Precambrian Research 103 (2000) 31 – 54
U – Pb geochronological constraints for Paleoproterozoic
evolution of the Core Zone, southeastern Churchill
Province, northeastern Laurentia
Donald T. James
a,*, Greg R. Dunning
baGeological Sur
6ey of Newfoundland and Labrador,P.O. Box8700,St. John’s Newfoundland, Canada A1B 4J6 bDepartment of Earth Sciences,Memorial Uni
6ersity 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 Paleoproterozoic lithotectonic domains assembled in the vise between obliquely colliding Archean Superior and North Atlantic cratons between ca. 1860 and 1810 Ma. Detailed geological and U – Pb geochronological studies of domains in the southern Core Zone highlight significant differences in Archean and Paleoproterozoic geology between domains, and constrain timing and models for Paleoproterozoic assembly. In the south-central part of the Core Zone, Crossroads and Orma domains consist of Late Archean granite – greenstone terrane crust. Crossroads domain also contains a significant number of granite – charnockite intrusions belonging to the 1840 – 1810 Ma De Pas batholith; magmatism pre-dates and 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 (De Pas related?) intrusions and appears to have mainly escaped Paleoproterozoic metamorphism and penetrative deformation. A model compatible with available data suggests the De Pas batholith is a continental magmatic arc formed above an east-dipping subduction zone. Paleoproterozoic metamorphism and deformation is concentrated in regions contiguous with arc magmatism. West of Crossroads domain, and separated from it by the Lac Tudor shear zone, McKenzie River domain is dominated by Archean orthogneiss (\80 m.y. older than gneisses in Crossroads and Orma domains), minor amounts of Paleoproterozoic supracrustal rocks (absent in domains to the east) and 1815 Ma tonalite. It does not contain De Pas intrusions. U – Pb geochronology suggests juxtaposition of the McKenzie River domain with domains to the east may have occurred around 1810 Ma. One possible model envisages the McKenzie River domain as a piece of reassembled Superior craton crust, incompletely rifted from the Superior margin during ca 2.17 Ga rifting. Domains occurring east of the Lac Tudor shear zone are interpreted to be exotic with respect to the Superior craton. © 2000 Elsevier Science B.V. All rights reserved.
Keywords:Northeastern Laurentia; Paleoproterozoic tectonics; U – Pb geochronology
www.elsevier.com/locate/precamres
* Corresponding author. Tel.: +1-709-7292774; fax: +1-709-7294270.
E-mail address:dtj@zeppo.geosurv.gov.nf.ca (D.T. James).
1. Objectives
There are few geochronological constraints on the intrusive and tectonothermal history of the southwestern Core Zone (James and Dunning, 1996; James et al., 1996) of the Southeastern Churchill Province (SECP), northeastern Lauren-tia (Fig. 1). A better understanding of the Core Zone (Fig. 2), a fundamental Paleoproterozoic tectonic division consisting of a tectonic collage of Archean and Paleoproterozoic rocks, is criti-cal because it contains not only a record of its internal construction but also of its assembly with bordering Archean cratons and their at-tached Paleoproterozoic supracrustal sequences. In this paper, we address some of the outstand-ing problems relatoutstand-ing to Paleoproterozoic con-struction of three domains, McKenzie River,
Crossroads and Orma domains, that make up the southwestern Core Zone. In particular, we investigate the timing and significance of granitic and mafic intrusions, their relationship to meta-morphic events and the development of major structures. 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 analysis of critical exposures containing unequivocal field relationships (James and Dunning, 1996). The results, when combined with available data from other areas in the southern SECP, which we summarize in section three of this paper, help to constrain a model for Paleoproterozoic evolution of the Core Zone. This model has broader im-plications for the overall development of north-eastern Laurentia.
D.T.James,G.R.Dunning/Precambrian Research103 (2000) 31 – 54 33
Fig. 2. Tectonic elements of Labrador and northeastern Que´bec. Mrd, McKenzie River domain; Crd, Crossroads domain; Od, Orma domain; M-Rd, Mistinibi-Raude domain; Kd, Kuujjuaq domain. Mesoproterozoic intrusions are indicated by the open triangle pattern.
2. Introduction
The SECP is a 300 km wide, north-trending composite tectonic belt of Archean and Pale-oproterozoic rocks that is one segment of a sys-tem of Paleoproterozoic orogens linking Archean cratons in northeastern Laurentia. It is principally a continuation of the Trans-Hudson Orogen, which can be traced around the western, northern and eastern margins of the Superior craton, but it also shares common elements with the Nagssug-toqidian Orogen of Greenland (Fig. 1). The SECP is exposed from Ungava Bay, where it disappears under Hudson Strait, to southern Labrador, where it is truncated by east – northeast-trending structures and tectonostratigraphic units that make up the Grenville Province. The SECP reap-pears north of Hudson Strait, on southern Baffin Island, although the precise correlation of major structures and tectonostratigraphic units between the two regions is speculative and the focus of
ongoing studies (see St-Onge et al., 1997; Scott and St-Onge, 1998; St-Onge et al., 1998, 1999).
de-veloped primarily in juvenile (B1.95 Ga) Pale-oproterozoic sediments and inferred to represent an accretionary complex along the suture between the Core Zone and the North Atlantic craton (Nain Province) (Fig. 2). Dextral (west) and sinis-tral (east) transcurrent shear zones, which are synchronous-to post-tectonic with respect to thrusting in the New Que´bec and Torngat oro-gens, respectively, separate the bordering ‘fore-land’ orogens from the Core Zone. The Core Zone is itself a mosaic of variably reworked Archean 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 Ga and B1.95 Ga supracrustal rocks (e.g. Van der Leeden et al., 1990; Girard, 1990; Scott and Gau-thier, 1996), and 1.84 – 1.81 Ga granitoid rocks belonging to the De Pas and Kuujjuaq batholiths (Perreault and Hynes, 1990; Dunphy and Skulski, 1996; James et al., 1996). Subsequent to Core Zone amalgamation, the Core Zone and border-ing orogens were overprinted by transcurrent shearing, which persisted locally to 1.74 Ga (Wardle and Van Kranendonk, 1996).
