Geology, NATMAP Shield Margin Project Area Flin Flon Belt, Manitoba/Saskatchewan Accompanying Notes

Geology, NATMAP Shield Margin Project Area Flin Flon Belt, Manitoba/Saskatchewan accompanying notes

Geological Survey of Canada Map 1968A Manitoba Energy and Mines Map A-98-2 Saskatchewan Energy and Mines Map 258A Manitoba ~

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Manitoba ~ Energy and Mines ~S Geological Services

E.G. Syme\ S.B. Lucas2 , H. V. Zwanzig1, A.H. Bailes1, K.E. Ashtod, and F.M. Haidl 3

Saskatchewa1

EN'1'1il'rTIlOEPARTENARIATSl . Saskatchewan l.UPLorTATlONMINDAlEl990. I\l' PARTNERSHIP ~ ASSOCIATIm

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Saskatchewan Energy and Mines • Saskatchewan Geological Surve~ I Manitoba Geological Services Branch 2 Geological Survey of Canada 3 Saskatchewan Geological Survey Table of Contents

1 Introduction 2 Comments on map presentation 3 Trans-Hudson Orogen overview 5 Precambrian geology 5 Introduction 7 Metamorphism 9 Superior Boundary Zone (unit A) 10 Paleoproterozoic and Archean rocks in the Pelican Window (units A2, A3, A4) 10 1.92-1.87 Ga volcanic, intrusive, and sedimentary rocks (units J, F, U, E) 11 Juvenile arc assemblages (unit J) 16 Flin Flon arc assemblage 17 Snow Lake arc assemblage 17 Fourmile Island arc assemblage 18 West Amisk arc assemblage 18 Birch Lake arc assemblage 18 Hanson Lake arc assemblage 18 Kisseynew Domain 'south flank' 19 Ocean-floor assemblages (unit F) 19 MORB-like basalts (units Fl, F2, F4, FS) 20 Synvolcanic mafic to ultramafic complexes (unit F6) 20 Ocean-island basalts (unit F3c) 20 Ocean-plateau basalts (units F3a, b, d) 21 Unknown geochemical affinity (unit U) 21 Isotopically evolved Proterozoic and Archean rocks 21 Archean crustal fragments (unit AI) 21 Evolved arc (unit E) 22 Tectonic setting of arc and ocean-floor assemblages 22 Neodymium-isotopic and geochronological evidence 23 1.88-1.83 Ga intrusive rocks (unit P) 24 Geochemistry and source constraints 24 Rocks of uncertain age (unit W) 24 Sub-Phanerozoic Precambrian geology 27 1.87-1.83 Ga sedimentary, volcanic and intrusive rocks (units S, M, B) 27 Schist-Wekusko assemblage (unit S) 28 Missi Group (unit M) 29 Burntwood Group (unit B) 30 Late intrusive rocks (unit L) 30 Mineral deposits 30 Volcanic-hosted massive-sulphide deposits 31 Gold deposits 32 Tectonic evolution

32 D 1: Accretion

32 D2 : Postaccretion magmatism, sedimentation, and deformation 33 D3: Collision 34 D4-DS: Postcollision 34 Phanerozoic bedrock geology 34 Paleozoic geology 36 Mesozoic geology 36 Quaternary geology 37 References Geology, NATMAP Shield Margin Project Area Flin Flon Belt, Manitoba/Saskatchewan accompanying notes

INTRODUCTION

The regional compilation maps of the Paleoproterozoic Flin Flon Belt, southern Kisseynew Domain and interpreted sub-Phanerozoic basement (six sheets, 1:100 000 scale), and adjacent Phanerozoic cover (Fig. la, 1:325000 scale), represent synthesis products stemming from the NATMAP Shield Margin Pro­ ject. These maps are the culmination of efforts by the project partners (Manitoba Geological Services Branch (MGSB), Saskatchewan Geological Survey (SGS), Geological Survey of Canada (GSC» and build on the rich history of geoscience mapping in the Flin Flon Belt. A companion set of surficial geology maps has been produced at 1: 100 000 scale (see below). A preliminary CD-ROM containing digital maps and an integrated geoscience knowledge base was released in 1993 (Broome et aI., 1993; Broome and Viljoen, in press). The final synthesis maps, more detailed maps, and most other geological, geophysical, and geo­ chemical information sets pertinent to the project area are planned for release in digital format in a final project CD-ROM in 1999.

The National Geoscience Mapping Program (NATMAP) aims to foster a multidisciplinary team approach to bedrock and surficial mapping and related research, combine the efforts offederal, provincial and university scientists, and utilize digital information technology for more efficient interdisciplinary research and map publication. The Shield Margin Project, initiated in March 1991, was an inaugural NATMAP venture (Lucas et aI., in press a, b). Its purpose was to study the Paleoproterozoic rocks of the Flin Flon Belt and their con­ tinuation below gently dipping Paleozoic carbonates and the Quaternary cover. The region was chosen in part because of its economic significance: the Flin Flon greenstone belt is the largest Paleoproterozoic volcanic-hosted massive-sulphide (VMS) district in the world (e.g. Galley, 1996, and references therein), and also contains producing gold deposits.

The Shield Margin Project is closely tied to a variety of initiatives, including the Partnership Agreements on Mineral Development (1990-95) in Manitoba and Saskatchewan, the Snow Lake Exploration Technology (EXTECH) project, LITHOPROBE's Trans-Hudson Orogen Transect (THOT), the International Metamor­ phic Map program, and the Industrial Partners Program (with Hudson Bay Mining and Smelting Co.). The NATMAP project team has brought together over 50 participants from government surveys, universities (Calgary, Saskatchewan, Regina, Manitoba, Queen's, Ottawa, UQAM, Concordia, McGill, UNB) and the minerals industry.

1 COMMENTS ON MAP PRESENTATION

The Shield Margin Project map set includes six 1: 100 000 scale Precambrian Geology maps, representing a compilation from more than 70 sources ranging in publication date from 1944 to 1997 (index shown on Sheet 7). Synoptic maps of the Shield Margin area are included as Figures 1a and 1b, presented at a scale of 1:325 000. Figure 1a depicts the bedrock geology of the Precambrian Shield and Western Canada Sedimen­ tary Basin, whereas Figure Ib depicts only Precambrian geology (including the sub-Phanerozoic base­ ment). These figures are annotated with locations of geological and topographic features described in these notes.

The scale and degree of complexity depicted on the approximately 70 original maps varies considerably, with the result that some recently mapped areas (e.g. Flin Flon area, Snow Lake area, Kisseynew Domain) are shown in considerably more detail than areas that are represented by older maps. Considerable effort was made to produce a seamless geology coverage, but the limitations inherent in any compilation apply.

All of the Precambrian supracrustal rocks and most of the intrusive rocks in the Shield Margin area have been metamorphosed. The prefix 'meta' has been omitted (in the legend and in these notes) in the interests of brevity. Volcanic rocks are subdivided according to silica content (e.g. Stem et aI., 1995a) as follows: basalt

«52 wt. % Si02), basaltic andesite (52-57 wt. % Si02), andesite (57-63 wt. % SiO), dacite (63-70 wt. %

Si02) and rhyolite (70-80 wt. % Si02). In the legend, the term 'mafic' is equivalent to the basalt - basaltic andesite compositional range, 'intermediate' is equivalent to the basaltic andesite - andesite compositional range, and 'felsic' is equivalent to the dacite - rhyolite compositional range. These terms are generally used where the precise composition of the rock is unknown (e.g. for breccias and tuffs, or gneisses derived from volcanic rocks).

The major groupings of rocks in the legend are arranged in temporal order (with ages defined), but the divi­ sions within those major groupings are lithological with no temporal significance implied. Unit numbers are constructed using a 3--4 character alphanumeric code. The upper-case letter portion of the code refers to its highest level grouping (e.g. J -Juvenile arc rocks). The numeric portion represents the next level of subdivi­ sion, in many instances compositional (e.g. 11 - basalt. basaltic andesite). The final, lower-case letter is the most detailed subdivision (e.g. J 1a - tholeiitic basalt, basaltic andesite, gabbro, derived amphibolite). Litho­ logical and geochemical classifications of volcanic, sedimentary, and plutonic rocks follow standard usage and are defined in referenced literature for the belt (e.g. Bailes and Syme, 1989; Stem et aI., 1995a, b; Whalen et aI., in press). Note that some major subdivisions (e.g. arc (unit J) vs. ocean floor (unit F) volcanic assemblages) are based on a combination ofgeochemical and empirical lithological characteristics, summa­ rized below and discussed at length in the referenced literature.

The selection of structural information to display on the maps was governed primarily by the need to ensure clarity of the geologicallinework. Given the variety of sources and level of geological unit detail (the work­ ing scale was 1 :50 000), the NATMAP working group determined that the simplest and most practical solu­ tion would be to include only a representative set of stratigraphic facing-direction symbols. The complex deformational and metamorphic history in this part of Trans-Hudson Orogen precludes more detailed

2 structural presentation (for example, multiple generations of foliations and lineations or metamorphic min­ eral assemblages) at 1: 100 000 scale. Note that detailed structural data are shown on original source maps, and are planned to be included on thematic digital maps and databases on the accompanying CD-ROM.

A significant research direction during the NATMAP project was a program of regional mapping of the Phanerozoic-covered basement south of the Shield margin (Leclair et al., 1997), where the mineral-rich rocks of the FIin FIon Belt are covered by Phanerozoic platformal rocks. Mapping this buried terrane using new information from the exposed Shield was anticipated to highlight new areas for mineral exploration. In order to map the covered Precambrian rocks, high-resolution geophysical data from detailed aeromagnetic and gravity surveys were integrated with an extensive geological data set derived from study of basement drill core. The potential field data served to identify basement domains of distinct physical properties and to establish continuity between these domains and tectonic elements in the exposed Shield. The drill core data­ base provided'ground-truth' constraints for the interpretation of aeromagnetic and gravity anomalies. The integration of the geological and geophysical data sets, combined with petrography, U-Pb geochronology, and geochemistry, led to the recognition of distinct lithotectonic domains in the sub-Phanerozoic basement (Leclair et al. , 1997). The delineation ofkey elements ofthe exposed FIin FIon Belt has been extended south­ ward into the subsurface on the basis of this work.

The principle sources for the information summarized in these notes are listed in the references. The reader is urged to refer to these original works for detail on all aspects of the geology and mineral deposits of the area. These accompanying notes are intended only to summarize in a most general manner the many journal papers and government maps and reports which contributed to the project. They are necessarily interpretive in order to provide a framework for understanding the geology in a manner that goes beyond the simple distribu­ tion of units.

TRANS-HUDSON OROGEN OVERVIEW

The Paleoproterozoic Trans-Hudson Orogen (THO; Fig. 2; Hoffman, 1989; Lewry and Stauffer, 1990) extends from South Dakota, through the exposed Shield in Saskatchewan and Manitoba, across Hudson Bay to northern Quebec. The orogen is part of a greater 'Pan-American' Paleo- to Mesoproterozoic system whose evolution involved assembly of dispersed Archean minicontinents and accreted juvenile Paleopro­ terozoic terranes during the main episode ofNorth American continental assembly (Hoffman, 1989). Time­ space relations of lithotectonic elements in the Trans-Hudson Orogen and the 'Pan-American' system are similar to those in younger orogens formed by subduction/accretion/collision at convergent plate bounda­ ries. In the Saskatchewan-Manitoba segment, four major lithotectonic zones are recognized:

• Superior Boundary Zone, a narrow, southeastern, ensialic foreland zone bordering Superior craton, comprising the Thompson Belt, Split Lake Block, and Fox River Belt.

