U–Pb geochronological constraints for Paleoproterozoic ...directory.umm.ac.id/Data Elmu/jurnal/P/Precambrian Research/Vol10… · The Core Zone of the Southeastern Churchill Province,

Precambrian Research 103 (2000) 31–54

U–Pb geochronological constraints for Paleoproterozoicevolution of the Core Zone, southeastern Churchill

Province, northeastern Laurentia

Donald T. James a,*, Greg R. Dunning b

a Geological Sur6ey of Newfoundland and Labrador, P.O. Box 8700, St. John’s Newfoundland, Canada A1B 4J6b Department of Earth Sciences, Memorial Uni6ersity of Newfoundland, St. John’s Newfoundland, Canada A1B 3X5

Received 9 September 1999; accepted 23 February 2000

Abstract

The Core Zone of the Southeastern Churchill Province, northeastern Laurentia, is a composite of Paleoproterozoiclithotectonic domains assembled in the vise between obliquely colliding Archean Superior and North Atlantic cratonsbetween ca. 1860 and 1810 Ma. Detailed geological and U–Pb geochronological studies of domains in the southernCore Zone highlight significant differences in Archean and Paleoproterozoic geology between domains, and constraintiming and models for Paleoproterozoic assembly. In the south-central part of the Core Zone, Crossroads and Ormadomains consist of Late Archean granite–greenstone terrane crust. Crossroads domain also contains a significantnumber of granite–charnockite intrusions belonging to the 1840–1810 Ma De Pas batholith; magmatism pre-datesand overlaps with ca. 1820–1775 Ma high-grade metamorphism and attendant deformation. In marked contrast,Orma domain, which occurs to the east of Crossroads domain, contains only local evidence of Paleoproterozoic (DePas related?) intrusions and appears to have mainly escaped Paleoproterozoic metamorphism and penetrativedeformation. A model compatible with available data suggests the De Pas batholith is a continental magmatic arcformed above an east-dipping subduction zone. Paleoproterozoic metamorphism and deformation is concentrated inregions contiguous with arc magmatism. West of Crossroads domain, and separated from it by the Lac Tudor shearzone, McKenzie River domain is dominated by Archean orthogneiss (\80 m.y. older than gneisses in Crossroadsand Orma domains), minor amounts of Paleoproterozoic supracrustal rocks (absent in domains to the east) and 1815Ma tonalite. It does not contain De Pas intrusions. U–Pb geochronology suggests juxtaposition of the McKenzieRiver domain with domains to the east may have occurred around 1810 Ma. One possible model envisages theMcKenzie River domain as a piece of reassembled Superior craton crust, incompletely rifted from the Superiormargin during ca 2.17 Ga rifting. Domains occurring east of the Lac Tudor shear zone are interpreted to be exoticwith respect to the Superior craton. © 2000 Elsevier Science B.V. All rights reserved.

Keywords: Northeastern Laurentia; Paleoproterozoic tectonics; U–Pb geochronology

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* Corresponding author. Tel.: +1-709-7292774; fax: +1-709-7294270.E-mail address: [email protected] (D.T. James).

0301-9268/00/$ - see front matter © 2000 Elsevier Science B.V. All rights reserved.

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D.T. James, G.R. Dunning / Precambrian Research 103 (2000) 31–5432

1. Objectives

There are few geochronological constraints onthe intrusive and tectonothermal history of thesouthwestern Core Zone (James and Dunning,1996; James et al., 1996) of the SoutheasternChurchill Province (SECP), northeastern Lauren-tia (Fig. 1). A better understanding of the CoreZone (Fig. 2), a fundamental Paleoproterozoictectonic division consisting of a tectonic collageof Archean and Paleoproterozoic rocks, is criti-cal because it contains not only a record of itsinternal construction but also of its assemblywith bordering Archean cratons and their at-tached Paleoproterozoic supracrustal sequences.In this paper, we address some of the outstand-ing problems relating to Paleoproterozoic con-struction of three domains, McKenzie River,

Crossroads and Orma domains, that make upthe southwestern Core Zone. In particular, weinvestigate the timing and significance of graniticand mafic intrusions, their relationship to meta-morphic events and the development of majorstructures. The timing relationships are estab-lished as the result of regional geological studies(James et al., 1993; James and Mahoney, 1994),and detailed field and geochronological analysisof critical exposures containing unequivocal fieldrelationships (James and Dunning, 1996). Theresults, when combined with available data fromother areas in the southern SECP, which wesummarize in section three of this paper, help toconstrain a model for Paleoproterozoic evolutionof the Core Zone. This model has broader im-plications for the overall development of north-eastern Laurentia.

Fig. 1. Principal tectonic elements of northeastern Laurentia including areas in northeastern North America and western Greenland.

D.T. James, G.R. Dunning / Precambrian Research 103 (2000) 31–54 33

Fig. 2. Tectonic elements of Labrador and northeastern Quebec. Mrd, McKenzie River domain; Crd, Crossroads domain; Od, Ormadomain; M-Rd, Mistinibi-Raude domain; Kd, Kuujjuaq domain. Mesoproterozoic intrusions are indicated by the open trianglepattern.

2. Introduction

The SECP is a 300 km wide, north-trendingcomposite tectonic belt of Archean and Pale-oproterozoic rocks that is one segment of a sys-tem of Paleoproterozoic orogens linking Archeancratons in northeastern Laurentia. It is principallya continuation of the Trans-Hudson Orogen,which can be traced around the western, northernand eastern margins of the Superior craton, but italso shares common elements with the Nagssug-toqidian Orogen of Greenland (Fig. 1). The SECPis exposed from Ungava Bay, where it disappearsunder Hudson Strait, to southern Labrador,where it is truncated by east–northeast-trendingstructures and tectonostratigraphic units thatmake up the Grenville Province. The SECP reap-pears north of Hudson Strait, on southern BaffinIsland, although the precise correlation of majorstructures and tectonostratigraphic units betweenthe two regions is speculative and the focus of

ongoing studies (see St-Onge et al., 1997; Scottand St-Onge, 1998; St-Onge et al., 1998, 1999).

The SECP formed as a result of relative north-ward movement and sequential collision ofArchean North Atlantic (Nain) and Superior cra-tons, and attached Paleoproterozoic supracrustalsequences, with an Archean craton(s) whichresided to the north (see Hoffman, 1990; VanKranendonk et al., 1993). Also involved in thecollisions were Archean crustal blocks of suspectparentage that are now confined to the interven-ing regions between the intact Archean cratons.The SECP has a broadly tripartite character con-sisting of (from west to east): (1) a west-vergingfold-and-thrust belt (New Quebec Orogen) devel-oped in 2.17–1.86 Ga sedimentary and volcaniccover rocks, and involving Superior craton base-ment, (2) a medial hinterland or composite ter-rane (the Core Zone) having Archean andPaleoproterozoic components, and (3) a doubly-verging, fan-shaped wedge (Torngat Orogen) de-


veloped primarily in juvenile (B1.95 Ga) Pale-oproterozoic sediments and inferred to representan accretionary complex along the suture betweenthe Core Zone and the North Atlantic craton(Nain Province) (Fig. 2). Dextral (west) and sinis-tral (east) transcurrent shear zones, which aresynchronous-to post-tectonic with respect tothrusting in the New Quebec and Torngat oro-gens, respectively, separate the bordering ‘fore-land’ orogens from the Core Zone. The CoreZone is itself a mosaic of variably reworkedArchean crustal blocks (Van der Leeden et al.,1990; Wardle et al., 1990; Nunn et al., 1990;James et al., 1996; Isnard et al., 1998), ca. 2.3 Gaand B1.95 Ga supracrustal rocks (e.g. Van derLeeden et al., 1990; Girard, 1990; Scott and Gau-thier, 1996), and 1.84–1.81 Ga granitoid rocksbelonging to the De Pas and Kuujjuaq batholiths(Perreault and Hynes, 1990; Dunphy and Skulski,1996; James et al., 1996). Subsequent to CoreZone amalgamation, the Core Zone and border-ing orogens were overprinted by transcurrentshearing, which persisted locally to 1.74 Ga(Wardle and Van Kranendonk, 1996).