Affinity of Archean crust in the Core Zone is an outstanding problem having significant impli-cations for developing Paleoproterozoic tectonic models for the region. Addressing this problem in a comprehensive way is outside of the scope of this paper, although one possible model proposes that the majority of Archean Core Zone crust is exotic with respect to the Superior and North Atlantic cratons. In contrast, other models sug-gest that Archean rocks in the Core Zone were part of the Superior craton prior to 2.2 Ga (James et al., 1998; Scott and St-Onge, 1998). Available geochronological data from the Core Zone indi-cates that Archean Core Zone rocks have broadly similar intrusive ages as rocks in the northeastern Superior craton, although this provides only cir-cumstantial evidence of their parentage. The Archean geochronological data from Core Zone rocks is non-unique and could be used to support either of the models. However, if a significant component of the Archean crustal blocks in the Core Zone did belong to the Superior Craton prior to 2.2 Ga, there is general consensus that these blocks acted as independent crustal units,
relative to the bounding Archean cratons, during Paleoproterozoic construction of the SECP (Wardle, 1998). There are no compelling geologi-cal or geophysigeologi-cal data to suggest the Core Zone includes a significant amount of North Atlantic craton crust, although minor amounts or variably reworked North Atlantic craton rocks may occur in regions adjacent to the boundary between the Core 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 with the 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, contained in Fig. 3, mainly consists of \2.7 Ga metapelitic gneiss, orthogneiss, metamorphosed mafic intru-sions, diatexite plutons of orthopyroxene-bearing monzogranite to granodiorite, and granite intru-sions. The sedimentary precursors of the metapelitic gneiss were intruded by plutons and related dykes of tonalite, gabbro and granite at ca. 2.7 Ga prior to a regional tectonothermal event at 2.68 – 2.65 Ga (Mortensen and Percival, 1987). Deformation was accompanied by gran-ulite – facies metamorphism and emplacement of the diatexite plutons, and followed by intrusion of late syn- to posttectonic plutons of biotite9 mus-covite granite. Of relevance to our discussion of the evolution of the Core Zone which follows, it should be noted that rocks belonging to the Ashuanipi Complex do not occur in the Core Zone.
3.2. New Que´bec Orogen (NQO)
D.T.James,G.R.Dunning/Precambrian Research103 (2000) 31 – 54 35
Paleoproterozoic supracrustal rocks, are con-tained in low-grade, thrust bound slices that make up a west-verging, Paleoproterozoic fold-and-thrust belt defined as the NQO. An understanding of the stratigraphic and structural relationships in the NQO is critical to understanding the Core Zone because stratigraphic (e.g. rifting) and struc-tural (e.g. initiation of thrusting) events in the foreland will be reflected or reflect events in the hinterland (i.e. Core Zone).
The Schefferville Zone, comprising the struc-turally lowest tectonostratigraphic unit in the southern NQO, consists of two sedimentary and volcanic cycles of the Knob Lake Group. In general, the two cycles record a transition from continental sedimentation and local alkaline vol-canism to progressively deeper water
sedimenta-tion and tholeiitic basaltic volcanism (see summaries in Skulski et al., 1993; Wardle and Van Kranendonk, 1996). The older sequence (Cy-cle one) of arkose, shale and dolomite records initiation of rifting and development of a passive margin sequence on extended Superior craton crust. Cycle one rocks are overlain discon-formably by a westwards-overstepping younger sequence (Cycle two) of quartzite, iron formation, sandstone – shale turbidites and arkose.
The Schefferville Zone is structurally overlain by the Howse Zone, which is dominated by thick sequences of siltstone – shale turbidites interbed-ded with tholeiitic basalt and gabbro – sill com-plexes. However, it also contains rocks equivalent to cycles one and two described in the preceding paragraph. A possible interpretation for the
bidites, basalts and related sills, is that they may have been deposited in a narrow, dextral transten-sional basin, perhaps analogous to the Gulf of California (Wardle et al., 1990; Skulski et al., 1993). Cycle one and two volcanic rocks and sills in the Howse Zone are dated at ca. 2.17 – 2.14 Ga and ca. 1.89 – 1.87 Ga (Birkett et al., 1991; Skulski et al., 1993; Rohon et al., 1993), respectively. These ages are interpreted to correspond with time of deposition for cycles one and two. If the narrow, transtensional basin model for Cycle two rocks is apropos, it would imply that rifted Supe-rior craton crust occurred to the east of Cycle two rocks at 1.87 Ga.
To the east, the Howse Zone is structurally overlain by the Doublet terrane composed of Doublet Group mafic pyroclastic rocks, turbiditic siltstones, mafic and ultramafic sill complexes. The rocks in the Doublet terrane lack unequivocal stratigraphic links with Cycle one or Cycle two rocks in the Schefferville and Howse zones, and on this basis they are inferred to be exotic with respect to the Superior craton (Wardle et al., 1995). Doublet terrane rocks may represent tran-sitional crust formed at the boundary between extended 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. It mainly consists of presumed Archean gneisses and metamorphosed Paleoproterozoic pelitic, arkosic and mafic volcanic rocks, and lesser amounts of quartzite, marble and gabbro. There are no obvi-ous correlations between supracrustal rocks in the Laporte terrane and Paleoproterozoic rocks oc-curring to the west, although the Archean rocks may represent reworked Superior craton crust. The supracrustal rocks may be the relics of an accretionary prism (Wardle, 1998).
3.3. The Core Zone
The Core Zone (James and Dunning, 1996; James et al., 1996) is informally defined as a composite Paleoproterozoic terrane consisting of lithotectonic 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 unequivocally linked with the Superior or North Atlantic cra-tons, or with Paleoproterozoic supracrustal se-quences which occur in the New Que´bec or Torngat orogens. The western boundary of the Core Zone is marked by the Ashuanipi River shear zone (James et al., 1996) in the south and the Lac Turcotte fault (see Perreault and Hynes, 1990) in the north. The western margin of the Abloviak shear zone defines the eastern boundary. This definition includes the Kuujjuaq domain and the Lac Lomier complex as part of the Core Zone; this is perhaps noteworthy as these are not included in Wardle’s definition (Wardle, 1998). In particular, tectonic affinity of the Kuujjuaq do-main (Perreault and Hynes, 1990) is somewhat controversial; some workers (e.g. Bardoux et al., 1998) consider the domain to be part of the NQO. The boundary between the NQO and the Core Zone is approximated by the inflection between a paired, negative (NQO side) and positive (Core Zone side) Bouguer gravity anomaly. On the basis of the gravity signature, which is consistent with a model involving west-directed transport of Core Zone rocks over the NQO, Thomas and Kearey (1980) proposed that this boundary is a Pale-oproterozoic suture marking a relic Andean-type margin. The offshore extension of the Core Zone underlying Ungava Bay has been traversed by the LITHOPROBE (ECSOOT) seismic reflection sur-vey, which showed the Core Zone to be domi-nated by 30° east-dipping reflectors that appear to penetrate the whole crust (Hall et al., 1995).