• The internal Reindeer Zone, a 400 km wide collage of Paleoproterozoic (1 .92-1 .83 Ga) arc volcanic rocks, plutons, volcanogenic sediments, and younger molasse, divisible into several lithostructural domains. Geochemical and Nd and Pb isotopic data indicate that most ofthese rocks evolved in an oceanic to transitional, subduction-related arc setting, with increasing

3 influence of Archean crustal components to the northwest. The Flin Flon-Snow Lake Domain, for example, is interpreted as an imbricated thrust wedge carried on a lower detach­ ment zone and overridden by higher grade Kisseynew gneisses (Lewry et ai., 1990; Lucas et aI., 1994, 1997). The Reindeer Zone overlies Archean basement exposed in structural win­ dows (Lewry et aI., 1990); this basement terrane is now termed the 'Sask craton' (Ansdell et aI., 1995).

• An Andean-type continental-margin, magmatic arc, represented by the Wathaman-Chipewyan Batholith emplaced at 1.86-1.85 Ga (Meyer et aI., 1992).

• A complexly deformed Northwestern Hinterland Zone, including the Peter Lake, Wollaston, and Seal River domains, and other parts ofthe Cree Lake Zone now included in Hearne Prov­ ince (Hoffman, 1989; Lewry and Stauffer, 1990).

, ' , , , :, ' :'. " : : :., " : : :" '. : : ' " " : '$as Phanerozoic .. ' . . ' , . ' ...... ' ...... ' .. .' .. ' . o Sedimentary rocks Paleoproterozoic o Sedimentary rocks Trans-Hudson Orogen o Continental arc + plutonic rocks Marginal basin/ O. collisional sedimentary and plutonic rocks

~ Arc plutons/ mixed gneisses Arc volcanic and plutonic rocks O~ ~ , Continental margin " " depOSits/reworked basement

o"I __100 , Archean km @ Archean cratons/ Pikwitonei Granulite Belt 52°25'N L...-______...... ;;...;.._____...... :.___...... u. Archean (exposed in Faults internal domains) ------LlTHOPROBE seismic reflection lines

Figure 2. Map ofthe Trans-Hudson Orogen, after Hoffman (1988), 'W'indicates location ofArchean base­ ment windows in the Reindeer Zone. FFB: Flin Flon Belt; GO: Glennie Domain HLB: Hanson Lake Block; LRD: La Ronge Domai; KD Kisseynew Domain; RD: Rottenstone Domain; TB: Thompson Belt; TF: Tabbernor Fault Zone; WB: Wathaman-Chipewyan Batholith; WD: Wollaston Domain.

4 PRECAMBRIAN GEOLOGY

Introduction

The Flin Flon Belt is a typical greenstone terrain and was once interpreted to be Archean in age (Harrison, 1951; Stockwell, 1961) based on its lithological, structural, and metamorphic similarities with greenstone belts in the Superior Province. The belt comprises polydeformed supracrustal and intrusive rocks, bounded to the north by metasedimentary gneisses of the Kisseynew Domain and to the south by flat-lying Paleozoic rocks of the Western Canada Sedimentary Basin.

The NATMAP Shield Margin Project and LITHOPROBE Trans-Hudson Orogen Transect built on an exten­ sive existing geological database to generate a much-improved understanding ofthe components and evolu­ tion of the southeastern Reindeer Zone, including the Flin Flon Belt (e.g. Lucas et al. 1996). These investigations have shown that, on the scale of the crust, the Flin Flon greenstone belt (and contained VMS deposits) is only one of three components in a northeast-dipping stack, juxtaposed during 1.84-1.80 Ga col­ lisional deformation (Fig. 3; Lucas et aI., in press b):

• at the lowest structural level (exposed in the Pelican Window): metaplutonic rocks and paragneisses (3.20-2.40 Ga) of the 'Sask craton'.

• at intermediate structural levels: Flin Flon Belt (now defined to include the Attitti Block and Paleoproterozoic rocks in the Hanson Lake Block) and Glennie Domain (together compris­ ing the 'Flin Flon-Glennie Complex'; Lucas et al. 1997).

• at the highest structural levels: marine turbidites (Burntwood Group; 1.85-1.84 Ga) and partly coeval distal facies of alluvial-fluvial sandstones (Missi Group) in the Kisseynew Domain.

Despite its location within a crustal-scale stack, the Flin Flon Belt contains a remarkably well preserved record of its earlier magmatic and tectonic history, crucial information to constrain the setting of contained mineral deposits in time and space.

Historically, the stratigraphy of the Flin Flon Belt has been described in terms of two stratigraphic groups, Amisk Group volcanic rocks and Missi Group continental sedimentary rocks (Fig. 4, top; Bruce, 1918; Harrison, 1951). The Flin Flon Belt, and in particular the Amisk Group, is now recognized to be a collage of distinct tectonostratigraphic assemblages that was assembled prior to the emplacement ofvoluminous granitoid plu­ tons and regional deformation related to the ca. 1.8 Ga Hudsonian Orogeny (Fig. 4, bottom). This is the basis for Lucas et al. (1996) terming the tectonic entity between the Sturgeon-weir River and Reed Lake as the 'Amisk collage', and for rejecting Amisk Group as the term to describe the 1.92-1.87 Ga volcano-plutonic rocks. In essence, the Amisk Group does not form a stratigraphic group in any sense of the term.

'Tectonostratigraphic assemblage' as used for the Shield Margin Project is not necessarily equated with 'terrane', nor is it implied that each assemblage is a fragment of a unique plate. However, each tectonostratigraphic assemblage does represent a distinct package of rocks in terms of its stratigraphy, geochemistry, isotopic

5 0')

3D Model of the Flin Flon Belt (Manitoba & Saskatchewan) (NATMAP Shield Margin Project, LITHOPROBE Trans-Hudson Orogen Transect) I

. ~~- ...... _.- -­ ~ I

Digital elevation models of the Quaternary, Phanerozoic and Precambrian surfaces (>2000 drillhole intersections)

I o ~ o ~~ j

_50000 L ' ( ~ 0'" 0'1)

~50000

:0\0 0'00 0 0 "?;,.

Precambrian crustal structure from seismic reflection profiles

Figure 3: Three-dimensional images of the NATMAP Shield Margin Project area from the LlTHOPROBE Trans-Hudson Orogen Transect (Lucas et aI., 1994). Conventional stratigraphy of signature, age, and inferred plate-tectonic the Flin Flon Belt setting (see below; Lucas et aI., 1996). The 1.92-1.87 Ga tectonostratigraphic assem­ Missi GP. blages recognized during the Shield Margin -1.845 Ga Project include those with juvenile arc, juvenile ocean-floor (including MORB-like basalt, ocean-plateau basalt and ocean­ Amisk GP. island basalt) and evolved arc affinities -1.9 Ga (Fig. 5; Syme and Bailes, 1993; Stem et aI., 1995a, b; David and Syme, 1994; Lucas et

'. _, ;;Ol~a~ic' }~ sandstone~ aI., 1996). This has important economic "." .; rocks :' ' conglomerate implications as not all of the tectonostra­ tigraphic assemblages are equally endowed New tectonostratigraphy of the Flin Flon Belt with mineral deposits. For example, all of the mined volcanic-hosted massive­ sulphide (VMS) base-metal deposits in the Flin Flon Belt are associated with the juve­ nile arc volcanic rocks (Syme and Bailes, 1993). Thus, knowledge of the physical and geochemical characteristics of the assem­ blages is crucial for effective·mineral explo­ ration in the Flin Flon Belt.

Amisk Collage Metamorphism

On the exposed Shield, peak regional meta­ 1 morphism at 1.82-1.80 Ga (David et aI., 1996) formed mineral assemblages in Flin Flon Belt rocks that range from prehnite­ pumpellyite to middle amphibolite facies in the east and upper amphibolite facies in the north and west (Froese and Gaspirini, 1975; Bailes, 1980a, b; Gordon, 1989; Bailes and Syme, 1989; Digel et aI., 1991; Ashton and Figure 4: Upper: Cartoon showing the conventional stratigra­ Digel, 1992; Digel and Gordon, 1995; phy of the Flin Flon Belt (Bruce, 1918; Harrison, 1951). Maxeiner et aI., 1995; Kraus and Menard, 1997; Menard and Gordon, in press). Lower: Cartoon showing the tectonostratigraphic framework for the NATMAP Shield Margin Project area (Lucas et aI., 1997). A general northward-increasing gradient is interrupted in the west by transitional upper amphibolite- to granulite-facies assemblages, marking a metamorphic culminationinrocks of the Pelican Window (Ashton et aI., in press).

7 co

Elbow-Athapap ocean floor assg. PHANEROZOIC 20km H.V. Zwanzia 1998

PRE-ACCRETION ASSEMBLAGES (1.87-1.92 Ga) SUCCESSOR-ARC and BASIN DEPOSITS _ Juvenile-arc and undivided metavolcanic rocks .c.. Missi Group (1.83-1.85 Ga) [B Ocean-floor (back arc) metabasalVsynvolcanic mafic intrusive Iwoi3X~ Continental sandstone / volcanics ~ Ocean-plateau metabasalt * Ocean-island metabasalt c=J Burntwood Group turbidites (1.84 - 1.85 Ga) [[[[[II] Tectonite L:~:~< I Schist-Wekusko Suite (1.85-1.88 Ga) FELSIC-MAFIC PLUTONS PELICAN WINDOW GNEISSES 1.76 - 1.82 Ga (Kisseynew Belt plutons) Archean charnockite ~ ~ ~ 1.83 - 1.84 Ga (late successor-arc plutons) Orthogneiss and pelitic gneiss ~ Ell! ~ 1.84 - 1.90 Ga (early juvenile-arc + early-middle [;iii VMS deposit successor-arc plutons) ~ Au deposit " FAULT ~ ca. 1.92 Ga ('evolved-arc' plutons) Sillimanite isograd

Figure 5: Tectonic assemblages in the NATMAP Shield Margin Project area, highlighting pre-accretion tectonostratigraphic assemblages and postaccretion (successor-arc) plutons, volcano-sedimentary basins and faults (Bailes and Syme, 1989; Syme and Bailes, 1993; Syme et al., 1995; Reilly et al., 1994; Lucas et al., 1996;Zwanzig, 1996, in press). B: Birch Lake arc assemblage; F: town of Flin Flon; FMI: Fourmile Island arc assemblage; ML: Mystic Lake 'evolved-arc' assemblage; S: town of Snow Lake; SB: Sandy Bay ocean-plateau assemblage. There is no obvious metamorphic break between rocks of the Flin Flon Belt and the middle to upper amphibolite-facies gneisses of the Kisseynew Domain (Bailes and McRitchie, 1978; Gordon, 1989; Gordon et aI., 1990; Nonnan et aI., 1995; Menard and Gordon, in press).

In Manitoba, the sillimanite-biotite isograd generally lies a few kilometres south ofthe northern limit ofpre­ dominantly volcanic-derived rock types, but in Saskatchewan it cuts across the Flin Flon Belt to the Sturgeon-weir River, where it bends southward and is obscured by Paleozoic cover (Fig. 5). Further west, it extends through Hanson Lake to the Tabbernor Fault Zone, where it bends sharply northward (Macdonald, 1981; Ashton and Balzer, 1995).

The cordierite-almandine-melt isograd (Bailes and McRitchie, 1978) lies within a few kilometres of the southern limit of purely sediment-derived migmatites and granites that constitute the central Kisseynew Domain. At least four generations of veining and granitic injection accompanied metamorphism on the south flank of the Kisseynew Domain; larger sheets and irregular plutons were intruded in the central part of the domain. Despite recrystallization during regional metamorphism, primary structures and textures are locally well preserved in both the Flin Flon Belt and Kisseynew Domain.

Metamorphic pressures range from less than 3 kbar in the prehnite-pumpellyite-facies rocks near Flin Flon (Digel and Gordon, 1995) to moderate values ofno more than about 6 kbar in the amphibolite-facies rocks to the north (Gordon, 1989; Bailes and McRitchie, 1978; Froese and Goetz, 1981; Digel et aI., 1991). Pressures approaching 7 kbar may have been reached in the vicinity the Pelican Window metamorphic culmina­ tion, which marks the deepest levels of crustal exposure (Digel et aI., 1991; Ashton and Digel, 1992; Ashton et aI., in press).