Affinity of Archean crust in the Core Zone isan outstanding problem having significant impli-cations for developing Paleoproterozoic tectonicmodels for the region. Addressing this problem ina comprehensive way is outside of the scope ofthis paper, although one possible model proposesthat the majority of Archean Core Zone crust isexotic with respect to the Superior and NorthAtlantic cratons. In contrast, other models sug-gest that Archean rocks in the Core Zone werepart of the Superior craton prior to 2.2 Ga (Jameset al., 1998; Scott and St-Onge, 1998). Availablegeochronological data from the Core Zone indi-cates that Archean Core Zone rocks have broadlysimilar intrusive ages as rocks in the northeasternSuperior craton, although this provides only cir-cumstantial evidence of their parentage. TheArchean geochronological data from Core Zonerocks is non-unique and could be used to supporteither of the models. However, if a significantcomponent of the Archean crustal blocks in theCore Zone did belong to the Superior Cratonprior to 2.2 Ga, there is general consensus thatthese blocks acted as independent crustal units,

relative to the bounding Archean cratons, duringPaleoproterozoic construction of the SECP(Wardle, 1998). There are no compelling geologi-cal or geophysical data to suggest the Core Zoneincludes a significant amount of North Atlanticcraton crust, although minor amounts or variablyreworked North Atlantic craton rocks may occurin regions adjacent to the boundary between theCore Zone and North Atlantic craton (e.g. Ryan,1990).

3. Tectonic elements of the southwestern SECP

3.1. Superior craton

The Archean Superior craton contiguous withthe southwestern SECP (Fig. 3) consists of high-grade Archean gneisses, part of the 90 000 km2

Ashuanipi Complex (Percival, 1991; James, 1997).The southeastern Ashuanipi Complex, containedin Fig. 3, mainly consists of \2.7 Ga metapeliticgneiss, orthogneiss, metamorphosed mafic intru-sions, diatexite plutons of orthopyroxene-bearingmonzogranite to granodiorite, and granite intru-sions. The sedimentary precursors of themetapelitic gneiss were intruded by plutons andrelated dykes of tonalite, gabbro and granite atca. 2.7 Ga prior to a regional tectonothermalevent at 2.68–2.65 Ga (Mortensen and Percival,1987). Deformation was accompanied by gran-ulite–facies metamorphism and emplacement ofthe diatexite plutons, and followed by intrusion oflate syn- to posttectonic plutons of biotite9mus-covite granite. Of relevance to our discussion ofthe evolution of the Core Zone which follows, itshould be noted that rocks belonging to theAshuanipi Complex do not occur in the CoreZone.

3.2. New Quebec Orogen (NQO)

The Superior craton is unconformably overlainby a succession of Paleoproterozoic sedimentaryand volcanic rocks, previously described as theLabrador Trough (see Dimroth, 1972; Wares andGoutier, 1990). These rocks, as well as variablyreworked Superior Province rocks and ‘exotic’


Paleoproterozoic supracrustal rocks, are con-tained in low-grade, thrust bound slices that makeup a west-verging, Paleoproterozoic fold-and-thrust belt defined as the NQO. An understandingof the stratigraphic and structural relationships inthe NQO is critical to understanding the CoreZone because stratigraphic (e.g. rifting) and struc-tural (e.g. initiation of thrusting) events in theforeland will be reflected or reflect events in thehinterland (i.e. Core Zone).

The Schefferville Zone, comprising the struc-turally lowest tectonostratigraphic unit in thesouthern NQO, consists of two sedimentary andvolcanic cycles of the Knob Lake Group. Ingeneral, the two cycles record a transition fromcontinental sedimentation and local alkaline vol-canism to progressively deeper water sedimenta-

tion and tholeiitic basaltic volcanism (seesummaries in Skulski et al., 1993; Wardle andVan Kranendonk, 1996). The older sequence (Cy-cle one) of arkose, shale and dolomite recordsinitiation of rifting and development of a passivemargin sequence on extended Superior cratoncrust. Cycle one rocks are overlain discon-formably by a westwards-overstepping youngersequence (Cycle two) of quartzite, iron formation,sandstone–shale turbidites and arkose.

The Schefferville Zone is structurally overlainby the Howse Zone, which is dominated by thicksequences of siltstone–shale turbidites interbed-ded with tholeiitic basalt and gabbro–sill com-plexes. However, it also contains rocks equivalentto cycles one and two described in the precedingparagraph. A possible interpretation for the tur-

Fig. 3. General geology of a transect across the southern SECP. The study area (this paper) is located in the southwestern part ofthe transect (see Fig. 4). Mg, Mesoproterozoic granitoid rocks; ARSZ, Ashuanipi River shear zone; LTSZ, Lac Tudor shear zone;GRSZ, George River shear zone. Paleoproterozoic thrust faults are decorated with widely spaced teeth. Grenvillian thrusts(southwest corner of Fig. 3) are decorated with closely spaced teeth.


bidites, basalts and related sills, is that they mayhave been deposited in a narrow, dextral transten-sional basin, perhaps analogous to the Gulf ofCalifornia (Wardle et al., 1990; Skulski et al.,1993). Cycle one and two volcanic rocks and sillsin the Howse Zone are dated at ca. 2.17–2.14 Gaand ca. 1.89–1.87 Ga (Birkett et al., 1991; Skulskiet al., 1993; Rohon et al., 1993), respectively.These ages are interpreted to correspond withtime of deposition for cycles one and two. If thenarrow, transtensional basin model for Cycle tworocks is apropos, it would imply that rifted Supe-rior craton crust occurred to the east of Cycle tworocks at 1.87 Ga.

To the east, the Howse Zone is structurallyoverlain by the Doublet terrane composed ofDoublet Group mafic pyroclastic rocks, turbiditicsiltstones, mafic and ultramafic sill complexes.The rocks in the Doublet terrane lack unequivocalstratigraphic links with Cycle one or Cycle tworocks in the Schefferville and Howse zones, andon this basis they are inferred to be exotic withrespect to the Superior craton (Wardle et al.,1995). Doublet terrane rocks may represent tran-sitional crust formed at the boundary betweenextended Superior continental crust and Pale-oproterozoic oceanic crust.

The Laporte terrane (Van der Leeden et al.,1990) is the eastern-most unit of the NQO. Itmainly consists of presumed Archean gneisses andmetamorphosed Paleoproterozoic pelitic, arkosicand mafic volcanic rocks, and lesser amounts ofquartzite, marble and gabbro. There are no obvi-ous correlations between supracrustal rocks in theLaporte terrane and Paleoproterozoic rocks oc-curring to the west, although the Archean rocksmay represent reworked Superior craton crust.The supracrustal rocks may be the relics of anaccretionary prism (Wardle, 1998).

3.3. The Core Zone

The Core Zone (James and Dunning, 1996;James et al., 1996) is informally defined as acomposite Paleoproterozoic terrane consisting oflithotectonic domains separated by Paleoprotero-zoic ductile high-strain zones (Fig. 3). The do-mains include Archean and Paleoproterozoic

(\1.8 Ga) rocks that cannot be unequivocallylinked with the Superior or North Atlantic cra-tons, or with Paleoproterozoic supracrustal se-quences which occur in the New Quebec orTorngat orogens. The western boundary of theCore Zone is marked by the Ashuanipi Rivershear zone (James et al., 1996) in the south andthe Lac Turcotte fault (see Perreault and Hynes,1990) in the north. The western margin of theAbloviak shear zone defines the eastern boundary.This definition includes the Kuujjuaq domain andthe Lac Lomier complex as part of the CoreZone; this is perhaps noteworthy as these are notincluded in Wardle’s definition (Wardle, 1998). Inparticular, tectonic affinity of the Kuujjuaq do-main (Perreault and Hynes, 1990) is somewhatcontroversial; some workers (e.g. Bardoux et al.,1998) consider the domain to be part of the NQO.The boundary between the NQO and the CoreZone is approximated by the inflection between apaired, negative (NQO side) and positive (CoreZone side) Bouguer gravity anomaly. On the basisof the gravity signature, which is consistent with amodel involving west-directed transport of CoreZone rocks over the NQO, Thomas and Kearey(1980) proposed that this boundary is a Pale-oproterozoic suture marking a relic Andean-typemargin. The offshore extension of the Core Zoneunderlying Ungava Bay has been traversed by theLITHOPROBE (ECSOOT) seismic reflection sur-vey, which showed the Core Zone to be domi-nated by 30° east-dipping reflectors that appear topenetrate the whole crust (Hall et al., 1995).