The southwestern Core Zone includes, from west to east, McKenzie River, Crossroads, Orma and Mistinibi-Raude domains (see Van der Lee-den et al., 1990; Nunn et al., 1990; Girard, 1990; Nunn, 1994; James et al., 1996). The McKenzie River domain (discussed in detail below, and see James et al., 1996) consists mainly of Archean tonalite gneiss and lesser amounts of inferred Paleoproterozoic supracrustal rocks, which are metamorphosed to upper amphibolite facies. The McKenzie River domain does not appear to share many Archean or Paleoproterozoic (\1815 Ma) features with domains to the east.
D.T.James,G.R.Dunning/Precambrian Research103 (2000) 31 – 54 37
high-grade Archean granite – greenstone terrane crust and Paleoproterozoic granitoid intrusions, which are part of the \500 km long De Pas batholith. The intrusions are variably deformed and recrystallized, demonstrating the domain has been overprinted by a Paleoproterozoic tec-tonothermal event which partially overlapped and postdated emplacement of the De Pas batholith. The western boundary with the McKenzie River domain is the Lac Tudor shear zone (Van der Leeden et al., 1990), which deforms Archean and Paleoproterozoic rocks in both domains, and has components of dextral transcurrent and reverse displacement (Van der Leeden et al., 1990; Bourque, 1991; James et al., 1996). The George River shear zone (see Van der Leeden et al., 1990; Girard, 1990), which forms the eastern boundary with the Orma domain, is a wide (several km), but very diffuse and poorly defined zone of heteroge-neous strain containing porphyroclastic protomy-lonitic and mylonitic rocks having dextral transcurrent kinematic indicators (James et al., 1996). Mylonitization in the Lac Tudor and George River shear zones was attendant with Paleoproterozoic amphibolite – facies metamorph-ism.
Archean geology of the Orma domain is similar to that of the Crossroads domain. The Orma domain contains relicts of Archean greenstone belt rocks, which are intruded by tonalite or-thogneisses having igneous crystallization ages be-tween 2682 and 2675 Ma, as determined by U – Pb age dating of zircons (Nunn et al., 1990). Titanite data from the same rocks suggest they were meta-morphosed to amphibolite – facies in the Late Archean. Notably, the U – Pb data show no evi-dence that the Archean rocks were overprinted by Paleoproterozoic thermal event(s) prior to the ca. 1720 – 1600 Ma Labradorian Orogeny; all Pb-loss in the titanites is younger than ca. 1640 Ma (Nunn et al., 1990). Based on this data, the Orma domain has been considered to have mainly es-caped the ca. 1820 – 1775 Ma Paleoproterozoic tectonothermal event (discussed in section four) which overprinted the Crossroads and McKenzie River domains. However, at possible variance with the U – Pb data, the domain also includes foliated granitoid rocks, which are undated, but
suspected to be Paleoproterozoic (Nunn, 1994) on the basis of lithologic correlations with dated intrusions in the Crossroads domain. The signifi-cance of this correlation will be discussed later.
The Orma domain also includes a sequence of wacke, quartz wacke, quartzite, tuffaceous rocks and metamorphosed basalt, named the Petscapiskau Group (Emslie, 1970). Petscapiskau Group rocks were not metamorphosed prior to being intruded by the Mesoproterozoic (ca. 1460 Ma) Michikamau Intrusion (Emslie, 1970). Age of the Petscapiskau Group is unknown, but there are two 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 River domains (i.e. \1820 Ma), it would support exist-ing U – Pb data from Archean rocks and confirm that the Orma domain escaped penetrative Pale-oproterozoic tectonothermal overprinting. If the Petscapiskau Group is younger than the Pale-oproterozoic event, it would leave open the possi-bility that the Orma domain has been overprinted by hitherto unrecognized Paleoproterozoic meta-morphism. It is remotely possible that the Petscapiskau Group is coeval with ca. 1650 Ma Labradorian supracrustal sequences (e.g. the Blueberry Lake group (James and Connelly, 1996), or MacKenzie Lake Group (Nunn, 1993)), which unconformably overly pre-1650 Ma rocks along the southeastern margin of pre-Labradorian Laurentia. The Labradorian volcanic rocks pre-dominately have felsic and intermediate composi-tions whereas Petscapiskau Group volcanic rocks are mafic. This may suggest that a Labradorian age for the Petscapiskau Group is unlikely, but it does not discount the possibility. There is an obvious need for more geochronological data from the Orma domain.
metabasic rocks that were intruded by pre- to syn-tectonic granitic sheets; these rock units and the high-grade metamorphism that overprints them are provisionally assigned to the Archean. The Archean rocks are intruded by the enigmatic ca. 2.3 Ga Pallatin Intrusive Suite (see Van der Leeden et al., 1990; Krogh, 1992) and overlain by presumed related sedimentary and volcanic rocks of the Ntshuku Complex (Girard, 1990). In the northern part of the domain, Ntshuku rocks are unconformably overlain by deformed sedimentary rocks of the Paleoproterozoic Hutte Sauvage Group (Van der Leeden et al., 1990). The Ntshuku and Hutte Sauvage rocks are metamor-phosed from greenschist to lower-amphibolite facies.
Mesoproterozoic (ca. 1460 Ma) intrusions of anorthosite and related granitoid rocks intrude the Orma and Mistinibi-Raude domains, and in the southeastern Core Zone, these obscure the Paleoproterozoic Abloviak shear zone (Wardle et al., 1990) and the boundary with the North At-lantic craton. The Abloviak shear zone marks the approximate eastern limit of penetrative Pale-oproterozoic deformation and defines the tectonic boundary with the North Atlantic craton. How-ever, in the area between the Mesoproterozoic Harp Lake Intrusive Suite and the Nain Plutonic Suite, the eastern limit of Paleoproterozoic defor-mation is marked by deformed Ingrid Group rocks, a sequence of conglomerates and volcanic rocks. Age of Ingrid Group volcanism is loosely constrained; U – Pb zircon and titanite ages deter-mined from two samples of felsic volcanic rocks give ages of ca. 1895 Ma and 1805 Ma (Wasteneys et al., 1996). Clasts in Ingrid Group conglomerate are derived from a variety of sources including
\3.7 Ga North Atlantic craton (presumed Saglek Block) crust and from ca. 1900 Ma granitoid rocks that may have been part of a Torngat magmatic arc (Wasteneys et al., 1996) constructed on the North Atlantic craton. The Early Archean clasts demonstrate that the sedimentary rocks were autochthonous with respect to North At-lantic craton crust.