Superior Boundary Zone (unit A)

Archean (3.2-2.7 Ga) basement gneisses and narrow belts of Paleoproterozoic siliciclastic and mafic­ ultramafic igneous rocks (Moak Lake gneiss and Ospwagan Group; Bleeker and Macek, 1988) are exposed 10 km east of the Shield Margin map area. They are adjacent to the southeastern part of the Kisseynew Domain and extend southwest under the Paleozoic cover as unit A (undivided) on the basis of prominent aeromagnetic and gravity trends (Green et aI., 1985). Seismic profiles (Lucas et aI., 1994) indicate that the Superior Boundary Zone forms the highest structural element in an easterly dipping crustal stack, with the exposed boundary with the Kisseynew Domain comprising a steeply dipping sinistral reverse fault and mylonite zone (Setting Lake Fault Zone; Zwanzig, 1997). The reworked basement and the supracrustal rocks feature high-grade metamorphism and intense deformation related to continental collision. Siliciclas­ tic rocks of the Moak Lake gneiss and Ospwagan Group have an Archean provenance (Brooks and Theyer, 1981) and associated mafic igneous rocks formed in an extensional tectonic setting that is unrelated to any within the internal zones of Trans-Hudson Orogen.

9 Paleoproteroloic and Archean rocks in the Pelican Window (units A2, A3, A4)

The Pelican Window, which represents one of only three known exposures ofthe Sask craton, comprises leu­ cocratic quartzofeldspathic gneisses (unit AI), migmatitic paragneisses (unit A2), and the Mirond Lake Igneous Suite, which includes heterogeneous, dominantly enderbitic rocks, ranging from granite to gabbro in composition (unit A3), and the Sahli chamockitic granite (unit A4) (Ashton et aI., in press). Lithological contacts are generally transposed, but rocks of the igneou-s suite appear to crosscut contacts between the quartzofeldspathic gneisses and migmatitic paragneisses (Ashtol1 and Shi, 1994). The metamorphic grade is transitional between upper amphibolite and granulite facies, with the latter essentially restricted to the Mirond Lake Igneous Suite (Ashton et aI., in press).

The leucocratic quartzofeldspathic gneisses (unit AI) contain 20-50% leucosome and generally less than 10% combined biotite±homblende. Based on their mineralogical and geochemical composition, most have been interpreted as calc-alkaline, arc-derived orthogneisses, although a magnetiferous variety containing elevated immobile elements is thought to represent a minor sedimentary component (Shi, 1995). A mini­ mum age of 2959 ± 13 Ma has been established for the igneous component (Ashton et aI., in press), which is consistent with previous 3.120-2.840 Ga depleted-mantle model Nd ages (Sun et aI., 1993).

The migmatitic paragneisses (unit A2) are graphitic gamet-cordierite-sillimanite rocks which have been interpreted as aluminous wackes (Ashton and Shi, 1994). Rare impure forsterite-phlogopite-spinel­ corundum-graphite-opaque-calcite-dolomite marble layers are thought to have originated as dolo stones.

The heterogeneous enderbitic rocks (unit A3) and Sahli chamockitic granite (unit A4) of the Mirond Lake Igneous Suite are variably retrogressed two-pyroxene-gamet granulites with no partial melt component. They display tholeiitic, within-plate chemical affinities and were emplaced into the quartzofeldspathic gneisses and migmatitic aluminous wackes at about 2.450 Ga (Ashton et aI., in press), although a model Nd age of3.280 Ga from the Sahli granite (Sun et aI., 1993) suggests earlier mantle derivation. A dioritic to gab­ broic component of the igneous suite mainly occurs as boudinaged mafic dykes up to 30 m thick, one of which has yielded a date of 2488 ± 12 Ma (Ashton et aI., in press). The Mirond Lake Igneous Suite is tenta­ tively thought to be part of a much more extensive, worldwide, ca. 2.450 Ga magmatic event. This 'Matach­ ewan Igneous Event' was characterized by mafic and associated magmas, which appear unrelated to any locally defined orogenic activity, and has been attributed to the breakup of a large Archean supercontinent (Heaman, 1997).

1.92-1.87 Ga volcanic, intrusive, and sedimentary rocks (units J, F, U, E)

Supracrustal rocks ranging in age between 1.92 and 1.87 Ga constitute most of the Flin Flon greenstone belt (Fig. 5). These 1.92-1.87 Ga greenstones vary in lithological and geochemical associations, in Sm/Nd iso­ topic signatures, and in U-Pb zircon ages, allowing subdivision into a series of arc and ocean-floor assem­ blages (Syme and Bailes 1993; Stem et aI. 1995a, b; Lucas et aI. 1996) which vary in potential to host the

10 Geological Survey of Canada Map 1968A Manitoba Energy and Mines Map A-98-2 Saskatchewan Energy and Mines Map 258A

Geology, NATMAP Shield Margin Project Area Flin Flon Belt, Manitoba/Saskatchewan accompanying notes

E.C. Syme, S.B. Lucas, H.V. Zwanzig, A.H. Bailes, K.E. Ashton, and F .M. Haidl

This table replaces Table 1: "Selected U-Pb geochronological information for the NATMAP Shield Margin Project area.", p. 12-15. ~ . Table 1: Selected U-Pb geochronological information for the NATMAP Shield Margin Project area.

Age Error UNIT Unit Rock Type Type (Ma) (Ma) Method Mineral Reference UTM-E UTM-N Zone No. Jan Lake Complex Quartzofeldspathic igneous 2959 ±13 U/Pb zircon Ashton et al. (in 641870 6101120 13 A2 gneiss press) Missi Group Sandstone lens in detrital 2721 ±4 Pb-Pb zircon Ansdell (1993) 323173 6055822 14 MIa conglomerate

Stroud Felsic Rhyolite breccia inherited 2715 Pb-Pb zircon David et al. (1996) 427600 6074000 14 J7a Breccia Mystic Lake Tonalite Tonalite inherited 2604 Pb-Pb zircon Stern et al. (in Ela press)

Wolf Lake Turbiditic Turbiditic wacke detrital 2523 ±12 Pb-evap zircon Ansdell and 663224 6053536 13' J9b Wacke Connors (1994)

Beaverhouse Lake Tonalite igneous 2518 +7/-4 U/Pb zircon David and Syme 327453 6069129 14 Ala Tonalite (1994)

Felsic Mylonite. NE Granodiorite igneous 2497 +3/-2 U/Pb zircon David and Syme 322420 6061324 14 Ala Arm Shear Zone (1994) Missi Group; Flin Sandstone (lower detrital 2494 ±2 U/Pb zircon Ansdell (1993) 702154 6072398 13' M3a Flon Basin; sequence) Beaverdam Member

Jan Lake Complex Gabbro igneous 2488 ±12 U/Pb zircon Ashton et al. (in 639800 6109140 13 Ala press)

Jan Lake Complex Chamockite igneous 2450 ±8 UlPb zircon Ashton et al. (in 636290 6104700 13 A4 press)

Burntwood Group Greywacke detrital 2397 Pb-Pb zircon David et al. (1996) 406450 6074500 14 Bla

Schist Lake Road Aplitic dyke inherited 2022 ±2 Pb-Pb zircon Stern et al. (1993) 315125 6058810 14 Pl0a Felsic Dyke

Jan Lake Complex Migmatitic detrital 2016 Pb-Pb zircon Ashton et al. (in 641390 6101560 13 A2 paragneiss press) Meridian Creek Tonalite igneous 1920 +18/-13 U/Pb zircon Stern and Lucas 309850 6067400 14 Ela Tonalite (1994)

Missi Group; Flin Conglomerate detrital 1914 ±6 Pb-evap zircon Ansdell et al. 697745 6076389 13" MIa Flon Basin (lower sequence) (1992)

Mystic Lake Tonalite igneous 1906 ±2 U/Pb zircon Heaman et al. 309165 6060575 14 Ela Intrusive Suite (1992)

Wo~ Lake Turbiditic Turbiditic wacke detrital 1906 ±20 Pb-evap zircon Ansdell and 663224 6053536 13' J9b Wacke Connors (1994)

Athapapuskow Diabase igneous 1904 ±4 U/Pb zircon Stern et al. 323550 6047350 14 F5a Diabase (1995b)

Mine Rhyolite Rhyolite igneous 1903 +7/-5 U/Pb zircon Stern et al. (in 314096 6073283 14 J4a press)

West Arm Tonalite Tonalite igneous 1903 +6/-4 U/Pb zircon Stern et al. (1993) 318378 6050352 14 Ela

Claw Lake Gabbro Gabbro pegmatite igneous 1901 +6/-5 U/Pb zircon Stern et al. 385286 6073221 14 F6a dyke (1995b)

Herblet Lake Gneiss Melanocratic gneiss igneous 1901 ±4 U/Pb zircon David et al. (1996) 445550 6102385 14 J14 Dome

Welsh Lake Greywacke detrital 1894 ±3 U/Pb zircon Heaman et al. 665131 6074326 13 J9b Assemblage (1993)

Stroud Felsic Rhyolite breccia igneous 1892 ±3 UlPb zircon Machado and 429129 6073267 14 J7a Breccia David (1992)

Herblet Lake Gneiss Granodiorite igneous 1890 +8/-6 U/Pb zircon Gordon et al. 443643 6089923 14 J14b Dome (1990)

Richard Lake Pluton Tonalite igneous 1889 +8/-6 U/Pb zircon Bailes et al. (1990) 426880 6081044 14 J12c

Bakers Narrows Porphyritic rhyolite igneous 1888 +11/-9 U/Pb zircon Stern et al. (in 330454 6067770 14 JI3c Rhyolite Dyke press)

Neagle Lake Rhyolite (?) igneous 1888 ±3 U/Pb zircon Heaman etal. 666713 6071812 13 J7b Rhyolite (1993)

Laurel Lake Rhyolite Rhyolite igneous 1887 ±3 U/Pb zircon Heaman etal. 676450 6065750 13 J4a • approximate location • I Table 1. (cont.)

Age Error UNIT Method Mineral Reference UTM·N Zone No. I Unit Rock Type Type (Ma) (Ma) UTM·E , (1993) Cliff Lake Pluton Quartz diorite igneous 1886 ±1 u/Pb zircon Stern et al. (in 317250 6076000 14 J12c press) Sneath Lake Pluton Tonalite igneous 1886 +17/-9 U/Pb zircon Bailes et al. (1990) 428914 6069820 14 J12c Two Portage Rhyolite crystal tuff igneous 1886 ±2 U/Pb zircon Gordon et al. 325860 6067880 14 J2a Rhyolite Crystal Tuff (1990) Vick Lake Tuff Shoshonitic tuff igneous 1885 ±3 Pb-Pb zircon Stern et al. (1993) 302670 6068240 14 J8a Herblet Lake Gneiss Melanocratic gneiss igneous 1884 ±6 U/Pb zircon David et al. (1996) 445550 6102385 14 J14 Dome Mystic Lake Tonalite-quartz igneous 1883 ±4 U/Pb zircon Heaman et al. 309110 6060240 14 Ela Intrusive Suite diorite (1992) West Amisk Andesitic tuff igneous 1882 ±3 U/Pb zircon Stern and Lucas 672055 6057596 13 J6a Andesitic Tuff (1994) Mikanagan Lake Sill Gabbro pegmatite igneous 1881 +3/-2 U/Pb titanite Stern et al. (in 327627 6071249 14 Jl0 press) Gants Lake Batholith Granodiorite igneous 1876 +7/-6 U/Pb zircon Whalen and Hunt 385740 6081260 14 P7a (1994) Puella Bay Dacite Dacite igneous 1876 ±2 U/Pb zircon Ansdell et al. (in 450960 6064440 14 S3b press) Hanson Lake Rhyolite igneous 1875 ±1 U/Pb zircon Heaman etal. 639500 6061000 13 J4a Rhyolite (1993) Jungle Lake Pluton Tonalite igneous 1874 ±2 u/Pb zircon Machado et al. (in 374075 6116600 14 P6a press) Finger Peninsula Granite igneous 1873 ±1.6 U/Pb zircon Heaman et al. 638130 6066280 13 J12a Sheared Granite (1994) Ragged Lake Pluton Granite gneiss igneous 1873 ±4 U/Pb zircon Hunt and Zwanzig 374303 6100053 14 P9a (1990) Hanson Lake Mine Granite igneous 1870 +7/-5 U/Pb zircon Heaman etal. 635740 6063500 13 J12a Road Granite (1994) Schist Lake Road Aplitic dyke igneous 1869 ±1 Pb-Pb zircon Stern et al. (1993) 315125 6058810 14 Pl0a Felsic Dyke Scheiders Bay Rhyolite crystal tuff detrital 1867 ±17 U/Pb zircon Stern et al. (in 327406 6059436 14 SIc Rhyolite Crystal Tuff press) Annabel Lake Pluton Granodiorite igneous 1866 ±3 U/Pb zircon Stern and Lucas 311450 6074850 14 P7a (1994) East Elbow Tonalite Quartz megacrystic igneous 1864 ±3 U/Pb zircon Whalen and Hunt 382630 6081518 14 P6a Stock tonalite (1994) Elbow Lake Pluton Tonalite igneous 1864 +5/-4 U/Pb zircon Whalen and Hunt 381225 6074488 14 P6a (1994)