The southwestern Core Zone includes, fromwest to east, McKenzie River, Crossroads, Ormaand Mistinibi-Raude domains (see Van der Lee-den et al., 1990; Nunn et al., 1990; Girard, 1990;Nunn, 1994; James et al., 1996). The McKenzieRiver domain (discussed in detail below, and seeJames et al., 1996) consists mainly of Archeantonalite gneiss and lesser amounts of inferredPaleoproterozoic supracrustal rocks, which aremetamorphosed to upper amphibolite facies. TheMcKenzie River domain does not appear to sharemany Archean or Paleoproterozoic (\1815 Ma)features with domains to the east.

Crossroads domain (also discussed in detail be-low, and see James et al., 1996) contains relicts of


high-grade Archean granite–greenstone terranecrust and Paleoproterozoic granitoid intrusions,which are part of the \500 km long De Pasbatholith. The intrusions are variably deformedand recrystallized, demonstrating the domain hasbeen overprinted by a Paleoproterozoic tec-tonothermal event which partially overlapped andpostdated emplacement of the De Pas batholith.The western boundary with the McKenzie Riverdomain is the Lac Tudor shear zone (Van derLeeden et al., 1990), which deforms Archean andPaleoproterozoic rocks in both domains, and hascomponents of dextral transcurrent and reversedisplacement (Van der Leeden et al., 1990;Bourque, 1991; James et al., 1996). The GeorgeRiver shear zone (see Van der Leeden et al., 1990;Girard, 1990), which forms the eastern boundarywith the Orma domain, is a wide (several km), butvery diffuse and poorly defined zone of heteroge-neous strain containing porphyroclastic protomy-lonitic and mylonitic rocks having dextraltranscurrent kinematic indicators (James et al.,1996). Mylonitization in the Lac Tudor andGeorge River shear zones was attendant withPaleoproterozoic amphibolite–facies metamorph-ism.

Archean geology of the Orma domain is similarto that of the Crossroads domain. The Ormadomain contains relicts of Archean greenstonebelt rocks, which are intruded by tonalite or-thogneisses having igneous crystallization ages be-tween 2682 and 2675 Ma, as determined by U–Pbage dating of zircons (Nunn et al., 1990). Titanitedata from the same rocks suggest they were meta-morphosed to amphibolite–facies in the LateArchean. Notably, the U–Pb data show no evi-dence that the Archean rocks were overprinted byPaleoproterozoic thermal event(s) prior to the ca.1720–1600 Ma Labradorian Orogeny; all Pb-lossin the titanites is younger than ca. 1640 Ma(Nunn et al., 1990). Based on this data, the Ormadomain has been considered to have mainly es-caped the ca. 1820–1775 Ma Paleoproterozoictectonothermal event (discussed in section four)which overprinted the Crossroads and McKenzieRiver domains. However, at possible variancewith the U–Pb data, the domain also includesfoliated granitoid rocks, which are undated, but

suspected to be Paleoproterozoic (Nunn, 1994) onthe basis of lithologic correlations with datedintrusions in the Crossroads domain. The signifi-cance of this correlation will be discussed later.

The Orma domain also includes a sequence ofwacke, quartz wacke, quartzite, tuffaceous rocksand metamorphosed basalt, named thePetscapiskau Group (Emslie, 1970). PetscapiskauGroup rocks were not metamorphosed prior tobeing intruded by the Mesoproterozoic (ca. 1460Ma) Michikamau Intrusion (Emslie, 1970). Age ofthe Petscapiskau Group is unknown, but there aretwo possible scenarios worthy of some discussion.If the Petscapiskau Group is older than the Pale-oproterozoic tectonothermal event that over-printed the Crossroads and McKenzie Riverdomains (i.e. \1820 Ma), it would support exist-ing U–Pb data from Archean rocks and confirmthat the Orma domain escaped penetrative Pale-oproterozoic tectonothermal overprinting. If thePetscapiskau Group is younger than the Pale-oproterozoic event, it would leave open the possi-bility that the Orma domain has been overprintedby hitherto unrecognized Paleoproterozoic meta-morphism. It is remotely possible that thePetscapiskau Group is coeval with ca. 1650 MaLabradorian supracrustal sequences (e.g. theBlueberry Lake group (James and Connelly,1996), or MacKenzie Lake Group (Nunn, 1993)),which unconformably overly pre-1650 Ma rocksalong the southeastern margin of pre-LabradorianLaurentia. The Labradorian volcanic rocks pre-dominately have felsic and intermediate composi-tions whereas Petscapiskau Group volcanic rocksare mafic. This may suggest that a Labradorianage for the Petscapiskau Group is unlikely, but itdoes not discount the possibility. There is anobvious need for more geochronological datafrom the Orma domain.

The Mistinibi-Raude domain (Van der Leedenet al., 1990) is in tectonic contact with the Ormadomain along a northwest-striking, subverticalhigh-strain zone containing a strong, horizontalmineral elongation lineation (Nunn, 1994). Defor-mation and attendant amphibolite-facies meta-morphism in the zone are undated but presumedto be Paleoproterozoic. The domain includesmetasedimentary gneiss, granitoid migmatite, and


metabasic rocks that were intruded by pre- tosyn-tectonic granitic sheets; these rock units andthe high-grade metamorphism that overprintsthem are provisionally assigned to the Archean.The Archean rocks are intruded by the enigmaticca. 2.3 Ga Pallatin Intrusive Suite (see Van derLeeden et al., 1990; Krogh, 1992) and overlain bypresumed related sedimentary and volcanic rocksof the Ntshuku Complex (Girard, 1990). In thenorthern part of the domain, Ntshuku rocks areunconformably overlain by deformed sedimentaryrocks of the Paleoproterozoic Hutte SauvageGroup (Van der Leeden et al., 1990). TheNtshuku and Hutte Sauvage rocks are metamor-phosed from greenschist to lower-amphibolitefacies.

Mesoproterozoic (ca. 1460 Ma) intrusions ofanorthosite and related granitoid rocks intrudethe Orma and Mistinibi-Raude domains, and inthe southeastern Core Zone, these obscure thePaleoproterozoic Abloviak shear zone (Wardle etal., 1990) and the boundary with the North At-lantic craton. The Abloviak shear zone marks theapproximate eastern limit of penetrative Pale-oproterozoic deformation and defines the tectonicboundary with the North Atlantic craton. How-ever, in the area between the MesoproterozoicHarp Lake Intrusive Suite and the Nain PlutonicSuite, the eastern limit of Paleoproterozoic defor-mation is marked by deformed Ingrid Grouprocks, a sequence of conglomerates and volcanicrocks. Age of Ingrid Group volcanism is looselyconstrained; U–Pb zircon and titanite ages deter-mined from two samples of felsic volcanic rocksgive ages of ca. 1895 Ma and 1805 Ma (Wasteneyset al., 1996). Clasts in Ingrid Group conglomerateare derived from a variety of sources including\3.7 Ga North Atlantic craton (presumed SaglekBlock) crust and from ca. 1900 Ma granitoidrocks that may have been part of a Torngatmagmatic arc (Wasteneys et al., 1996) constructedon the North Atlantic craton. The Early Archeanclasts demonstrate that the sedimentary rockswere autochthonous with respect to North At-lantic craton crust.