The Ingrid Group is overthrust to the west by granulite – facies granitoid migmatite determined by U – Pb zircon geochronology to have an
em-placement age of 2870922 Ma (Wasteneys et al., 1996). The affinity of these Archean gneisses is uncertain; 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 Archean gneisses are tectonically interleaved with Tasiuyak gneiss, a distinctive metasedimentary gneiss, which can be traced for at least 500 km along the entire eastern boundary with the North Atlantic craton, although it’s total strike length may ex-ceed 1300 km (Scott, 1998). The high-strain zone containing deformed Tasiuyak gneiss and Archean gneisses may be related to and coeval with the sinistral transcurrent Abloviak shear zone. Detrital zircon populations and field rela-tions, which bracket the depositional age of Tasi-uyak sediments between 1940 and 1895 Ma (Scott and Machado, 1993; Scott and Gauthier, 1996), and Nd-isotopic data (Kerr et al., 1993; The´riault and Ermanovics, 1993) suggest that sedimentary precursors of the Tasiuyak gneiss were exotic with respect to the North Atlantic craton (Wardle and Van Kranendonk, 1996). Thus, Archean rocks that occur to the west of the Tasiuyak gneiss might never have been part of the North Atlantic craton. In conflict with this interpretation, Ryan (1990) has suggested that Archean gneisses and anorthositic rocks occurring more than 50 km west of Tasiuyak gneiss and Paleoproterozoic shear zones, in the area north of the Mistastin batholith (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 craton which contains tectonically interleaved North At-lantic craton rocks and Archean rocks which are exotic with respect to the North Atlantic craton (e.g. Van Kranendonk et al., 1993).
D.T.James,G.R.Dunning/Precambrian Research103 (2000) 31 – 54 39
Fig. 4. Geology of the southwestern core zone in the area around the Smallwood Reservoir, Labrador (southern part of NTS map area 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 Smallwood Reservoir. ARSZ, Ashuanipi River shear zone; LTSZ, Lac Tudor shear zone; GRSZ, George River shear zone.
4. U – Pb geochronological studies in the southwestern Core Zone
4.1. McKenzie Ri6er domain
The McKenzie River domain (Fig. 4) is sepa-rated from the NQO and the Crossroads domain by the Paleoproterozoic Ashuanipi River and Lac Tudor shear zones, respectively. Both shear zones are inferred to be transpressive (dextral) mylonite zones having components of east-over-west
re-verse displacement. Field relations demonstrate they were developed concomitant with amphibo-lite – facies metamorphism. The Ashuanipi River shear zone corresponds approximately to the infl-ection between a regionally persistent, paired Bouguer gravity anomaly; negative on the NQO side and positive on the Core Zone side (see James et al., 1996, Fig. 2, page 218).
(James et al., 1996). o Nd values from the Flat
Point gneiss from −0.5 to−5, at 2800 Ma (Kerr et al., 1994), suggest the rocks incorporated a component of older crust. The rocks are meta-morphosed to upper-amphibolite facies and have a gneissosity which is folded into relatively open and flat-lying superposed, mainly dome-and-basin, folds. The age of the gneissosity and the folding are unknown, but are assumed to be Archean. The gneissosity and superposed folds are overprinted by a steep, north-striking and east-dipping foliation accompanied by a pervasive recrystallization that locally obliterates the gneis-sosity and preexisting structure. The north-strik-ing foliation is interpreted to be a Paleoproterozoic fabric for three reasons: (1) it has the same attitude as the foliation in ca. 1815 Ma tonalite (see discussion of sample MR1 be-low), (2) it is defined by the peak-assemblage metamorphic minerals and amphibolite – facies metamorphism was attained by ca. 1805 Ma (see discussion of sample MR2 below), and (3) it becomes progressively more intense in areas near the major Paleoproterozoic high-strain zones (e.g. Lac Tudor shear zone), which define the domain boundaries.
Metamorphosed supracrustal rocks of un-known depositional age occur as thin (B1 km) tectonically bound units within the Flat Point gneiss. Informally named the Lobstick group, they include metasedimentary (pelitic) migmatite and lesser amounts of quartzite, marble, calcsili-cate derived from impure siliceous carbonate, and amphibolite of uncertain protolith. These supracrustal rocks may be correlative with litho-logically similar, Laporte group rocks, which oc-cur to the north of the study area. Geochronology samples were collected from an outcrop contain-ing migmatitic Lobstick group rocks, a tonalite dyke which cross cuts the metamorphic leuco-some, and meta-tonalite which is in tectonic con-tact with the Lobstick group rocks. The supracrustal rocks are steeply foliated and have steeply plunging folds of relict sedimentary bed-ding. They do not contain the superposed fold structures contained in the Flat Point gneiss. On the basis of orientation, the foliation in the supracrustal gneisses is correlated with the folia-tion in Paleoproterozoic tonalite (see discussion of sample MR1 below).
Sample MR1 is from a unit of strongly foliated and recrystallized tonalite that is in tectonic con-tact with Lobstick group supracrustal gneiss. MR1 tonalite differs significantly from the Flat Point gneiss in that it lacks metamorphic layering and the superposed fold structures that character-ize the latter. The sample yielded three fractions of concordant and nearly concordant zircons (Fig. 5, Table 1), and the two overlapping concordant analyses give an age of 181593 Ma, interpreted to represent the igneous crystallization age. These data represent the first indication of Paleoprotero-zoic granitoid intrusive rocks in the McKenzie River domain, although the extent of ca. 1815 Ma tonalite is unknown.
