur-ally overlying Paleoproterozoic rocks by high-strain zones(Chiarenzelli 1989; Ashton et al. 1999). The seismic datashows that the Archean rocks are isolated from the Archeanrocks of the Hearne and Superior provinces, and extend intothe subsurface as a crustal-scale anticlinorium. These Archeanrocks form part of the Sask craton (Ansdell et al. 1995),which now represents an important component in the devel-opment of the MS-THO, as initially predicted by Lewry(1981b) and Lewry et al. (1990). The presence of Archeanrocks in the Pelican window was initially identified by Belland Macdonald (1982), and more recently further U–Pb geo-chronology has shown that the Pelican window and the areasin the Glennie domain have a distinct Archean and earliestPaleoproterozoic history (Chiarenzelli et al. 1998; Ashton etal. 1999; Rayner et al. 2005).

The overall geometry and extent of the Sask craton hasbeen estimated from Lithoprobe and COCORP (Consortiumfor Continental Reflection Profiling) seismic sections (Lucaset al. 1993; Baird et al. 1996; Hajnal et al. 2005), Nd and Pbisotope analyses of post-collisional pegmatites in the exposedMS-THO (Bickford et al. 2005; Ansdell and Stern 1997;

Prokopiuk and Ansdell 2000), and Precambrian rocks inter-sected in drill holes through the Phanerozoic cover to thesouth (Collerson et al. 1988). Together, these methods showthat the Sask craton extends over 1000 km southwards underthe Phanerozoic cover and tapers northwards towards thenorthern end of the Glennie domain. In fact, it appears asthough the volume of juvenile Paleoproterozoic crust belowthe Phanerozoic cover is insignificant (Collerson et al. 1988).In addition, the presence of diamondiferous kimberlites atFort à la Corne indicates that the Sask craton has a significantlithospheric root.

The Sask craton exposed in the structural windows in theGlennie domain consists mainly of 2.4–2.5 Ga granitic toferrodioritic orthogneisses, which are separated from Paleopro-terozoic rocks by the mylonitic Nistowiak Thrust (Chiarenzelliet al. 1998; Rayner et al. 2005). A variety of rocks havebeen mylonitized including Archean and Paleoproterozoicgranitoid and mafic rocks, and metasedimentary rocks thatcontain 2.8–2.95 Ga detritus and that may have been depositedon the margin of the Sask craton. The Pelican window consistsof ca. 2.45 Ga orthopyroxene-bearing granitoid and maficrocks, which intrude migmatitic paragneisses and leucocraticorthogneisses that are older than 2.96 Ga (Ashton et al.1999; Rayner et al. 2005). The age and character of therocks exposed in the structural windows and the structuralcharacter and history of the bounding Nistowiak and Pelicanthrusts are remarkably similar; this together with the seismicdata convincingly support the correlation between these twomylonitic high-strain zones and thus the suggestion that theArchean rocks represent exposures of the same Archean block.The small areal extent of the exposed Sask craton has madeit difficult to determine the heritage of the craton. However,the age of igneous and metamorphic events in the Saskcraton do not correlate well with the age of rocks exposed inthe Hearne and Superior cratons (Ansdell et al. 2005). Bickfordet al. (2005) suggest that the Pb isotopic composition ofSask craton rocks is more similar to parts of the Wyomingcraton. Thus, the Sask craton may represent a fragment ofthe Wyoming craton that rifted and migrated northwards toits present position. Unfortunately, no paleomagnetic datasupport this speculation. In addition, some of the Neoarcheanfragments preserved in the Flin Flon belt may actually befragments of the Sask craton.

Continental collisions: the tale of threecratons

The geometrical relationships observed today and imagedby Lithoprobe seismic lines within the MS-THO are essen-tially a result of continental collisions that occurred in thePaleoproterozoic and which led to the final closure of theManikewan Ocean. The structural complexity of the regionresults from the diachronous collision between three Archeancratons, namely the Hearne, Sask, and Superior cratons. Thesecratons also preserve significant, albeit variably developed,Paleoproterozoic structural and thermal events.

Superior–Reindeer zone interactionThe orientation of seismic reflectors in the vicinity of the

Superior Boundary Zone suggests that the rocks of theReindeer Zone underlie the margin of the Superior craton

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(White et al. 1999, 2002), which was a surprise when con-sidering that the Superior craton had been interpreted tohave formed the lower plate during collision with the Rein-deer Zone, based on the lack of arc rocks on the craton(Lewry 1981b; Green et al. 1985a; Bleeker 1990). However,the apparent conflict between the dip of seismic reflectorsand the evidence for the polarity of subduction has now beenreconciled by comparing the seismic signature exhibited alongline 3B with those from lines 1 and 2 (White et al. 2002).The northernmost seismic lines (1 and 2) preserve predomi-nantly east-dipping structures, resulting from intensive reori-entation of earlier west-dipping structures in a retroshearzone during long-lasting collision. The preservation of low-

grade metamorphic rocks and west-dipping structures in thevicinity of line 3B and intensely deformed and high-grademetamorphic rocks in the Thompson area is related to longerlasting and greater convergence in the vicinity of theThompson promontory as opposed to the flanking reentrants(White et al. 2002). White et al. (2002) convincingly showthat the low metamorphic grade supracrustal rocks of theSuperior Boundary Zone, which are buried below the Phanero-zoic, are preserved in an east-verging fold and thrust belt.