The Ingrid Group is overthrust to the west bygranulite–facies granitoid migmatite determinedby U–Pb zircon geochronology to have an em-

placement age of 2870922 Ma (Wasteneys et al.,1996). The affinity of these Archean gneisses isuncertain; they resemble North Atlantic (Nain)craton rocks, although Ermanovics and Ryan(1990) note that west of the Ingrid Group,Archean gneisses lack Paleoproterozoic Kikker-tavak dykes, diagnostic of adjacent North At-lantic craton crust. Farther west, the Archeangneisses are tectonically interleaved with Tasiuyakgneiss, a distinctive metasedimentary gneiss,which can be traced for at least 500 km along theentire eastern boundary with the North Atlanticcraton, although it’s total strike length may ex-ceed 1300 km (Scott, 1998). The high-strain zonecontaining deformed Tasiuyak gneiss andArchean gneisses may be related to and coevalwith the sinistral transcurrent Abloviak shearzone. Detrital zircon populations and field rela-tions, which bracket the depositional age of Tasi-uyak sediments between 1940 and 1895 Ma (Scottand Machado, 1993; Scott and Gauthier, 1996),and Nd-isotopic data (Kerr et al., 1993; Theriaultand Ermanovics, 1993) suggest that sedimentaryprecursors of the Tasiuyak gneiss were exotic withrespect to the North Atlantic craton (Wardle andVan Kranendonk, 1996). Thus, Archean rocksthat occur to the west of the Tasiuyak gneissmight never have been part of the North Atlanticcraton. In conflict with this interpretation, Ryan(1990) has suggested that Archean gneisses andanorthositic rocks occurring more than 50 kmwest of Tasiuyak gneiss and Paleoproterozoicshear zones, in the area north of the Mistastinbatholith (north of the area shown in Fig. 3),correlate with North Atlantic (Nain) craton rocks.It may be possible that there is a wide (50 km?)region contiguous with the North Atlantic cratonwhich contains tectonically interleaved North At-lantic craton rocks and Archean rocks which areexotic with respect to the North Atlantic craton(e.g. Van Kranendonk et al., 1993).

Locally, the Mesoproterozoic intrusions andCore Zone rocks are unconformably overlain byclastic sedimentary rocks and gabbro sills of the1275–1225 Ma Seal Lake Group (Romer et al.,1995). Seal Lake Group rocks obscure theboundary between the Core Zone and the NorthAtlantic craton in the region south of the HarpLake Intrusive Suite.


Fig. 4. Geology of the southwestern core zone in the area around the Smallwood Reservoir, Labrador (southern part of NTS maparea 23I). Geochronological data summarized from James et al. (1996)(zr, zircon; ti, titanite; ig, igneous crystallization age; mt,metamorphic age). Ellipses show sample locations (this study). Most samples were collected from island outcrops in the SmallwoodReservoir. ARSZ, Ashuanipi River shear zone; LTSZ, Lac Tudor shear zone; GRSZ, George River shear zone.

4. U–Pb geochronological studies in thesouthwestern Core Zone

4.1. McKenzie Ri6er domain

The McKenzie River domain (Fig. 4) is sepa-rated from the NQO and the Crossroads domainby the Paleoproterozoic Ashuanipi River and LacTudor shear zones, respectively. Both shear zonesare inferred to be transpressive (dextral) mylonitezones having components of east-over-west re-

verse displacement. Field relations demonstratethey were developed concomitant with amphibo-lite–facies metamorphism. The Ashuanipi Rivershear zone corresponds approximately to the infl-ection between a regionally persistent, pairedBouguer gravity anomaly; negative on the NQOside and positive on the Core Zone side (seeJames et al., 1996, Fig. 2, page 218).

The domain is dominated by the Flat Pointgneiss, which has an emplacement age of 277695Ma as determined by U–Pb dating of zircon


(James et al., 1996). o Nd values from the FlatPoint gneiss from −0.5 to −5, at 2800 Ma (Kerret al., 1994), suggest the rocks incorporated acomponent of older crust. The rocks are meta-morphosed to upper-amphibolite facies and havea gneissosity which is folded into relatively openand flat-lying superposed, mainly dome-and-basin, folds. The age of the gneissosity and thefolding are unknown, but are assumed to beArchean. The gneissosity and superposed foldsare overprinted by a steep, north-striking andeast-dipping foliation accompanied by a pervasiverecrystallization that locally obliterates the gneis-sosity and preexisting structure. The north-strik-ing foliation is interpreted to be aPaleoproterozoic fabric for three reasons: (1) ithas the same attitude as the foliation in ca. 1815Ma tonalite (see discussion of sample MR1 be-low), (2) it is defined by the peak-assemblagemetamorphic minerals and amphibolite–faciesmetamorphism was attained by ca. 1805 Ma (seediscussion of sample MR2 below), and (3) itbecomes progressively more intense in areas nearthe major Paleoproterozoic high-strain zones (e.g.Lac Tudor shear zone), which define the domainboundaries.

Metamorphosed supracrustal rocks of un-known depositional age occur as thin (B1 km)tectonically bound units within the Flat Pointgneiss. Informally named the Lobstick group,they include metasedimentary (pelitic) migmatiteand lesser amounts of quartzite, marble, calcsili-cate derived from impure siliceous carbonate, andamphibolite of uncertain protolith. Thesesupracrustal rocks may be correlative with litho-logically similar, Laporte group rocks, which oc-cur to the north of the study area. Geochronologysamples were collected from an outcrop contain-ing migmatitic Lobstick group rocks, a tonalitedyke which cross cuts the metamorphic leuco-some, and meta-tonalite which is in tectonic con-tact with the Lobstick group rocks. Thesupracrustal rocks are steeply foliated and havesteeply plunging folds of relict sedimentary bed-ding. They do not contain the superposed foldstructures contained in the Flat Point gneiss. Onthe basis of orientation, the foliation in thesupracrustal gneisses is correlated with the folia-tion in Paleoproterozoic tonalite (see discussion ofsample MR1 below).

Sample MR1 is from a unit of strongly foliatedand recrystallized tonalite that is in tectonic con-tact with Lobstick group supracrustal gneiss.MR1 tonalite differs significantly from the FlatPoint gneiss in that it lacks metamorphic layeringand the superposed fold structures that character-ize the latter. The sample yielded three fractionsof concordant and nearly concordant zircons (Fig.5, Table 1), and the two overlapping concordantanalyses give an age of 181593 Ma, interpretedto represent the igneous crystallization age. Thesedata represent the first indication of Paleoprotero-zoic granitoid intrusive rocks in the McKenzieRiver domain, although the extent of ca. 1815 Matonalite is unknown.

To determine the age of metamorphism, a sam-ple (MR2) of K-feldspar+biotite+garnet-bear-ing leucosome contained in pelitic migmatite wascollected. Two fractions of monazite (Fig. 5) areconcordant (M2) and nearly concordant (M1) andindicate that monazite crystallized in the leuco-some at 180593 Ma, and demonstrate that up-per-amphibolite facies metamorphism wasattained by this time. The data only loosely con-

Fig. 5. Schematic diagram of an outcrop in the McKenzieRiver domain showing field relationships and U–Pb concordiadiagrams for samples MR1 to MR3. In the field, the distancebetween MR1 and MR3 sample locations is less than 50 m.

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Table 1U–Pb analytical resultsa

Age (Ma)Corrected atomic ratiosConcentration

206 Pb/204/Pb 208 Pb/206 Pb 206 pb/208 U 207 Pb/235 U 207 Pb/206 Pb 206 Pb/238 UPb(c) 207 Pb/235 UFraction description 207 Pb/206 PbPb(r)UWeight(mg) (pg)(ppm)

MR1 meta-tonalite (366167E, 60102040N)37Z1 3339Large clear 0.2451 0.32534 (124) 4.9747 (202) 0.1109 (14) 1826 1815 18140.121 50 19.2

prisms6490 0.2839 0.32295(122) 4.9282(182) 0.11067 (20) 180420.8 180754 181110Z2 Clear, euhedral 0.06

Z3 12546Large, euhedral 0.2439 0.32584 (142) 4.9908 (208) 0.11109 (24) 1818 1818 18030.083 86 32.8 12

MR2 leucosome, msed migmatite (3861617E, 601204N)9633 39.246 0.32337 (190) 4.9287 (296) 0.11054 (12) 1806 1807 1808M1 Large, clear 0.091 293 3317.3 5610619 45.062 0.32294 (132) 4.9069 (206) 0.1102 (14) 1804 1803 180348M2 Clear, yellow 0.1 252 3253.3

MR3 tonalite dyke (386167E, 6010204N)8918 0.1183 0.32667 (124) 5.0751 (200) 0.11268 (14)Z1 1822Small, clear 1832 18430.021 512 177.1 243890 0.1461 0.32617 (100) 5.0618 (168) 0.11255 (14) 182038 1830Z2 1841215.66100.012Clear prisms