To determine the age of metamorphism, a sam-ple (MR2) of K-feldspar+biotite+ garnet-bear-ing leucosome contained in pelitic migmatite was collected. Two fractions of monazite (Fig. 5) are concordant (M2) and nearly concordant (M1) and indicate that monazite crystallized in the leuco-some at 180593 Ma, and demonstrate that up-per-amphibolite facies metamorphism was attained by this time. The data only loosely
D
Fraction description Weight U Pb(r) 207 Pb/206 Pb (mg) (ppm) (pg)
MR1meta-tonalite(366167E,60102040N)
37
Z1 Large clear 0.121 50 19.2 3339 0.2451 0.32534 (124) 4.9747 (202) 0.1109 (14) 1826 1815 1814 prisms
6490 0.2839 0.32295(122) 4.9282(182) 0.11067 (20) 1804
20.8 1807
54 10 1811
Z2 Clear, euhedral 0.06
Z3 Large, euhedral 0.083 86 32.8 12 12546 0.2439 0.32584 (142) 4.9908 (208) 0.11109 (24) 1818 1818 1803
MR2leucosome,msed migmatite(3861617E,601204N)
9633 39.246 0.32337 (190) 4.9287 (296) 0.11054 (12) 1806 1807 1808 M1 Large, clear 0.091 293 3317.3 56
10619 45.062 0.32294 (132) 4.9069 (206) 0.1102 (14) 1804 1803 1803 48
M2 Clear, yellow 0.1 252 3253.3
MR3tonalite dyke(386167E,6010204N)
8918 0.1183 0.32667 (124) 5.0751 (200) 0.11268 (14)
Z1 Small, clear 0.021 512 177.1 24 1822 1832 1843 3890 0.1461 0.32617 (100) 5.0618 (168) 0.11255 (14) 1820
38 1830
Z2 Clear prisms 0.012 610 215.6 1841
27
Clear, cracks 0.023 429 152.5 7645 0.1368 0.33024 (110) 5.2562 (188) 0.11544 (12) 1840 1862 1887 Z3
18072 0.1344 0.32499 (112) 5.0302 (186) 0.11226 12) 1814 1824 1836 16
0.034 410 142.9 Clear prisms
Z4
CR1tonalite orthogneiss(386602E,601415N)
Large, cleaar, 0.054 41.4 7 17440 0.1387 0.47972 (154) 11.5395 (400) 0.17446 (16) 2526 2568 2601
Z1 76
euhedral
7
Z2 Small, brown 0.071 85 46.9 26173 0.1319 0.48777 (150) 11.9289 (400) 0.17737 (16) 2561 2599 2628 prisms
18 13576 0.1339 0.48608 (164) 11.8186 (428) 0.17634 (16) 2554 2590
82 2619
0.098 Large, clear
Z3 45.1
euhedral
CR2leucogranite(386602E,601451N)
2843 0.2202 0.31757 (118) 4.8755 (188)
Z1 Brown prisms 0.009 251 91.3 15 0.11135 (16) 1778 1798 1821 6 5907 0.2364 0.32289 (162) 4.9749 (226) 0.11174 (30) 1804
121.3 1815
Z2 Small, brown 0.005 324 1828
prisms
Small, brown 0.033 111 48 4209 0.2231 0.32588 (94) 5.1139 (166) 0.11381 (10) 1818 1838 1861
Z3 296
needles
CR3diorite dyke(366002E,601415N)
12075 0.2525 0.32278 (86) 4.919 (290)
Best clear prisms0.034 1833 694.2 0.11053 (10) 1803 1806 1808
Z1 105
471 174.3 23 18209 0.2262 0.32139 (134) 4.8995 (214) 0.11057 (10) 1797 1802 1809 Z2 Small, brown 0.045
prisms
3867.3 218 4777 10.996 0.32161 (152) 4.8551 (240) 0.10949 (10) 1798 1795 1791 M1 Clear, pale yellow0.045 1146
CR4pegmatite(386602E,6014151N)
24977 0.1889 0.33642 (1556) 5.4943 (268) 0.11845 (10)
0.192 1869
Z1 Large, pale pink 160 60.4 26 1900 1933
56.4 7 19898 0.1909 0.3244 (96) 4.9989 (162) 0.11176 (12) 1811 1819 1828 0.046
Z2 1 large prism 155
73 16946 39.429 0.32165 (186) 4.8867 (286) 0.10968 (12)
M1 Large, dark 0.117 524 5926.6 1798 1797 1796 yellow
18625 29.371 0.32222 (108) 4.873(176) 0.10968 (12) 1801
5786.6 1798
Clear, pale yellow 57 1794
M2 0.077 679
37
D
Table 1 (Continued)
Age (Ma)
4640.4 167 5236 24.456 0.32083 (138) 4.8429 (218) 0.10948 (10) 1794 1792 1791 Clear
M4 0.066 652
CR5spotted diorite duke(39138E,6014188N)
7363 0.2571 0.31986 (120) 4.8956 (192) 0.11101 (12)
0.038 196 1789 1802 1816
Clear, cracks
Z1 73.9 20
8
Large, cracks 0.058 271 106.5 37935 0.3131 0.3206 (112) 4.9098 (184) 0.11107 (12) 1793 1804 1817 Z2
40810 0.3295 0.32262 (110) 4.9396 (180) 0.11105 (10) 1802
23 1809
277 110.9 1817
Z3 Prisms, cracked 0.168
12
Elongate prisms 0.023 224 90.2 8498 0.3297 0.32406 (112) 4.9618 (180) 0.11105 (14) 1810 1813 1817 Z4
3
Clear prisms 0.013 53 20.8 5190 0.2768 0.3253 (130) 4.9847 (196) 0.11114 (20) 1816 1817 1818 Z5
381 0.098 0.31651 (104) 4.7448 (198) 0.10872 (20) 1773
1191 1775
T1 Clear brown 0.525 42 13.7 1778
801 458 0.0951 0.31696 (98) 4.7592 (184) 0.1089 (18) 1775 1778
T2 Clear, medium 0.349 51 16.8 1781
brown
533 474 0.1191 0.31821 (108) 4.7822 (200) 0.109 (22)
57 1781
0.217 19 1782 1783
T3 Medium brown, prisms
T4 15 large brown 0.139 42 13.7 341 355 0.0916 0.31717 (104) 4.7466 (194) 0.10854 (22) 1776 1776 1775 prisms
CR6granite dyke(386167E,6010304N)
122.5 34 4420 0.0951 0.31967 (116) 4.8567 (182) 0.11019 (16) 1788 1795 1803 Z1 2 large grains 0.02 370
9459 0.1036 0.30827 (118) 4.6515 (176) 0.10944 (18) 1732
14 1759
Z2 Small grain 0.018 394 126.7 1790
4 32919 0.1246 0.32029 (142) 4.8659 (222) 0.11018 (12) 1791 1796 1802 Z3 Small, clear, 0.025 261 88.9
euhedral
28133 0.098 0.3165 (136) 4.8557 (212) 0.11017 (14) 1788
103.5 4 1795 1802
312 0.02 Clear, euhedral Z4
Z5 3 large grains 0.022 113 42.5 4 11193 0.2565 0.31813 (186) 4.8344 (262) 0.11021 (30) 1781 1791 1803
CR7De Pas monzogranite(389070E,6016195N) 15
Z1 Sharp, elongate 0.023 240 85.9 7451 0.1797 0.3225 (114) 4.9203 (176) 0.11065 (18) 1802 1806 1810 prisms
7724 0.1896 0.32281 (156) 4.925 (228) 0.11065 (24) 1803
87.3 1807
242 11 1810
Z2 Sharp, elongate 0.018
aThe 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 UTM
D.T.James,G.R.Dunning/Precambrian Research103 (2000) 31 – 54 43
Fig. 6. Lobstick group supracrustal rocks including pelitic migmatite (right), calc – silicate gneiss (centre) and quartzite (left). The metasedimentary rocks are cut by a late syn-tectonic tonalite dyke (folded). Sample MR3 was collected from a similar dyke.