Interaction between the Superior craton and rocks of theReindeer Zone may have started as early as ca. 1.89 Ga(Fig. 2D), as mafic dykes, thought to be similar to the Molsondykes (1883 Ma; Heaman et al. 1986), crosscut the earliest

Fig. 5. Schematic cross-section across the Superior Boundary Zone. The section is perpendicular to the overall motion of the Saskcraton and the Superior craton, so all convergent plate boundaries are likely oblique. Between 1865 and 1840 Ma, the Flin Flon –Glennie complex collides with the Superior craton, whereas in F the Superior boundary fault is a sinistral strike–slip fault with theSuperior craton moving northwards with respect to the Reindeer Zone. See text for details.

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deformation event (Figs. 5B, 5C). White et al. (1999)emphasized that the character of the terrane that collidedwith the Superior craton at this time is not known, as therocks to the west of the Superior Boundary Zone are buriedby younger sedimentary rocks of the Reindeer Zone and thecolliding terrane may also have rifted away during theeruption of 1864 Ma komatiites (Fig. 5D; Hulbert et al.1994), which have been intersected below the Phanerozoiccover. These komatiites are incorporated in the thrust slices,and so the best maximum age estimate for the start of colli-sion with the Superior craton is 1864 Ma. The minimum agefor the start of a close spatial association between the Supe-rior craton and the Reindeer Zone is probably ca. 1.84 Ga(Fig. 5E), based on the age of granitoid and mafic rocks(e.g., Bleeker et al. 1995) that were emplaced in both tectonicelements and the assumed age of the terrestrial Grass RiverGroup (Zwanzig 1999). During collision, the Superior cratoncontinued to march northwards as a result of a convergentplate boundary (Figs. 2F–2I; Figs. 6B, 6C) above whichsome of the arc magmatic rocks preserved within the easternReindeer Zone may have been generated (e.g., White et al.1999). The northeastern portion of the Reindeer Zone ispoorly understood, although Archean crust is more extensivethan originally appreciated (Zwanzig and Bohm 2002). TheSuperior craton is now thought to extend as far northwest asthe Owl River shear zone and includes Mesoarchean crustthat may have been accreted to the Superior craton during

the Trans-Hudson orogeny (Fig. 6B; Bohm et al. 2000). Thus,many of the granitoid gneisses may be Archean in age ormelts originally generated in the Paleoproterozoic that wereextensively contaminated by Archean crust. Nevertheless, itis likely that Superior craton crust may have significantlyunderthrust the southern margin of the Hearne craton(Figs. 6C, 6D).

The northward movement of the Superior craton duringthe final stages of continental collision generated sinistraland dextral strike–slip faulting on the western and easternsides of the Thompson promontory, respectively, and likelyresulted in strike–slip displacement along the Owl River shearzone (e.g., Gibb 1975). Deformational features interpreted tobe related to collision with the Superior were synchronouswith peak metamorphism at ca. 1.81 Ga in the eastern ReindeerZone, but postdate peak metamorphism farther to the west(e.g., Connors et al. 1999). The northeast–southwest-trendingregional-scale folds in the Reindeer Zone were generated bythe collision with the Superior craton and led to folding ofearlier structures, such as the high-strain zones separatingPaleoproterozoic rocks from the Sask craton (Figs. 2I, 2J;e.g., Lewry et al. 1990; Ashton et al. 1999).

Collisions in the Reindeer Zone (Flin Flon – Glenniecomplex – Sask craton)

The Kisseynew domain (Fig. 1A) consists of metamor-phosed turbiditic sedimentary rocks that were deposited in a

Fig. 6. Schematic cross-sections showing the interaction between the Superior and Hearne cratons across the Thompson promontory ofthe Superior Boundary Zone. See text for details.