Missi Group Cross-bedded detrital 1864 ±7 Pb-Pb zircon Ansdell et al. (in M3a sandstone press) West Arm Turbiditic sandstone detrrtal 1863 ±3 Pb-Pb zircon Stern et al. (in 320782 6071821 14 SIb Sandstone press) Pluton intruding Quartz-feldspar igneous 1861 +4/-3 U/Pb zircon Heaman et al. 639327 6065954 13 Pl0b Hanson Lake porphyry (1997) metasediments Missi Group; Flin Conglomerate detrital 1861 ±17 Pb-evap zircon Ansdell et al. 697745 6076389 13" MIa Flon Basin (lower sequence) (1992) Puffy Lake Tonalite Tonalite igneous 1860 ±2 U/Pb zircon Machado et al. (in 370150 6099500 14 P7e press) Batty Lake Intrusive Tonalite gneiss igneous 1859 +14/-7 U/Pb zircon Hunt and Zwanzig 391400 6114890 14 P8c Complex (1993) " approximate location \ . Table 1. (cont.)

Age Error UNIT Unit Rock Type Type (Ma) (Ma) Method Mineral Reference UTM-E UTM-N Zone No. Burntwood Group Greywacke detrital 1859 ±7 Pb-Pb zircon David et al. (1996) 442400 6075300 14 Bla Missi Group Cross-bedded detrital 1859 ±8 Pb-Pb zircon Ansdell et al. (in M3a sandstone press) Neso Lake Pluton Quartz diorite igneous 1858 ±3 UlPb zircon Syme et al. (1991) 332482 6059654 14 P5a Schist Lake Trachyandesite igneous 1858 ±4 U/Pb zircon Stern et al. (in 320374 6058788 14 Sla Conglomerate press) Schist Lake Qz-fp phyric rhyolite igneous 1858 ±1 UlPb zircon Stern et al. (in 320782 6058583 14 Sla Conglomerate cobble press) Missi Group; Flin Sandstone (lower detrital 1857 ±2 U/Pb zircon Ansdell (1993) 702154 6072398 13' M3a Flon Basin; sequence) Beaverdam Member East Wekusko Rhyolite igneous 1856 ±1 UlPb zircon Ansdell et al. (in 449180 6069140 14 S3a Rhyolite press) Kaminis Lake Pluton Granodiorite igneous 1856 ±2 U/Pb zircon Stern and Lucas 313750 6059750 14 P7a (1994) Mari Lake Pluton Biotite granodiorite igneous 1855 +4/-3 U/Pb zircon Heaman etal. 308025 6096090 14 P8b (1992) Kakinagimak Lake Biotite leucotonalite igneous 1852 +6/-4 U/Pb zircon Heaman et al. 669565 6104335 13 P8b Leucotonalite (1993) Reynard Lake Porphyritic igneous 1850 ±3 U/Pb zircon Stern et al. (1993) 689775 6062125 13 P7a Pluton granodiorite Wekach Lake Pluton Gabbro igneous 1850 ±2 UlPb zircon Heaman etal. 311400 6063775 14 P2a (1992) Missi Island Trondhjemite igneous 1848 ±11 Pb-evap zircon Ansdell and Kyser 678000 6064250 13 P8a Trondhjemite (1991)

Puffy Lake Dacite igneous 1848 ±4 UlPb zircon Machado et al. (in 372150 6100100 14 M5d Metadacite press) Lynx Lake Pluton GranodiOrite igneous 1847 ±4 U/Pb zircon Gordon et al. 333749 6052482 14 P7a (1990) Syn-Meridian-West Plagioclase­ igneous 1847 ±2 U/Pb zircon Stern and Lucas 309400 6069012 14 Pl0b Arm Shear Zone porphyry dyke (1994) Dyke MissiGroup Pebbly sandstone detrital 1847 ±2 Pb-Pb zircon Ansdell (1993) 317480 6073800 14 M2a Neagle Lake Pluton Granodiorite to igneous 1846 +14/-6 Pb/Pb titanite Ashton (1992) 667191 6073948 13 P6c quartz monzodiorite Missi Group Sandstone lens in detrital 1846 ±6 Pb-Pb zircon Ansdell (1993) 323173 6055822 14 Mia conglomerate Big Rat Lake Pluton Granodiorite igneous 1845 ±3 Pb-Pb zircon Whalen and Hunt 377307 6077068 14 P7a (1994) Hanson Lake Pluton Granodiorite igneous 1844 ±2 UlPb zircon Heaman et al. 642921 6064730 13 P7a (1993)

Hanson Lake Pluton Gabbro igneous 1843 ±2 UlPb zircon Heaman etal. 640833 6060206 13 P2a (1994) Pelican Quartz monzodiorite igneous 1843 ±2 U/Pb zircon Ashton et al. (in 652450 6086575 13 P9c Decollement Zone press) Boundary Intrusion Metagabbro igneous 1842 ±3 U/Pb zircon Heaman etal. 315770 6065500 14 P2g (1992) Burntwood Group Psammite detrital 1842 ±2 U/Pb zircon Machado et al. (in 390650 6116675 14 Bib press) Post-Meridian-West M icrocline-porphyry igneous 1839 ±3 U/Pb zircon Stern and Lucas 309475 6068919 14 Pl0b Arm Shear Zone dyke (1994) Dyke Boot Lake Intrusion Monzodiorite igneous 1838 ±2 U/Pb zircon Heaman et al. 315770 6065500 14 P4c (1992) , approximate location Table 1. (cont.)

Age Error UNIT Unit Rock Type Type (Ma) (Ma) Method Mineral Reference UTM-E UTM-N Zone No.

Phantom Lake Granodiorite igneous 1838 ±2 U/Pb zircon Heaman et al. 315770 6065500 14 P7a Granodiorite Dyke (1992)

Tramping Lake Granite igneous 1837 +8/-6 U/Pb zircon David et al. (1996) 436260 6072890 14 P9a Pluton J Missi Group Metasandstone detrital 1837 ±4 Pb-Pb zircon Machado et al. (in Mlb press)

Bujarski Lake Pluton Quartz diorite igneous 1836 +4/-3 U/Pb zircon Bailes et al. (1990) 425128 6067841 14 P5a

Chickadee Rhyolite Rhyolite igneous 1836 ±1 U/Pb zircon Ansdell et al. (in 452840 6071083 14 M5a I press) Wekusko Granite Granite igneous 1834 +8/-6 U/Pb zircon Gordon et al. 438981 6068838 14 P9a (1990)

Herb Lake Felsic Rhyolite igneous 1833 +6/-2 U/Pb zircon Ansdell et al. (in 450860 6071110 14 M5a Volcanic press)

Nelson Bay Gneiss Granodiorite gneiss igneous 1832 ±2 U/Pb zircon David et al. (1996) 413226 6085967 14 P7e Dome

Rex Lake Plutonic Granodiorite igneous 1832 +4/-3 UlPb zircon Gordon et al. 450113 6076491 14 P7a Complex (1990) Ham Lake Pluton Granodiorite igneous 1830 +27/-19 U/Pb zircon Gordon etal. 417475 6083464 14 P7a (1990)

Mirond Lake Homogeneous calc­ igneous 1830 ±1 U/Pb zircon Ashton et al. (in 645450 6103970 13 P5d Enderbite alkaline enderbite press) Touchbourne Enderbite igneous 1830 +11 /-5 U/Pb zircon Gordon et al. 414000 6130900 14 P5d Intrusive Suite (1990)

Little Swan Lake Granodiorite igneous 1828 ±6 Pb-Pb titanite Whalen and Hunt 338380 6073200 14 P9b Pluton (1994)

Missi Group Quartz porphyry 1 igneous 1826 +11/-5 U/Pb zircon Hunt and 353834 6097861 14 M6b pink gneiSS Schledewitz (1992)

Anderson Rhyolite Rhyolite metamorphic 1812 ±15 U/Pb titanite David et al. (1996) 438778 6078845 14 J4a

Hanson Lake Psammopelitic metamorphic 1808 ±2.6 U/Pb monazite Heaman etal. 639450 6066280 13 J9b Psammopelitic schist (1994) Schist

Pelican Quartz monzodiorite metamorphic 1808 ±3 U/Pb titanite Ashton et al. (in 652450 6086575 13 P9c Decollement Zone press)

Attitti Lake Felsic Felsic gneiss metamorphic 1807 +3/-2 U/Pb zircon Heaman et al 664360 6104205 13 Ulb Volcanic (volcanic ?) (1992)

Herblet Lake Gneiss Tonalite gneiss metamorphic 1807 ±3 U/Pb zircon David et al. (1996) 446780 6102510 14 J14a Dome

Pre-Sturgeon-weir Granitoid pegmatite igneous 1806 ±2 UlPb zircon Ashton (1992) 658530 6095230 13 L4 Shear Zone Pegmatite

Hanson Lake Mine Rhyodacite metamorphic 1804 +5/-4 U/Pb titanite Heaman et al. 634400 6064980 13 J4a Road Rhyodacite (1994)

Pegmatitic Granite Granite igneous 1799 ±2 U/Pb monazite Ansdell and 337814 6097681 14 Ll Nonman (1993)

Mystic Lake Porphyritic rhyolite igneous 1797 Pb-Pb zircon Heaman etal. 308910 6059950 14 L 1 Intrusive Suite dyke (1991) Pegmatite Dyke Pegmatite igneous 1796 ±1 Pb-Pb monazite Hunt and Zwanzig 388430 6116480 14 l1 (1993)

Kakinagimak Lake Biotite leucotonalite metamorphic 1789 ±3 U/Pb titanite Heaman et al. 669565 6104335 13 P8b Leucotonalite (1993)

Jan Lake Granite Aplitic-pegmatitic igneous 1773 ±9 UlPb zircon Bickford et al. 633396 6086198 13 L3a Suite granite (1987)

Post Sturgeon-weir Granitoid pegmatite metamorphic 1767 ±1 U/Pb monazite Ashton (1992) 653620 6091600 13 L4 Shear Zone Pegmatite I • approximate location

:t.-.-. Z) The tholeiitic rocks are similar to modem island-arc tholeiites, having low high-field'-strength-element (HFSE) and rare-earth-element (REE) abundances relative to MORE, and chondrite-normalized light REE depletion to slight enrichment (Stem et aI., 1995a). The calc-alkaline andesite-rhyolite and shoshonite are more strongly LREE-enriched and have comparatively higher HFSE abundances. These calc-alkaline and alkaline-series rocks have trace-element signatures (high Th/Nb, La/Nb) that are almost identical to those forming in modem intra-oceanic arcs (Stem et aI., 1995a). The extreme extent ofthe HFSE depletion exhib­ ited by arc tholeiites at Flin Flon is observed in the island-arc tholeiites of the Tonga-Kermadec arc (Ewart and Hawkesworth, 1987) and Fiji (Gill, 1987). For the most primitive Flin Flon arc assemblage rocks, these values are only 1-2.5 times their abundances in the estimated mantle source ofmodernN-MORBs, probably due to their derivation from a highly refractory (depleted) mantle source (Stem et al. 1995a). This is consis­ tent with the arc tholeiites representing primitive arc segments built on oceanic lithosphere. Neodymium isotopic and trace-element data indicate that the Flin Flon arc assemblage volcanic and plutonic rocks are

predominantly juvenile (i.e. positive initial ENd values of +2 to +5, sirnilar to the contemporaneous depleted mantle; Stem et aI., 1995a, b; Fig. 6).

