The Margaree orthogneiss: an Ordovician, peri-Gondwanan, mafic–felsic igneous complex insouthwestern Newfoundland1

Pablo Valverde-Vaquero, Greg R. Dunning, and Cees R. van Staal

Abstract: The igneous protoliths of the Margaree orthogneiss intruded the peri-Gondwanan rocks of the Port-aux-Basques Complex, in the southwestern corner of the Hermitage Flexure, prior to regional deformation and metamor-phism. Field relationships and U–Pb geochronology indicate that the Margaree orthogneiss represents a 20 km long,mafic–felsic intrusive complex formed by amphibolite, dioritic orthogneiss,474 4

14−+ Ma hornblende-bearing tonalitic

orthogneiss with mafic enclaves, 472 ± 2.5 Ma and 465 ± 3 Ma biotite-bearing granitic orthogneiss, and minor ultra-mafic rocks. Most amphibolite have the chemistry of volcanic-arc tholeiite and are interpreted to be coeval with the in-trusion of tonalitic and granitic magma with volcanic-arc geochemical signatures. Locally, mafic magmatism continuedafter the intrusion of the 465 Ma granite. The Margaree orthogneiss and its country rock were overprinted by upperamphibolite-facies metamorphism and deformation associated with the final closure of the Iapetus Ocean. A late-synkinematic granite dates the late stages of the high-temperature regional deformation at417 4

7−+ Ma, while metamor-

phic titanite (411 ± 2 Ma) is interpreted to date extensive recrystallization under amphibolite-facies conditions as EarlyDevonian. The Margaree orthogneiss is broadly coeval with the plutons that postdate the Early OrdovicianPenobscottian arc – passive margin imbrication in central and southern Newfoundland. It also coincides with the exten-sive late Arenig – early Llanvirn back-arc rifting event along the entire peri-Gondwanan margin of the northern Appa-lachians. The external position of the Port-aux-Basques Complex with respect to the back-arc elements in theHermitage Flexure, coupled with the rock types and geochemistry of the orthogneiss, suggest formation of theMargaree orthogneiss in an arc and (or) back-arc transitional setting.

Résumé: Les roches du complexe péri-Gondwana de Port aux Basques, dans le coin sud-ouest de la flexure Hermi-tage, ont subi l’intrusion des protolithes ignés de l’orthogneiss Margaree avant la déformation et le métamorphisme ré-gional. Les relations de terrain et la géochronologie U–Pb indiquent que l’orthogneiss Margaree représente uncomplexe intrusif mafique felsique de 20 km de long formé par de l’amphibolite, de l’orthogneiss dioritique, del’orthogneiss tonalitique à hornblende de474 4

14−+ Ma avec des enclaves mafiques, des orthogneiss granitiques à biotite

de 472 ± 2,5 Ma et de 465 ± 3 Ma ainsi qued’autres roches mineures ultramafiques. La plupart des amphibolites ontla composition chimique d’une tholéiite d’arc volcanique et on les croit contemporains de l’intrusion d’un magma gra-nitique et tonalitique présentant des signatures géochimiques d’arc volcanique. Localement, il y eut continuation dumagnétisme mafique après l’intrusion du granite à 465 Ma. L’orthogneiss Margaree et sa roche hôte ont été surimpri-més par de la déformation et du métamorphisme au faciès amphibolite élevé associés à la fermeture finale de l’océanIapetus. Un graphite tardi-syncinématique donne une date de417 4

7−+ Ma pour les phases tardives de la déformation ré-

gionale à haute température alors qu’une titanite métamorphique (411 ± 2 Ma) donnerait le Dévonien précoce commedate d’une recristallisation généralisée sous des conditions de faciès de l’amphibolite. L’orthogneiss Margaree est large-ment contemporain des plutons qui postdatent l’imbrication, au Ordovicien précoce, de l’arc penobscottien et de la demarge passive, au centre et au sud de Terre-Neuve. Il coïncide également avec un événement de distension importantede l’arrière arc, à l’Arénigien tardif – Llanvirnien précoce, le long de toute la marge péri-Gondwana des Appalachesdu Nord. La position externe du complexe de Port aux Basques par rapport aux éléments d’arrière arc dans la flexureHermitage, jumelée aux types de roches et à la géochimie de l’orthogneiss, suggère la formation de l’orthogneiss Mar-garee dans un environnement d’arc et (ou) d’arrière arc transitionnel.

[Traduit par la Rédaction] Valverde-Vaquero et al. 1710

Can. J. Earth Sci.37: 1691–1710 (2000) © 2000 NRC Canada

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Received May 25, 1999. Accepted June 5, 2000. Published on the NRC Research Press Web site on November 20, 2000.Paper handled by Associate Editor L. Corriveau.

P. Valverde-Vaquero2 and G.R. Dunning. Department of Earth Sciences, Memorial University of Newfoundland, St. John’s,NF A1B 3X5, Canada.C.R. van Staal. Continental Geosciences Division, Geological Survey of Canada, 601 Booth Street, Ottawa, ON K1A 0E8, Canada.

1Geological Survey of Canada Contribution 1998263.2Corresponding author. Present address: Institut für Geowissenchaften und Lithosphärenforschung, Justus-Liebig Universität,Senckenbergstrasse 3, Giessen, D-35390, Germany (e-mail: pablo.valverde@geolo.uni-giessen.de).

Introduction

The classic lithotectonic zones of the Newfoundland Ap-palachians (Williams et al. 1988) converge in southwesternNewfoundland at the western corner of the Hermitage Flex-ure (Fig. 1A; Williams et al. 1970). This convergence is theresult of the final closure of the Iapetus Ocean and theSiluro-Devonian continental collision between the Lauren-tian and peri-Gondwanan margins of the Iapetus Ocean (e.g.,Williams et al. 1988; Dunning et al. 1990; Lin et al. 1994).The present-day expression of the suture zone is the Red In-dian Line and its continuation in southern Newfoundland,the Cape Ray Fault (Fig. 1B; e.g., Dubé et al. 1996). TheCape Ray Fault juxtaposes against the peri-Laurentian rocksof the Dashwoods subzone, a package of gneiss, the Port-aux-Basques gneiss (Brown 1977). This gneiss is currentlydivided into the Grand Bay and Port-aux-Basques complexesand the Harbour le Cou Group (Fig. 1B; van Staal et al.1996a, 1996b). These rocks are considered part of the peri-Gondwanan margin of the Iapetus Ocean (e.g., Williams etal. 1988; Lin et al. 1994). A straightforward lithological cor-relation of these gneissic rocks with the low-grade rocks ofthe early Paleozoic miogeocline of the Gander Zone, and thearc- and back-arc-derived rocks and ophiolite suites of theExploits subzone, is hampered by the intensity of the Siluro-Devonian metamorphism and deformation (e.g., Dunning etal. 1990; Burgess et al. 1995), and the lack of precise timeconstraints on their pre-Silurian history, particularly in thecase of the Grand Bay and Port-aux-Basques complexes.This study of the Margaree orthogneiss (van Staal et al.1996a, 1996b) is directed at unravelling part of this pre-Silurian evolution. This orthogneiss predates regional defor-mation and metamorphism and contains the oldest units tointrude into the Port-aux-Basques Complex (Brown 1977;van Staal et al. 1996a, 1996b). These new data provide cru-cial details on the Ordovician evolution of the Port-aux-Basques Complex and illustrate the geological complexity ofthe rocks that form the southern edge of the peri-Gondwananmargin of the Newfoundland Appalachians.