45% probability of fit which yields a lower inter-cept age of 1802 +9/ −14 Ma, interpreted to represent the igneous crystallization age of the dyke. The upper intercept is 2850 Ma with a large uncertainty. This brackets formation of the steeply east-dipping foliation in the host rocks to be between 1815 Ma, the emplacement age of the foliated tonalite (MR1) and 1802 Ma.
The field and geochronological data demon-strate that Archean Flat Point gneiss and \1815 Ma Lobstick group rocks were intruded by tonalite, metamorphosed to upper-amphibolite – facies and deformed in the interval between 1815 and 1802 Ma. The data indicate that the gneissos-ity and superposed folds in the Flat Point gneiss are older than 1815 Ma. It is tacitly assumed that these features are Archean, although this has not been confirmed by a geochronological test.
4.2. Crossroads domain
Field relations demonstrate that the oldest rocks in Crossroads domain are high-grade supracrustal gneisses, informally named the Overflow group, consisting principally of metasedimentary (pelitic) migmatite (Fig. 7), mi-nor amounts of mafic and felsic metavolcanic rocks and associated chert – magnetite iron forma-tion. The precise age of the supracrustal rocks is undetermined, although they are constrained to be \2704 Ma on the basis of U – Pb geochrono-logical data from younger intrusive units (de-scribed below).oNd values for two samples of the
metasedimentary rocks are approximately +1 and +2 at 2700 Ma, and they have depleted-mantle model ages (T dm ages) of approximately 2800 Ma (Kerr et al., 1994). The Nd data indicate the rocks were probably not derived from erosion of Middle or Early Archean crust, and suggest that their depositional ages are between 2700 and 2800 Ma. One possibility is that they could have been derived from erosion of coeval Overflow group volcanic rocks. These supracrustal rocks are provisionally correlated with similar Archean high-grade metasedimentary and metavolcanic rocks in the Orma domain (see Nunn and Noel, 1982; Nunn, 1993).
Fig. 7. Archean pelitic migmatite (Overflow group) containing superposed folds of Archean age and a deformed Archean amphibolite dyke, Crossroads domain.
strain the metasedimentary rocks to be older than 1805 Ma. They are probably older than 1815 Ma, the age of the MR1 tonalite, although field rela-tions do not unequivocally prove this. A sample of paleosome from the pelitic migmatite has a depleted-mantle Nd-model age of 2310 Ma (Kerr et al., 1994).
Overflow group rocks are intruded by variably deformed and metamorphosed plutonic units in-cluding tonalite and granite orthogneisses that contain several phases of metamorphic leuco-some, granitic rocks belonging to the ca. 1840 – 1810 Ma De Pas batholith, deformed granitic plutons, some of which are probably related to the De Pas batholith, and several ages of variably deformed and metamorphosed mafic dykes. To better understand age relationships between the various 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 tonalite orthogneiss which intrudes Overflow group supracrustal gneisses. Three zircon fractions from the sample define a discordia line (Fig. 8, Table 1) with an upper intercept of 2704915 Ma, inter-preted to represent the igneous crystallization age of the rock. This age is significantly older than the ca. 2620 Ma emplacement age determined for a monzogranite orthogneiss from the southeastern Crossroads domain (James et al., 1996), although it is consistent with 2682 – 2675 Ma emplacement ages of tonalite intrusions in the Orma domain
(see Nunn et al., 1990). The lower intercept of the discordia line is 1815 Ma, and is thought to represent incipient Pb-loss during a ca. 1815 Ma metamorphic event. Crossroads domain or-thogneisses have o Nd values of between 0 and +2 at 2650 Ma, and T dm model ages between 2800 and 2650 Ma (Kerr et al., 1994). The Nd and U – Pb data indicate that these intrusions are juve-nile, Late Archean additions to the crust. Nd data from Orma domain tonalite orthogneiss are simi-lar; rocks have o Nd values of +1 at 2675 Ma
and T dm ages of approximately 2770 Ma (Kerr et al., 1994).
The gneissosity in CR1 orthogneiss is cut by a pink, recrystallized leucogranitic dyke, which is deformed by a locally intense foliation that also overprints the host orthogneiss. A discordia defined by two fractions of zircon collected from the 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 Pas batholith granite collected from the southern Crossroads domain. On this basis, the dyke is interpreted to be related to De Pas magmatism.
D.T.James,G.R.Dunning/Precambrian Research103 (2000) 31 – 54 45
Fig. 9. Metamorphosed diorite dyke (CR3 sample location; left side of photograph) cutting CR1 tonalite orthogneiss.
The data constrain the age of the gneissosity in the host rocks (i.e. in sample CR1) to be \1836 Ma, and as Archean rocks in the domain do not contain any isotopic evidence of thermal events between ca. 2620 and 1836 Ma, we interpret the gneissosity to be an Archean feature.
