Archaean cratonization and deformation in the northern …€¦ · the batholithic roots of Andean-type plate margins and intra-oceanic arcs. Existing horizontal-tectonic models propose

Precambrian Research 127 (2003) 61–87

Archaean cratonization and deformation in the northernSuperior Province, Canada: an evaluation of plate

tectonic versus vertical tectonic models

Jean H. Bédarda,∗, Pierre Brouillettea, Louis Madoreb, Alain Berclazca Geological Survey of Canada, Division Québec, 880, ch.Ste-Foy, Quebec City, Que., Canada G1S 2L2

b Géologie Québec, Ministère des ressources naturelles du Québec, 5700, 4e Avenue Ouest, Charlesbourg, Que., Canada G1H 6R1c Géologie Québec, Ministère des ressources naturelles du Québec, 545 Crémazie Est, bureau 1110, Montreal, Que., Canada H2M 2V1

Accepted 10 April 2003

Abstract

The Archaean Minto Block, northeastern Superior Province, is dominated by tonalite–trondhjemite, enderbite (pyroxenetonalite), granodiorite and granite, with subordinate mafic rocks and supracrustal belts. The plutons have been interpreted asthe batholithic roots of Andean-type plate margins and intra-oceanic arcs. Existing horizontal-tectonic models propose thatpenetrative recrystallization and transposition of older fabrics during terrane assembly at∼2.77 and∼2.69 Ga produced aN-NW tectonic grain. In the Douglas Harbour domain (northeastern Minto Block), tonalite and trondhjemite dominate theFaribault–Thury complex (2.87–2.73 Ga), and enderbite constitutes 50–100 km-scale ovoid massifs (Troie and Qimussinguatcomplexes, 2.74–2.73 Ga). Magmatic muscovite and epidote in tonalite–trondhjemite have corroded edges against quartz+plagioclase, suggesting resorption during ascent of crystal-charged magma. Foliation maps and air photo interpretation show thecommon development of 2–10 km-scale ovoid structures throughout the Douglas Harbour domain. Outcrop and thin-section scalestructures imply that many plutons experienced a phase of syn-magmatic deformation, typically followed by high temperaturesub-magmatic overprints. Thermobarometric data for plutons indicate near-solidus recrystallization at 4–6 kbar pressures. Thecommon preservation of syn-magmatic fabrics in plutons of different ages seems incompatible with the origin of these fabricsthrough superimposed regional orogenesis. The broad uniformity of intrusion ages and lithologies throughout the Minto Block,and the rarity of shallowly-dipping planar fabrics, also seem inconsistent with accretion of disparate older terranes, each of whichshould preserve distinct histories. A possible alternative explanation for these features is provided by vertical tectonic models,whereby buoyant felsic magmas ascended as crystal slurries, while dense supracrustal rocks (and solidified felsic intrusionsemplaced into them) subsided as cold fingers (10–20 km-scale instabilities). Shear between upwelling and downwelling limbswould have concentrated in the weak intrusions, generating steeply-plunging syn-magmatic fabrics, and producing ductileoverprints in solidified rocks.© 2003 Elsevier B.V. All rights reserved.

Keywords: Archaean; Orogeny; Vertical tectonics; Minto Block; Superior Province; Cratonization

∗ Corresponding author. Tel.:+1-418-654-2671; fax:+1-418-654-2615.E-mail addresses: [email protected] (J.H. Bedard), [email protected] (L. Madore), [email protected]

(A. Berclaz).

0301-9268/$ – see front matter © 2003 Elsevier B.V. All rights reserved.doi:10.1016/S0301-9268(03)00181-5

62 J.H. Bedard et al. / Precambrian Research 127 (2003) 61–87

1. Introduction and regional framework

Most recent Archaean crustal growth models in-voke a near-uniformitarian process of subduction,arc maturation, and lateral accretion of oceanic arcsand plateaux (e.g.Condie, 1986; Card, 1990; de Wit,1998; Percival et al., 2001). However, recent workin a number of cratons has favored vertical tectonicmodels (e.g.Chardon et al., 1996; Collins et al.,1998), leading to a revival of this classic debate. TheMinto Block (Fig. 1) is the largest plutonic-dominatedterrane of the Superior Craton, and is an idealplace to investigate how Archaean crust formed andstabilized.

The Minto Block was poorly known prior to pio-neering work byPercival et al. (1992, 1994, 2001),who subdivided it into lithotectonic domains (Fig. 1),and developed tectonic models. The oldest MintoBlock rocks are mafic and ultramafic lavas and felsictuffs (3.825 Ga) from the Porpoise Cove greenstonebelt, embedded in younger (2.75 Ga) tonalitic rocksof the Inukjuak domain (David et al., 2002). Thenext oldest Minto Block rocks are tonalite and trond-hjemite with embedded supracrustal belts from theGoudalie and Douglas Harbour domains (3–2.8 Ga:Stern et al., 1994; Madore et al., 1999; Percival and

Fig. 1. Simplified map of Minto Block adapted fromPercivalet al. (1997). PN = Pelican Nantais belt, Ko= Kogaluc belt,Vz = Vizien belt.

Skulski, 2000; Percival et al., 2001), which were inter-preted to represent cratonic nuclei. Juvenile isotopicsignatures of Qalluviartuuq belt lavas (2.84–2.83 Ga)were considered evidence for an intra-oceanic arcsetting (Skulski et al., 1996). A 2.81 Ga shear zonein the Qalluviartuuq belt was interpreted as anintra-oceanic accretionary thrust (D1:Percival andSkulski, 2000). However, analogous D1 structuresare younger in the Vizien belt (<2.718 Ga), and soD1 structures cannot record a unique tectonic event(Percival and Skulski, 2000). Percival et al. (2001)proposed that a composite cratonic basement wasformed when oceanic and continental terranes dockedat about 2.77 Ga (Fig. 2), with crustally contaminated

Fig. 2. Cartoon illustrating existing plate-tectonic scenarios(Percival et al., 2001) for origin and assembly of Minto Blockdomains.

J.H. Bedard et al. / Precambrian Research 127 (2003) 61–87 63

calc-alkaline volcanics in the Kogaluc belt, Lac Mintodomain, representing development of a successorarc at 2.77–2.76 Ga (Skulski et al., 1996). Younger(<2.748 Ga), unconformable greywacke and iron for-mations that contain ancient detritus (Percival et al.,1995), and quartzites and ultramafic lavas in the Farib-ault belt (Douglas Harbour domain), were interpretedas a continental overlap sequence (Percival et al.,1997), linked to intra-arc extension bySkulski et al.(1994).

This older composite basement was intruded byvoluminous granodiorite and granite (subordinatetonalite, enderbite, pyroxene tonalite and mafic intru-sions) of the Leaf River suite (2.73–2.72 Ga:Percivalet al., 1994; Stern et al., 1994), which constitutesmost of the Lac Minto and Utsalik domains. Relictzircon cores and Nd-isotopic signatures indicate recy-cling of older continental crust, and so the Leaf Riversuite was interpreted to be the plutonic root zonesof Andean-type continental arcs (Stern et al., 1994;Percival et al., 1994). The great areal extent of syn-chronous magmatism and heterogeneous Nd-isotopicsignatures were explained in terms of two simultane-ously active subduction zones (Fig. 2; Percival et al.,2001).

Lin et al. (1996)originally proposed that collisionand amalgamation of the Lake Minto, Goudalie andUtsalik arcs generated the dominant N-NW foliation ofa regionally distributed, tectono-metamorphic episode(D2). Subsequently,Percival and Skulski (2000)dateda D2 fabric at 2.693–2.675 Ga and re-interpretedD2 as an overprint related to overthrusting of the2.71–2.70 Ga Tikkerutuk continental arc onto theamalgamated Lac Minto+ Goudalie+ Utsalik +Douglas Harbour proto-craton (Fig. 2, at 2.7 Ga).They further proposed that deformation and amphibo-lite to granulite facies metamorphism in supracrustalbelts of the west-central Minto Block resulted fromthis crustal thickening event, with transposition andrecrystallization of older fabrics into the dominantN-NW grain, and production of crustally-derived plu-tons and diatexites (2.69 Ga). Subsequent activity ismanifested as a series of less penetrative deformationevents (D3–D5:Lin et al., 1996; Percival and Skulski,2000) and late- to post-tectonic granitoid and syen-ite intrusions (2.696–2.645 Ga:Stern et al., 1994;Skulski et al., 1996; Percival and Skulski, 2000).Since the dominant tectono-metamorphic pulse of the

Minto Block (D2) is interpreted as a 2.69 Ga over-print by Percival and Skulski (2000), then how isthe pre-Leaf River suite accretion event (∼2.77 Ga)manifested in terms of fabric development and meta-morphism, and what was the impact of the 2.69 Gaevent on the older fabrics?