The Margaree orthogneiss is a composite unit of tonaliticto granitic gneiss, amphibolite, and minor ultramafic rocks.This gneiss crops out in Channel Island, near Port-aux-Basques, and along the coastline of Margaree and FoxRoost, where it forms a 2 km wide band that can be mappedat least 15 km inland beyond Grandys Brook (Fig. 2). Theoutcrop pattern of the Margaree orthogneiss coincides withthe trend of a well-defined positive regional magnetic anom-aly (Kilfoil 1993), which highlights its geophysical contrastwith the surrounding paragneiss. Mineralogical differenceswith the surrounding, sedimentary-derived gneiss are the ab-sence of garnet, Al-silicate, and muscovite. Its metaluminouscharacter also differentiates the Margaree orthogneiss fromthe locally garnet–muscovite-bearing Port-aux-Basquesgranite (van Staal et al. 1996a, 1996b) and the garnet–gedrite-bearing Kelby Cove orthogneiss (Schofield et al.1998), in the Port-aux-Basques and nearby Grand Bay com-plexes. Previous work on the Margaree orthogneiss is re-stricted to the description of some of its rock units in Brown(1977) and Burgess et al. (1995) and the geochemical studyof some of its members within a wider regional context byWhalen et al. (1997) and Schofield et al. (1998).

Geological setting

The Port-aux-Basques Complex (Figs. 1B and 2), includ-ing the Margaree orthogneiss, is part of a package ofgneissic rocks that crop out in southwestern Newfoundlandbetween the peri-Laurentian rocks of the Dashwoodssubzone (Williams 1995), including the Windsor PointGroup and the peri-Gondwanan, Exploits subzone equiva-lent, rocks of the Bay du Nord Group (see Tucker et al.1994). The gneissic rocks are divided into three lithologicunits bounded by shear zones, from west to east: Grand BayComplex, Port-aux-Basques Complex, and Harbour le CouGroup (Fig. 2; van Staal et al. 1996a,1996b; Schofield et al.1998). All three units share the same regional metamor-phism and three major phases of regional deformation (vanStaal et al. 1992; Lin et al. 1993). Peak metamorphism wasreached during the D2 deformation, and metamorphic gradeincreases, according to a Barrovian series, from lower am-phibolite-facies in the Grand Bay Complex to upper amphi-bolite-facies in the eastern portion of the Port-aux-BasquesComplex and the Harbour le Cou Group. (e.g., Owen 1992;Burgess et al. 1995).

The Grand Bay ComplexThe Grand Bay Complex (Fig. 2) is the westernmost unit

and is tectonically bound to the Dashwoods subzone by theCape Ray Fault. Dubé et al. (1996) demonstrated Late Silu-rian to Early Devonian, high-angle, oblique thrusting of theGrand Bay Complex over the low-grade Windsor PointGroup along this fault (Fig. 1B). This complex comprisesgedrite-bearing schist, metapsammite, abundant intercala-tions of amphibolite, minor coticule beds, orthogneiss (vanStaal et al. 1996a), and hosts a significant Zn–Pb–Cu–Agmassive sulphide deposit (Isle-aux-Morts prospect; O’Neill1985). The orthogneiss (Fig. 2) include undated, pre-D1,tonalitic to granitic orthogneiss with associated amphibolite(Kelby Cove orthogneiss; van Staal et al. 1996a) and syn-D1granite (Port-aux-Basques granite; van Staal et al. 1996a).Schofield et al. (1998) differentiated a package of tholeiiticmafic and ultramafic rocks, the Big Barachois assemblage,which they interpreted to have volcanic-arc affinities. Thisset of rocks was metamorphosed in the staurolite and kyanitezones, and peak metamorphic conditions were reached at ca.580°C within the Ky + St stability field (e.g., Burgess et al.1995). Metamorphism in this complex is dated by a meta-morphic titanite206Pb/238U age of 412 ± 2 Ma(Dunning etal. 1990). Burgess et al. (1995) and Dubé et al. (1996) ob-tained40Ar/39Ar cooling ages for hornblende (406–393 Ma)and biotite (403 Ma). This complex is also intruded by thelate-tectonic Isle-aux-Morts granite (386 ± 3 Ma, U–Pb Zrn;Dubé et al. 1996), which locally stitches the Cape Ray Fault.The contact with the Port-aux-Basques Complex, to the east,is a late syn- to post-D2, thrust-sense shear zone that isknown as the Grand Bay Fault (Burgess et al. 1992).

The Port-aux-Basques ComplexThe Port-aux-Basques Complex (Fig. 2) comprises

quartzo-feldspathic paragneiss, amphibolite, minor meta-pelite, very rare coticule, and abundant orthogneiss: the Port-aux-Basques granite and the Kelby Cove and Margareeorthogneiss (van Staal et al. 1996a, 1996b). These rocks are

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Fig. 1. (A) Location of the Hermitage Flexure and the area of study (rectangle) in the northern Appalachians (modified after Barr andWhite 1996) showing the lithotectonic zones of the Newfoundland Appalachians with peri-Gondwanan affinities (patterned) and the ex-tent of the boundary between peri-Gondwanan and peri-Laurentian rocks. (B) Generalized geological map of the western corner of theHermitage Flexure, southwest Newfoundland (modified after Chorlton and Dallmeyer 1986; Lin et al. 1994; Tucker et al. 1994; Dubéet al. 1996) showing the location of the Port-aux-Basques Complex (rectangle, area in Fig. 2).

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Fig. 2. Geological map of the area between Port-aux-Basques and Garia Bay showing the distribution of the Margaree orthogneiss (compiled after van Staal et al. 1996a,1996b and Dubé et al. 1996). Geochronological data after Chorlton and Dallmeyer (1986), Dunning et al. (1990), van Staal et al. (1994), Burgess et al. (1995), and Dubé etal. (1996). The U–Pb samples sites (this study) are indicated.

bound to the west by the Grand Bay Fault and to the east bythe Isle-aux-Morts fault. The Port-aux-Basques Complexwas partially metamorphosed in the kyanite zone, thesillimanite zone, and the transition between the sillimaniteand the sillimanite + K-feldspar zones. The highest peakmetamorphic conditions were reached in the eastern part ofthe complex (Owen 1992; Burgess et al. 1995). Field rela-tionships indicate that the igneous protoliths of the KelbyCove and the Margaree orthogneiss are pre-D1 and olderthan the Port-aux-Basques granite (Brown 1977, van Staal etal. 1996a). The Kelby Cove orthogneiss crops out along thewestern side of the complex, whereas the Margareeorthogneiss occupies the eastern portion of the complexalong the contact with the Harbour le Cou Group. Van Staalet al. (1994) reported a U–Pb age of ca. 450 Ma for the syn-D1 Port-aux-Basques granite in this complex. The U–Pbmonazite ages of 420–415 Ma from migmatitic paragneissare interpreted to date peak metamorphism (van Staal et al.1994). Burgess et al. (1995) obtained 407–399 Mahornblende and 394–391 Ma muscovite40Ar/39Ar coolingages, and Lowdon et al. (1962) and Wanless et al. (1965) re-ported K–Ar muscovite ages of 415 and 400 ± 20 Ma fromthese gneissic rocks.

The Harbour le Cou GroupThe Harbour le Cou Group (Fig. 2) does not contain

orthogneiss with pre-D1 protoliths and has been divided intwo formations (van Staal et al. 1996b): the Otter Bay andGrandy’s formations. The Otter Bay Formation containssiliciclastic rocks interleaved with mafic pillow lavas anddykes with tholeiitic affinities and has a gradational contactwith the overlying metapelitic rocks and interleavedcoticules of the Grandy’s Formation (Schofield et al. 1998).The eastern part of the group contains abundant sheets ofsyn-D2 two-mica granite (Rose Blanche granite; e.g., Bennet al. 1993). Peak metamorphic conditions were in excess of700°C and 6.6–5.2 kbar (1 kbar = 100 mPa; Burgess et al.1995). Burgess et al. (1995) obtained a U–Pb titanite age of418 ± 9 Ma and40Ar/39Ar hornblende (419–404 Ma) andmuscovite (391 Ma) cooling ages from the western portionof the Harbour le Cou Group. This group is bound to thewest by the Isle-aux-Morts fault. To the east, it is bound tothe greenschist-facies rocks of the Bay du Nord Group bythe Bay le Moine shear zone (Lin et al. 1993). Inland, thecontact is uncertain as the two groups are separated by largemasses of synmetamorphic granitoid rocks (Rose Blanche-like granite) and late-kinematic granite (416 ± 4 Ma, LaPoile granite; Chorlton and Dallmeyer 1986). According toLin et al. (1993), lithological correlations suggest that theHarbour Le Cou Group is a higher grade metamorphicequivalent of the Bay du Nord Group, and therefore part ofthe Exploits subzone.