The Flin Flon Belt contains six geographically sepa­ rate juvenile arc assemblages, each of which is FUN 20-50 km across (Hanson Lake, West Amisk, Birch +4 FLON Lake, Flin Flon, Fourmile Island, and Snow Lake MANTLE ------~~~77~!r assemblages; Fig. 1b, 5). These are separated by +2 Missi Group major faults or intervening ocean-floor rocks (unit (fluvial sediments) CHUR F), Burntwood Group turbidites (unit B), plutons o~------+~~~ (unit P), or a combination ofthese rock types. The arc Postaccretion Rocks (Successor Arc & Basins) Pre-Accretion assemblages are internally complex, comprising -2 Assemblages "C (Amisk Collage) numerous fault-bounded and folded volcanic suites Z c:: (e.g. Bailes and Syme, 1989), rendering correlation ..2 -4 Evolved 'iii Arc of volcanic stratigraphy within and between the 0. (Mystic Lake w Assem blage) assemblages nearly impossible. It is unclear whether -6 the segments represent the fragmented parts of a for­ Sask Craton merly single are, or were generated in completely -8 different arcs (e.g. Syme et aI., 1995; Lucas et aI., 1996).

-12 Postcollisional leucogranites Flin Flon arc assemblage and pegmatites -14 Flin Flon arc assemblage contains mostly mafic vol­ canic rocks that were deposited in a subaqueous envi­ 1770 1790 1810 1830 1850 1870 1890 1910 Age (Ma) ronment (Bailes and Syme, 1989). Basalt and basaltic andesite flows dominate the assemblage, and Figure 6: CNd VS. time plot for units from the Flin Flon Belt. Neodymium-isotopic data from Stern et al. (1992, 1993, belong mainly to tholeiitic (unit J1a) and calc­ 1995a, b, unpub. data), Whalen et al. (in press), Bickford alkaline (unit J1 b) suites (Stem et aI., 1995a, b). et al. (1992) and Mock et al. (1993) . Locally, the calc-alkaline rocks are stratigraphically associated with terrigenous turbidites (unit J9),

16 MORB-like rift basalts (unit J2), and turbidites derived solely from shoshonitic volcano( es) (unit J8) (Bailes and Syme, 1989; Stern et aI., 1995a), sequences attributed to episodes of arc rifting and the development of intra-arc basins (Lucas et aI., 1996; Syme et aI., in press). The majority of arc rocks contain primary struc­ tures indicating that they were deposited in a subaqueous environment, but there is clear morphologic evi­ dence (e.g. presence ofbubble-wall shards, pumice) principally in the younger calc-alkaline and shoshonitic sequences that resedimented pyroclastic rocks may have been erupted in a very shallow marine or subaerial setting (Bailes and Syme, 1989; Syme and Bailes, 1993). Synvolcanic intrusions form a calcic gabbro­ diorite-quartz diorite-tonalite series (Whalen et aI., in press) and occur as high-level, discrete sills, dykes, and plutons (Bailes and Syme, 1989), such as the 1.886 Ga Cliff Lake pluton (Stem, pers. comm., 1996). Stra­ tigraphic sequences are complex and typically display a wide variety of rock types with interfingering relationships, lenticular units, and abrupt facies variations.

The dominance of basaltic andesite and basalt in the Flin Flon arc assemblage contrasts with the apparently greater abundance of andesite and rhyolite in the West Amisk arc assemblage (below; Walker and Watters, 1982) and the greater proportion of felsic rocks in the Snow Lake arc assemblage (below). Note that these variations are not incompatible with the geochemical and isotopic evidence that all of the arc assemblages formed in a dominantly oceanic regime.

Snow Lake arc assemblage

The 1.892 Ga Snow Lake arc assemblage (David et aI., 1996) is similar to the Flin Flon segment except for more extensive hydrothermal alteration, a higher proportion of volcaniclastic rocks, and higher metamorphic grade (Bailes and Galley, 1996). Much of the Snow Lake assemblage is well preserved despite polyphase deformation and regional metamorphism from middle greenschist to middle amphibolite facies. The arc volcanic rocks were deposited under subaqueous conditions and, like the Flin Flon assemblage, the sequence also includes some material derived from shallow marine to subaerial pyroclastic deposits. The thick (>6 km) juvenile oceanic-arc sequence at Snow Lake records in its stratigraphy and geochemistry (Bailes and Galley, 1996) a temporal evolution in geodynamic setting from a primitive arc, to a mature arc, to a rifted arc. Shallow synvolcanic, multiphase tonalite intrusions (unit 112) and associated high-temperature alteration zones in the VMS-hosting primitive and mature portions of the Snow Lake arc assemblage (1.886 Ga Sneath Lake pluton, 1.889 Ga Richards Lake pluton; cf. Bailes and Galley, 1996) are interpreted to be the products of high heat flow and increased fluid circulation accompanying arc-rifting processes (Bailes and Galley, in press). The Snow Lake arc assemblage is bounded to the southwest and northeast by low-angle faults interpreted as thrusts, and by tectonic slices of the 1.84 Ga Burntwood Group turbidites (unit B1a).

Fourmile Island arc assemblage

Greenschist-facies volcanic rocks of the 5.5 km thick Fourmile Island arc assemblage occur on western Reed Lake (Syme et aI., 1995). They are separated from the Reed Lake mafic-ultramafic complex (unit F6) to the west by a wide zone of heterogeneous tectonite and sheet-like bodies of felsic-intermediate intrusive

17 rocks (unit W6b), and are bounded on the east by a fault-bound slice of Burntwood Group turbidites (unit B la). The Fourmile Island assemblage ranges from basaltic andesite to rhyolite in composition, and has trace-element characteristics similar to arc rocks elsewhere in the Flin Flon Belt (Syme and Bailes, 1996).

West Amisk arc assemblage

The West Amisk arc assemblage stratigraphy is marked bya lower tholeiitic sequence (unit 11 a), including a high-level, mafic to felsic volcanic complex interpreted as an emergent volcano (Ayres et aI., 1991), overlain by greywacke turbidites (unit J9; ca. 1.887 Ga; Heaman et aI., 1993) which interfinger with shallow-water felsic complexes (unit 17; 1.888-1.887 Ga; Heaman et aI., 1992, 1993), and which are in turn overlain by andesitic flows and volcaniclastic rocks (units J6, 13; 1.882 Ga; Stern and Lucas, 1994, 1995). The stratigra­ phy is attributed to an episode of arc rifting and the development of an intra~arc basin prior to the resumption of arc magmatism (Lucas et aI., 1996).

Birch Lake arc assemblage

The Birch Lake arc assemblage (Reilly et aI., 1994, 1995), host to the Konuto, Flexar, Birch, and Coronation VMS deposits, comprises massive plagioclase-phyric and amygdaloidal mafic flows with minor ash and lapilli tuffs. It has tholeiitic arc affinities and is probably a tectonic slice ofeither the Flin Flon or West Amisk assemblage.

Hanson Lake arc assemblage

Supracrustal rocks between the Sturgeon-weir River and Sarginson Lake represent an amphibolite-facies, juvenile arc assemblage (Maxeiner and Sibbald, 1995; Maxeiner et aI., 1995, 1996, in press; Slimmon, 1995) that shows evolution from primitive arc tholeiites to evolved calc-alkaline arc rocks (1.875 Ga rhyolite; Heaman et aI., 1993). Rocks at Hanson Lake comprise a mixed suite of subaqueous to subaerial, dacitic to rhyolitic and intercalated minor mafic volcanic rocks, overlain by greywackes. East of Sarginson Lake, the Northern Lights volcanics comprise tholeiitic, arc, pillowed mafic flows and felsic to intermediate volcani­ clastic rocks and greywackes, that can be traced as far west as Wapawekka Lake in the south-central part of the Glennie Domain. The arc assemblage is intruded by synvolcanic granitic plutons (ca. 1.873 Ga; Heaman et aI., 1994), subvolcanic alkaline porphyries (ca. 1.860 Ga; Heaman et aI., 1997), and the younger Hanson Lake Pluton (ca. 1.844 Ga; Heaman et aI., 1993).

Kisseynew Domain 'south flank'

Juvenile arc and ocean-floor units extend into the gneissic rocks in the southern part of the Kisseynew Domain where they retain many of their geochemical characteristics in spite of the high-grade metamor­ phism. Felsic gneisses derived from volcanic precursors (units J4b, J4c), which are very prominent in the Sherridon-Batty Lake area, contain extensive alteration (Jl5c, part of J4c) and associated Cu-Zn deposits

18 that are structurally flattened, but otherwise similar to VMS deposits in lower grade rocks (Zwanzig and Schledewitz, 1992). These units are structurally overlain and locally underlain by the Bumtwood and Missi Group metasedimentary rocks, and refolded into domes, basins, and complex sheets.

Ocean-floor assemblages (unit F)

The juvenile ocean-floor assemblages are composed principally of MORB-like basalts (units Fl, F2) and related kilometre-scale, layered, mafic-ultramafic plutonic complexes (unit F6) (Fig. 5; Syme and Bailes 1993; Stem et al. 1995b). Other distinct ocean-floor assemblages, in areal extent an order of magnitude smaller than the MORB-like basalt/mafic-ultramafic intrusive complexes, include ocean-island basalt (unit F3c) and ocean-plateau basalt (units F3a, b, d).

MORB-like basalts and mafic-ultramafic complexes are always tectonically juxtaposed or separated by younger intrusions (Syme 1995). Stem et al. (1995b) suggested there are sufficient lithological and geo­ chemical grounds to consider the largest domain of ocean-floor rocks (Elbow-Athapapuskow assemblage, Fig. 1b, 5) as a stratigraphically fragmented back-arc ophiolite despite the apparent absence ofsheeted dyke complexes or harzburgite tectonite. Uranium-lead zircon ages for synvolcanic hypabyssal sills within basalts (1904 ±4 Ma) and gabbro pegmatites in the cumulate complexes (1901 +6/-5 Ma; Stem et al. 1995b) clearly indicate that ocean-floor magmatism was coeval with tholeiitic arc volcanism at Flin Flon (1903 +9/-4 Ma; David and Machado 1993).

MORB-like basalts (units F1 , F2, F4, FS)

Ocean-floor volcanic sequences comprise thick units of pillowed and massive basalt (units Fl, F2; Syme 1995) that were emplaced in a setting far removed from coeval arcs or Archean continents (Stem et aL 1995b). Interbedded arc-derived volcaniclastic sedimentary rocks are absent, and the rare fragmental units that do occur include mafic volcaniclastic rocks (unit F4), basaltic flow top breccias, and reworked hyalo­ clastite. Phreatomagmatic fragmental deposits are absent, suggesting water depths below that at which such eruptions can occur. Synvolcanic diabasic dykes and sills (unit F5) are common and locally abundant (Syme et aL 1993); rhyolitic rocks are rare. The basalts are mapped as laterally coherent 'formations' (named units in Fl and F2), 4 to more than 60'km in strike length with stratigraphic thickness of 0.3-3.0 km, each having characteristic weathering colour, flow morphology, alteration assemblage, and geochemistry (Syme 1995; Stem et aL 1995b; Syme et aL 1995). Some basalt 'formations' display abundant evidence for low­ temperature seafloor hydrothermal activity (e.g. epidosite domains and veins, interpillow chert; Syme 1991, 1992), consistent with eruption in a ridge setting (Stem et aI., 1995b).