27Clear, cracks 7645 0.1368 0.33024 (110) 5.2562 (188) 0.11544 (12) 1840 1862 18870.023 429 152.5Z318072 0.1344 0.32499 (112) 5.0302 (186) 0.11226 12) 1814 1824 1836160.034 410 142.9Clear prismsZ4

CR1 tonalite orthogneiss (386602E, 601415N)Large, cleaar, 41.4 7 17440 0.1387 0.47972 (154) 11.5395 (400) 0.17446 (16) 2526 2568 26010.054Z1 76euhedral

7Z2 26173Small, brown 0.1319 0.48777 (150) 11.9289 (400) 0.17737 (16) 2561 2599 26280.071 85 46.9prisms

18 13576 0.1339 0.48608 (164) 11.8186 (428) 0.17634 (16) 2554 259082 26190.098Large, clearZ3 45.1euhedral

CR2 leucogranite (386602E, 601451N)2843 0.2202 0.31757 (118) 4.8755 (188)Z1 0.11135 (16)Brown prisms 1778 1798 18210.009 251 91.3 15

6 5907 0.2364 0.32289 (162) 4.9749 (226) 0.11174 (30) 1804121.3 1815Z2 18283240.005Small, brownprismsSmall, brown 111 48 4209 0.2231 0.32588 (94) 5.1139 (166) 0.11381 (10) 1818 1838 18610.033Z3 296needles

CR3 diorite dyke (366002E, 601415N)12075 0.2525 0.32278 (86) 4.919 (290)Best clear prisms 0.11053 (10)0.034 1803 1806 18081833 694.2Z1 105

471 174.3 23 18209 0.2262 0.32139 (134) 4.8995 (214) 0.11057 (10) 1797 1802 1809Z2 0.045Small, brownprisms

3867.3 218 4777 10.996 0.32161 (152) 4.8551 (240) 0.10949 (10) 1798 1795 1791M1 0.045Clear, pale yellow 1146

CR4 pegmatite (386602E, 6014151N)24977 0.1889 0.33642 (1556) 5.4943 (268) 0.11845 (10)0.192 1869Z1 1900 1933Large, pale pink 160 60.4 26

56.4 7 19898 0.1909 0.3244 (96) 4.9989 (162) 0.11176 (12) 1811 1819 18280.046Z2 1551 large prism73 16946 39.429 0.32165 (186) 4.8867 (286) 0.10968 (12)M1 1798Large, dark 1797 17960.117 524 5926.6

yellow18625 29.371 0.32222 (108) 4.873(176) 0.10968 (12) 18015786.6 1798Clear, pale yellow 179457M2 6790.077

37Clear, pale yellow 15446 21.575 0.32121 (112) 4.8485 (180) 0.10948 (12) 1796 1793 17910.048 595 3759.8M3

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Table 1 (Continued)

Age (Ma)Concentration Corrected atomic ratios

Pb(r) Pb(c) 206 Pb/204/Pb 208 Pb/206 Pb 206 pb/208 U 207 Pb/235 U 207 Pb/206 Pb 206 Pb/238 U 207 Pb/235 U 207 Pb/206 PbWeightFraction description U(mg) (pg)(ppm)

4640.4 167 5236 24.456 0.32083 (138) 4.8429 (218) 0.10948 (10) 1794 1792 1791ClearM4 0.066 652

CR5 spotted diorite duke (39138E, 6014188N)7363 0.2571 0.31986 (120) 4.8956 (192) 0.11101 (12)0.038 1789196 1802 1816Clear, cracksZ1 73.9 20

8Large, cracks 37935 0.3131 0.3206 (112) 4.9098 (184) 0.11107 (12) 1793 1804 18170.058 271 106.5Z240810 0.3295 0.32262 (110) 4.9396 (180) 0.11105 (10) 180223 1809277 1817110.9Z3 Prisms, cracked 0.168

12Elongate prisms 8498 0.3297 0.32406 (112) 4.9618 (180) 0.11105 (14) 1810 1813 18170.023 224 90.2Z43Clear prisms 5190 0.2768 0.3253 (130) 4.9847 (196) 0.11114 (20) 1816 1817 18180.013 53 20.8Z5

381 0.098 0.31651 (104) 4.7448 (198) 0.10872 (20) 17731191 1775T1 177813.7420.525Clear brown801 458 0.0951 0.31696 (98) 4.7592 (184) 0.1089 (18) 1775 1778T2 1781Clear, medium 0.349 51 16.8

brown533 474 0.1191 0.31821 (108) 4.7822 (200) 0.109 (22)57 17810.217 1782 178319T3 Medium brown,

prismsT4 34115 large brown 355 0.0916 0.31717 (104) 4.7466 (194) 0.10854 (22) 1776 1776 17750.139 42 13.7

prisms

CR6 granite dyke (386167E, 6010304N)122.5 34 4420 0.0951 0.31967 (116) 4.8567 (182) 0.11019 (16) 1788 1795 1803Z1 0.022 large grains 370

9459 0.1036 0.30827 (118) 4.6515 (176) 0.10944 (18) 173214 1759Z2 1790126.73940.018Small grain4 32919 0.1246 0.32029 (142) 4.8659 (222) 0.11018 (12) 1791 1796 1802Z3 Small, clear, 0.025 261 88.9

euhedral28133 0.098 0.3165 (136) 4.8557 (212) 0.11017 (14) 1788103.5 17954 18023120.02Clear, euhedralZ4

Z5 111933 large grains 0.2565 0.31813 (186) 4.8344 (262) 0.11021 (30) 1781 1791 18030.022 113 42.5 4

CR7 De Pas monzogranite (389070E, 6016195N)15Z1 7451Sharp, elongate 0.1797 0.3225 (114) 4.9203 (176) 0.11065 (18) 1802 1806 18100.023 240 85.9

prisms7724 0.1896 0.32281 (156) 4.925 (228) 0.11065 (24) 180387.3 1807242 181011Z2 Sharp, elongate 0.018

a The analytical methods followed by Dunning at Memorial University of Newfoundland are described in Dube et al. (1996) and references therein. UTM co-ordinates for each sample are shown in parentheses. All UTMco-ordinates are for Grid Zone 20, NAD 1927, and NTS map area 23I (Woods Lake map sheet). Z-zircon. T-titanite. M-Monazite. Pb(r) is the total radiogenic lead after correction for blank, common lead and spike. Pb(c)represents the picograms of common lead in the analysis. Corrected atomic ratios are corrected for fractionation, spike, 4–10 pg laboratory blank, initial common lead (calculated using the model of Stacey and Kramers(1975) for the age of the sample), and 1 pg uranium blank. Numbers in parentheses after the corrected atomic ratios refer to uncertainties of 2 s on the final digits of the isotope ratios. The uncertainties were calculatedusing an unpublished error propagation program, as reported in Dube et al. (1996).


Fig. 6. Lobstick group supracrustal rocks including peliticmigmatite (right), calc–silicate gneiss (centre) and quartzite(left). The metasedimentary rocks are cut by a late syn-tectonictonalite dyke (folded). Sample MR3 was collected from asimilar dyke.

45% probability of fit which yields a lower inter-cept age of 1802 +9/−14 Ma, interpreted torepresent the igneous crystallization age of thedyke. The upper intercept is 2850 Ma with a largeuncertainty. This brackets formation of thesteeply east-dipping foliation in the host rocks tobe between 1815 Ma, the emplacement age of thefoliated tonalite (MR1) and 1802 Ma.

The field and geochronological data demon-strate that Archean Flat Point gneiss and \1815Ma Lobstick group rocks were intruded bytonalite, metamorphosed to upper-amphibolite–facies and deformed in the interval between 1815and 1802 Ma. The data indicate that the gneissos-ity and superposed folds in the Flat Point gneissare older than 1815 Ma. It is tacitly assumed thatthese features are Archean, although this has notbeen confirmed by a geochronological test.