Local settingThe metasedimentary rocks around the Margaree

orthogneiss (Fig. 2) are dominated by quartzo-feldspathicparagneiss alternating with amphibolite dykes and sills, am-phibole-rich paragneiss, and, locally, pelitic micaschist. Allthese rocks are part of the Port-aux-Basques Complex. Inpelitic rocks, leucosomes of granitic composition and restiteselvedges of Bt + Sil ± Grt provide evidence for partialanatexis. These leucosomes make up to 30% of the outcrop.

According to Burgess et al. (1995), these anatecticmigmatite rocks were produced by muscovite-out reactionsduring peak metamorphism at conditions of 700°C and 8–10 kbar. In the vicinity of Isle-aux-Morts (Fig. 2). TheMargaree orthogneiss is intruded by an undated, strongly fo-liated, pre-D2, garnet-bearing granitic gneiss, which is corre-lated with the Port-aux-Basques granite (van Staal et al.1996b). Locally, the Margaree orthogneiss (Fig. 3A) and thesurrounding paragneiss were intruded by syn- and late-kinematic granitic dykes (Fig. 3B), coeval with the D2 andD3 phases of regional deformation, and by pre- and (or) syn-D3 and post-D3 pegmatites (Fig. 3C; Valverde-Vaquero1997).

The Margaree orthogneiss and the surrounding paragneisspreserve evidence of the three major phases of regional de-formation and a local D4 phase (van Staal et al. 1992;Valverde-Vaquero 1997). The gneissosity in the orthogneisswas produced during D1 and was partially transposed by theD2 deformation. The complicated fold interference patterns(Fig. 3C) and folded boudins observed in the gneiss are theproduct of the D2 and D3 phases of deformation. The maingneissosity (S0–S1) was steepened along a southwest–northeast trend as a result of the F3 folding. D2 coincidedwith peak metamorphism at upper amphibolite-facies condi-tions (Burgess et al. 1995); local formation of granitic peg-matitic melts in the necks of D3 boudins (Fig. 3D) indicatethat high-temperature conditions were maintained duringpart of D3. Microstructural data (Valverde-Vaquero 1997) in-dicates that recrystallization was partially to almost com-pletely achieved after D3. The latter D4 deformationoccurred at greenschist facies and was associated with localshear zones.

The map pattern (Fig. 2) shows that the Margareeorthogneiss is interleaved with the paragneiss. The field con-tacts between the country rock and the Margaree orthogneissare straight due to the intense regional deformation or are re-worked by local brittle–ductile shear zones. Only theGrandys Brook section has a noticeable intrusive relation-ship between a granitic orthogneiss, assigned to theMargaree orthogneiss, and the country rock paragneiss.

The Margaree orthogneiss: lithological units

The Margaree orthogneiss consists of hornblende-bearingtonalitic orthogneiss, granitic orthogneiss, amphibolite, andultramafic rocks. These units were studied in detail in thecoastal section around Margaree and Fox Roost and 15 kminland, in the Grandys Brook section, and are recognizablealong the entire trace of the Margaree orthogneiss. No traceof primary minerals or textures are preserved in these rocks.

Tonalitic orthogneissThis unit makes up more than 60% of the Margaree

orthogneiss. Hornblende-bearing tonalitic rock types withsubordinate dioritic and granodioritic rocks have beengrouped in this unit. It includes homogeneous, leucocratic,tonalitic orthogneiss with mafic enclaves and minormelanocratic, tonalitic, and (or) dioritic orthogneiss. Theserocks are locally intruded by the granitic orthogneiss and byamphibolite dykes. These relationships suggest that theserocks are the oldest members of the Margaree orthogneiss.

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Homogeneous, hornblende-bearing, leucocratic, tonaliticorthogneiss with abundant mafic enclaves is well exposed inquarries along the old road to Margaree, the Dolphin road,and the Grandys Brook section. These rocks include most ofthe Margaree orthogneiss. They are fine- to medium-grained(~1 mm) and have 35–30% plagioclase (An20–30), 30% greenhornblende, 10–15% pale–dark brown biotite, 20% quartz,and >10% K-feldspar. Epidote, apatite, zircon, opaque min-erals, and rare titanite form accessory minerals. Thegneissosity (Fig. 3E) is defined by millimetre-widediscontinous hornblende and biotite-rich layers alternatingwith felsic layers, folded and transposed hornblende-bearingfelsic veins, and flattened mafic enclaves. Alignedhornblende defines a mineral lineation in the foliation planeof the melanocratic and leucocratic, tonalitic gneiss.

The melanocratic, tonalitic, and (or) dioritic orthogneissoccurs in the coastal section of Fox Roost as compositesheets of dioritic orthogneiss, amphibolite, and small ultra-mafic pods intruded by felsic centimetre- to metre-wide gra-nitic veins (Fig. 3A). These granitic veins coalesce with thenearby granitic orthogneiss and mingle with the mafic rocksto form a network of veins (Fig. 3F). The resultant compos-ite gneiss exhibits compositional domains of amphibolite,tonalite–granodiorite, and granite (Fig. 3C). Melanocratic,dioritic orthogneiss is also present in the Grandys Brooksection as decametre-wide, massive bodies of diorite hostinga network formed by minor, centimetre-scale, felsic irregularpatches and intrusive veins. As in the Fox Roost section(Fig. 3B), this gneiss is intruded by late-synkinematic granite.

Granitic orthogneissThis rock type forms 30–40% of the outcrop of the

Margaree – Fox Roost section. In this coastal section and in-land, in the Grandys Brook section, the granitic orthogneissintrudes into the tonalitic gneiss and contains discrete am-phibolite enclaves and (or) dykes (Figs. 3A, 3D, 3F, 3G).

The granitic orthogneiss is pale grey, medium grained(1 mm), and composed of approximately 40% quartz, 25%plagioclase (An15–25), 15% K-feldspar, and 20% brown bio-tite. Epidote, opaque minerals, apatite, zircon, and rare titan-ite form accessory phases. Discontinuous, 1–3 mm thick,biotite-rich layers alternating with granoblastic layers ofquartz, plagioclase, and K-feldspar define the gneissic folia-tion. Coarse-grained felsic veins (<5 cm thick) with diffusecontacts are folded and transposed parallel to the gneissosity.

AmphiboliteAmphibolite occurs as folded, stretched, and boudinaged,

30–60 cm thick and up to several metres long, bodies. Dueto their scale, they can not be separated as a regionalmappable unit. At Fox Roost, decametre-long sheets of am-phibolite mingled with granite can make up to 50% of thecomposite tonalitic gneiss (Fig. 3A). Most amphibolites arecentimetre- to metre-scale enclaves in the leucratic, tonaliticgneiss and granitic orthogneiss and dykes. The amphiboliteshave developed mullions during the boudinage associatedwith the solid-state D2 and D3 deformations, therefore, thecuspate-lobate shape of the enclaves can not be used to inferintrusive relationships. Complicated refolded patterns makeit difficult to assess if the amphibolites are mafic xenoliths,synplutonic enclaves, or boudinaged dykes. Locally, in lowstrain areas in Fox Roost, evidence for partial hybridizationbetween the mafic enclaves and the host granitic orthogneissis observed. That is the case of the mafic dyke in Fig. 3G.This dyke is disrupted and injected by leucocratic granite,but the enclaves at the front of the dyke are directly in con-tact with the granitic orthogneiss and have diffuse bound-aries grading into hornblende-bearing schlieren in the host.Therefore, some of the amphibolite represent mafic meltscoeval with the intrusion of the granite. Whereas in the FoxRoost section most amphibolite are intruded by the graniticorthogneiss (Fig. 3A, 3F) or represent synplutonic dykes(Fig. 3G), in the Grandys Brook section, one generation ofamphibolite dykes intrudes and postdates the graniticorthogneiss (Fig. 3D). This granitic orthogneiss in GrandysBrook intrudes hornblende-bearing, leucocratic, tonaliticgneiss with mafic, amphibolitic enclaves. This demonstratesthe presence of at least two generations of amphibolite.