Ocean-floor-assemblage basalts are exclusively tholeiitic, with MgO contents typical of modem MORBs, falling mostly in the range 6-10 wt.%. They can be readily distinguished from the arc rocks by their higher Ti and Zr contents at given MgO (Stem et aI., 1995b) and lower Th/Nb ratios. On the basis oftheir trace-element geochemical characteristics relative to modem ocean-ridge basalts, Stem et aL (1995b) subdivided basalts of the Elbow-Athapapuskow assemblage into N- (normal) and E- (~nriched, or plume-related) types. N -type basalts resem ble modem N -M ORB sand Mariana-type back -arc basin basalts (BAB B), having depleted to flat REE

19 patterns, high Zr/Nb, variaole Th/Nb, and initial ENd =+3.3 to +5.4 (Fig. 6; Stem et aI., 1995b). The ocean­ floor basalts with high Th/Nb are thought to be derived, in part, from metasomatized arc mantle similar to that which produced the arc basalts (Stem et aI., 1995b). The E-type basalts resemble modem transitional and plume MORBs (Stem et aI., 1995b), with slightly enriched REE, lower Zr/Nb, and initial ENd = +3.1 to

+4.5. The variation in initial ENd compositions (+3 to +5 at 1.9 Ga) is attributed to mixing of depleted and enriched MORB-like mantle sources and not to contamination by older crust.

Synvolcanic mafic to ultramafic complexes (unit F6)

Mafic-ultramafic intrusive rocks in the ocean-floor assemblages occur in kilometre-scale sequences that are either bound by faults or intruded by plutons, masking primary stratigraphic relations with the ocean-floor basalts (Syme 1988, 1992). The best-exposed sequence (Claw Lake complex: Syme 1992; Williamson 1993; Williamson and Eckstrand 1995) is composed of 1) an older layered series (unit F6c,d: decimetre- to metre­ scale layered gabbro and lesser pyroxenite, peridotite, and anorthosite); and 2) a younger, isotropic to wispy layered gabbro (unit F6a) with locally abundant pegmatitic gabbro veins. The abundances ofREEs in the gab­ bros and peridotites are systematically lower than in the MORB-like basalts, but the range ofLa/Yb values is similar, consistent with the subparallel increase of REEs with fractionation in the basalts (Stem et al. 1995b). The mafic-ultramafic intrusive complexes are interpreted as oceanic gabbros and crustal cumulates (Layer 3).

Ocean-island basalts (unit F3c)

An isolated occurrence of conglomerates (Long Bay basalt, unit F3c; Syme, 1991) consisting principally of basaltic detritus, has been mapped in contact (unconformable?) with arc-assemblage rocks in the Elbow Lake area. The basaltic clasts are scoriaceous to strongly amygdaloidal, commonly display vesicle banding, and preserve ropy or crenulated internal contacts in some clasts. These features are consistent with subaerial eruption of the basalts, although conglomeratic turbidite bedforms suggest that the conglomerates were deposited subaqueously. Most of the samples of this unit are subalkaline, but they span the Macdonald and

Katsura (1964) tholeiite/alkali basalt dividing line. The basalt clasts have high MgO and low Al20 3 contents (9.5-13.5 wt. %), and modest overall LREE enrichment, with concave-downward LREE profiles and HREE that are strongly fractionated. Stem et ai. (1995b) concluded that the geochemical characteristics of these rocks are similar to tholeiitic ocean-island basalts (e.g. Hawaii). Initial ENd values (at 1.90 Ga) for three sam­ ples range from +2.2 to +3.4, which, coupled with primitive Th/Nb ratios «0.1), characterize a ca. 1.90 Ga enriched mantle source. Stem et ai. (1995b) speculated that the ocean-island basalts may be derived from Fiji-like, post-subduction, hot-spot magmatism.

Ocean-plateau basalts (units F3a, b, d)

Sandy Bay ocean-plateau assemblage (Fig. la, 5; unit F3a) is a ca. 3 km thick, monotonous sequence of subaqueous basalt flows and synvo1canic sills of unknown age (Reilly et aI., 1994; Slimmon, 1995). The Sandy Bay basalts are uniformly tholeiitic (Stem et aI., 1995b) and plot in the E-MORB/tholeiitic ocean­ island basalt field on aZr-Th-Nb diagram. However, the basalts are geochemically distinct from those ofthe arc and ocean-floor assemblages (Stem et aI., 1995b). Their trace-element characteristics include strong

20 enrichment in high-field-strength elements (Nb, Zr, Ti), LREE enrichment ([La/Yb]N = 1.3 to 4.5), high TiN, and low Zr/Nb (Stem et aI., 1995b). An important feature of these basalts is their content of fraction­ ated heavy REEs, which suggests the involvement of residual garnet during melting (Stem et aI., 1995b; Watters et aI., 1994) and contrasts with the other basalt types (arc, ocean floor). Two samples of Sandy Bay basalt that bracket the sequence's trace-element compositional range yielded identical initial ENd values of +4.5 (Stem et aI., 1995b). Stem et ai. (1995b) proposed an ocean-plateau or ocean-island origin for the basalts, on the basis of their physical and geochemical characteristics coupled with their juvenile Nd­ isotopic signature and absence of crustal contamination.

Unknown geochemical affinity (unit U)

Metavolcanic (and derived gneissic) rocks whose geochemical affinities are completely unknown, and that cannot be grossly correlated with rocks of known affinity, have been placed in unit U. These rocks are assumed to be part ofthe 1.92-1.87 Ga assemblages. They include parts of the sub-Phanerozoic extension of the Flin Flon Belt in Saskatchewan and gneisses on the exposed Shield interpreted to have metavolcanic pro­ toliths. Although units of volcanic rocks can be continuously traced from the Flin Flon area into the more highly metamorphosed terrains, extreme structural shortening and attenuation preclude the discrimination of distinct assemblage types.

Isotopically evolved Proterozoic and Archean rocks

Archean crustal fragments (unit A 1)

Archean crustal fragments represent a minor but important component in the Flin Flon-Glennie Complex, comprising «1% by area. There is no evidence of significant Archean crust at depth at the time ofits forma ­ tion (i.e. 1.88-1.87 Ga; Stem and Lucas, 1994). Granitoid rocks, dated at 2.497 and 2.518 Ga (David and Syme, 1995) and with an average initial ENd composition of -6.9 (Stem et aI., 1995a), occur as fault-bound lozenges (10-100 m wide x hundreds ofmetres long) within the Northeast Arm shear zone on eastern Schist Lake (Lucas et aI., 1996). The Archean rocks are intruded by mafic dykes that may be related to a sequence of tholeiitic (arc-rift?) basalts with relatively evolved initial ENd compositions (-1 to +2 at 1.9 Ga; cf. Stem et aI., 1995b; Fig. 6) found immediately east ofthe shear zone (Scotty Lake section; Bailes and Syme, 1989).

Evolved arc (unit E)

The Flin Flon Belt also contains an isotopically evolved arc sliver (,Mystic Lake assemblage', Fig. 5) that contains 1.920-1.903 Ga, calc-alkaline, massive to layered, amphibolite-grade tonalite, granodiorite, and diorite orthogneiss (unit E1a; Table 1; Reilly et aI., 1994; Syme, 1991) with initial ENd values of -3.1 to -6.1 and evidence for xenocrystic Archean zircons (2.56-2.67 Ga; Stem et aI., 1992, 1993; Stem and Lucas, 1994; Fig. 6). The Mystic Lake assemblage is marked by LREE enrichment and elevated Th/Nb ratios (Stem, unpub. data). Coupled with the evidence from xenocrystic zircons, these results suggest that Archean

21 basement was involved in the generation ofthese plutonic rocks (Stem, un pub. data). Lucas et aI. (1996) sug­ gested that the Mystic Lake evolved-arc assemblage represents a tectonic slice of the middle crust of an arc built on Archean crust, possibly a microcontinental fragment.

- Tectonic setting ofarc and ocean-floor assemblages

Two principal factors suggest that the ocean-floor-assemblage basalts may have formed intimately with the Plin Plon arc assemblage, independent of their current spatial proximity: 1) some basalt formations show

dual features of MORB-like major-element geochemistry and arc trace-element signature (e.g. lower Ti02 and higher Th/Nb ratios), characteristics of some basalts in modem intra-oceanic back-arc basins (Tamey et aI. 1981; Sinton and Fryer 1987); and 2) the identical crystallization ages (Ga. 1.9 Ga) ofthe basalts, related gabbro-peridotite, and Flin Plon arc tholeiites. Stem et aI. (1995b) proposed-that the Elbow-Athapapuskow ocean-floor assemblage formed in an intra-oceanic back-arc basin setting similar to the modem-day Mari­ ana Trough or Lau Basin. Individual basalt formations within the Elbow-Athapapuskow assemblage differ by virtue ofthe interplay between depleted (N-MORB), more enriched (OIB-like) sources, and subduction­ modified (arc) mantle associated with adjacent rifted arcs.

Neodymium-isotopic and geochronological evidence

Neodymium-isotopic and trace-element data indicate that the Plin Plon-assemblage arc rocks are predomi­ nantly juvenile (above) and show only limited contributions from older crustal sources. Stem et aI. (1995a) suggested that the contributions from older crustal sources were best explained by recycling of small amounts «10%) of Archean and/or older Proterozoic crust via sediment subduction or possibly intracrustal contamination. This is supported by U-Pb geochronological study ofdetrital zircons in greywackes associa­ ted with arc-rift basins that principally contain 1.92-1.887 Ga zircons, interpreted as associated with arc vol­ canism, but also zircons at ca. 2.5 Ga and older (Heaman et aI., 1993; Ansdell and Stem, 1997).

Although the juvenile oceanic rocks at Snow Lake are similar in age to the ca. 1.90-1.88 Ga, VMS-hosting,

juvenile oceanic-arc rocks at Flin Plon, they display distinctly lower ENd (-0.4 to +3) than most of their Plin Plon equivalents (+2 to +5) and likely evolved as an independent system (Stem et aI., 1993, 1995a). The Snow Lake arc assemblage also contains direct evidence for interaction of juvenile arc magmas with older crustal materials in xenocrystic zircons of 2.65-2.82 Ga age in an 1.892 Ga rhyolitic breccia unit (David et aI., 1996). Evidence for limited contamination by older crust in stratigraphically overlying, geochemically

'evolved-arc', basaltic andesite (cf. Bailes and Galley, 1996) is indicated in initial ENd values of -0.4 to +2.4 (Stem et aI., 1995a).

In a broad context, the Flin Plon 'arc' and Elbow-Athapapuskow 'back-arc' system probably occurred in a peri-continental setting, with rifted fragments ofcontinental crust locally forming the basement to arcs (e.g. Mystic Lake assemblage and possibly part ofthe Snow Lake arc assemblage) and available for incorporation in the accretionary collage atca. 1.87-1.88 Ga (cf. Lucas etaI., 1996). The 2.5 Ga signature ofthe older crust may be related to the 'Sask craton' (Ansdell et aI., 1995), an Archean block that has a characteristic 2.45-2.50 Ga signature as well as older ages (to >3.2 Ga, Table 1; Heaman et aI., 1995; Chiarenzelli et aI., 1996). Further evidence of this is found in the detrital-zircon record of turbidites that were deposited in

22 basins developed on or adjacent to arc assemblages (e.g. 'Welsh Lake' turbidites in the West Amisk arc assemblage; cf. Reilly et aI., 1995). Ansdell and Stem (1997) reported detrital zircons with SHRIMP U-Pb ages of 2.426,2.431, and 3.003 Ga in a sample with the dominant population of grains ranging in age from 1.887-1.968 Ga, consistent with results of Heaman et aI. (1993) on a sample of the same turbidite sequence. The ages of ca. 2.43 and 3.03 Ga are virtually identical to ages from the 'Sask craton' in the Pelican Window, suggesting that it may have been exposed and supplying detritus at the time of ca. 1.90--1.88 Ga arc volcan­ ism (cf. Ansdell and Stem, 1997).