To clarify these issues, we present field, structural,and petrographic data from the poorly-known north-eastern part of the Minto Block, recently mapped ona 1:250,000 scale (Figs. 3 and 4) by the Ministèredes Ressources Naturelles du Québec. After com-paring our data with observations from elsewhere inthe Minto Block, we discuss the relative merits ofdifferent tectonic models to explain the characteristicfabrics and lithological assemblages associated withArchaean cratonization in this area.

2. Douglas Harbour domain

We divide the Douglas Harbour domain into theFaribault–Thury, Troie and Qimussinguat plutoniccomplexes (FTC, TC and QC) (Figs. 3–5), each ofwhich contains supracrustal belts. In the west, the FTCis cut by granodiorite, granite, pyroxene tonalite andenderbite of the Lepelle complex (2.729–2.727 Ga),which envelops screens and enclaves of rocks equiv-alent to the FTC (Fig. 3, Kapijuq and Bottequinsuites). To the northeast, Diana complex gneisses(2.78–2.76 Ga) were juxtaposed against the DouglasHarbour rocks during Proterozoic dextral compres-sion (Madore and Larbi, 2000), which overprints theeastern FTC (Figs. 1, 3 and 5). To the south, Troiecomplex rocks are intruded by Utsalik domain graniteand granodiorite (Berclaz et al., 2001; Leclair et al.,2001a).

The FTC is dominated by hornblende tonalite andbiotite trondhjemite (Fig. 6a), with subordinate dior-ite, granodiorite and granite. Western FTC tonalitesyield 2.88–2.86 Ga crystallization ages, while easternFTC tonalites yield younger ages between 2.81 and2.77 (Madore et al., 1999; Madore and Larbi, 2000;Percival et al., 2001). Inherited zircon cores yield agesup to 3 Ga (Percival et al., 2001), indicating recy-cling of older sialic material. The age of FTC pluton-ism overlaps with ages from embedded supracrustalbelts (∼2.82–2.78 Ga:Madore et al., 1999; Madoreand Larbi, 2000).


Fig. 3. Simplified geological map of the Douglas Harbour domain, adapted fromMadore et al. (1999, 2001), Madore and Larbi (2000), andCadieux et al. (2002). K = Kapijuq and Bottequin suite, strongly deformed tonalite and trondhjemite, equivalent to the Faribault–Thurycomplex. L= Lepelle complex, syn-kinematic sheet-like intrusions of granite, granodiorite and tonalite, with steep dips. M= McMahonsuite, enderbite, diorite and gabbro-norite, which straddle the Lepelle/Utsalik boundary. Among the late granitic suites, we distinguish themonzonitic rocks from the Troie complex and a series of late granite and granodiorite from the extreme northeast of the Faribault–Thurycomplex (Madore and Larbi, 2000). Note that associated granites are included with the monzonitic suite rocks in this figure, unlikeFig. 5.All other late granitoid rocks have been grouped together since there has yet to be a systematic investigation of their chemistry andpetrography. Cu= Curotte belt, Fa= Faribault belt, Ha= Hamelin belt, Ki= Kimber belt, Ta= Tasiaalujjuaq belt, Th= Thury belt.


Fig. 4. Simplified map of structural subdivisions of the Douglas Harbour domain, with stereonets showing mineral lineations (dots) andcontoured plots (Schmidt) of mineral foliation and layering. In the case of the Troie complex, filled dots represent the core of the massif,while open dots are from the periphery. Areas to the northeast have been overprinted by Proterozoic orogenesis.

The TC and QC are dominated by enderbite (py-roxene tonalite and trondhjemite), with subordinatepyroxene granodiorite, pyroxene granite, hornblendetonalite and biotite trondhjemite (Fig. 6a). Ages of

2.74–2.73 Ga were obtained from TC enderbites(inherited cores 2.83–2.8 Ga:Madore et al., 1999;Percival et al., 2001), and a QC trondhjemite (2.73 Ga,inherited cores 2.81–2.80 Ga:Madore et al., 1999).


Fig. 5. Map of the Lac Peters mapsheet (SE Douglas Harbour), showing the trend of foliations defined from field observations andinterpretation of air photographs. Plotted strike and dips are averaged values for each sub-area. Supracrustal belts also define a similar,swirly pattern, whereas late- to post-kinematic monzonitic intrusions strike N-S. Inset shows schematic cross-section.


Fig. 6. Thematic maps of the Lac Peters area (same area asFig. 5) showing the distribution of: (a) the nature of felsic plutonic rocksdetermined in thin section, using stained slabs, or where field determinations were unambiguous. TTG= tonalite, trondhjemite, granodiorite.Granite and monzonite are not shown. (b) Metamorphic assemblages in metabasalts. (c) The distribution of eastern and western FTC tonaliteand trondhjemite (TT) and diorite. Granodiorite and granite are not represented. W-FTC rocks have distinct geochemical signatures (higherU–Th-HREE) and it is from this domain that the oldest age dates originate. (d) Presence of igneous epidote and of (e) igneous(?) muscovite.


Poorly exposed contacts between the enderbitic TCor QC, and the tonalitic FTC, are marked by zonesof ductile deformation and strongly deformed gran-ite and granodiorite intrusions. The QC and TCare separated by a corridor of FTC-type rocks, theThury Deformation Zone (TDZ,Figs. 3–5). Rocksfrom the TDZ are younger than the rest of the FTC,yielding a 2.734 Ga crystallization age, with inher-ited cores at 3.013–2.762 Ga (Percival et al., 2001).Thus, to a first approximation, emplacement of theQC, TC and TDZ magmas (2.74–2.73 Ga) is coevaland nearly age-correlative with the Leaf River suite(∼2.73–2.72 Ga:Percival et al., 1994; Stern et al.,1994). However, tonalite and trondhjemite bodiesin the QC are significantly older (2.77 Ga, inher-ited cores 2.83 Ga:Madore et al., 1999). Since this

Fig. 7. (A) Layered trondhjemite to leuco-trondhjemite with segregations of leuco-trondhjemite (Thury Deformation Zone). A structuralcontrol on segregation is apparent. Lenscap 55 mm for scale. Black spots are lichen. (B) Folds in banded iron formation (quartz/magnetite)(E-FTC). Lenscap for scale. Irregular black and white spots are lichen. (C) Trondhjemite 98-3230A (QC/E-FTC). Note the euhedral an-tiperthitic plagioclase phenocryst, and the high-energy feldspar–feldspar grain boundaries. Crossed polarizers, 5 mm across. (D) Melatonalite98-5064A (W-FTC). Note the serrated, high-energy feldspar–feldspar and feldspar–quartz grain boundaries, and presence of subgrains andneoblasts at grain contacts. Crossed polarizers, 5 mm across.

is the age of the eastern FTC rocks, we infer thattonalite and trondhjemite within the TC and QC aremostly septa of FTC rocks engulfed by enderbiticmagmas.

The TC and QC are invaded by gabbronorite,pyroxene+ mica diorite, granodiorite, monzonite,quartz monzonite and granite intrusions (Figs. 3 and5). Among the younger intrusions, a hornblendegabbronorite from the TC yields an average Pb–Pbage of 2.73 Ga (Madore et al., 1999), similar to thatof TC enderbites. A quartz monzonite sheet fromthe TC yields a 2.697 Ga U/Pb zircon age (inher-ited cores 2.73–2.72 Ga:Madore et al., 1999). Apyroxene+ mica diorite dyke cuts this quartz mon-zonite, implying that these diorites are younger thanthe hornblende gabbronorites.


3. Field relationships and lithological assemblages

3.1. The Faribault–Thury complex

Tonalite and trondhjemite in the FTC are commonlyinterlayered on a centimeter- to meter-scale, and con-tain deformed enclaves of amphibolite, paragneiss,and pyroxenite. The absence of melanosome sheathsseparating tonalite from trondhjemite, or of preservedmetatexite domains, suggests that these are probablynot in-situ migmatites. Outcrop-scale structures indi-cate that magmatism and deformation overlapped intime: e.g. trondhjemite may occupy strain shadowsand fill shear bands, and trondhjemitic breccia con-tains abundant tonalitic to dioritic enclaves in variousstates of transition from angular blocks to elongateschlieren (Fig. 7A). Layered tonalite–trondhjemite do-mains alternate on a 1–5 km-scale with more homoge-neous trondhjemite or tonalite intrusions, which alsocommonly display penetrative fabrics.

Mineral foliation generally parallels layering andis typically steeply dipping. Regionally, the foliationstrikes N-NW (Fig. 4), but in detail the foliation andlayering define complexly swirling, 2–10 km-scaleovoid patterns (Figs. 5 and 8). Lineations charac-teristically plunge steeply (Fig. 4), and L-tectonitesare locally prominent in the vicinity of supracrustalbelts. In the TDZ (Figs. 3–5), N-NW-trending fabricsare reoriented to an E-W grain with sub-vertical lin-eations. Proterozoic orogenesis affected the easternFTC (Fig. 4), with partial preservation of Archaeanfabrics and textures between deformation corridors(Madore and Larbi, 2000).