The modal composition of the amphibolites varies from 80–90% green hornblende + 10–20% plagioclase (An35) to 50–60% green hornblende + 5–10% brown biotite + ~30%plagioclase (An30) + (minor) quartz. Rutile, apatite, and opaqueminerals are accessory phases in the hornblende-rich varieties,whereas titanite, apatite, and opaque minerals are accessory inthe biotite–quartz-bearing amphibolite. Epidote with allanitecores occurs locally associated with retrogression of greenhornblende. An L–S fabric, defined by recrystallizedhornblende and biotite, is present in most amphibolite bodies.

Ultramafic rocksThese rocks are common (<5% of the outcrop) in the

Margaree-Fox Roost section, but were not observed in the

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Fig. 3. (A) Contact between the granitic orthogneiss (U–Pb sample 93PV5) and the melanocratic, tonalitic orthogneiss dominated byamphibolite. The amphibolite is intruded by a network of granitic veins, which coalesce with the granitic orthogneiss (L. Hall forscale; location: Fox Roost). (B) Late-synkinematic dyke (U–Pb sample 92GD11) intrusive into tonalitic gneiss with migmatitic aspect(pen for scale; location: Fox Roost). (C) Margaree orthogneiss, folding interference pattern crosscut by late pegmatite in melanocratictonalitic gneiss. Compositional domains in the tonalitic gneiss: (1) amphibolite; (2) tonalite–granodiorite; (3) granite (location: FoxRoost). (D) Amphibolite dykes intrusive into granitic orthogneiss (U–Pb sample 94PV2). The dyke on the right has been boudinagedduring D3 and pegmatite patches have formed in the neck of the boudins; white arrow, mullion structure in the amphibolite (notebookfor scale is 12 cm wide; location: Grandys Brook). (E) Detail of the gneissosity in the hornblende-bearing leucocratic tonaliticorthogneiss, hornblende porphyroblasts in the leucocratic vein (U–Pb sample 93PV3; location: abandoned quarry, old road toMargaree). (F) Detail of the net-veining in the amphibolite. The granitic veins show sharp contacts and coalesce with the graniticorthogneiss (location: same as 3A). (G) Low-strain zone preserving a partially mingled amphibolite dyke intrusive into graniticorthogneiss. The dyke is disrupted and injected by leucocratic veins of granite. Broken enclaves at the front of the dyke are in contactwith the surrounding granitic orthogneiss (white arrow) and show partial hybridization, suggesting that the dyke is synplutonic. Left:high-strain, ductile D3 shear zone (location: Fox Roost).

Grandys Brook section. They crop out as unfoliated, 5–10 mdiameter pods composed of coarse-grained (0.3–1 cm) maficminerals, primarily amphibole. These ultramafic pods are as-sociated with the mafic-rich areas in the melanocratic,tonalitic orthogneiss. This association indicates that theymight represent igneous cumulate.

These rocks have a decussate texture defined by 70%colourless tremolite and pale-green Mg-rich clinoamphibole(cummingtonite?), and 30% partially aligned aggregates of amicaceous mineral (Ø≈ 0.2 mm), optically identified as Mg-rich chlorite. Apatite, rutile, and opaque minerals are acces-sory minerals. Burgess et al. (1992) described another vari-

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ety of ultramafic pod with actinolite–spinel–clinopyroxene,which was not observed in this study. They also reportedclinopyroxene–talc replacements.

Other lithologies

Layered gneissThis lithology is located in areas of high-strain, ductile

shearing in the coastal section around Fox Roost. Similarlayered gneiss was also identified and mapped as Margareeorthogneiss in a high-strain, ductile shear zone in the upperpart of the Grandys Brook section. This gneiss consists ofparallel, discontinuous, metre-scale layers most of which arederived by ductile stretching of tonalitic, granitic orthogneissand amphibolite. Locally, the presence of quartz-rich layersand garnet-bearing leucocratic layers suggests derivationfrom the country rock quartzo-feldspathic paragneiss.

Migmatitic gneissThis rock type is limited to the coastal section around Fox

Roost and resembles a diatexitic migmatite (sensu Mehnert1968). Field relationships and petrography (Valverde-Vaquero 1997) indicate that it comprises a set of aplitic gra-nitic dykes intruded into tonalitic rocks, both of which havebeen ductily deformed (i.e., a solid-state migmatite).

Geochemistry

Major and trace element geochemistry was gathered fromrepresentative samples of the main rock types of theMargaree orthogneiss in the area between Margaree andIsle-aux-Morts to elucidate a possible tectonic scenario. Thedata set includes five amphibolites, two ultramafic rocks,two leucocratic tonalitic orthogneiss, and two graniticorthogneiss (Table 1).

Analytical techniquesRock samples weighing between 10–20 kg were quartered

after crushing and prior to pulverization to obtain represen-tative samples. Resulting sampling error on the basis ofgrain size and variability of mineral phases is estimated at5%. Major (fused glass pellets) and trace (press pellets) ele-ments were analyzed by X-ray fluorescence (XRF) spec-trometry at Memorial University of Newfoundland.Analytical procedures, limits of detection, precision, and ac-curacy are described in Longerich (1995). Lanthanides (orrare earth elements, REE) and Th were determined by Na2O2sinter digestion, followed by solution and inductively cou-pled plasma – mass spectrometry (ICP-MS) analysis at Me-morial University. Details of the technique can be found inJenner et al. (1990). The precision and accuracy is generallybetween 3 and 10% for the REE. Limits of detection for theREE are at the sub-parts per million level. The internalcheck with Y analyzed by XRF and ICP-MS indicates thatthere were no dissolution problems with the REE.

Geochemistry

Ultramafic rocksThese have low SiO2 (45.6%) and TiO2 (~0.5%), high

MgO (~18.5%), significant Fe2O3* (~10.5%), total alkalis(Na2O + K2O = 1.62–1.31%), and Al2O3 (~15%), and theirNi (452–382 ppm) and Cr (888–846 ppm) contents are not

particularly high when compared with the amphibolite (Ta-ble 1). Ultramafic samples with similar geochemistry werealso reported by Schofield et al. (1998). The ultramaficrocks are REE poor, show a light REE (LREE) enrichment,and a small positive Eu anomaly (Fig. 4). Mid-ocean ridgebasalt (MORB; Pearce 1983) normalized, multielement pat-terns show a general depletion in high field strength ele-ments (HFSE), a relative enrichment in large ion lithophileelements (LILE), and the presence of a small negative Nbanomaly (Fig. 4). This geochemical signature could be com-patible with the primary mineralogy of a clinopyroxene–plagioclase cumulate (Green 1994) and is consistent with theclose field association of the ultramafic rocks with the am-phibolite-rich areas at Fox Roost. Therefore, the ultramaficrocks could represent a cumulate phase that separated fromthe mafic magma that was responsible for the amphibolite.