1.88-1.83 Ga intrusive rocks (unit P)

The 1.92-1.87 Ga tectonostratigraphic assemblages (above) were amalgamated to form an accretionary col­ lage prior to the emplacement of voluminous 1.88-1.83 Ga granitoid plutons and deposition of younger sedimentary and volcanic rocks (Lucas et aI., 1996; below). The plutons and coeval volcanic rocks are associated with younger arc(s) imposed on the collage, resulting in the development of a microcontinent by 1.85-1.84 Ga that probably included the adjacent Hanson Lake Block and Glennie Domain. We use the term 'successor' to describe the temporal context ofthe postaccretion intrusive and sedimentary rocks (below), in that they are younger than (and thus succeed) the 1.92-1.87 Ga tectonostratigraphic assemblages (Lucas et aI., 1996).

Granitoid magmatism in the Flin Flon Belt spans three evolutionary stages: 'evolved' arc (in unit E; -1.920 Ga) plus early juvenile arc (in unit J; 1.904-1.880 Ga) plutonism during intra-oceanic arc/back-arc formation (discussed above), early (1.878-1.860 Ga) and middle (1.860--1.844 Ga) successor-arc plutonism following accretion and successor-arc(s) development, and late (1.843-1.826 Ga) successor-arc plutonism accompanying successor-basin formation and waning arc magmatism.

Each stage comprises a wide compositional spectrum; accordingly, all granitoids in stages 2 and 3 above, the majority of which remain undated, are included together (unit P) on the 1: 100000 compilation. Within this large-scale grouping the units are arranged from mafic (unit PI) to felsic (unit P9); hypabyssal intrusions (unit PlO) and tectonites derived from plutonic rocks (unit PIt) complete the subdivision. The following discussion of the nature and geochemistry of granitoid rocks in the Flin Flon Belt is from Whalen et aI. (in press).

Flin Flon Belt plutonic rocks are lithologically and compositionally variable, even at the outcrop scale. They range in texture from fine to coarse grained, and equigranular to porphyritic, and in composition from gab­ broic and dioritic to tonalitic and granodioritic. Homogeneous, blue quartz-eye-porphyritic to megacrystic tonalite is the predominant felsic rock type in the early juvenile-arc and early successor-arc pluton group. In contrast, both normal and reverse compositional zonation, from diorite to granodiorite, frequently accom­ panied by variation in grain-size, is a common feature of middle successor-arc plutons. Granites, sensu stricto, and intrusive phases in which K -feldspar was an early liquidus phase are rare and restricted to the late successor-arc group of plutons. Fine- to medium-grained, angular to ovoid, hornblende-rich gabbro and diorite inclusions, which probably represent quenched samples of coexisting mafic magmas, are present, and locally abundant, in plutons of all ages. Metasedimentary inclusions are rare or absent, and

23 metavolcanic inclusions are usually angular and restricted to near intrusive contacts. Mafic dykes are also common within plutons of all ages; some exhibit features, such as cuspate margins and irregular sinuous contacts, suggestive of emplacement prior to complete solidification of their host plutons.

Geochemistry and source constraints

Amphibole-bearing mineralogy, metaluminous compositions and igneous micro granitoid enclaves indicate that Flin Flon Belt granitoids are derived from infracrustal sources. The predominance of intermediate calc­

alkaline compositions and negative Nb' anomalies on normalize a patterns over a 46-77 wt. % Si02 range indicate an arc setting. Basaltic end members indicate important contributions directly from the mantle. The

ENd (T) values are predominately in the range 0 to +4.3 (Fig. 6), reflecting mixing between depleted mantle melts and an Archean crustal component preserved in evolved-arc plutons (~J.9 to -6). Temporal variations includel) early juvenile-arc plutons are low-K, high-field-strength-element (HFSE) depleted, with rela­ tively flat rare-earth element (REE) patterns and negative Eu anomalies, indicative of low-pressure partial melting/fractionation in the mantle wedge, with residual pyroxene and plagioclase; 2) early and middle successor-arc plutonism is medium-K, with steep REE patterns and no Eu anomalies, indicative of input from melting of basaltic sources (likely subducted back-arc oceanic crust) under high-pressure conditions with residual garnet and/or amphibole and no plagioclase; 3) late successor-arc plutons are high-K, more HFSE-enriched, with both variable REE pattern slopes and Eu anomalies, indicative of a significant petro­ genetic role of recycling of pre-existing juvenile arc/accretionary complex crust.

Rocks of uncertain age (unit W)

Unit W contains a disparate assemblage of tectonites (unit W6), gneisses (unit W5), metasediments (units W2, W3, W4) and metavolcanic rocks (unit WI). The common thread for all is that an unequivocal age can­ not be determined with presently available data. Rocks of the greywacke, derived gneisses, and migmatite unit (unit W2) are considered correlative with either the broadly syn-volcanic volcaniclastic rocks of unit J9b or the Burntwood Group (unit B I), but structural complexities preclude this distinction on the map-unit scale in the absence of detrital zircon studies. The interlayered sandstones, feldspathic quartzites, and minor pelites (unit W3) generally occur at the boundary between rocks of the Burntwood (unit B) and Missi (unit M) groups, and are interpreted as representing a depositional transition. The carbonate and calc-silicate rocks (unit W4) probably include both calcareous sedimentary rocks and volcanic rocks which were sub­ jected to carbonatization prior to metamorphism.

The largest domain of unit W rocks lies in the Sub-Phanerozoic portion of the Shield Margin area (Fig. lb), which is described below, including rocks other than unit W.

Sub-Phanerozoic Precambrian geology

The northern edge of Phanerozoic platformal rocks of the Western Canada Sedimentary Basin overlies the Flin Flon Belt in Manitoba and Saskatchewan. Regional mapping of the Phanerozoic-covered basement, involving the integration of high-resolution aeromagnetic and gravity data with extensive drill core

24 information (Leclair et al., 1997), resulted in the recognition of several major domains in the buried base­ ment, each with a distinct lithotectonic character and potential field-anomaly pattern (Fig. 7). Three lithotec­ tonic domains in the buried basement (Clearwater, Athapapuskow, and Amisk Lake domains) are characterized by northerly trending, positive gravity and aeromagnetic anomalies (Fig. 7) and correlate with the 1.92-1.83 Ga volcanic and plutonic rocks of the exposed Flin Flon Belt (Clearwater = Snow Lake arc assemblage; Athapapuskow = Elbow-Athapapuskow ocean-floor assemblage; Amisk Lake = West Amisk arc assemblage).

An upper-amphibolite-grade orthogneiss complex (Namew Gneiss Complex, Fig. 7), containing calc­ alkaline intrusive rocks (various unit P) ranging in age from 1.88 to 1.83 Ga and screens derived from the older volcano-sedimentary rocks (unit U2), is interpreted as the middle crust of a 1.88-1.84 Ga arc exposed in the Flin Flon Belt. Discordant intrusive complexes, such as the 1.830 Ga Cormorant Batholith (unit P9; Fig. 1b) are centered on magnetic-gravity lows (Fig. 7) and truncate the structural trend ofadjacent lithotec­ tonic domains. Correlation of Flin Flon Belt geology with that beneath the Phanerozoic cover shows that its constituent lithotectonic elements have north-south strikes ofup to 150 km, and form a predominantly east­ dipping crustal section, consistent with LITHOPROBE seismic-reflection profiles (Leclairet al., 1997). The Namew Gneiss Complex forms the structurally deepest part ofthe east-dipping crustal section imaged along LITHOPROBE line 3 (Fig. 5) (Lucas et al., 1994; White et al., 1994; Leclair et al., 1997).

The Namew Gneiss Complex is unique to the Flin Flon Belt in terms of its high metamorphic grade (upper amphibolite facies) and predominantly orthogneissic character (Leclair et al., 1997). In general, it consists of variably deformed granitoid rocks (tonalite, quartz diorite, granodiorite, and diorite) with metre- to kilometre-scale enclaves of mixed mafic, calcic, psammitic, and pelitic gneisses (unit U2), and plutons (various unit P) ofless deformed granodiorite, tonalite, and diorite. These rocks display curvilinear to tightly folded magnetic anomaly patterns. Although hornblende-biotite tonalite and quartz diorite (unit P6) are the predominant plutonic units, biotite-hornblende granodiorite (units P8, P9), mafic-ultramafic rocks (unit P 1), and biotite monzogranite/syenogranite (unit P4) are also important components ofthe complex. A penetrative foliation defined mainly by biotite and hornblende in the plutonic units parallels the composi­ tionallayering. Detailed drill core and underground examinations (at Namew Lake Mine) indicate the pres­ ence ofmUltiple intrusive phases (Leclair et al., 1997). Uranium-lead zircon dating ofdrill-core samples has yielded the following igneous crystallization ages: 1) 1880 ±2 Ma for quartz diorite gneiss, and 2) 1850 ±2 Ma for weakly foliated tonalite (Leclair et al., 1997). Mafic-ultramafic rocks (units P2, PI) occur as sub­ map-scale bands within the layered sequence and as discrete, subcircular plutons marked by positive aero­ magnetic highs (Fig. 7). Monzogranite (unit P9) occurs as layer-parallel to discordant veins and rare, dis­ crete, kilometre-scale plutons which cut all other components of the gneiss complex. These late granitic phases may be related to the Cormorant Batholith (below).

The most conspicuous geophysical feature of the buried basement is a roughly ovoid-shaped (about 60 x 25 km) magnetic and gravity low (Fig. 7) that appears to be superimposed on the anomaly patterns of adjacent rocks and is termed the Cormorant Batholith (unit P9; Leclair et al., 1997). The batholith is charac­ terized mainly by moderate-intensity aeromagnetic relief (Fig. 7), and the negative gravity anomaly has a minimum value at -65 mGal. Its internal magnetic fabric is either weak or absent and is transected by a series

25 99 0 103 0 102 0 lor 1000

55 0

540 54· 990 102 0 lor 1000

1030 99°

550

54" 0 0 99 103 102 0 101 0 1000

55 '

54'

26 oflinear north-northwest-striking structures, possi­ Figure 7: Upper: Shaded relief representation of bly representing fractures or faults. Seven widely the total field aeromagnetic data for the NATMAP spaced drillholes in the Cormorant Batholith Shield Margin Project area (cf. Broome and Viljoen, magnetic-gravity low have all intersected unde­ in press). formed monzogranite, two of which have yielded Middle: Bouger anomaly gravity map of the U-Pb zircon ages of 1.830 Ga (Leclair et aI., 1997). NATMAP Shield Margin Project area (cf. Broome and Viljoen, in press).

Lower: Major lithotectonic domains and regional 1.87-1.83 Ga sedimentary, volcanic tectonic framework of the exposed and buried Flin and intrusive rocks (units S, M, B) Flon Belt (Leclair et al. 1997). The Hanson Lake Block, eastern Glennie Domain, Tabbernor Fault Zone, Kisseynew Domain, and the West Amisk, Volcanic, volcaniclastic, and sedimentary rocks Elbow-Athapap and Snow Lake assemblages that are younger than the 1.92-1.87 Ga tectonostra­ extend southward into the subsurface. Abbreviations tigraphic assemblages (Fig. 5) have been docu­ are: BCF - Berry Creek Fault; CBF - Crowduck Bay Fault; ELSZ - Elbow Lake Shear Zone; NLS­ mented across the central part of the Flin Flon belt, Namew Lake Structure; SASZ - South and are termed 'successor' basin deposits, in paral­ Athapapuskow Shear Zone; SLF - Suggi Lake lel with terminology for the 1.88-1.83 Ga mag­ Fault; SRSZ - Spruce Rapids Shear Zone; SWSZ - Sturgeon-weir Shear Zone. matic rocks. These sedimentary rocks may represent the remnants of depositional basins that formed alluvial aprons, fluvial systems, and marine turbidite basins. Successor basin deposits fall into two contrasting types: older (>1.85 Ga) marine to subaerial volcaniclastic and epiclastic deposits (unit S), and younger «1.85 Ga) subaerial (unit M) to marine (unit B) deposits derived from erosion of successor-arc volcanic and plutonic rocks as well as the older tec­ tonostratigraphic assemblages.