Tonalite contains 10–30% hornblende, biotite,and Fe–Ti-oxides, with accessory epidote, allan-ite, titanite and apatite (Fig. 6d). Feldspar-phyricfacies are rare. Trondhjemite has<10% ferromag-nesian minerals, principally biotite, with subordi-nate muscovite, and the same accessory phases astonalite (Fig. 6e). Trondhjemite pegmatites asso-ciated with meta-pelitic supracrustal rocks containtourmaline+ garnet+ muscovite+ biotite. Graniteand granodiorite (<15% of the FTC) occur as dis-crete bodies that are distinctly more massive and lessdeformed than tonalite–trondhjemite hosts, as smallbodies of augen gneiss associated with the edges ofthe TC and QC, or as widely-distributed, essentiallyundeformed, granitic pegmatite dykes. Granite and

granodiorite may contain quartz and alkali feldsparphenocrysts, while muscovite, epidote, allanite, ap-atite, titanite and rare garnet are concentrated in maficschlieren with biotite and hornblende.

Mafic ‘dioritic’ schlieren are common, but largerbodies (∼100–200 m) are rare, and are engulfed bytonalite–trondhjemite. Diorite is massive to weaklyfoliated, and heterogeneous in mode and grain size.Feldspars are lath-shaped. Hornblende, biotite andminor clinopyroxene relicts constitute 30–50%. Ac-cessory apatite, titanite, magnetite and ilmenite areubiquitous and can be abundant. Minor allanite andquartz may be present.

3.2. The Troie and Qimussinguat complexes

Enderbite contains plagioclase+ quartz+ ortho-pyroxene± clinopyroxene+ biotite + magnetite+apatite+zircon±minor interstitial hornblende. Pyrox-ene granodiorite and granite are uncommon. Ender-bite may form massive intrusions showing little modalor textural variation over many kilometers, but morecommonly, mela- and leuco-enderbite are interlay-ered on every scale (like FTC tonalite–trondhjemite).Veins of unfoliated or weakly foliated leuco-enderbitemay form reticulated networks that crosscut com-positional layering, occupy strain shadows aroundxenoliths, or are injected along shear zones (seeBédard, 2003). These features are interpreted tosignify that deformation was syn-magmatic, withstructural control on migration of evolved residualleuco-enderbite melt. Melanosome sheaths sepa-rating mela- and leuco-enderbite have never beenobserved, suggesting that these are probably not insitu migmatites. Swarms of pyroxene diorite andmela-enderbite enclaves in leuco-enderbite matricesare common. Calc-silicate and partially-reacted am-phibolite enclaves with pyroxene-rich reaction rimsalso occur. Steeply-dipping foliations generally strikeN-NW (Fig. 4), but in detail, they define swirling,2–10 km-scale ovoid patterns (Fig. 5). Lineationsgenerally plunge steeply, but plunges are slightlyshallower near contacts with the FTC (Fig. 4).

In the east and north of the TC, extensive domainsof homogeneous, massive granite with blue quartz areattributed to the monzonitic suite. Small, strongly de-formed and heterogeneous bodies of granite, granodi-orite, tonalite and trondhjemite are common near the


Fig. 8. Detail map of the Hamelin belt, located in the W-FTC.

edges of the TC and QC (Fig. 6a). The tonalite andtrondhjemite are geochemically and petrographicallysimilar to FTC equivalents, and may be cut by ender-bite intrusions, which suggests that many are engulfedslivers of the FTC. However, some tonalites and trond-hjemites may be intrusions from a younger phase ofmagmatism similar to that of the TDZ.

3.3. Troie complex monzonites, quartz monzonites,granites, gabbronorites and diorites

North-trending, steeply-dipping, intrusive sheets ofporphyritic monzonite to quartz monzonite (5–10 kmwide, up to 40 km long) constitute∼15% of the Troiecomplex. Similar rocks occur in the Quimussinguatcomplex, but have not been systematically examined,so that we will focus on the Troie complex as a typeexample of the late intrusive rocks. The monzonitesand quartz monzonites are typically massive andhomogeneous, and are crowded with euhedral phe-

nocrysts (1–10 cm) of alkali feldspar, plagioclase andlocally quartz. Most intrusions are undeformed, butin some bodies, phenocrysts have ovoid shapes anddefine a contact-parallel foliation. Hornblende, biotiteand Fe–Ti-oxides are the dominant mafic minerals.Minor apatite, titanite and zircon are common. Thereare isolated igneous microgranular enclaves, andthin (∼1 m), micro-monzonite dykes. Clinopyroxenemonzodiorite and biotite granite have sharp intru-sive contacts against monzonite or quartz monzonite.Granular textured, fine- to medium-grained biotitegranite forms extensive, homogeneous bodies. It hasfew phenocrysts, and commonly contains blue quartz.

Small intrusions of pyroxene+ mica diorite andhornblende gabbronorite are dispersed throughout theTroie complex and also occur in the Quimussinguatcomplex (Madore et al., 1999). Abundant (∼10–15%)Fe–Ti-oxides are diagnostic. Hornblende gabbronoriteis more strongly deformed than pyroxene+mica dior-ite, with fabrics being similar to those in adjoining


supracrustal and plutonic rocks. Pyroxene+ micadiorite contains plagioclase and biotite phenocrystsaligned parallel to chilled contacts, indicating mag-matic flow. Isolated gabbronorite to diorite bodiesmay be massive and homogeneous, with the grainsize coarsening inwards. Agmatitic complexes arecommon, with gradations from dykes with chilledmargins to swarms of cm- to m-scale ovoid or lobatebodies embedded in enderbite, tonalite or monzonite.The agmatites are interpreted to be disaggregatedintra-plutonic dykes (Hibbard and Watters, 1985;Pitcher, 1991). Rare potassic pyroxenite and harzbur-gite form discrete lensoid bodies (sheared dykes ormetalavas?) in felsic hosts, or are associated withsupracrustal belts (feeders or sills?).

3.4. Supracrustal belts

All three Archaean complexes contain remnants ofsupracrustal belts, generally<3 km wide and<15 kmlong (Fig. 3). Decameter-scale fold closures are ap-parent (Fig. 8), with ubiquitous sub-vertical fold axesand stretching lineations. Centimeter-scale folds arevisible in the laminated quartz–magnetite ironstone(Fig. 7B). Fabric orientations in supracrustal and ad-joining plutonic rocks are similar. A coherent stratig-raphy is rarely preserved in supracrustal belts, andthere are marked along-strike contrasts in the domi-nant lithology. For example, in the Thury belt (easternFTC:Fig. 3) a section of 100% paragneiss+ ironstonechanges within<20 m into a section dominatedby metabasalt. This variability implies significantinternal deformation and probable tectonic sliver-ing. Detail mapping of the Hamelin belt (westernFTC) reveals a complex fold interference pattern(Fig. 8) and ubiquitous injection of syn-kinematictonalite–trondhjemite sheets (cf.Pawley and Collins,2002).

Tholeiitic metabasalts form sequences hundreds ofmeters in thickness (Fig. 8), or are interbedded on ameter-scale with paragneiss and komatiitic metalava.Pillow lavas are preserved locally, but more typically,FTC metabasalt is converted to massive or lami-nated amphibolite, commonly with a well-developed,steeply-plunging, hornblende lineation. Commonleuco-trondhjemitic, foliation-parallel leucosomes(1 mm–1 cm) have millimeter-thick hornblenditerinds, suggesting an anatectic origin (Kriegsman,

2001). Associated coarse to pegmatitic leuco-trondhjemite pods probably represent segregations ofanatectic melts. In the TC and QC, metabasalt is trans-formed into mafic granulite with individual foliation-parallel laminae having different feldspar/pyroxeneand hornblende/pyroxene ratios. Laminae are stronglytransposed and may define steeply plunging isoclinalfold closures. Hornblende is commonly concentratedinto anastomozing amphibolitic gabbro laminae withinthe hornblende-poor mafic granulite. Metabasalt inthe TDZ varies from amphibolite to granulite grade(Fig. 6b).

Komatiitic flows, subvolcanic sills and peridotitic,pyroxenitic, and gabbroic cumulates are interlay-ered with metabasaltic lavas. Rhyodacitic to daciticmetatuffs are thin (<1 m) and laterally continu-ous (>20 m). Coarser blocky tuffs and pyrite-richgossans occur locally. Metapelite is the most com-mon metasedimentary rock, forming 1–10 m thickbands of migmatitic paragneiss dominated by quartz,garnet and biotite± sillimanite. Well-preservedmeta-greywacke, meta-quartzite and pebbly sandstoneoccur in the Hamelin belt (Fig. 8). Bands of ironformation are typically 1–10 m thick, with up to 50 mof apparent thickness in fold closures. Rare marbleand calc-silicate rocks form recessive layers 1–10 mthick.