The tonalite orthogneiss, leucogranite dyke and it’s contained foliation are cross-cut by a grey-weathering, recrystallized and weakly deformed diorite dyke (CR3, Figs. 8 and 9), which is in turn cut by an undeformed granitic pegmatite (CR4). The igneous crystallization age of the diorite dyke is interpreted to be 180992 Ma based on two fractions of concordant and nearly concordant zircons. A single monazite fraction from CR3 was dated at 179593 Ma and interpreted to represent the time of metamorphism. The pegmatite dyke (CR4, Fig. 8) is interpreted to have a crystalliza-tion age of 180092 Ma based on data from zircon and three concordant monazite analyses.
To provide additional constraints on the age of the supracrustal rocks, their included metamor-phic and structural features, and ages of intru-sions, samples of a diorite dyke (CR5) and a granite dyke (CR6) were also dated. The diorite dyke (Fig. 10), which has a similar mineralogy and texture to the CR3 dyke, cross-cuts relict primary layering, gneissosity and foliation in host Overflow group mafic metavolcanic rocks. How-ever, the diorite dyke is itself metamorphosed, it contains 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 crystallization age of CR5. Four fractions of titanite from the same rock define an age of 1775 Ma. The titanite data may represent a metamorphic cooling age, or they may indicate renewed thermal activity and crystallization of new titanite at 1775 Ma. Sample CR6 is from an undeformed, white-weathering granite dyke that is discordant to gneissosity in host Overflow group metasedimentary migmatite. Five fractions of zircon define a discordia line (85% probability of fit) with an upper intercept age of 180695 Ma (Fig. 12) and interpreted to be the crystallization age of the rock.
The Crossroads domain contains intrusions of variably foliated and recrystallized K-feldspar
Fig. 10. Metamorphosed (‘spotted’) diorite dyke (CR5 sample location, lower left) cutting Archean mafic and felsic volcanic rocks (left side of photo) of the Overflow group.
megacrystic granite, granodiorite and charnockite belonging to the main De Pas batholith, sensu stricto, and presumed satellite intrusions of the De Pas batholith, which occur to the east of the main batholith. The satellite intrusions are corre-lated with the batholith on the basis of lithology. Two samples of De Pas batholith megacrystic granite, from the main part of the batholith in the southern Crossroads domain, have emplacement ages of 183195 Ma (James et al., 1996) and 181193 Ma (Krogh, 1986), as determined by U – Pb dating of zircon. One of the presumed satellite intrusions, consisting of strongly foliated megacrystic granite, is dated at 182395 Ma (James et al., 1996). From farther north in the batholith, Dunphy and Skulski (1996) have deter-mined that a foliated De Pas batholith tonalite
has an emplacement age of 1840 Ma on the basis of preliminary U – Pb dating of zircon.
In an attempt to obtain minimum ages of em-placement for De Pas K-feldspar megacrystic granite, and to constrain the timing of deforma-tion that overprints these rocks, a unit of isotropic to very weakly foliated, pink, biotite monzogran-ite containing xenoliths of strongly foliated K-feldspar megacrystic granite was sampled. On the basis of lithology and structure, the xenoliths are correlated with foliated De Pas K-feldspar megacrystic granite. Field relations suggest that the biotite monzogranite (CR7) is late syn-tec-tonic with respect to the deformation in the in-cluded megacrystic granite. Two fractions of concordant zircons from sample CR7 (Fig. 13) yield an age of 181093 Ma, interpreted to repre-sent the crystallization age of the rock. This age is within error of the youngest emplacement ages from the De Pas batholith. The strong foliation in the batholith is inferred to have formed between 1810 and 1823 Ma.
5. Discussion
5.1. Archean e6olution
Field, geochronological and Nd-isotopic data indicate that the Crossroads and Orma domains contain a significant component of Late Archean granite – greenstone terrane crust. Supracrustal rocks in both domains are undated but are con-strained to be older than 2704 Ma. They are intruded by tonalite to granite plutons in the interval between ca. 2704 and 2620 Ma; these are inferred to be pre- to syn-tectonic with respect to high-grade metamorphism and attendant defor-mation. Granitic intrusions correlated with the De Pas batholith, and mafic dykes correlated with the 1817 – 1809 Ma dykes, are discordant to super-posed folds of metamorphic leucosome in the supracrustal rocks in the Crossroads domain, sug-gesting that these features are Archean. A possible interpretation of the data is that Archean rocks in Crossroads and Orma domains have a similar Archean depositional, intrusive and tectonother-mal evolution.
Fig. 12. U – Pb concordia diagram for sample CR6.
D.T.James,G.R.Dunning/Precambrian Research103 (2000) 31 – 54 47
The McKenzie River domain lacks Archean supracrustal rocks and the ca. 2776 Ma Flat Point gneiss is significantly older than intrusive units in the Crossroads and Orma domains. In contrast to the 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 the rocks incorporated a component of older crust. Archean rocks in the McKenzie River domain were probably not part of the same granite – greenstone terrane that makes up the Crossroads and Orma domains.
Parentage of Archean rocks that occur in the Core Zone of the southwestern SECP is uncertain. Archean data from these rocks is non-unique and could be used to support radically different mod-els. However, the fact that the entire exposed margin of the Superior craton shows evidence of a significant rifting event, initiated between ca. 2.2 and 2.0 Ga, and reflected in the development of Cycle one rocks in the southern Labrador Trough, strongly suggests that even if Archean rocks in the Core Zone have a Superior craton affinity, they acted as independent crustal blocks after ca. 2.0 Ga. Moreover, if the De Pas batholith is a subduction-related magmatic arc, then Paleoproterozoic ocean basins must have separated distinct cratons or blocks consisting of Archean 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 De Pas batholith. These granitoid rocks are variably deformed and recrystallized demonstrating they predate and overlap with a tectonothermal event, which persisted to at least 1775 Ma on the basis of U – Pb titanite data from other rocks the do-main. Paleoproterozoic metamorphism in the Crossroads domain reached amphibolite facies, but it did not result in the production of meta-morphic leucosome. Major- and trace-element geochemistry of De Pas batholith rocks define calc-alkaline trends (Van der Leeden et al., 1990;
Dunphy and Skulski, 1996), and are compatible with an interpretation for the batholith as a conti-nental magmatic arc formed above an east-dip-ping subduction zone (Martelain, 1989; Van der Leeden et al., 1990; Dunphy and Skulski, 1996). In apparent contradiction to this interpretation, Kerr et al. (1994) have noted that De Pas batholith rocks do not have all of the geochemical signatures characteristic of modern subduction-re-lated continental magmatic arcs; they note, as one example, that most samples show a strong enrich-ment in Zr. However, some of the geochemical signatures, which are not characteristic of modern arcs, may in part be related to the nature of the host Archean crust and to how much of the host was assimilated during emplacement of the De Pas batholith. The batholith rocks are character-ised by negative o Nd values between −3 and −7, calculated at 1830 Ma, and T dm ages between 2.24 and 2.64 Ga indicating that De Pas magma was significantly contaminated by the sur-rounding Archean crust (Kerr et al., 1994; Dun-phy and Skulski, 1996). The geochemical signatures of De Pas batholith rocks are inter-preted to reflect the mixing of juvenile, subduction related magma and Archean crust of the Core Zone.