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marine basin at broadly the same time (ca. 1850 Ma) asfluvial-alluvial sediments were being deposited on the FlinFlon – Glennie complex (Fig. 2F; Zwanzig 1990; Ansdelland Norman 1995; Ansdell et al. 1995; Machado et al. 1999).However, turbiditic material (e.g., Levesque Bay assemblage)derived from adjacent arcs was being deposited in basinswithin the Manikewan Ocean much earlier than this and arenow structurally interleaved with the rocks of the Kisseynewdomain. The Kisseynew basin, sensu stricto, is interpretedto have developed after the Flin Flon – Glennie complex hadcollided with the southern margin of the La Ronge – LynnLake belts at ca. 1850 Ma (Fig. 4F; Ansdell et al. 1995;White et al. 2000). Obducted remnants of ocean-floor rocksare preserved in the Granville Lake structural zone. Thesubduction zone marking the northern boundary of the Saskplate retreated, leading to extension of the Flin Flon – Glen-nie complex in the upper plate and development of a basinwhich collected detritus from adjacent arc terranes (Figs. 4E,4F). The basin was rapidly thickened and collapsed as theSask craton started to underthrust the Flin Flon – Glenniecomplex (Fig. 4G). Arc plutons, formed between 1840 Maand 1830 Ma, crosscut the Kisseynew domain turbidites andwere generated by broadly north-dipping subduction either ofthe Sask plate (Fig. 4G; Hollings and Ansdell 2002) or theSuperior plate (Fig. 6C; White et al. 2000) underneath theFlin Flon – Glennie complex and the Hearne margin.

Ashton et al. (1999) suggest that collision between theSask craton and the Flin Flon – Glennie complex was initiatedbefore 1826 Ma (Fig. 2H), the age of a diorite dyke that cutsthe Pelican Thrust. The youngest successor arc magmaticrocks in the Flin Flon – Glennie complex are about 1830 Ga(Ansdell et al. 1999a; Hollings and Ansdell 2002) and maydocument the ultimate clogging of a subduction zone as a resultof underthrusting by the Sask craton. However, the presenceof Neoarchean detrital zircons in ca. 1880 Ma turbidites(Ansdell and Stern 1997) and ca. 1845 Ma fluvial-alluvialsedimentary rocks (Ansdell 1993) in the Flin Flon belt mayimply that Sask craton crust was attempting to subduct andthus collide with the margin of the Flin Flon – Glenniecomplex much earlier than envisaged. Also, the presence ofNeoarchean cores in zircons from Rottenstone domain granitoids(Clarke et al. 2005) may also suggest that material derivedfrom the Sask craton was available in the magma sourcealong the western margin of the THO. However, the Neo-archean zircons may also have been derived from fragmentsof “Sask craton” crust that were swimming in the ManikewanOcean. In general, the older igneous rocks of the Flin Flon –Glennie complex have isotopic signatures which indicate thatArchean crust was not involved in their development, andthus there was likely a significant distance between theseterranes and the nearest Archean blocks.

Chiarenzelli et al. (1998) and Ashton et al. (1999, 2005)provide geological, structural, and geochronological evidencethat the Sask craton was severely deformed and metamor-phosed during the Trans-Hudson orogeny. The boundarybetween the Sask craton and the Paleoproterozoic rocks ofthe Flin Flon – Glennie complex is a high-strain zone (e.g.,Ashton et al. 2005), emphasizing that they are allochthonouswith respect to each other. The Sask craton migrated north-wards, underthrusting the younger rocks, and extended atleast as far north as the southern end of Reindeer Lake

(Figs. 2G–2J; Figs. 4G, 4H). The extent of the Sask cratonis, however, difficult to determine from seismic data, butmay lie below a significant portion of the Reindeer Zone(Hajnal et al. 2005).

The western and eastern boundaries of the Sask plate areconsidered to be transform boundaries, or at least highlyoblique, to allow the northward movement of the Sask platewith respect to the Hearne plate and the northward movementof the Superior plate with respect to the Sask plate (Fig. 2).The North American Central Plains conductivity anomaly(Camfield and Gough 1977), the origin of which is still contro-versial, may be the geophysical signature of the western marginof the Sask plate (e.g., Jones et al. 2005).

Hearne–Rae hinterlandThe Hearne and Rae cratons in the western Churchill

Province (Fig. 1B) were juxtaposed during the Archean(Hanmer et al. 1994), but were extensively reworked duringthe Paleoproterozoic (e.g., Sandeman 2001). Seismic profilesacross the western margin of the Hearne domain as part ofthe Alberta Basement Lithoprobe record east-dipping reflec-tivity, whereas those recorded by Trans-Hudson OrogenTransect exhibit west-dipping structures. The geometry ofthese structures indicates that the Hearne domain formed theupper plate of two collision zones with opposing vergence,namely the Taltson–Alberta orogen and the THO. The west-dipping reflectors along the eastern Hearne margin may beconsistent with long-lived westerly-dipping subduction (Figs. 2,4). The evidence in the crust for southeasterly directed sub-duction has thus been destroyed, if it ever existed.