Considering the nature of the lithologies present,an intra-cratonic or peri-continental environmentseems likely for most of these supracrustal belts (cf.Ayres and Thurston, 1985; Thurston and Chivers,1990; Lowe, 1994; Percival et al., 1997, p. 211).This is consistent with the observation of mafic andultramafic dykes cutting tonalites (Percival et al.,1992) and of basal unconformities in other MintoBlock belts (Moorhead, 1989; Skulski and Percival,1996).

3.5. Proterozoic rocks and overprints

Proterozoic dykes in the eastern FTC are affectedby amphibolite-grade regional metamorphism anddeformation. Proterozoic deformation is concentratedinto corridors hundreds of meters wide, characterizedby dextral shear (Madore and Larbi, 2000). Brittle,west-verging reverse faults near the eastern TC/FTCcontact may represent distal manifestations of theProterozoic deformation (Fig. 5; Madore and Larbi,


2000). Proterozoic supracrustal packages are klippe,with kinematic indicators implying that they werethrust towards the south.

4. Petrographic and mineral-chemistryconstraints

Many rocks show partial retrogression to thegreenschist grade of original igneous, or high-grademetamorphic assemblages. Hydrothermal greenschistfacies metamorphism is more intense in the vicinityof major greenschist facies fractures. These fracturesare spaced several kilometers apart, on average, andextend as much as 100 km, but are of unknown ageand origin.

4.1. Supracrustal rocks

Metabasaltic rocks in the FTC are dominatedby hornblende and plagioclase, with minor quartz,biotite, Fe–Ti-oxides, pyrite, titanite, epidote andgarnet. Relict clinopyroxene and olivine phenocrystpseudomorphs are uncommon. Textures range fromgranoblastic polygonal, to nematoblastic, to por-phyroclastic. Generally, elongate hornblende grainsin a thin section share the same pleochroism andextinction angle (see Fig. 3e inBédard, 2003), im-plying extensive syn-deformational recrystallization.Plagioclase feldspar forms a polygonal mosaic withstraight grain boundaries and 120◦ triple junctions.Plagioclase may also be concentrated with quartzinto foliation-parallel anatectic veinlets. Amphiboliticmetabasalt affected by lower-temperature deformationhas a porphyroclastic texture, with feldspar and horn-blende augen, deformation twins, marginal granula-tion of feldspar into a fine-grained quartzo–feldspathicmatrix, oblique shear bands, and concentrations ofquartz + chlorite in pressure shadows and frac-tures.

In granulitic metabasalt of the TC and QC, pyrox-ene and feldspar have straight or gently curved grainboundaries and 120◦ triple junctions, indicating com-plete re-equilibration. Some rocks preserve evidenceof prograde granulitization of amphibolite. Wherepresent, hornblende typically makes up a polygo-nal mosaic with pyroxenes and feldspar, with rarerims on pyroxene that suggest a replacement origin.

Hornblende may also be concentrated into centimet-ric lamellae or veins that alternate with granuliticassemblages. Retrogression to amphibolite gradecaused by channelized penetration of water seemslikely.

Ultramafic lavas contain porphyroclastic metamor-phic(?) olivine (Fo77) and minor chromite, surroundedby clinochlore (Mg-chlorite) needles in a matrix ofiddingsite, tremolitic hornblende, and epidote. As-sociated ultramafic cumulates are harzburgite (sub-ordinate wehrlite and lherzolite) containing olivine(Fo88.5) and either green spinel or brown chromite,both surrounded by poikilitic pyroxene. Pyroxenitecontains both ortho- and clino-pyroxenes, olivine,Fe–Ti-oxides, amphibole, mica, plagioclase and al-kali feldspar. Pale brown, interstitial phlogopitic micais ubiquitous, and may be abundant. With increas-ing proportions of poikiloblastic hornblende, rocksgrade towards hornblendite and feldspathic horn-blendite.

4.2. Gabbronorite and diorite of the Troie complex

Hornblende gabbronorite has a mosaic texture, andis dominated by pyroxene and plagioclase (An26–56).It is distinguished from metabasaltic granulite by hav-ing abundant apatite (1–3%), magnetite and ilmenite(∼10–15%). Green hornblende (15–30%) replaces py-roxene or forms oikocrysts. Pyroxene+ mica dior-ite commonly preserves euhedral, zoned, antiperthitic,plagioclase phenocrysts (≤15%, An24–32; see Fig. 12in Bédard, 2003). Abundant red biotite (5–15%) oc-curs as phenocrysts or small flakes. Hornblende issubordinate (<5%). The groundmass is a texturallyequilibrated, polygonal mosaic of two pyroxenes andplagioclase, together with magnetite, ilmenite, minorsulphide, titanite and quartz.

4.3. Monzonite, quartz monzonite, monzodiorite,and granite of the Troie complex

Monzonite and quartz monzonite have large(1–5 cm), generally euhedral, perthitic alkali feldsparphenocrysts, commonly with inclusions of euhedralplagioclase (An25) and/or quartz. Quartz phenocrystsmay be embayed. Myrmekite is common. Monzodior-ite contains relics of clinopyroxene within hornblende,and has little quartz or alkali feldspar.


4.4. Felsic plutonic rocks in the Faribault–Thurycomplex

Tonalite and trondhjemite are generally not por-phyritic, though a few contain sub-euhedral, an-tiperthitic plagioclase phenocrysts (1–5 cm, 5–15%),or micro-phenocrysts (Fig. 7C), and deformed phe-nocrysts (phenoclasts) are equally rare. Plagio-clase composition ranges from An15 to An30, withthe tonalite/trondhjemite transition at ca. An25.Plagioclase–plagioclase contacts are typically ser-rated high-energy boundaries showing subgrain de-velopment (Fig. 7C and D). Weakly-deformed rockscontain quartz with deformation bands and undula-tory extinction, yet which locally retains an inter-stitial habit. More typically, quartz is recrystallizedinto neoblasts or annealed ribbons. Microcline is rare(<5%), and occurs interstitially, as exsolutions, or assmall neoblastic subgrains. Micro-perthite to micro-cline exsolutions are generally undeformed (Fig. 7C),and myrmekitic plagioclase+ quartz intergrowths arecommon. In rocks affected by submagmatic, solid-state deformation, plagioclase forms twinned augen,mortar textures develop, and the matrix is neoblastic.

Tonalite contains subequal proportions of horn-blende and biotite, typically concentrated intoschlieren with oxides, titanite, apatite, allanite andepidote. Hornblende grains in a thin section do nothave common extinction angles and pleochroism (un-like the amphibolites), suggesting substantial grainrotation or translation during deformation, or simplyan absence of dynamic recrystallization. Some ofthe larger hornblende grains may be recrystallizedphenocrysts. Trondhjemite has<10% ferromagne-sian minerals, principally biotite, with subordinatemuscovite, and traces of the same accessory phasesas tonalite. Apatite (∼1%) is commonly included inmagnetite and ilmenite (1–4%), which occur as inter-growths or as discrete grains. Titanite is ubiquitous (1–5%), as rims on Fe–Ti-oxides and as anhedral grainswith complex internal zoning. Euhedral to roundedzircons contain inherited cores and multiple igneousovergrowths separated by dissolution surfaces.

Epidote is nearly ubiquitous in FTC tonalite andtrondhjemite (0.5–5%,Fig. 9A). It is pale green in thinsection, commonly contains a core of allanite (Figs.3b–d and 5 inBédard, 2003), has a narrow, typicallyigneous, compositional range (Ps23–29), and may show

internal growth zoning that is truncated by corrosionstructures, features to which we attribute an igneousorigin (Zen and Hammarstrom, 1984, 1986; Evans andVance, 1987; Dawes and Evans, 1991). Epidote is eu-hedral when completely embedded in biotite or mus-covite, but shows embayments and wormy textureswhen exposed to the quartzo–feldspathic matrix (Fig.9A; cf. Fig. 3a–d inBédard, 2003), which we interpretas the result of magmatic corrosion or dissolution. Ig-neous epidote in contact with chlorite is replaced bydeeper green (Fe-rich) metamorphic epidote rims.

Allanite is generally metamict, but rare grains pre-serve fine, concentric lamellar growth zoning. Otherallanite grains have complex internal zonations thatprobably reflect multiple growth/dissolution events,locally associated with the development of Th- andrare-earth minerals. Allanite also forms isolatedprismatic grains with scalloped edges suggestingmagmatic resorption (seeBédard, 2003). With pro-gressive solid-state deformation, the relatively rigidepidote, allanite and titanite grains tend to break upinto domino structures (Kruse and Stünitz, 1999)and micro-boudins, and eventually evolve into sliversenclosed within biotite or muscovite fish.

Muscovite has the high FeO∗ (4.5–6 wt.%) andTiO2 (0.5–1.7 wt.%) contents typical of igneousmuscovite (Miller et al., 1981; Speer, 1984; Zen,1988). Muscovite intergrown with biotite occurs aswell-crystallized, faceted crystals. Muscovite that isin direct contact with quartz or plagioclase, on theother hand, has irregular terminations, with wormytextures (Fig. 9B) that Bédard (2003)attributed tomagmatic resorption.