AmphiboliteThese rocks have basaltic compositions (47.7–50.2%

SiO2), variable Al2O3 (17.6–14%), MgO (11.2–5.4%), TiO2(2–0.8%), Ni (283–24 ppm), and Cr (772–62 ppm) concen-trations. The Ni and Cr contents show a positive correlationwith MgO. A negative correlation is observed between MgOand the LREE, with the highest MgO concentrations corre-sponding with the LREE depleted patterns (Fig. 4). The bio-tite ± quartz-bearing amphibolites show LREE enrichedpatterns (Fig. 4), which are similar to those reported bySchofield et al. (1998). The HFSE (i.e., Zr, Nb) increasewith decreasing MgO. The MORB (Pearce 1983) normal-ized, multielement patterns show an enrichment in LILE, anegative Nb anomaly, an overall depletion in HFSE, and apoorly defined Ti negative anomaly. These trace elementcharacteristics are shared with the amphibolite with flat-REEpatterns reported by Whalen et al. (1997). Secondary ele-ment mobility is probably responsible for the relative varia-tions of K2O, Rb, and Ba among samples. However, theincrease in Sr, Th, and HFSE contents with decreasing MgOobserved in the least evolved samples suggests that the rela-tive enrichment in LILE and Th with respect to MORB is aprimary feature.

The major element contents of the amphibolite, althoughpotentially modified by metamorphic processes, are indica-tive of a magma of basaltic composition with tholeiitic affin-ities, as pointed out by Schofield et al. (1998). With theexception of sample MA-8 (LREE enriched), all samplesshow LREE-poor to flat-REE patterns and MORB-normalized multielement patterns indicative of volcanic-arctholeiite: LILE enrichment, HFSE and REE depletion withrespect to MORB and negative Nb anomaly (Fig. 4 ; Pearce1983). The high MgO, Ni, and Cr concentrations of theLREE-poor amphibolite suggest that these are primitive orlittle evolved samples. Ternary tectonic discrimination dia-grams based on immobile elements (Fig. 5) also confirm thetholeiitic character of these samples and show a shift withdecreasing MgO towards the field of volcanic-arc basalts.Schofield et al. (1998) reported LREE-enriched amphibolitelike sample MA-8. The geochemical signature of this sampleresembles a volcanic-arc basalt. However, its low Mg num-ber (24) indicates that this is an evolved sample, and there-fore, this geochemical signature might not necessarily reflectthat of its mantle source.

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Leucocratic tonalitic and granitic orthogneissThe hornblende-bearing, leucocratic, tonalitic orthogneiss

has an intermediate composition (62.2–64.8% SiO2),whereas the granitic orthogneiss ranges from 71.1 to 74.1%SiO2. Although the sample set is limited, there is a remark-able trend of decreasing TiO2, Al2O3, Fe2O3*, MgO, MnO,CaO, and P2O5 with increasing SiO2. Only the alkali ele-ments do not seem to show a systematic variation with de-creasing SiO2. The REE patterns (Fig. 6) are enriched inLREE and show a minor Eu anomaly. The tonaliticorthogneiss shows an enrichment in REE with respect to the

granitic orthogneiss and has a better defined Eu anomalyand flatter HREE. Primitive mantle-normalized (Sun 1980),multielement patterns (Fig. 6) are characterized by enrichedLILE and negative Nb, P, and Ti anomalies, with thetonalitic orthogneiss showing an enrichment over the gra-nitic orthogneiss. Whalen et al. (1997) reported a tonalite(65.79% SiO2) with the same trace element characteristics asthe tonalite reported here and anεNd(0.47Ga) value of +1.1.

Despite the wide range of SiO2 contents and the REEvariations, the geochemical signatures of the tonalite and thegranite are strikingly similar. When classified, according to

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Table 1. Major and trace element analysis, Margaree orthogneiss.

the schemes of Maniar and Piccoli (1989) and Pearce et al.(1984), they correspond to volcanic-arc granitoid (VAG;Fig. 6).

Petrogenetic processes and tectonic signaturesThe Mg-rich amphibolite have a geochemical signature

typical of tholeiitic basaltic magma. The high MgO, Ni, Cr,and low SiO2 concentrations and the depletion of the HFSEwith respect to the MORB suggest that there was no signifi-cant crustal contamination (Pearce 1983). It is feasible thatthe tholeiitic LREE-depleted patterns are related byplagioclase–olivine–clinopyroxene fractionation. But, the ab-sence of enhancement of the Ti anomaly with increasingLREE suggests that magnetite or amphibole fractionation,which otherwise could explain the small Nb anomaly (e.g.,Pearce and Parkinson 1993), did not take place. This and thelack of evidence of crustal contamination argues against theinterpretation of Schofield et al. (1998) that the Nb negativeanomaly in the MgO-rich amphibolite of the Margareeorthogneiss was due to crystal fractionation and crustal con-tamination. Instead, we interpret that the geochemical signa-tures of the MgO-rich amphibolite reflect the mantle sourceof these volcanic-arc tholeiites. The LREE-depleted patternsare typical of primary magma produced by a high degree ofpartial melting from a depleted-mantle source (e.g., Pearce1983). This can be explained by shallow melting of a spinel–lherzolite mantle source (asthenosphere?) slightly affectedby a subduction zone component (Wilson 1989; Ellam 1992;Pearce and Parkinson 1993).

The trace element signature of the tonalitic orthogneiss issimilar to that of a volcanic-arc andesite (Gill 1981). The

εNd(0.47Ga) value of +1.1 reported by Whalen et al. (1997)suggests that the tonalites probably represent hybrid melts.The higher contents of trace elements in the tonaliticorthogneiss, with respect to the granitic orthogneiss (Fig. 6),indicate that the tonalitic magmas are not the product ofsimple binary mixing between granitic and basaltic endmembers in the orthogneiss. Such trace element depletion inthe granite could be attributed to crystal fractionation fromthe tonalitic melt. However, the geochemical signature ofthe granite could also be the result of pure crustal anatexisof an immature crustal source (Pearce et al. 1984) or hybrid-ization with tholeiitic basalt (see Patiño Douce 1995)

Geochronology

U–Pb geochronology was carried out to obtain crystalliza-tion ages of most units of the Margaree orthogneiss and toconstrain their tectonothermal overprint. The leucocratic,tonalitic orthogneiss was sampled in a quarry outsideMargaree (sample 93-PV-3; Fig. 3E); it represents the oldestmember of the Margaree orthogneiss and contains mafic en-claves and strained mafic dykes. The granitic orthogneisswas sampled along the coast (at Fox Roost; sample 93-PV-5), where it is the youngest member of the Margareeorthogneiss and intrudes the mafic members of themelanocratic, tonalitic orthogneiss associated with the ultra-mafic rocks (Fig. 3A, 3F). The amphibolite with tholeiiticaffinities and LREE-depleted signatures described in the pre-vious section were also sampled at this locality, where theyare intruded by and mingled with this granitic orthogneiss(Fig. 3A, 3F). A granitic orthogneiss was sampled inland atGrandys Brook (sample 94-PV-2; Fig. 3D), where it is alsothe youngest member of the Margaree orthogneiss. This gra-nitic orthogneiss intruded the leucocratic, tonaliticorthogneiss with mafic enclaves and the melanocratic,tonalitic–dioritic orthogneiss. This sample (94-PV-2) iscrosscut by a later generation of mafic dykes. The maficmembers of the Margaree orthogneiss (sample 93-PV-6)were sampled at Fox Roost. The timing of the tectono-thermal overprint of the Margaree orthogneiss is constrainedby a late-synkinematic, granitic dyke sampled at Fox Roost(sample 92GD11; Fig. 3B). This dyke provides the youngestage limit for the early stages of the D2 deformation. Further-more, metamorphic titanite (Tc < 650°C; Heaman andParrish 1991) from these rocks will provide a younger agelimit for the regional upper amphibolite facies metamor-phism and deformation.