The underplating of young oceanic crust from the westand continental crust from the east would have provided thedrive for thickening and uplift, with the Hearne domain con-sidered to be a collisional plateau (Ross et al. 2000). Therewere numerous geological consequences of these collisions.Paleoproterozoic granulite-grade metamorphism (Crocker etal. 1993) indicates significant exhumation. Ross et al. (2000)suggest that extensional structures are lacking in the southernHearne Province, although the development of the BakerLake basin, which is broadly synchronous with collisionsalong the southern Hearne domain (Rainbird et al. 2002),and the eruption and intrusion of the Christopher IslandFormation rocks (Cousens et al. 2001) may suggest other-wise. In addition, the basement rocks of the Hearne domainwere incorporated into thick-skinned thrust sheets, and wereinvolved in crustal-scale buckling (Tran 2001; Aspler et al.2002). Ross et al. (2000) suggested that part of the Hearnelithospheric mantle was removed. This led to an influx ofasthenosphere which likely underwent decompression melting,and the resulting mantle melts infiltrated and metasomatizedthe remaining lithospheric mantle, sparking the productionof potassic magmas and lower crustal melts (Cousens et al.2001; Peterson et al. 2002). The generation of these magmaswas initiated at about 1830 Ma and is temporally related tothe last vestiges of subduction-related magmatism in the MS-THO (Figs. 2H, 4G, 6C). By this time, subduction of thickcontinental crust (Sask and (or) Superior cratons) was beinginitiated, albeit unsuccessfully, below the Hearne margin.The denser leading edges of these plates, consisting ofeclogitized oceanic crust, probably broke off leading toasthenospheric upwelling and melting of metasomatized litho-

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spheric mantle and lower crust. It is interesting to note thatthe youngest intrusive rocks along the Hearne margin arealkalic, potassic, and carbonatitic (e.g., Halden and Fryer1999; Mumin 2002) and may have been generated by theprocess described above. The concentration of Paleoproterozoicdeformation into discrete zones within the Hearne craton,the development of sedimentary basins, and the generationof alkalic and potassic magmatism would suggest that thecrust north of the Reindeer Zone is broadly comparable tothe present-day Tibetan plateau and Tien Shan in Asia. Theinternides of the THO would thus represent the erodedremnants of a Himalayan-type continent–continent collisionzone, an idea first suggested by Dewey and Burke (1973).

Post-collisionOn a global scale, the MS-THO represents a small section

of the Paleoproterozoic orogenic belts that welded the conti-nents of Laurentia and Baltica, and which may have formedparts of a significant (super?)continent that was probablyextant until about 1.27 Ga (Columbia; Rogers and Santosh2002; Zhao et al. 2002). The extension of the THO intoScandinavia is well accepted, although the extension south-wards is more difficult to unravel as the rocks are overlainby Phanerozoic sedimentary basins. The Great Falls TectonicZone, which preserves Archean and juvenile Paleoproterozoicrocks that were overprinted by ca. 1820 Ma metamorphismand deformation as a result of collision between the Wyomingcraton and the Medicine Hat block (Mueller et al. 2002;Roberts et al. 2002), may have linked with the southern partof the MS-THO. Similar magmatic and metamorphic ageshave also led to the suggestion that the Halls Creek orogenin northern Australia may represent a continuation of theMS-THO by way of the Great Falls Tectonic Zone (Tyler etal. 2002). Hill and Bickford (2001) also suggest that theTHO and similar age rocks of the Penokean orogeny mayextend much further southwestwards into Colorado andArizona.

However, since the final stages of continental collision ca.1.8 Ga, the MS-THO has remained within the internal partsof North America. Nevertheless, many structures have beenreactivated episodically during the development of the intra-cratonic Athabasca basin, the Williston basin, and the WesternCanada sedimentary basin (e.g., Ramaekers and Catuneanu2002; Elliott 1996) with concomitant movement of hydro-thermal fluids and control of sedimentary depocentres. Inaddition, localized intraplate magmatic activity has occurred,namely the intrusion of MacKenzie dykes and sills at 1.27 Ga,which may have helped drive fluid flow within the uranium-rich Athabasca basin, and by the emplacement and eruptionof diamondiferous kimberlites during the Cretaceous.

Acknowledgments

This study represents an attempt to compile and interpreta significant volume of work performed over the last 30 yearsby the Manitoba and Saskatchewan geological surveys, theGeological Survey of Canada, and numerous universityresearchers. It is dedicated to the memory of John Lewry,whose many ideas have generally stood the test of time.Funding provided by Natural Sciences and EngineeringResearch Council of Canada and Lithoprobe research grants.

Reviews by Ken Ashton, Pat Bickford, Ken Collerson, andRon Clowes are greatly appreciated.

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