Apart from a higher proportion of alkali feldspar,granodiorite resembles tonalite in most respects (in-cluding the types of trace phases), while graniteresembles trondhjemite. Granite and granodioritemay contain minor Fe–Mn-garnet with biotite inclu-sions, and/or minor prismatic tourmaline. The garnetfalls in (or near) the igneous garnet field ofClarke(1981). Trondhjemitic pegmatite is associated withmetapelite-dominated supracrustal packages, andcontains almandine-rich garnet, muscovite, biotiteand tourmaline. Dichroic tourmaline has blue cores(schorl) and more dravitic green rims. The presenceof tourmaline implies a metasedimentary component(London, 1999), supporting local derivation from, orcontamination by, paragneiss.


Fig. 9. (A) Resorbed magmatic epidote from tonalite 98-3094A (E-FTC). Note the absence of secondary magnetite from the spongy domains,and the presence of U–Th-LREE enriched high-reflectibity veinlets, which suggest a late deuteric(?) fluid-migration event (Bea, 1996). (B)Trondhjemite 98-2200A (E-FTC). B= biotite. Note that muscovite (M) is corroded when in contact with the quartzo–feldspathic matrix,but has euhedral faces when protected by biotite. Plane polarized light, 5 mm across. (C) Enderbite 98-3266 (TC). Euhedral antiperthiticplagioclase (P) with apatite (A), zircon (z) and magnetite (mt) inclusions. Interstitial quartz (Q). Crossed polarizers, 5 mm across. (D)Enderbite 98-1064D (TC). Euhedral orthopyroxene (O) and magnetite (black). Biotite (B) is abundant, as is interstitial hornblende (H).Quartz and feldspar are white. Plane polarized light, 5 mm across.

Rocks of the eastern FTC are typically more de-formed than those in the west as a result of Proterozoicorogenesis. Strong, solid-state deformation causeddevelopment of plagioclase and quartz neoblasts, re-crystallization of biotite, concentration of muscoviteinto phacoidal cleavages, and development of pris-matic epidote. There is also extensive retrogressionto the greenschist facies. As a result, igneous texturesand minerals are less well preserved in the easternFTC, though still recognizable in low-strain domains.Some of the metamorphic epidote may representrecrystallized igneous phases. The occurrences of ig-neous epidote shown inFig. 6drepresent cases wherean igneous origin could be established based on tex-tural considerations (growth zoning, igneous allanitecores, marginal resorption structures).

4.5. Felsic plutonic rocks in the Troie andQimussinguat complexes

Textures of feldspars and quartz in TC and QC en-derbites do not differ significantly from those in theFTC. Plagioclase (An27–34) generally has high-energy,high-temperature, serrated grain boundaries and iscommonly antiperthitic. In relatively undeformedrocks, plagioclase may be euhedral and quartz inter-stitial (Fig. 9C; cf. Fig. 6c in Bédard, 2003), unam-biguously indicating an igneous origin. Phenoclastsare rare. With progressive deformation, quartz de-velops deformation bands, sub-grains, and finally isreduced to a neoblastic mosaic.

Orthopyroxene generally forms stubby, euhe-dral, pleochroic prisms (Fig. 9D). Clinopyroxene is


commonly euhedral and prismatic, but may also beinterstitial to orthopyroxene. Fine exsolution lamellaeoccur in either pyroxene. Minor green hornblendeis interstitial (Fig. 9D) and rims clinopyroxene.Fe–Ti-oxide granules and TiO2-rich biotite (brightred in thin section) are ubiquitous and abundant. Mi-nor apatite and zircon are common (Fig. 9C), butaccessory igneous titanite and epidote are absent.Hornblende tonalite and biotite trondhjemite in theTC and QC are identical to those found in the FTC.

5. Geothermobarometric constraints

Bédard (2003)presented geothermobarometric cal-culations on these rocks. The Al-in-hornblende geo-barometer (Hammarstrom and Zen, 1986; Schmidt,1992) yielded 3–6.4 kbar pressures for tonalite of theFTC, QC and TC in agreement with other estimatesfrom Minto Block plutons (3.5–5.6 kbar:Percivalet al., 1992; Percival and Mortensen, 2002), andsupracrustal belts (generally 2–5 kbar:Percival andBerman, 1996; Madore et al., 1999; Percival andSkulski, 2000). Granulite-grade supracrustal rocksmay show fragmentary preservation of a higherpressure history (8–12 kbar:Percival and Skulski,2000; Cadéron et al., 2003). The QUILF thermome-ter (Andersen et al., 1993) gave 697–911◦C forhornblende gabbronorite, pyroxene+ mica diorite,and enderbite of the TC. TheBlundy and Holland(1990) plagioclase–hornblende geothermometeryields near-solidus temperatures (711–745◦C) forFTC felsic plutons. A hornblende-bearing TC en-derbite records a higher temperature (790◦C), asdo TC gabbronorite and pyroxene+ mica diorite(760–826◦C). Reconstructed pre-exsolution compo-sitions of antiperthite grains (Bédard, 2003) yieldminimum ternary plagioclase liquidus tempera-tures of 550–870◦C in tonalite–trondhjemite, and810–1045◦C in enderbite.

6. Discussion

Existing plate tectonic models for the Minto Block(Percival et al., 2001) have great predictive value,since the geology can be compared directly withwell-studied Phanerozoic environments. In the follow-

ing, we discuss the implications of our observationswith regard to the interplay of igneous, metamorphicand deformational processes active in the DouglasHarbour domain and other parts of the Minto Block.We then compare the geology of the Douglas Harbourdomain and the rest of the Minto Block with An-dean/Cordilleran accretionary orogens. The compari-son highlights problems with application of actualisticmodels to Archaean craton formation, leading us toconsider vertical tectonic models in the final section.

6.1. Metamorphic or plutonic origin of minerals andfabrics in plutons

Minto Block supracrustal rocks have tectono-metamorphic fabrics and assemblages (Percival andSkulski, 2000; Bédard, 2003) and display leucosome–melanosome pairs and diatexitic/metatexitic struc-tures that indicate anatexis. Existing genetic modelsfor the Minto Block attribute deformation and meta-morphism of supracrustal rocks to overprinting,regional-scale orogenic/metamorphic events thatcompletely transposed and recrystallized pre-existingfabrics into the dominant N-NW grain (Fig. 2). How-ever, during a regional-scale orogenic overprint, theplutonic rocks in which the deformed and metamor-phosed supracrustal slivers were embedded must haveexperienced the same events, and so should exhibitmicrostructures and mineral assemblages consistentwith transposition and prograde recrystallization attemperatures and pressures similar to those of thesupracrustal rocks. Do the plutonic rocks record thisoverprinting tectono-metamorphic event?

Many Douglas Harbour (and other Minto Block)plutonic rocks preserve relict magmatic textures andstructures, such as: (1) variably deformed tonalitic(or mela-enderbitic) enclaves in trondhjemitic (orleuco-enderbitic) breccia complexes; (2) mingledmelatonalite–trondhjemite intra-plutonic dykes (cf.Fig. 2b inBédard, 2003) showing varying degrees ofdeformation, with fabric orientations similar to thoseobserved in adjacent tonalitic hosts; (3) euhedraltwinned laths of plagioclase in diorite (cf.Vernon,2000); (4) quartz with interstitial habits (Fig. 9C); (5)enderbite with coarse, massive, homogeneous textures(Percival and Skulski, 2000), locally with alignedphenocrysts indicating magmatic flow (Percival et al.,1997); (6) epidote (Fig. 9A) and allanite with textures


and compositions (Bédard, 2003) identical to thosefrom undeformed and unmetamorphosed Phanerozoicplutons (e.g.Zen and Hammarstrom, 1984, 1986;Zen, 1985); (7) an absence of coronitic structures(compare withSt-Onge and Ijewliw, 1996); (8) oc-casional euhedral plagioclase phenocrysts, which areonly rarely transformed into phenoclasts.

The presence of leucocratic segregations occupy-ing strain shadows and shear zones within tonalite ormela-enderbite (Figs. 2, 6 and 7a inBédard, 2003; cf.Sawyer, 2000; Pawley and Collins, 2002) implies thatdeformation was syn-magmatic. The fact that maficminerals generally do not share crystallographic axes,unlike the hornblende grains in metabasalt from ad-joining supracrustal belts, is consistent with this in-ference, as is the homogeneity of the fabrics, whichsuggests a viscous (mushy) medium (Gapais, 1989).

At the granulite to amphibolite grade condi-tions preserved in supracrustal rocks (peaks at∼800◦C from Percival and Skulski, 2000), biotite+muscovite-bearing granitoids should be migmatized(e.g.Thompson, 2001). However, felsic plutons of theDouglas Harbour domain (and elsewhere in the MintoBlock: Bédard, 2003) show few of the hallmark traitsof migmatites (e.g.Kriegsman, 2001; Brown, 2001),and commonly record temperatures higher than thoseof adjacent supracrustal belts (Bédard, 2003).