4.2. Crossroads domain

Field relations demonstrate that the oldestrocks in Crossroads domain are high-gradesupracrustal gneisses, informally named theOverflow group, consisting principally ofmetasedimentary (pelitic) migmatite (Fig. 7), mi-nor amounts of mafic and felsic metavolcanicrocks and associated chert–magnetite iron forma-tion. The precise age of the supracrustal rocks isundetermined, although they are constrained tobe \2704 Ma on the basis of U–Pb geochrono-logical data from younger intrusive units (de-scribed below). o Nd values for two samples of themetasedimentary rocks are approximately +1and +2 at 2700 Ma, and they have depleted-mantle model ages (T dm ages) of approximately2800 Ma (Kerr et al., 1994). The Nd data indicatethe rocks were probably not derived from erosionof Middle or Early Archean crust, and suggestthat their depositional ages are between 2700 and2800 Ma. One possibility is that they could havebeen derived from erosion of coeval Overflowgroup volcanic rocks. These supracrustal rocksare provisionally correlated with similar Archeanhigh-grade metasedimentary and metavolcanicrocks in the Orma domain (see Nunn and Noel,1982; Nunn, 1993).

Fig. 7. Archean pelitic migmatite (Overflow group) containingsuperposed folds of Archean age and a deformed Archeanamphibolite dyke, Crossroads domain.

strain the metasedimentary rocks to be older than1805 Ma. They are probably older than 1815 Ma,the age of the MR1 tonalite, although field rela-tions do not unequivocally prove this. A sampleof paleosome from the pelitic migmatite has adepleted-mantle Nd-model age of 2310 Ma (Kerret al., 1994).

To place minimum constraints on the age of thefoliation and leucosome formation in the Lobstickgroup rocks, sample MR3 was collected from aweakly deformed, late syn-tectonic tonalite dyke(Fig. 6) which cross-cuts the supracrustal rocks,their foliation and included leucosomes. Fourfractions of zircons (Fig. 5) define a line having a


Overflow group rocks are intruded by variablydeformed and metamorphosed plutonic units in-cluding tonalite and granite orthogneisses thatcontain several phases of metamorphic leuco-some, granitic rocks belonging to the ca. 1840–1810 Ma De Pas batholith, deformed graniticplutons, some of which are probably related tothe De Pas batholith, and several ages of variablydeformed and metamorphosed mafic dykes. Tobetter understand age relationships between thevarious units, samples from one outcrop contain-ing unequivocal contact relationships (Fig. 8)were collected for U–Pb geochronological studies.

Sample CR1 is from the paleosome of a tonaliteorthogneiss which intrudes Overflow groupsupracrustal gneisses. Three zircon fractions fromthe sample define a discordia line (Fig. 8, Table 1)with an upper intercept of 2704915 Ma, inter-preted to represent the igneous crystallization ageof the rock. This age is significantly older than theca. 2620 Ma emplacement age determined for amonzogranite orthogneiss from the southeasternCrossroads domain (James et al., 1996), althoughit is consistent with 2682–2675 Ma emplacementages of tonalite intrusions in the Orma domain

(see Nunn et al., 1990). The lower intercept of thediscordia line is 1815 Ma, and is thought torepresent incipient Pb-loss during a ca. 1815 Mametamorphic event. Crossroads domain or-thogneisses have o Nd values of between 0 and+2 at 2650 Ma, and T dm model ages between2800 and 2650 Ma (Kerr et al., 1994). The Nd andU–Pb data indicate that these intrusions are juve-nile, Late Archean additions to the crust. Nd datafrom Orma domain tonalite orthogneiss are simi-lar; rocks have o Nd values of +1 at 2675 Maand T dm ages of approximately 2770 Ma (Kerret al., 1994).

The gneissosity in CR1 orthogneiss is cut by apink, recrystallized leucogranitic dyke, which isdeformed by a locally intense foliation that alsooverprints the host orthogneiss. A discordiadefined by two fractions of zircon collected fromthe dyke (CR2, Fig. 8) suggest an igneous crystal-lization age of 1836910 Ma. This age is coeval,within error, of the 183195 Ma age (James et al.,1996) determined from a sample of De Pasbatholith granite collected from the southernCrossroads domain. On this basis, the dyke isinterpreted to be related to De Pas magmatism.

Fig. 8. Sketch map of an outcrop in Crossroads domain showing field relationships and U–Pb concordia diagrams for samples CR1to CR4. The fresh, ENE-striking mafic dyke (Mesoproterozoic?) is undeformed and does not contain minerals suitable for U–Pbdating. The diagram represents an area of approximately 600 m2.


Fig. 9. Metamorphosed diorite dyke (CR3 sample location;left side of photograph) cutting CR1 tonalite orthogneiss.

The data constrain the age of the gneissosity inthe host rocks (i.e. in sample CR1) to be \1836Ma, and as Archean rocks in the domain do notcontain any isotopic evidence of thermal eventsbetween ca. 2620 and 1836 Ma, we interpret thegneissosity to be an Archean feature.

The tonalite orthogneiss, leucogranite dyke andit’s contained foliation are cross-cut by a grey-weathering, recrystallized and weakly deformeddiorite dyke (CR3, Figs. 8 and 9), which is in turncut by an undeformed granitic pegmatite (CR4).The igneous crystallization age of the diorite dykeis interpreted to be 180992 Ma based on twofractions of concordant and nearly concordantzircons. A single monazite fraction from CR3 wasdated at 179593 Ma and interpreted to representthe time of metamorphism. The pegmatite dyke(CR4, Fig. 8) is interpreted to have a crystalliza-tion age of 180092 Ma based on data fromzircon and three concordant monazite analyses.

To provide additional constraints on the age ofthe supracrustal rocks, their included metamor-phic and structural features, and ages of intru-sions, samples of a diorite dyke (CR5) and agranite dyke (CR6) were also dated. The dioritedyke (Fig. 10), which has a similar mineralogyand texture to the CR3 dyke, cross-cuts relictprimary layering, gneissosity and foliation in hostOverflow group mafic metavolcanic rocks. How-ever, the diorite dyke is itself metamorphosed, itcontains a distinctive ‘spotted’ hornblende–por-phyroblastic texture, and has a weak foliation.Five fractions of zircon (Fig. 11) define a discor-dia line with an upper intercept of 181792 Ma,interpreted to represent the igneous crystallizationage of CR5. Four fractions of titanite from thesame rock define an age of 1775 Ma. The titanitedata may represent a metamorphic cooling age, orthey may indicate renewed thermal activity andcrystallization of new titanite at 1775 Ma. SampleCR6 is from an undeformed, white-weatheringgranite dyke that is discordant to gneissosity inhost Overflow group metasedimentary migmatite.Five fractions of zircon define a discordia line(85% probability of fit) with an upper interceptage of 180695 Ma (Fig. 12) and interpreted to bethe crystallization age of the rock.

The Crossroads domain contains intrusions ofvariably foliated and recrystallized K-feldspar

Fig. 10. Metamorphosed (‘spotted’) diorite dyke (CR5 samplelocation, lower left) cutting Archean mafic and felsic volcanicrocks (left side of photo) of the Overflow group.

Fig. 11. U–Pb concordia diagram for sample CR5.


megacrystic granite, granodiorite and charnockitebelonging to the main De Pas batholith, sensustricto, and presumed satellite intrusions of theDe Pas batholith, which occur to the east of themain batholith. The satellite intrusions are corre-lated with the batholith on the basis of lithology.Two samples of De Pas batholith megacrysticgranite, from the main part of the batholith in thesouthern Crossroads domain, have emplacementages of 183195 Ma (James et al., 1996) and181193 Ma (Krogh, 1986), as determined byU–Pb dating of zircon. One of the presumedsatellite intrusions, consisting of strongly foliatedmegacrystic granite, is dated at 182395 Ma(James et al., 1996). From farther north in thebatholith, Dunphy and Skulski (1996) have deter-mined that a foliated De Pas batholith tonalite

has an emplacement age of 1840 Ma on the basisof preliminary U–Pb dating of zircon.