Analytical techniquesU–Pb (zircon and titanite) dating was carried out in the

isotopic laboratory of Memorial University of Newfound-land (Canada) following procedures modified after Krogh(1973; zircon), Manhès et al. (1979; titanite), and Krogh(1982). Isotopic analysis was done using a MAT262 multi-collector mass spectrometer equipped with a SEM (second-ary electron multiplier) ion-counting system. Procedureblanks for the zircon and titanite analyses were less than 10and 20 pg Pb, respectively. The U procedure blanks wereless than 1 pg. Regression lines were calculated using themethod of Davis (1982). The uncertainties in the isotopic

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Fig. 4. Margaree orthogneiss, amphibolites, and ultramafic rocks:Chondrite (Nakamura 1974) normalized rare earth element (REE) pat-terns and MORB (Pearce 1983) normalized multielement patterns.

ages are reported at the 95% confidence level. Further de-tails of the procedure can be found in Dubé et al. (1996).

U–Pb results

Sample 93-PV-3: tonalitic orthogneissFour zircon fractions define a collinear discordance pat-

tern (Table 2; Fig. 7A). Fraction Z1, subconcordant, corre-sponds to air-abraded, euhedral prisms. Fractions Z2 and Z3of euhedral, needle-like, crack-free zircon, although air-abraded, are more disturbed by Pb-loss than fraction Z1, dueto their larger surface/volume ratio. Fraction Z4 of cracked,needle-like zircon was not air abraded. The resultingdiscordia line has a 90% probability of fit (see Davis 1982)with an upper intercept of474 5

14−+ Ma and a lower intercept

of 142 165161

−+ Ma. The large positive error of the upper intercept

is due to the shallow angle of intersection between thediscordia line and the concordia curve (Davis 1982). Thisupper intercept at474 5

14−+ Ma is the crystallization age of the

rock.

Sample 93-PV-5: granitic orthogneiss, Fox RoostFive fractions of euhedral zircon prisms with a

width/length ratio of 1:5–1:7 were selected from this rockfor analysis (Table 2; Fig. 7B). Fractions Z1, Z4, and Z5 de-fine a discordia line (12.2% probability of fit) with an upperintercept of473 6

16−+ Ma and a lower intercept of 137 Ma with

a large associated error. Fractions Z2 and Z3 plot to the rightof the discordia line due to minor inherited zircon coupledwith secondary Pb-loss. Fraction Z4 plots slightly below thediscordia line and is responsible for the low probability of fitof the discordia line. Fraction Z1 is concordant and providesa crystallization age of 472 ± 2.5 Ma.

Sample 94-PV-2: granitic orthogneiss, Grandys BrookThree fractions of needle-like zircon crystals were ana-

lyzed from this rock (Table 2; Fig. 7C). Fraction Z1, Z2, andZ3 are collinear and define a discordia line (29% probabilityof fit) with an upper intercept of464.5−

+2 718 7.. Ma and a nega-

tive lower intercept at –221 Ma with a large associated error.If the lower intercept is fitted at 1 ± 50 Ma, theupper inter-cept has an age of466.52.5

+4 Ma. Fraction Z1 is sub-concordant, plotting slightly above concordia. The U–Pbages from this fraction provide an age of 465 ± 3 Ma. Thisage of 465 ± 3 Ma can beconsidered the best estimate of thecrystallization age, and it coincides within error with the207Pb/206Pb ages from fractions Z2 and Z3, as well as withthe upper intercept of the discordia line.

Inclusion-free, clear, brown titanite (T1 and T2; Table 2)was analyzed and provides a duplicated concordant age of411 ± 2 Ma(Fig. 7C). This age reflects closure of the U–Pbsystem after the upper amphibolite facies metamorphic over-print.

Sample 93-PV-6: quartz-bearing amphibolite dyke,Margaree – Fox Roost

This dyke, intrusive into the leucocratic tonalitic ortho-gneisses, has green hornblende, plagioclase (oligoclase), andbrown biotite. Quartz, opaque minerals, and titanite are inminor proportions. Apatite and epidote are the main acces-sory minerals. All these minerals, except the apatite, arerecrystallized, and the amphibole minerals define a minerallineation parallel to the plunge of the minor F3 folds.

Only titanite was recovered from this sample, and onefraction of brown crystals provided a concordant age of410 ± 2 Ma (Table 2; Fig. 7D). This is a metamorphic agethat provides the youngest age limit for the amphibolitedykes.

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Fig. 5. Ternary tectonic discrimination diagrams for the amphibolites (circles; enriched sample, square) of the Margaree orthogneiss.TiO2–MnO–P2O5 diagram (Mullen 1983). CAB, calc-alkaline basalt; IAT, island-arc tholeiite; MORB, mid-ocean ridge basalt; OIT,ocean-island tholeiite; OIA, ocean-island alkaline basalt. Ti–Zr–Y diagram (Pearce and Cann 1973): WPB, within-plate (alkaline) ba-salt; LKT, low-K tholeiite; OFB, ocean-floor basalt; CAB, calc-alkaline basalt. Nb–Zr–Y diagram (Meschede 1986): WPAB, within-plate alkaline basalt; WPTB, within-plate tholeiitic basalt; E-MORB, enriched-MORB; N-MORB, normal-MORB; VAB, volcanic-arcbasalt. Field of modern back-arc basalts from the southwest Pacific Eastern Lau spreading centre (Lau Basin; Pearce et al. 1995) andNgatoro Basin (Gamble et al. 1995) plotted for comparison.

Sample 92GD11: late-synkinematic granitic dyke, FoxRoost

This is a grey, fine- to medium-grained, equigranular, bio-tite-bearing, granitic dyke. This dyke crosscuts the steepgneissosity in the host country rock and is gently folded(Fig. 3B). The gneissosity is folded by tight, upright, F3folds (Fig. 3C). But, since the dyke does not cut through

both limbs of the F3 folds, a late syn-D2 or D3 time of intru-sion cannot be ruled out.

Six zircon and one titanite fraction have been analyzed(Table 2; Fig. 7E). The titanite fraction (T1) is clear andvery pale, and is 8% discordant due to the effects of thecommon Pb correction used. U–Pb ages of 407 Ma and412 Ma for this titanite are, however, in good agreement and

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Fig. 6. Margaree orthogneiss, tonalitic and granitic orthogneiss. Chondrite-normalized REE and primitive mantle (Sun 1980) normal-ized multielement patterns. Tectonic discrimination diagrams: Rb vs. Y + Nb diagram (Pearce et al. 1984). COLG, collisionalgranitoid; VAG, volcanic-arc granitoid; WPG, within-plate granitoid; ORG, ocean-ridge granite. A/NK (Al2O3/(Na2O + K2O) vs.A/CNK (Al 2O3/(CaO + Na2O + K2O) diagram (Maniar and Piccoli 1989); IAG, island-arc granite; CAG, continental-arc granite.

©2

00

0N

RC

Ca

na

da

Valverde-Vaqueroet

al.1703

Table 2. U–Pb data, Margaree orthogneiss.

are consistent with the titanite age from the nearby amphibo-lites, previously presented.

Most zircon grains extracted from this rock are heavilycracked, even the best quality prisms and needles (fractionsZ1, Z2, Z3, Z4, and Z5) show cracks. All fractions werestrongly abraded, except fraction Z5 (unabraded). FractionZ6, comprising the largest zircon prisms and prism frag-ments, has the oldest207Pb/206Pb age and is the most discor-dant fraction (5% discordancy; Fig. 5E). This suggests thepresence of inherited zircon. Fractions Z1, Z2, Z3, Z4, andZ5 are closely clustered defining a discordia line with an up-per intercept of 417.3 Ma and a lower intercept of 18.9 Ma.If the lower intercept is fitted at 1 ± 50 Ma (Davis1982), the

resultant discordia line has a 64% probability of fit and anupper intercept age of417 4

7−+ Ma (Fig. 7E). This upper inter-

cept is in agreement with the207Pb/206Pb ages of the five in-dividual zircon fractions (Table 2). Therefore, this age of417 4

7−+ Ma is taken as the best estimate of the intrusion age of

the dyke.