The common preservation of outcrop- andmicro-scale structures indicating that many fabrics(mafic schlieren, lineation, foliation, preferred crystalorientation) began to form at the magmatic or submag-matic stage, the preservation of delicate exsolutiontextures (Figs. 7C and 9C) recording near-liquidustemperatures; and the rarity of phenoclasts imply thatmost plutons did not experience penetrative, prograde,tectono-metamorphic recrystallization. Most of thetextural and mineralogical characteristics of the Dou-glas Harbour felsic plutons are more consistent withslow cooling of a syn-kinematic igneous system, withgradual development of solid-state fabrics as plutonscontinued to deform in the sub-magmatic stage. Thefact that feldspar behaved as plastically as quartz inthe felsic plutons suggests that this solid-state de-formation occurred at high (>900◦C) temperatures(Gapais, 1989; Dell’Angelo and Tullis, 1996). Theubiquitous serrated, high-energy, feldspar–feldspargrain boundaries and subgrains (Fig. 7C and D) sug-gest recrystallization by fast grain boundary migration

at high temperatures (Tullis and Yund, 1985; Lafranceet al., 1996; Vigneresse et al., 1996; Rosenberg,2001). These high-temperature feldspar microstruc-tures in the plutons contrast with the straight, wellequilibrated feldspar grains (120◦ triple junctions) inmetabasalt, which indicate a very different, and moreclearly ‘metamorphic’, prograde recrystallizationhistory (Kretz, 1966; Vernon, 1999). Since geother-mometry yields near-solidus temperatures for mostmineral phases in the plutonic rocks (Bédard, 2003),we infer that the plastic deformation textures gener-ally represent a continuation of syn-magmatic defor-mation beyond the rheological locking point into thesub-subsolidus, sub-magmatic regime (e.g.Hibbard,1987; Paterson et al., 1989; Vigneresse et al., 1996),rather than an overprinting orogenic/metamorphicevent. Later overprinting fabrics exist, but tend to beconcentrated in relatively narrow corridors with steepfoliations and sub-horizontal lineations (Cadieuxet al., 2002), and are not obviously associated withthe development of the penetrative N-NW fabricscharacterizing the Douglas Harbour domain and mostof the central-western Minto Block.

6.2. Melt ascent and emplacement in the DouglasHarbour domain

Recognition of relict igneous textures and of thesyn-kinematic nature of Douglas Harbour magmatismallows partial reconstruction of the melt ascent andemplacement history. Textures in igneous epidote andmuscovite from FTC rocks are important, becausethe stability of these minerals is sensitive to pressure.Epidote commonly has embayed and spongy marginswhen in contact with the quartzo–feldspathic matrix(e.g.Fig. 9A), but faceted faces when armored by bi-otite or muscovite. Identical partial magmatic resorp-tion structures in epidote from Phanerozoic tonalitesare attributed to syn-magmatic decompression (Zenand Hammarstrom, 1984, 1986; Zen, 1985; Schmidtand Thompson, 1996). A 6 kbar threshold is com-monly suggested (Zen and Hammarstrom, 1984, 1986;Zen, 1985), but Schmidt and Thompson (1996)con-cluded that epidote could remain stable to 3 kbar inoxidized magmas. Thermobarometric calculations forDouglas Harbour rocks (Bédard, 2003) suggest that a6 kbar breakdown pressure may be applicable. Mus-covite also exhibits evidence of magmatic resorption


(Fig. 9B, alsoBédard, 2003). Experimental data indi-cate a 2–4 kbar lower limit for the stability of igneousmuscovite (Chatterjee and Johannes, 1974; Andersonand Rowley, 1981), though variations in oxygen andvolatile activity may shift this threshold (Miller et al.,1981; Weidner and Martin, 1987; Zen, 1988).

Despite uncertainty about the absolute pressureof phase destabilization, the common presence ofcorroded epidote and muscovite in Douglas Harbourtonalite–trondhjemite (Fig. 6d and e) indicates thatthese magmas ascended rapidly in the form of crys-tal slurries from a depth where these phases werestable on the liquidus (cf.Zen, 1985; Brandon et al.,1996). Identical igneous epidote and muscovite oc-cur in tonalite–trondhjemite from most Minto Blockdomains (Bédard, 2003), suggesting that these con-clusions are generally applicable. Ascent of melt isan efficient heat transport mechanism, andBédard(2003) proposed that the metamorphism observedin Minto Block supracrustal rocks is pluton-driven(DeYoreo et al., 1989), which removes the need fora regional, Barrovian metamorphic episode linked toorogenesis.

6.3. Do voluminous tonalite–trondhjemite–granodiorite–granite (TTG) plutons implysubduction under an Andean-type margin?

It has been proposed that the closest modern ana-logues to Archaean TTG suites are Andean mar-gin cordilleran-type batholiths constructed on apre-existing sialic crust (Weaver and Tarney, 1981;Condie, 1986; Stern et al., 1994; Martin, 1999).However, the differences between Minto Block TTGand Andean margins leads us to question whetherthis proposal is correct (cf.Maaløe, 1982; Hamilton,1998; Smithies, 2000).

Firstly, the broad areal distribution of TTG in Ar-chaean cratons (∼500 km wide distribution of the2.73 Ga Leaf River suite in the Minto Block,Fig. 1)and large volumes produced, are probably incon-sistent with a single, linear source of arc magma(Reymer and Schubert, 1984; Hamilton, 1998). Notethat the plutonic belts of the North and South Amer-ican Cordilleras are only∼150 km wide (Chardonet al., 1999; Scheuber and Gonzalez, 1999; Brownet al., 2000), even though subduction operated nearlycontinuously for 200 Ma, a timescale similar to that

required to generate the bulk of the Minto Block.Percival et al. (1994)circumvented this problem byproposing the existence of two synchronous subduc-tion zones. Slab rollback may generate extensionalaccretionary orogens characterized by wider plutonicbelts (Collins, 2002), and represent another possibleanalogue for the Minto Block plutonism. However,the basalts associated with extensional accretionaryorogens have prominent arc-related geochemical sig-natures, unlike typical Minto Block tholeiites (Madoreet al., 1999), and their closure is characterized bythin-skinned thrusting and duplexing (see referencesin Collins, 2002), which is not seen in the MintoBlock.

Secondly, the structural pattern in the Minto Blockis very different from that of Phanerozoic accretionaryorogens. Terrane accretion in the North and SouthAmerican Cordilleras is characterized by the devel-opment of synchronous, shallowly-dipping thrustsand steeply-dipping strike-slip faults (Chardon et al.,1999; Scheuber and Gonzalez, 1999; Brown et al.,2000). This contrasts with the extreme rarity ofshallowly-dipping planar fabrics in the Minto Block(e.g.Fig. 4). Domains characterized by syn-magmaticfabrics with steep plunges exist in Phanerozoic accre-tionary orogens (McClelland et al., 2000), but formnarrow belts (10’s of km) embedded within dominantshallowly-dipping structures. While a pure flatteningfabric associated with terrane accretion or collisioncould produce steeply-dipping and plunging L–Sand S–L fabrics on a local scale (Choukroune et al.,1995, 1997), it is difficult to imagine a collisionalor accretionary process that repeatedly generatedsuch fabrics without also generating coeval forelandthrust-and-fold belts.

Another problem is the scattered distribution ofsupracrustal belts (Figs. 3 and 5), a pattern whichseems inconsistent with a marginal accretion or‘obduction’ zone, as proposed for the southern Supe-rior Province (Card, 1990; Kimura et al., 1993; Calvertand Ludden, 1999). The absence of thrust faults withconsistent vergence also seems inconsistent with suchmodels. Furthermore, many Minto Block supracrustalrocks do not resemble oceanic crust, being moresimilar to platform sequences (Thurston and Chivers,1990; Lowe, 1994; Percival et al., 1997). One possi-ble alternative to obduction is subcretion, or lateralaccretion of soft, largely melted slabs to the margins


of the continent (Kusky and Polat, 1999). However,this type of lateral accretion should also produceobliquely dipping fabrics with a consistent vergence(Beaumont et al., 1996; Barr et al., 1999).

6.4. Origin of the dominant North-Northwest-trending fabric of the Minto Block

In the terrane accretion model originally proposedfor the Minto Block (Percival et al., 1992; Sternet al., 1994; Lin et al., 1996), the amalgamation ofseveral distinct oceanic and continental arcs gener-ated N-NW-striking foliations and steeply dippinglineations at, or near, the metamorphic peak. In therevised model (Percival and Skulski, 2000), a firstamalgamation event at ca. 2.77 Ga (Fig. 2) createda composite cratonic basement, while collision andoverthrusting of the Tikkerutuk arc onto the westernand central Minto (2.70–2.69 Ga) thickened the crustand produced the dominant N-NW fabric throughtransposition and recrystallization synchronous withmetamorphism.Percival and Skulski (2000)alsosuggested that subsequent exhumation and erosionremoved most of this overthrust terrane, except forthe topmost lavas in the Vizien belt.