In an attempt to obtain minimum ages of em-placement for De Pas K-feldspar megacrysticgranite, and to constrain the timing of deforma-tion that overprints these rocks, a unit of isotropicto very weakly foliated, pink, biotite monzogran-ite containing xenoliths of strongly foliated K-feldspar megacrystic granite was sampled. On thebasis of lithology and structure, the xenoliths arecorrelated with foliated De Pas K-feldsparmegacrystic granite. Field relations suggest thatthe biotite monzogranite (CR7) is late syn-tec-tonic with respect to the deformation in the in-cluded megacrystic granite. Two fractions ofconcordant zircons from sample CR7 (Fig. 13)yield an age of 181093 Ma, interpreted to repre-sent the crystallization age of the rock. This age iswithin error of the youngest emplacement agesfrom the De Pas batholith. The strong foliation inthe batholith is inferred to have formed between1810 and 1823 Ma.

5. Discussion

5.1. Archean e6olution

Field, geochronological and Nd-isotopic dataindicate that the Crossroads and Orma domainscontain a significant component of Late Archeangranite–greenstone terrane crust. Supracrustalrocks in both domains are undated but are con-strained to be older than 2704 Ma. They areintruded by tonalite to granite plutons in theinterval between ca. 2704 and 2620 Ma; these areinferred to be pre- to syn-tectonic with respect tohigh-grade metamorphism and attendant defor-mation. Granitic intrusions correlated with the DePas batholith, and mafic dykes correlated with the1817–1809 Ma dykes, are discordant to super-posed folds of metamorphic leucosome in thesupracrustal rocks in the Crossroads domain, sug-gesting that these features are Archean. A possibleinterpretation of the data is that Archean rocks inCrossroads and Orma domains have a similarArchean depositional, intrusive and tectonother-mal evolution.




The McKenzie River domain lacks Archeansupracrustal rocks and the ca. 2776 Ma Flat Pointgneiss is significantly older than intrusive units inthe Crossroads and Orma domains. In contrast tothe isotopically juvenile intrusions in the Cross-roads and Orma domains, negative o Nd values,at 2800 Ma, and T dm ages\3.0 Ga (Kerr et al.,1994) from the Flat Point gneiss suggest that therocks incorporated a component of older crust.Archean rocks in the McKenzie River domainwere probably not part of the same granite–greenstone terrane that makes up the Crossroadsand Orma domains.

Parentage of Archean rocks that occur in theCore Zone of the southwestern SECP is uncertain.Archean data from these rocks is non-unique andcould be used to support radically different mod-els. However, the fact that the entire exposedmargin of the Superior craton shows evidence of asignificant rifting event, initiated between ca. 2.2and 2.0 Ga, and reflected in the development ofCycle one rocks in the southern LabradorTrough, strongly suggests that even if Archeanrocks in the Core Zone have a Superior cratonaffinity, they acted as independent crustal blocksafter ca. 2.0 Ga. Moreover, if the De Pasbatholith is a subduction-related magmatic arc,then Paleoproterozoic ocean basins must haveseparated distinct cratons or blocks consisting ofArchean crust prior to ca. 1.84 Ga.

5.2. Paleoproterozoic e6olution

The oldest Paleoproterozoic rocks in the south-ern Crossroads domain (Fig. 14) are ca. 1835–1810 Ma K-feldspar megacrystic granite, granite,granodiorite and charnockite intrusions of the DePas batholith. These granitoid rocks are variablydeformed and recrystallized demonstrating theypredate and overlap with a tectonothermal event,which persisted to at least 1775 Ma on the basisof U–Pb titanite data from other rocks the do-main. Paleoproterozoic metamorphism in theCrossroads domain reached amphibolite facies,but it did not result in the production of meta-morphic leucosome. Major- and trace-elementgeochemistry of De Pas batholith rocks definecalc-alkaline trends (Van der Leeden et al., 1990;

Dunphy and Skulski, 1996), and are compatiblewith an interpretation for the batholith as a conti-nental magmatic arc formed above an east-dip-ping subduction zone (Martelain, 1989; Van derLeeden et al., 1990; Dunphy and Skulski, 1996).In apparent contradiction to this interpretation,Kerr et al. (1994) have noted that De Pasbatholith rocks do not have all of the geochemicalsignatures characteristic of modern subduction-re-lated continental magmatic arcs; they note, as oneexample, that most samples show a strong enrich-ment in Zr. However, some of the geochemicalsignatures, which are not characteristic of modernarcs, may in part be related to the nature of thehost Archean crust and to how much of the hostwas assimilated during emplacement of the DePas batholith. The batholith rocks are character-ised by negative o Nd values between −3 and−7, calculated at 1830 Ma, and T dm agesbetween 2.24 and 2.64 Ga indicating that De Pasmagma was significantly contaminated by the sur-rounding Archean crust (Kerr et al., 1994; Dun-phy and Skulski, 1996). The geochemicalsignatures of De Pas batholith rocks are inter-preted to reflect the mixing of juvenile, subductionrelated magma and Archean crust of the CoreZone.

There is a need for additional field and isotopicdata from the Orma domain, although existingU–Pb geochronological data collected by Nunnet al. (1990) suggest that Archean rocks in thedomain have not been overprinted by a Pale-oproterozoic tectonothermal event. The analysedminerals do not show any significant post-Archean Pb loss, nor do the samples contain new(Paleoproterozoic) zircon or titanite. However wesuggest that the domain probably contains Pale-oproterozoic intrusions petrogenetically relatedto, and part of, the De Pas batholith. In particu-lar, we provisionally correlate a unit of foliated,pyroxene-bearing K-feldspar megacrystic granitemapped by Nunn, (1994) (Nunn’s Unit 5, page433) with the De Pas batholith. That these rocksare foliated and metamorphosed; pyroxene isrimmed by hornblende and the rocks containgarnet, suggests that at least locally the Ormadomain contains Paleoproterozoic structures andamphibolite–facies minerals.


Fig. 14. Summary diagram of U–Pb geochronological data (Paleoproterozoic ages) for the study area. Evidence for coevalmetamorphism and deformation in McKenzie River and Crossroads domains after ca. 1810 Ma, and different magmatic historiesprior to that, suggests the domains were juxtaposed by that time. Included are U–Pb ages for De Pas batholith rocks from 1 (Jameset al., 1996) 2 (Krogh, 1986) and 3 (Dunphy and Skulski, 1996).

The proposed correlation in the preceding para-graph is significant because it implies that move-ment along the George River shear zone, whichpostdated emplacement of the De Pas batholith isprobably not significant in the southern part ofthe Core Zone. This interpretation is consistentwith field and geochronological data indicatingthat deformation in the George River shear zone,170 km north of our study area, was occurringbut waning at 1825 Ma (Dunphy and Skulski,1996). Furthermore, the fact that the Crossroadsand Orma domains have very similar Archean

geology suggests that in the study area, theGeorge River shear zone does not have a signifi-cant total-finite transcurrent displacement, andthat the two domains have acted as an essentiallyintact Archean block throughout their history.Paleoproterozoic (ca. 1825–1775 Ma) amphibo-lite–facies metamorphism and deformation ismore pervasive in areas contiguous with the DePas batholith (i.e. in the Crossroads domain),consistent with models for metamorphism in amagmatic arc setting. This may explain why theOrma domain, which is mainly lacking in De Pas


intrusive rocks, effectively escaped Paleoprotero-zoic tectonothermal effects. Absence of a late syn-to post-De Pas batholith metamorphism in theOrma domain is also consistent with tectonicmodels for the batholith which predict that it isexposed as a tilted, oblique section having deeper,higher-grade rocks in the west and lower-graderocks in the east (Dunphy and Skulski, 1996).

The McKenzie River domain does not containDe Pas batholith granitic rocks, suggesting theMcKenzie River and Crossroads domains werenot in their present configuration until after ca.1810 Ma, the youngest emplacement age for DePas granite in the region. An occurrence of ca.1815 Ma tonalite indicates igneous activity in theMcKenzie River domain at this time, althoughthe significance of this tonalite is uncertain. Leu-cosome from amphibolite–facies Lobstick groupmetasedimentary rocks, dated at ca. 1805 Ma,indicates that the peak of metamorphism in theMcKenzie River domain falls within the range ofmetamorphic monazite and titanite ages from theCrossroads domain. These data and the fact thatdeformation in the Lac Tudor shear zone wasattendant with amphibolite–facies metamorphismsuggests that the McKenzie River and Crossroadsdomains were linked by ca. 1805 Ma. A recrystal-lized tonalite dyke, which cross-cuts the metamor-phic leucosome and strong foliation and is datedat 1802 +9/−14 Ma, indicates that metamor-phism outlasted deformation in the McKenzieRiver domain. Ages of metamorphic monaziteand titanite from the Crossroads and McKenzieRiver domains (this study) are consistent with theages of metamorphic monazite (179395 and178392 Ma) and titanite (177495 and 1783911 Ma) in orthogneisses from the Kuujjuaq do-main (Machado et al., 1989), 300 km north of thestudy area, and support widespread metamor-phism at this time.