Fig. 14. Summary diagram of U – Pb geochronological data (Paleoproterozoic ages) for the study area. Evidence for coeval metamorphism and deformation in McKenzie River and Crossroads domains after ca. 1810 Ma, and different magmatic histories prior to that, suggests the domains were juxtaposed by that time. Included are U – Pb ages for De Pas batholith rocks from 1 (James et 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, which postdated emplacement of the De Pas batholith is probably not significant in the southern part of the Core Zone. This interpretation is consistent with field and geochronological data indicating that deformation in the George River shear zone, 170 km north of our study area, was occurring but waning at 1825 Ma (Dunphy and Skulski, 1996). Furthermore, the fact that the Crossroads and Orma domains have very similar Archean
D.T.James,G.R.Dunning/Precambrian Research103 (2000) 31 – 54 49
intrusive rocks, effectively escaped Paleoprotero-zoic tectonothermal effects. Absence of a late syn-to post-De Pas batholith metamorphism in the Orma domain is also consistent with tectonic models for the batholith which predict that it is exposed as a tilted, oblique section having deeper, higher-grade rocks in the west and lower-grade rocks in the east (Dunphy and Skulski, 1996).
The McKenzie River domain does not contain De Pas batholith granitic rocks, suggesting the McKenzie River and Crossroads domains were not in their present configuration until after ca. 1810 Ma, the youngest emplacement age for De Pas granite in the region. An occurrence of ca. 1815 Ma tonalite indicates igneous activity in the McKenzie River domain at this time, although the significance of this tonalite is uncertain. Leu-cosome from amphibolite – facies Lobstick group metasedimentary rocks, dated at ca. 1805 Ma, indicates that the peak of metamorphism in the McKenzie River domain falls within the range of metamorphic monazite and titanite ages from the Crossroads domain. These data and the fact that deformation in the Lac Tudor shear zone was attendant with amphibolite – facies metamorphism suggests that the McKenzie River and Crossroads domains were linked by ca. 1805 Ma. A recrystal-lized tonalite dyke, which cross-cuts the metamor-phic leucosome and strong foliation and is dated at 1802 +9/ −14 Ma, indicates that metamor-phism outlasted deformation in the McKenzie River domain. Ages of metamorphic monazite and titanite from the Crossroads and McKenzie River domains (this study) are consistent with the ages of metamorphic monazite (179395 and 178392 Ma) and titanite (177495 and 17839
11 Ma) in orthogneisses from the Kuujjuaq do-main (Machado et al., 1989), 300 km north of the study 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 and obliquely converging North Atlantic and Superior cratons. Convergence of the cratons resulted in
sequential closure of intervening Paleoproterozoic basins, and their northward movement terminated in their ultimate collision with Archean terranes along their northern margins (see Hoffman, 1990). Core Zone collision and deformation occurred first along it’s eastern margin; accretion of Tasi-uyak sediments to the North Atlantic craton and the onset of deformation in the Torngat Orogen and eastern Core Zone occurred at 1.86 – 1.85 Ga (Van Kranendonk et al., 1993; Wardle and Van Kranendonk, 1996), and preceded tectonothermal and magmatic events in the western Core Zone and in the New Quebec orogen. The tectonic setting of formation and accretion of the 1.84 – 1.82 Ga Lac Lomier Complex (subduction-related arc?), occurring along the boundary between the Core Zone and the North Atlantic craton are uncertain (see discussion in Wardle, 1998).
have an affinity with the Superior craton; they were incompletely rifted from the craton after ca. 2.2 Ga.
McKenzie River domain was not terminally juxtaposed to the Crossroads domain until ca. 1810 Ma; a consequence of this collision was the shutting down of the subduction zone and cessa-tion of De Pas magmatism. U – Pb
geochronologi-cal data (this study) demonstrates ca. 1815 Ma emplacement of tonalite in the McKenzie River domain overlaps in time with De Pas magmatism. Tectonic setting for this magmatism, which is slightly younger than emplacement ages from the Kuujjuaq batholith, is uncertain. Amphibolite – fa-cies metamorphism and production of ca. 1805 Ma metamorphic leucosome in the domain may
D.T.James,G.R.Dunning/Precambrian Research103 (2000) 31 – 54 51
be related to crustal thickening as a result of overthrusting by the Crossroads domain along the Lac Tudor shear zone. The youngest structures in the shear zone are related to dextral transcurrent displacement. A switch from mainly convergent (thrust) to strike-slip displacement may have oc-curred when shortening in the southwestern Core Zone could no longer be accommodated by crustal thickening.
The onset of deformation in the New Que´bec Orogen is inferred to be related to collision of the western Core Zone with the Superior craton mar-gin. Our model predicts that this did not occur until ca. 1810 Ma. There are insufficient data from the New Que´bec Orogen to constrain the timing of initial thrusting and to test this model, although Machado et al. (1997) claim that thrust-ing occurred locally before 1813 Ma. Evidence from the Cape Smith Belt along the northern Superior margin suggests that final collision oc-curred there at ca. 1800 Ma (St-Onge et al., 1998). The lack of detailed geochronological data to constrain the timing of thrusting in the New Que´bec Orogen and the timing and kinematic
history of shear zones separating it from the Core Zone remain outstanding problems in understand-ing foreland – hinterland relationships in the SECP.
Acknowledgements
The U – Pb geochronological work summarized here was funded by LITHOPROBE (ECSOOT Transect), Canada. Geological field studies were funded 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 the geochronology laboratory at Memorial University of Newfoundland in St. John’s. Richard Wardle is thanked for his comments on an earlier version of this manuscript. Our paper benefitted from re-views by N. Machado and M. St-Onge. This paper is published with the permission of the Director of the Geological Survey of Newfound-land and Labrador.
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