Discussion

The Margaree orthogneiss: an Early Ordovician mafic–felsic intrusive complex

The U–Pb data indicates that the 474 Ma tonaliticorthogneiss is coeval within error with the 472 Ma granitic

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Fig. 7. U–Pb concordia diagrams for the samples from the Margaree orthogneiss. Z, zircon; T, titanite: (A) tonalitic orthogneiss (93-PV-3), Margaree; (B) granitic orthogneiss (93-PV-5), Fox Roost; (C) granitic orthogneiss (94-PV-2), Grandys Brook; zircon; (D) gra-nitic orthogneiss (94-PV-2), Grandys Brook titanite; (E) amphibolite (93-PV-6), Fox Roost; (F) late-synkinematic granitic dyke(92GD11), Fox Roost.

orthogneiss (Fig. 8). The mafic members of the Margareeorthogneiss could not be directly dated. However, contem-poraneous intrusion of felsic and mafic magma at 474–472 Ma is supported, at least locally, by the following fieldrelationships between the amphibolite and the 472 Ma gra-nitic orthogneiss at Fox Roost: (1) mafic dykes (now amphi-bolite) disrupted and dispersed as mafic enclaves in the472 Ma granitic orthogneiss (Fig. 3G); and (2) net-veiningof the amphibolite and mafic tonalite–diorite by the 472 Magranite (Figs. 3A, 3F). The widespread net-veining of themafic members of the gneiss by the 472 Ma granite is one ofthe most striking features of the Margaree orthogneiss atFox Roost. This net-veining could be interpreted as perva-sive injections of granite into mafic xenoliths. This, how-ever, seems unlikely given the fact that there are noxenoliths derived from the country-rock paragneiss, not evenrestitic ones, in the granite. Also none of this net-veining isobserved at the contact between the granite and the meta-sedimentary country rocks. The alternative is that the net-veining represents deformed, 472 Ma granite mingled withsynplutonic mafic pillows produced by a coeval injection ofmafic magma. Evidence for coeval mafic magmatism is pro-vided by the dyke in Fig. 3G. Although the surroundingleucocratic material could be attributed to back-veining, thepartially hybridized mafic enclaves in contact with the472 Ma granite indicate that this is a synplutonic dyke.These features, including the presence of ultramaficcumulates associated with the mafic-rich areas, are com-monly displayed in ultramafic–mafic–silicic intrusive com-plexes (Sutcliffe et al. 1990; Sha 1995; Fernandez andGasquet 1994) or during coeval intrusion of mafic magmainto felsic magma chambers (Fernandez and Barbarin 1991;Wiebe 1991; Bateman 1995). This interpretation implies thatthe mafic protoliths of the amphibolite with volcanic-arctholeiitic signatures, described in the geochemical section,were coeval with the 472 Ma granite that intrudes them. Ad-ditionally, the presence of abundant mafic enclaves in thetonalitic rocks and the absence of metasedimentary xenolithscould be interpreted to indicate that basic magma intruded atthe same time as the 474 Ma tonalitic magma.

Mafic magmatism in the area, however, continued afterthe intrusion of the 474–472 Ma tonalite and granite, asdemonstrated by the intrusion of the 465 ± 3 Ma granite atGrandys Brook. This granite provides a younger age limitfor the tonalite and most of the mafic members of theMargaree orthogneiss (Fig. 8) in this section and is locallyintruded by mafic dykes of unknown geochemical signature.This indicates that mafic magmatism in the Port-aux-Basques Complex continued after 465 Ma and before the on-set of deformation.

Tectonothermal overprintThe intrusion of the late-synkinematic granitic dyke into

the Margaree orthogneiss at417 47

−+ Ma was either coeval with

the late stages of the D2 deformation or the regional D3 de-formation. The U–Pb titanite ages of 411 ± 2 and 410 ±2 Ma could be interpreted as cooling ages due to opening ofthe isotopic system by thermally induced volume diffusion(Tc for titanite≈ 650–550°C; Heaman and Parrish 1991). Al-ternatively, these ages might represent opening and closureof the isotopic system as a result of deformation-induced,

dynamic recrystallization (see Gromet 1991) during the am-phibolite-facies D3 deformation (Burgess et al. 1995;Valverde-Vaquero 1997). Either of these two interpretationssuggest that the titanite ages possibly date the final stages ofthe D3 deformation and predate the local, D4, greenschist-facies overprint of the Margaree orthogneiss (Fig. 8).

These deformational, intrusive, and metamorphic eventsare not exclusive to the Margaree orthogneiss, but are part ofa large-scale Siluro-Devonian tectonic event that affected theentire Hermitage Flexure (Dunning et al. 1990). The intru-sion of the late-synkinematic granite (417 4

7−+ Ma) is coeval

within uncertainty with the intrusion of the late syn-D2 RoseBlanche granite (Dunning and van Staal, unpublished data)into the upper amphibolite facies Harbour le Cou Group andthe intrusion of the late-kinematic La Poile batholith (416 ±4 Ma; Chorlton and Dallmeyer 1986) into the nearby,greenschist-facies (Lin et al. 1993) Bay du Nord Group(Fig. 2). This dyke is also coeval within error with monaziteages (420–414 Ma), which, according to van Staal et al.(1994), date peak metamorphism and D2 deformation in thePort-aux-Basques and Grand Bay complexes, and the Har-bour le Cou Group. The titanite ages (411–410 Ma) are co-eval within uncertainties with a titanite age in the Grand BayComplex (Fig. 2; 412 ± 2 Ma; Dunning et al. 1990) andslightly younger than (but overlapping with) a titanite age of418 ± 9 Ma in theHarbour le Cou Group (Fig. 2), obtainedby Burgess et al. (1995). This indicates that the medium-and high-grade rocks of the Grand Bay Complex, Harbour leCou Group, and Port-aux-Basques Complex, including theMargaree orthogneiss, were overprinted by the same Siluro-Devonian regional deformation and metamorphism, as a re-sult of the final suturing of the peri-Laurentian and peri-Gondwanan margins of the Iapetus Ocean in southwesternNewfoundland (Dunning et al. 1990; Lin et al. 1994; Dubéet al. 1996).

Tectonic significance of the Early Ordovician Margareeorthogneiss

The data from the Margaree orthogneiss indicates that thisis a 474–465 Ma mafic–felsic igneous complex dominatedby tonalitic rocks with abundant mafic members and subor-dinate granite. The preponderance of tonalites with volcanic-arc-like geochemical signatures could be attributed to a vol-canic-arc origin. The tholeiitic, volcanic-arc chemistry of theamphibolites that are mingled with the 472 Ma granite is in-dicative of shallow melting of asthenospheric material in asuprasubduction-zone environment. This requires astheno-spheric upwelling and a relatively large degree of partialmelting. A transitional volcanic-arc – back-arc setting on at-tenuated continental crust, such as the Taupo volcanic zonein northern New Zeland (Gamble et al. 1995), would be con-sistent with these requirements.

Regional significanceThe late Arenig – early Llanvirn intrusion of the Margaree

orthogneiss is contemporaneous with the main back-arc rift-ing event along the peri-Gondwanan margin of the CanadianAppalachians (Swinden et al. 1990; van Staal et al. 1991).The oldest tonalite (474 4

14−+ Ma) and granite (472 ± 2.5 Ma)

of the Margaree orthogneiss are contemporaneous within er-ror with the intrusion of Baggs Hill granite (477 ± 1 Ma;

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Valverde-Vaquero et al. 1705

Tucker et al. 1994) in the nearby Bay du Nord Group(Fig. 1B) and the intrusion of the Partridgeberry Hills gran-ite (474 4

6−+ Ma; Colman-Sadd et al. 1992) in central New-

foundland (Fig. 9). The Baggs Hill granite postdates theimbrication of "ophiolitic" gabbro with the oldest volcano-sedimentary members of the Bay du Nord Group in southernNewfoundland (Tucker et al. 1994); while the PartridgeberryHills granite stitches the contact between the allochthonousPipestone Complex (Exploits subzone Tremadocian ophiolite)

and the autochthonous Mount Cormack subzone of the Gan-der Zone in central Newfoundland (Colman-Sadd et al.1992). Therefore, the intrusion of the Margaree orthogneisspostdates the Tremadoc – early Arenig Penobscotian amal-gamation of the Early Ordovician Exploits subzone (arc)with the Gander Zone (passive margin; cf. Colman-Sadd etal. 1992; van Staal et al. 1996c).