Terranes accreted to a margin or entrapped betweencolliding blocks in Phanerozoic orogens typicallyretain distinctive lithological, deformational, andpressure–temperature histories, with terrane bound-aries being recognizable either as high-strain zones,or zones where oceanic crustal fragments and ac-cretionary prisms occur (e.g.Williams and Hatcher,1983; Coney, 1989; Bluck and Dempster, 1991;Monger, 1993; Huang et al., 2000). Preservation ofterrane boundaries in the Proterozoic Grenville oro-gen (Rivers, 1997; Indares et al., 2000) suggests thatthese boundaries should still be recognizable, despitedeep burial and intense deformation. The limitedlithological variations, and apparent lack of tectonicterrane boundaries in the Minto Block, suggests to usthat this was not an accretionary orogen.

In the context of a collisional or accretionary oro-genic scenario, homogeneous and orthogonal ENE-WSW compression of the Minto Block would berequired to generate a uniform, N-NW-striking grain,with extensive crustal thickening to generate sub-vertically-plunging lineations. To keep fabrics uni-form during repeated terrane accretion, each terrane

must have traveled in exactly the same direction asthe preceding one, with near-orthogonal convergencein every case. While not impossible, this appearstoo fortuitous. Finally, the common preservation ofsyn-magmatic fabrics in all Minto Block domainssuggests that amalgamation must have occurred whileeach terrane was still magmatically active, a conclu-sion at odds with the proposed very young age of theD2 event.

In their revised model,Percival and Skulski (2000)proposed that metamorphic effects (seen mostlyin supracrustal rocks) decrease eastward from agranulite-facies peak at 8 kbar and 850◦C, down tothe amphibolite facies in the Goudalie domain, withonly weak evidence of metamorphism in the Utsalikdomain (zircon overgrowths). Comparison with otherexamples of crustal thickening and exhumation (e.g.Frisch et al., 2000; Streepey et al., 2000; Hynes, 2002)suggests that prominent exhumation-related struc-tures should separate structural panels with distinctpressure–temperature histories. Furthermore, the east-ern limit of the reactivated orogen should correspondto a marked contrast in metamorphic grade, structuralstyle, and depth of exhumation, and be characterizedby prominent structures (thrust-and-fold belt and/ormajor strike-slip faults). However, recent mapping inthe Minto Block has not discovered evidence of suchexhumation-related structures, or of an eastern struc-tural boundary to the deformation and metamorphismcaused by a Tikkerutuk accretion event. Finally, asnoted previously, textures of plutons in the centralMinto Block (Bédard, 2003) do not seem to recordthis orogenic event.

6.5. Vertical tectonic models?

The preceding discussion underscores the problemsinvolved in applying uniformitarian plate-tectonicmodels to Archaean craton formation. Could verticaltectonic models provide an alternative framework inwhich to interpret our observations? Greenstone beltsin many Archaean cratons occur as anastomizing syn-forms between large domiform TTG bodies, a patternoriginally interpreted in terms of diapiric ascent ofgranitoids into pre-existing or pene-contemporaneoussupracrustal sequences (e.g.Macgregor, 1951;Blackburn, 1981; Hickman, 1983; Ayres and Thurston,1985). These vertical tectonic models fell into disfavor


principally because repetition of supracrustal pack-ages, and the presence of nappe- and thrust-likestructures, were interpreted as being due to marginalaccretion processes (Myers, 1976; de Wit, 1982;Williams, 1990; Kusky and Kidd, 1998; Polat andKerrich, 1999; the ‘D1’ event ofPercival and Skulski,2000); while the ‘dome-and-basin’ structures werereinterpreted as fold interference patterns (Snowdenand Bickle, 1976; Drury et al., 1984; Myers andWatkins, 1985).

Vertical tectonic models have been revived becausemany studies have revealed a general absence of fab-rics compatible with fold interference in dome cores(Schwerdtner et al., 1979; Chardon et al., 1996, 1998;Choukroune et al., 1995, 1997; Collins et al., 1998),and showed that there are no structural repetitionsin some of these so-called ‘accretionary’ sequences,many of which are now recognized as being essen-tially autochthonous (Chardon et al., 1998; Bleekeret al., 1999; Van Kranendonk, 2001; Thurston, 2002;Ayer et al., 2002; Van Kranendonk et al., 2002,2003, this issue). Many objectors to vertical tectonicmodels have invoked the absence of radial ‘diapiric’structural patterns in the domal massifs. However,natural and experimental diapirs show complex in-ternal structures, particularly where viscosity con-trasts between diapir and host are small (Jackson andTalbot, 1989). In any case, petrologic studies implythat ascent of material within TTG ‘domes’ mostly in-volved movement of magma in large conduits (Ayreset al., 1991; Collins et al., 1998; this paper), so thattheir lack of resemblance to simple diapiric structuresshould come as no surprise (see alsoVan Kranendonket al., 2003, this issue; andPawley et al., 2003,this issue).

Proponents of vertical tectonic models arguethat structures in the synclinal troughs, including‘accretionary’ or thrust faults, can also be explainedas a consequence of the complex flow pattern ex-pected in a downwelling limb (e.g.Bouhallier et al.,1995; Chardon et al., 1996, 1998; Collins et al.,1998; Van Kranendonk, 2001; Van Kranendonk et al.,2003, this issue). In this context, it seems proba-ble that complex overprinting structures and intensecompression would develop as supracrustal packagesflowed into convergence zones above downwellinglimbs (Fig. 10; cf. Dixon and Summers, 1983), andprominent decollement structures would be expected

at the base of the migrating supracrustal sequences(e.g. Chardon et al., 1996). Further increments ofdeformation might be expected as material floweddown into constrictions between impinging upwellingzones, and as sheath folds developed in the deeper,ductile portion of the downwelling limbs.

We now discuss how vertical tectonic modelscould be used to interpret the formation of fabricsin the Douglas Harbour rocks. The steep fabricsand near-ubiquitous∼2–10 km-scale ‘swirly’ foli-ation patterns found within the TC, QC, and FTC(Figs. 4 and 5) can be interpreted either as: (1) aninterference pattern generated by near-orthogonalcompressional deformation events (horizontal tec-tonics); or (2) as a section through steeply-plungingL-tectonite or diapiric bodies (vertical tectonics). Theintimate interpenetration and similar fabric orienta-tions of felsic plutonic and supracrustal rocks impliesthat both were deformed together, with many ofthe plutonic facies recording a supra- to sub-soliduscontinuum. The absence of orthogonal, overprintingcleavages, and the preservation of syn-magmatic tex-tures in most plutons, are difficult to reconcile withsuperimposed, mutually perpendicular, orogenic de-formation events, and so we favor a vertical tectonicsmodel.

In a vertical tectonic model, repeated deformationof supracrustal rocks would occur as supracrustalpackages migrated towards, and then into, the zone(s)of downwelling (Fig. 10B and C). The steeply dippingand plunging syn-magmatic fabrics in the plutonicrocks, on the other hand, could originate in a varietyof ways. Given the evidence for rapid melt ascent inlarge conduits (resorbed epidote), and for commonsyn-magmatic deformation in Douglas Harbour intru-sions, it seems probable that most of these fabrics wereacquired deep within the crust. Newly emplaced intru-sions would tend to concentrate strain (down-dip shear,Fig. 10B, ‘a–b’) between downwelling and upwellingzones, since in rheological terms they would be theweakest part of the crust (Pavlis, 1996; Vauchez et al.,1997; McCaffrey et al., 1999). Subsequent intrusionswould create new weak zones, shifting the locus ofmaximum strain and preserving the fabrics devel-oped in older intrusions (Schwerdtner et al., 1979).Alternatively, some of these fabrics may have beenacquired due to magmatic overpressure from youngerintrusions emplaced into partially consolidated older



plutons (cf.McNulty et al., 1996). A more radical hy-pothesis involves slow convective overturn of the en-tire lower crust (Fig. 10B and C), which is supportedby fluid dynamic models that imply bulk convectionof the largely molten lower part of the crust duringthe peak magmatic flux (Ridley and Kramers, 1990;Ridley, 1992). However, convective overturn need notbe complete (Collins et al., 1998; Van Kranendonket al., 2003, this issue), and only modest amountsof shear might be sufficient to develop the observedfabrics. Detailed mapping and structural studies areneeded to verify these suggestions and discriminatebetween mechanisms.