6. Tectonic model

The Core Zone is a composite terrane assem-bled in the vise between northward moving andobliquely converging North Atlantic and Superiorcratons. Convergence of the cratons resulted in

sequential closure of intervening Paleoproterozoicbasins, and their northward movement terminatedin their ultimate collision with Archean terranesalong their northern margins (see Hoffman, 1990).Core Zone collision and deformation occurredfirst along it’s eastern margin; accretion of Tasi-uyak sediments to the North Atlantic craton andthe onset of deformation in the Torngat Orogenand eastern Core Zone occurred at 1.86–1.85 Ga(Van Kranendonk et al., 1993; Wardle and VanKranendonk, 1996), and preceded tectonothermaland magmatic events in the western Core Zoneand in the New Quebec orogen. The tectonicsetting of formation and accretion of the 1.84–1.82 Ga Lac Lomier Complex (subduction-relatedarc?), occurring along the boundary between theCore Zone and the North Atlantic craton areuncertain (see discussion in Wardle, 1998).

It is inferred that collision between the easternmargin of the Core Zone and the North Atlanticcraton resulted in basin closure, initiation of east-dipping subduction under the western Core Zoneby 1840 Ma, and formation of the earliest intru-sions of the De Pas magmatic arc between 1840Ma and 1830 Ma (Dunphy and Skulski, 1996;James et al., 1996) (Fig. 15). Initiation of De Pasmagmatism pre-dates the peak of metamorphismand the onset of deformation in the southwesternCore Zone; these are estimated to be B1820 Ma.Absence of De Pas intrusions in the McKenzieRiver domain suggests that it was west of thetrench and on the subducting plate. Moreover, thesignificant differences in Archean geology betweenthe McKenzie River domain and Core Zone do-mains to the east, and the fact that Paleoprotero-zoic Lobstick group sediments are restricted tothe McKenzie River domain, suggests that do-mains on either side of the Lac Tudor shear zoneare exotic with respect to each other. This inter-pretation implies that ca. 1845–1833 Ma arc-re-lated (?) intrusive rocks in the Kuujjuaq domain(Poirier et al., 1990; Perreault and Hynes, 1990),occurring west of the Lac Tudor shear zone in thenorthwestern Core Zone (Fig. 2), do not correlatewith the De Pas batholith. We propose thatArchean rocks occurring west of the Lac Tudorshear zone, including those in the McKenzie Riverdomain, Laporte terrane and Kuujjuaq domain


have an affinity with the Superior craton; theywere incompletely rifted from the craton after ca.2.2 Ga.

McKenzie River domain was not terminallyjuxtaposed to the Crossroads domain until ca.1810 Ma; a consequence of this collision was theshutting down of the subduction zone and cessa-tion of De Pas magmatism. U–Pb geochronologi-

cal data (this study) demonstrates ca. 1815 Maemplacement of tonalite in the McKenzie Riverdomain overlaps in time with De Pas magmatism.Tectonic setting for this magmatism, which isslightly younger than emplacement ages from theKuujjuaq batholith, is uncertain. Amphibolite–fa-cies metamorphism and production of ca. 1805Ma metamorphic leucosome in the domain may

Fig. 15.


be related to crustal thickening as a result ofoverthrusting by the Crossroads domain along theLac Tudor shear zone. The youngest structures inthe shear zone are related to dextral transcurrentdisplacement. A switch from mainly convergent(thrust) to strike-slip displacement may have oc-curred when shortening in the southwestern CoreZone could no longer be accommodated bycrustal thickening.

The onset of deformation in the New QuebecOrogen is inferred to be related to collision of thewestern Core Zone with the Superior craton mar-gin. Our model predicts that this did not occuruntil ca. 1810 Ma. There are insufficient datafrom the New Quebec Orogen to constrain thetiming of initial thrusting and to test this model,although Machado et al. (1997) claim that thrust-ing occurred locally before 1813 Ma. Evidencefrom the Cape Smith Belt along the northernSuperior margin suggests that final collision oc-curred there at ca. 1800 Ma (St-Onge et al., 1998).The lack of detailed geochronological data toconstrain the timing of thrusting in the NewQuebec Orogen and the timing and kinematic

history of shear zones separating it from the CoreZone remain outstanding problems in understand-ing foreland–hinterland relationships in theSECP.

Acknowledgements

The U–Pb geochronological work summarizedhere was funded by LITHOPROBE (ECSOOTTransect), Canada. Geological field studies werefunded by a Canada–Newfoundland Mineral De-velopment Agreement contract carried by the Ge-ological Survey of Newfoundland and Labrador.Our thanks to Robbie Hicks for assistance in thegeochronology laboratory at Memorial Universityof Newfoundland in St. John’s. Richard Wardle isthanked for his comments on an earlier version ofthis manuscript. Our paper benefitted from re-views by N. Machado and M. St-Onge. Thispaper is published with the permission of theDirector of the Geological Survey of Newfound-land and Labrador.

Fig. 15. Series of cartoons showing sequence of assembly of the southern Core Zone with the Superior and North Atlantic cratons.(1) \1.91 Ga: Initial configuration. McKenzie River and Kuujjuaq domains represent Superior craton rocks that were incompletelyrifted from the craton at ca. 2.17 Ga. Archean rocks in the central and eastern parts of the Core Zone may have been part of theSuperior craton prior to 2.2 Ga, but were independent crustal blocks by 1.91 Ga and separated from the Superior and NorthAtlantic cratons by Paleoproterozoic ocean basins. (2) 1.91–1.88 Ga. Convergence of the eastern Core Zone with the North Atlanticcraton and formation of an east-dipping subduction zone. A continental magmatic arc (Burwell domain; not shown) is formedlocally on the North Atlantic craton margin. There are no relicts of arc rocks along the southwestern margin of the North Atlanticcraton; the arc(s) may have been structurally excised or may never have existed there. An accretionary prism (Tasiuyak sediments)receives input from the arc and from emerging Paleoproterozoic terranes to the north (e.g. Cumberland batholith). (3) 1.88–1.85 Ga.Rifting along the Superior margin and deposition of Cycle two rocks, mafic volcanic and intrusive complexes in a narrow basin.Initial collision of the eastern Core Zone with the North Atlantic craton, and accretion of arc(s) to the North Atlantic cratonmargin. (4) 1.84–1.81 Ga. Convergence between the eastern and central Core Zone and the Superior margin. Subduction-relatedcontinental arc magmatism occurs in the McKenzie River–Kuujjuaq domains, and in Crossroads domain (De Pas batholith).Crustal thickening, high-grade metamorphism and deformation in regions contiguous with arc magmatism. The eastern Core Zonelacks intrusions and, at least locally, escapes Paleoproterozoic metamorphism. Continued convergence in the Torngat Orogen andsinistral shearing in the Abloviak shear zone (ASZ). (5) 1.81–1.78 Ga. Collision between the Core Zone (including McKenzie Riverand Kuujjuaq domains) and the Superior craton. Collision results in formation of a fold and thrust belt (New Quebec Orogen,NQO). Dextral transpressive kinematics on the Ashuanipi River (ARSZ) and Lac Tudor (LTSZ) shear zones. The George Rivershear zone is a dextral transcurrent structure, although it may have a component of late extension (e.g. Bardoux et al., 1998).Continued convergence, thrusting and sinistral shearing in the Torngat Orogen. The George River shear zone may have a pre- andpost-1.8 Ga history, and may have influenced localization of subsequent Mesoproterozoic events. We note, for example, that allMesoproterozoic intrusions and Seal Lake Group rocks occur to the east of the shear zone.


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