The intrusion of the youngest granite (465 ± 3 Ma) of theMargaree orthogneiss is coeval with volcanism in the com-

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Fig. 8. Proposed sequence and timing of events recorded in the Margaree orthogneiss. Time scale after Tucker and McKerrow (1995)and Tucker et al. (1998).

posite Bay du Nord (466 ± 3; Dunning et al. 1990) and Bayd’Espoir groups (468 ± 2 Ma; Colman-Sadd et al. 1992) ofthe Exploits subzone in the central and eastern HermitageFlexure (Fig. 9). This also coincides with low pressure –high temperature anatexis at 465 ± 2 Ma andintrusion of464 Ma anatectic granites in the Meelpaeg and MountCormack subzones of the Gander Zone in central Newfound-land (Fig. 9; Colman-Sadd et al. 1992). This metamorphismin the autochthon of Exploits subzone indicates the presenceof a high geothermal gradient under the amalgamated arc –passive margin, as would be expected in a continental back-arc environment.

Early Ordovician events are also well documented in theExploits subzone in northern Newfoundland (Fig. 9), wherethe outboard Tremadoc – early Arenig Tea Arm arc was jux-taposed along the peri-Gondwanan margin of the IapetusOcean in the Arenig (O’Brien et al. 1997; MacLachlan andDunning 1998a). Volcanic-arc magmatism at ca. 472 Maabove the amalgamated Tremadocian arc (MacLachlan andDunning 1998b) suggests an arc polarity reversal and east-ward-dipping subduction under the peri-Gondwanan marginof the Newfoundland Appalachians as a result of the late

Tremadoc – early Arenig (Penobscottian) amalgamation.This is followed shortly after by ca. 470 Ma eruption of en-riched tholeiite and alkalic basalt with no subduction zonecomponent and interpreted to represent arc rifting and for-mation of back-arc basins (O’Brien et al. 1997; MacLachlanand Dunning 1998b). According to O’Brien et al. (1997), arcmagmatism in the northern Exploits subzone was renewed inthe Llanvirn. Van Staal (1994) described a similar tectonicscenario based on rocks of the Exploits subzone and GanderZone in New Brunswick: arc–passive margin amalgamation,subduction polarity reversal, slab roll-back, Arenig–Llanvirnback-arc rifting, and mafic magmatism up into the Caradoc(Sullivan and van Staal 1995). Therefore, the data along thenorthern Appalachians in the late Arenig indicate east-dipping subduction under the peri-Gondwanan margin of thenorthern Appalachians (Fig. 1A) and associated back-arcrifting.

The Bay du Nord and Bay d’Espoir groups in southernNewfoundland are time equivalents of the Exploits Group innorthern Newfoundland. Hence, the back-arc rifting volca-nism in southern Newfoundland may be represented by theHarbour le Cou Group (Schofield et al. 1998), the ca.

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Fig. 9. Arenig to early Llanvirn magmatism along the peri-Gondwanan margin of the Newfoundland Appalachians (Exploits subzoneand Gander Zone) and interpretative tectonic setting for the intrusion of the Margaree orthogneiss. Subduction polarity based on vanStaal (1994).

466 Ma volcanosedimentary sequences of the Bay du NordGroup (Fig. 9; Dunning et al. 1990; Tucker et al. 1994), andthe 468 ± 2 Mavolcanic rocks of the Bay d’Espoir Group(Fig. 9; Colman-Sadd et al. 1992) in the western HermitageFlexure. The western position of the Margaree orthogneisswith respect to the Harbour Le Cou and Bay du Nord groups(Fig. 1B) suggests that this igneous complex occupied amore external, outboard position, in relation to these ele-ments of the peri-Gondwanan margin of the Iapetus Ocean.Oceanward migration of the volcanic-arc front as a result ofthe widening of the back-arc area during slab roll-back is anattractive explanation for the late Arenig – early Llanvirntransitional-arc – back-arc setting proposed herein (Fig. 9)for the Margaree orthogneiss.

Conclusions

(1) The Margaree orthogneiss represents a 20 km longEarly Ordovician mafic–felsic intrusive complex withinthe Port-aux-Basques Complex. This composite gneissis formed by amphibolites;474 4

14−+ Ma, hornblende-

bearing, tonalitic orthogneiss with mafic enclaves;472 ± 2.5 Ma and 465 ± 3 Magranitic orthogneiss; lo-cal melanocratic tonalitic–dioritic orthogneiss; and mi-nor ultramafic rocks. Despite the Silurian upperamphibolite-facies overprint, field relationships suggestthat 472 Ma granite was coeval with widespread maficmagma. Locally, a second generation of mafic dykes in-truded the 465 Ma granite.

(2) Most amphibolite have basaltic compositions and traceelement signatures of volcanic-arc tholeiite. The474 Ma tonalite and the 472 Ma granite have similartrace element geochemical signatures that correspond tovolcanic-arc granitoids.

(3) Timing of the regional tectonometamorphic overprint isconstrained by a417 4

7−+ Ma late-synkinematic, granitic

dyke and 411 ± 2 and 410 ± 2 Matitanite (Early Devo-nian) U–Pb ages. The granitic dyke dates the late stagesof the high-temperature, regional, D2 and D3 deforma-tion in the Margaree orthogneiss. The titanite ages areinterpreted to date the amphibolite faciesrecrystallization during the late stages of the D3 defor-mation. This Siluro-Devonian overprint is associatedwith the final closure of the Iapetus Ocean during theLate Ordovician – Silurian.

(4) The intrusion of the igneous protoliths of the Margareeorthogneiss at 474–465 Ma coincided within error withthe plutons that postdate the Penobscotian ophioliteobduction, arc – passive margin imbrication, and amal-gamation of the Exploits subzone and the Gander Zone(477 ± 1 Ma Baggs Hills granite, Tucker et al. 1994;474 4

6−+ Ma Partridgeberry Hills granite, Colman-Sadd et

al. 1992), local anatexis and granite intrusion at ca.465 Ma in central Newfoundland, and extensiveArenig – early Llanvirn back-arc volcanism along theeastern margin of the Iapetus Ocean in the northern Ap-palachians. It is suggested that the Margaree ortho-gneiss occupied an oceanward arc – back-arctransitional setting with respect to the major Arenig –early Llanvirn back-arc rifts along the peri-Gondwananmargin of the Newfoundland Appalachians.

Acknowledgments

These data form part of the Ph.D. work of P.V.-V. at Me-morial University of Newfoundland, which was primarily fi-nanced by a Natural Sciences and Engineering ResearchCouncil of Canada grant to G.R. Dunning. Fieldwork wascarried out within a Geological Survey of Canada mappingproject part of the Canada–Newfoundland co-operationagreement on mineral development, 1990–1995. We grate-fully acknowledge Shoufa Lin, Lindsay Hall, and DaveSchofield for field assistance and discussion. P.V.-V. alsothanks S.J. O’Brien (Newfoundland Department of Mines)and A. Indares (Memorial University) for general discus-sions. Wouter Bleeker (Geological Survey of Canada) isthanked for reviewing an earlier version of the manuscript.Constructive reviews by Dave Schofield and an anonymousreferee are appreciated.

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