We suspect that the broad domical structures of theTroie and Quimussinguat complexes (Figs. 3 and 5)formed in a late warping event, since the presence ofshear zones at the margins of the domes indicates thatthe crust was rather stiff and cold at this time. Concen-trations of granite at sheared contacts to the TC andQC domes suggest that the bounding structures alsoguided the ascent of late K2O-rich melts and fluids(cf. Smit and van Reenen, 1997; Kramers et al., 2001).The zone of E-W-trending Archaean fabrics that sepa-rates the TC and QC (Thury Deformation Zone) couldrepresent a crush zone related to upward movement of

�

Fig. 10. Vertical tectonics scenario for the Douglas Harbour domain, which may also be applicable to the central and western Minto Block.Dotted line is sea level. Note the progressive horizontal extension of the crust from (A) to (C). (A) Extensive mantle melting producesthick oceanic plateaux. Iceland or Ontong-Java would be good analogues (e.g.Maaløe, 1982; Kröner, 1991; Tejada et al., 2002). Collisionsbetween plateaux may cause local thickening and compressional tectonics (D1 ofPercival and Skulski, 2000). Different mantle domainsyield lavas (tholeiites and komatiites) with distinct trace element and isotopic signatures, such as volcanic terranes 1 and 2 (VT1 and VT2).Small dark arrows indicate movement of mafic–utramafic melt out of the mantle residuum. Continued under- and intra-plating by basaltand komatiite (UKT) eventually causes the base of the lava plateau to melt, yielding a first generation (TTG1) of tonalite/trondhjemitemagmatism. Inheritance from the lavas (VT1 vs. VT2) gives the resulting TTGs distinct isotopic and trace element signatures. (B)Underplated cumulates (UKTC) and restites delaminate (Jull and Kelemen, 2001; Zegers and van Keken, 2001), allowing fresh mantle towell up, decompress, and generate another major pulse of mafic–ultramafic magmatism, much of which underplates/intraplates the crust.As a result, TTG1 (and host rocks) melt to give a second generation of TTG magmas. When the proportion of TTG2 magma reachesa critical threshold, the overlying volcanics become unstable and density-driven convective instabilities develop. Compressional forces inthe downgoing limbs cause folding and thrusting that may correspond to some of the D1 events seen in Minto Block supracrustal rocks.Ascent of TTG magma is dominantly as intrusions (‘a’ and ‘b’), which focus shear between the downwelling limb and the ascending TTGintrusions, producing the steep D2 fabrics. Diapiric ascent (‘d’), and thermally-driven convective instabilities (‘c’) may also occur, bothof which would contribute to generation of steep, D2-like fabrics. (C) Zones where underplated cumulates and restites delaminate anewreceive a fresh pulse of mafic–ultramafic magmatism, generating a third TTG event which may correspond to domains dominated by theLeaf River suite (2.73 Ga, Utsalik, Lac Minto). Instabilities such as those illustrated in part B (a–d) would be repeated, reinforcing theD2 structural overprint. Domains where the underplated material does not, or only partly, delaminates would only develop minor amountsof late magmas that could fill ‘tensional’ fractures or shear zones (e.g. Troie complex monzonitic suite), with lower magmatic fluxesgenerating only a broad ‘domal’ movement, rather than pervasive vertical instabilities. Note that the supracrustal relicts in TTG1 and TTG2have been omitted for clarity. The central trough structure includes subsiding lavas and sediments, as well as intruded, or adjoining, olderTTGs, and would represent the older, more tectonized domains (Goodalie, Faribault–Thury).

these two domical masses (cf.Bouhallier et al., 1995;Chardon et al., 1996, 1998; Collins et al., 1998; VanKranendonk, 2001).

Given that fabrics developed in the Douglas Har-bour domain plutons (Figs. 4 and 5) were originallysyn-magmatic, then they must reflect the stress fieldthat existed at the time of intrusion. Since this N-NWgrain is shared by the western FTC (2.85–2.87 Ga), theeastern FTC (2.8–2.77 Ga), the enderbites of the TCand QC (2.73–2.74 Ga), and the TC monzonitic suite(2.69 Ga), it seems as though this stress field persistednearly unchanged in the Douglas Harbour domain foralmost 200 million years. What was the nature of thisstress field, how was it produced, and how could it re-main so uniform, so long? As pointed out previously,Phanerozoic accretionary orogens commonly exhibitstrike-slip faults and shallowly-dipping fabrics, fea-tures which appear to be inevitable consequences ofnon-orthogonal subduction and terrane accretion. Con-sequently, it is difficult to account for the long-liveduniformity of the Douglas Harbour stress field (and itsfabrics) in terms of multiple compressional orogenicpulses associated with docking of a variety of terranesagainst a proto-cratonic core. On the other hand,differential vertical movement (shear) in an overall


extensional stress field would explain the absence ofshallowly-dipping thrusts and transpressive structures(cf. Barros et al., 2001) and simplifies the space prob-lem associated with generation and emplacement ofthe huge volumes of magma involved. Although mostPhanerozoic Cordilleran-type batholiths are unde-formed (Zen, 1985), emplacement-related structuresare locally prominent, preserving sub-vertical fabrics(e.g.McNulty et al., 1996; Miller and Paterson, 2001),some of which have been related to emplacementof successive sub-vertical sheets during a magmaticextensional phase (Campbell-Stone et al., 2000). Inprinciple, the continental magmatic extension scenarioproposed here (Fig. 10) resembles seafloor-spreading,where the crust forms by repetitive intrusion per-pendicular to the principal extensional stress. Wespeculate that the driving force was extensional col-lapse of the weak continental crust (Rey et al., 2001;Liu, 2001), a scenario that would be favored by on-going mafic–ultramafic underplating and the partiallymolten nature of the lower crust (e.g.Sandiford, 1989;Fountain, 1989; Laube and Springer, 1998).

Thus, the bulk of the evidence is compatible withan origin of the dominant ‘D2’ fabric of the DouglasHarbour domain as a result of diachronous verticaltectonic processes associated with cratonization. Thebroad domical structures (TC and QC) could repre-sent later, cooler manifestations of the same processesacting on a more rigid crust (Fig. 10C, right side). Avertical tectonic model can account for the regionalextent and temporal persistence of the N-NW-trendingfabric, the common occurrence of steeply-plunginglineations, the rarity of shallowly-dipping planarfabrics, the syn-magmatic nature of the early de-formation, the scattered distribution of supracrustalbelts and their pluton-driven metamorphism. We hy-pothesize that vertical tectonic models might explainmany of the problematic features of the western andcentral Minto Block also. In this context, changesin the locus of peak magmatic flux in the MintoBlock between 2.9 and 2.7 Ga (Leclair et al., 2001b)would reflect shifts in the locus of convective insta-bility, rather than the presence of subduction zones.This hypothesis is also consistent with the provin-ciality of intrusion and inherited zircon ages (Leclairet al., 2001b), and explains why domain bound-aries are principally intrusive contacts, rather thantectonic ones.

7. Conclusions

Archaean tonalite, trondhjemite, enderbite, gran-odiorite and granite of the Minto Block, northeast-ern Superior Province, have been interpreted as adeeply-eroded collage of Andean and intra-oceanicarcs, with the dominant N-NW-trending foliationresulting from orogenic transposition and metamor-phism associated with terrane accretion. However,outcrop scale structures and microstructures im-ply that deformation was initially syn-magmatic inmost Minto Block plutons. Magmatic muscoviteand epidote grains have resorbed edges that sug-gest ascent of magma as crystal slurries from higherpressures. Thermobarometry indicates final equi-libration at ∼4−6 kbar during near-solidus defor-mation, with several intrusive phases preservinghigh-temperature, near-liquidus signatures. Preserva-tion of syn-magmatic fabrics and of fossil liquidustemperatures in the plutons is probably not consis-tent with development of the dominant, penetrative,N-NW fabrics of the Minto Block during a late,orogenic, tectono-metamorphic overprint. Further-more, the broad uniformity of these high-temperaturefabrics in rocks of widely different ages, the over-all lithological similarity of rocks from differentMinto Block domains, and apparent absence of ma-jor domain-limiting structures, also seem inconsistentwith terrane accretion models. Instead, we infer anoverall extensional regime, with dense supracrustalassemblages and associated subsolidus intrusions sub-siding as cold fingers, while felsic magmas ascendedinitially as syn-kinematic intrusions. Shear betweenascending and descending limbs, or partial, possiblycyclical, convective overturn, would have generatedthe dominant steeply-dipping and plunging fabrics.Broader domical structures would have developed asthe crust stiffened, with bounding shear zones guidingascent of late granitoids and fluids.

Acknowledgements

This study was funded by the Geological Surveyof Canada (contribution # 2002-145) and Ministèredes Ressources Naturelles du Québec (contribution #2002-5130-04). Patrice Rey, Steve Sheppard, AlainLeclair, Léopold Nadeau and John Percival provided


many valuable comments and kept us on the straightand narrow path. Marc Choquette assisted with themicroprobe analyses. We also wish to thank the fieldcrews, pilots, mechanics, cooks and expediters whomade this research possible.

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Archaean cratonization and deformation in the northern …€¦ · the batholithic roots of Andean-type plate margins and intra-oceanic arcs. Existing horizontal-tectonic models propose