Basement to the older basins and younger marine deposits has yet to be found, whereas younger subaerial sandstones and conglomerates unconformably overlie 1.92-1.87 Ga assemblages and 1.88-1.85 successor arc plutons (Bailes and Syme, 1989; Holland et aI., 1989; Stauffer, 1990).

Schist-Wekusko assemblage (unit S)

The Schist-Wekusko assemblage includes 1) 1.867 Ga (Stem et aI., in press) grey wacke turbidites, rhyolitic tuffs, and associated high-level differentiated tholeiitic sills (Scheiders Bay sequence, Athapapuskow Lake; Syme, 1988); 2) trachyandesite marine conglomerate (McCafferty Liftover sequence) and overlying 1.876 Ga rhyolite flows, tuff, and tuff-breccia, (Wekusko Lake, Ansdell et aI., in press); 3) 1.858 Ga trachyandesite marine conglomerate-sandstone sequence on southern Schist Lake (Syme, 1988; Stem et aI., in press); and 4) 1.856 Ga East Wekusko rhyolite (Ansdell et aI., in press). Defined contacts between the Schist-Wekusko assemblage and Burntwood or Missi Groups are invariably faulted.

27 Missi Group (unit M)

Missi Group deposits are characterized by thick packages (>2 km) of conglomerate, pebbly sandstone, and sandstone (units M1-M3) interpreted to have been deposited in alluvial and fluvial environments (Bailes and Syme, 1989; Syme, 1988; Stauffer, 1990), similar to that of the Temiskaming sequences in the Superior Province. Uranium-lead analysis of detrital zircon populations in the sandstones (youngest zircon is 1.846-1.847 Ga; Ansdell et al., 1992; Ansdell, 1993) and crosscutting intrusions (1.842 Ga; Heaman et al., 1992) has bracketed sedimentation to approximately 1.845 Ga in the central Flin Flon Belt (Table 1). In the eastern Flin Flon belt, Ansdell and Connors (1994) and Ansdell et al. (in press) have bracketed sedimentation between 1.832 Ga (youngest detrital zircon) and 1.826 Ga (crosscutting intrusion), suggesting that Missi Group sedimentation is diachronous. In the southern Kisseynew Domain, a spread in zircon ages from 1.846 Ga to 1.837 Ga has been recognized between the base and top of a section of Missi Group paragneisses at Puffy Lake (Machado et al., in press; below). Detrital zircon ages in all these Missi Group rocks indicate provenance from Flin Flon-Glennie Complex sources (1.92-1.85 Ga) as well as older (2.2-2.6 Ga) crustal sources (Ansdell et al., 1992; Ansdell, 1993).

Key features of the Missi Group siliciclastic rocks where they occur within the Flin Flon Belt are: 1) uncon­ formable deposition on 1.92-1.87 Ga volcanic assemblages and 1.88-1.85 Ga plutons; 2) development of an oxidized paleosol (regolith) at the unconformity; 3) removal of significant (ca. 2 km) stratigraphic section along the (angular) unconformity; 4) abundant clasts of metavolcanic and metasedimentary rock types derived from the 1.92-1.87 Ga assemblages, medium- to coarse-grained successor-arc plutonic rocks, regolith, and jasper; and 5) locally significant sections of mafic-felsic calc-alkaline and tholeiitic volcanic rocks (units M4-M5) and rare trachyandesite sills (unitM6a) (Bailes and Syme, 1989; Syme, 1987; Stauffer, 1990; Holland et al., 1989). Together, these features suggest that the Missi Group sedimentation occurred during postaccretion (successor) arc magmatism on an uplifted and deeply incised terrain (e.g. Bailes and Syme, 1989; Stauffer, 1990). Depositional environments included alluvial fans and braided river systems. It is likely that earlier structures (e.g. folds, steep belts, shear zones) and associated topography controlled the pattern of fluvial drainage systems and associated Missi Group sedimentation at 1.85-1.84 Ga (Lucas et al., 1996).

Missi Group volcanic rocks (units M4-M5) include basalt, andesite, trachyandesite, dacite, rhyolite, and derived orthogneisses. The largest and least recrystallized sections, east of Wekusko Lake, include mafic volcanic rocks that generally lack pillows, and felsic fragmental rocks (1836 ± 2 Ma, Gordon et al. , 1990) with flattened shards and fragments (Gordon and Gall, 1982; Shanks and Bailes, 1977), tentatively inter­ preted as subaerial deposits (Gordon and Gall, 1982; Bailes, 1985). Basalts and andesites at Wekusko Lake are enriched in Zr, Y, and Ti02 relative to juvenile arc volcanic rocks that predate 1.88 Ga (Connors and Ansdell, 1994a, b). Potential gneissic equivalents of the east Wekusko subaerial felsic rocks have been reported throughout the southern Kisseynew Domain (Bailes, 1975, 1980a, b; Zwanzig, 1996).

In the southern Kisseynew Domain, the Missi Group comprises magnetite-bearing quartzofeldspathic gneiss with local crossbedding (units M3f-i), flattened conglomerate beds, layers of fine-grained gneiss derived from subaerial volcanic rocks, felsic gneiss with relict quartz phenocrysts (units M6b, M7b and parts of M3g), and mafic gneiss with flattened amygdaloidal zones (unit M4c). The Missi Group gneisses are structurally stacked with imbricates of Burntwood Group turbidite-derived gneisses and gneissic

28 equivalents of Amisk collage and successor-arc plutons (Zwanzig and Schledewitz, 1992) in large-scale recumbent folds and small nappes. The recumbent structures are refolded into overturned domes and basins to form the complex interference patterns in the northern part of the NATMAP Shield Margin area.

In the lower part ofthis structural pile, the basal Missi unconformity extends for at least 10 km north of Sherridon and probably all along the south flank of the Kisseynew Domain. This part of the unconformity is dated directly as 1848 ± 4 Ma from felsic gneiss interpreted as a bed of tuff within the basal conglomerate (Machado et aI., in press). A higher part of the same stratigraphic section contains 1837 ±4 Ma detrital zir­ cons (Machado et aI., in press) to suggest over 10 Ma ofMissi sedimentation and volcanism in the Kissey­ new Domain. The volcanic rocks are bimodal, including moderate- to high-K rhyolite and tholeiitic to calc-alkaline basalt and basaltic andesite. Their geochemistry is distinctive from arc volcanic rocks that pre­ date 1.88 Ga; they contain relatively high Kp, Zr, and HFSE, comparable to type Missi volcanic rocks (above) and Aegean-arc volcanic rocks (Zwanzig, 1996).

In the upper part of the structural pile, the Missi Group contains no conglomerate, but grades downward in the Burntwood Group through a unit of interbedded protoquartzite and pelite (unit W3a) interpreted as a regressive shallow-marine succession. The entire succession is considered to be a distal marine and alluvial plane facies of the Missi Group 'basin' at Flin Flon (cf. Stauffer, 1990).

Burntwood Group (unit B)

In the Flin Flon Belt (Reed Lake and Snow Lake areas; Fig. 5), the Burntwood Group includes greywacke, siltstone, mudstone, and rare conglomerate (unit B la), with bedforms and sedimentary structures consistent with deposition by turbidity currents. Within the low-grade Flin Flon Belt these rocks are generally in fault contact with other units. In Saskatchewan, Burntwood Group rocks are interpreted locally to pass gradation­ ally into broadly synvolcanic volcaniclastic rocks (unit J9b), without obvious structural or stratigraphic break (Hartlaub et aI., 1996). Burntwood Group psammitic to pelitic gneisses (unit Bib) with upper almandine­ amphibolite-facies mineral assemblages and derived migmatitic gneiss (unit Blc) are the major components of the Kisseynew Domain (Bailes and McRitchie, 1978; Bailes, 1980a).

The sedimentological (Bailes, 1980a, b), stratigraphic (Zwanzig, 1990), and geochronological (David et aI., 1996; Machado and Zwanzig, 1995) constraints on the Burntwood Group suggest that deposition was at about 1.855-1.84 Ga, partly coeval with the Missi Group, in prograding submarine fans fed by braided river systems draining from adjacent mountain range(s) and active successor arc(s). Southward coarsening of the turbidites (Bailes, 1980a; Syme et aI., 1995) and northward fining of conglomeratic Missi Group rocks (Harri­ son, 1951; Bailes, 1971; Bailes and Syme, 1989; Ansdell et aI., 1995) suggest that the two groups may have been sedimentary facies equivalents (e.g. Syme et aI., 1995). Gradational to disconformable sedimentary contacts between the groups occur in the higher part of the structural pile in the southern Kisseynew Domain (i.e. in the distal facies (Zwanzig and Schledewitz, 1992; Zwanzig, 1995»).

Uranium-lead ages for calc-alkaline to alkaline intrusions in Flin Flon-Glennie Complex and the La Ronge-Lynn Lake Belt (Bickford et aI., 1990; Heaman et aI., 1992; Stem and Lucas, 1994) overlap the age of Burntwood and Missi Group sedimentation, suggesting that arc magmatism was sustained during

29 sedimentation (Lucas et aI., 1996) and providing an explanation for the abundance of immature, volcanic­ derived detritus in the turbidites (Bailes, 1980b). A back-arc (Mediterranean-style) basin setting has been proposed to explain 1.85-1.84 Ga sedimentation, magmatism, deformation, and metamorphism associated with the Kisseynew Domain (Ansdell et aI., 1995; Zwanzig, 1996).

Late intrusive rocks (unit L)

Fine-grained to peg mati tic granitoid rocks postdating arc-derived magmatism fall into two broad categories. The oldest are synkinematic melts (units 1.1, L2), which were emplaced as sheets, variably transposed dykes, and small plutons when metamorphic temperatures exceeded minimum melting conditions. One pegmatite, displaying a tectonic stretching lineation related to deformation in the Pelican Decollement Zone, yielded an 1806 ± 2 Ma age (Ashton et aI., 1992), consistent with 1.815-1.805 Ga estimates for 'peak' metamorphic conditions (Gordon et aI., 1990; Ashton et aI., in press). Syn- to late-tectonic leucogranite (unitL1a) in the Jungle Lake area of the Kisseynew 'south flank' yielded magmatic monazite ages of 1.80 to 1.79 Ga (Parent et aI., 1995).

Later intrusive rocks occur as straight-sided dykes and rare sheets of fine-grained to pegmatitic, dominantly biotite±gamet granite (units L3, L4). The majority of these appear undeformed, but in the Pelican Window they exhibit a weak tectonic foliation which is axial planar to north-south folding, and are elsewhere folded by a later northeast-trending phase. In the Hanson Lake Block, these intrusions have been grouped under the term Jan Lake Granite (unit L3), and dated at about 1.770 Ga (Bickford et aI., 1987; Ashton et aI., 1992), which appears representative of other late pegmatite dates in the region (Krogh et aI., 1985; Machado et aI., 1987; Chiarenzelli, 1989). Neodymium isotopic studies have shown that at least some of these ca. 1.770 Ga granitic rocks in the Hanson Lake Block were derived by partial melting of Archean material from the Sask Craton at depth (Bickford et aI., 1990, 1992; Mock et aI., 1993).

MINERAL DEPOSITS

Volcanic-hosted massive-sulphide deposits

The Flin Flon Belt is one of the largest Proterozoic volcanic-hosted massive-sulphide (VMS) districts in the world, in which more than 118 700 000 t of sulphide has already been mined from 25 deposits (Fig. 1b, 5), with a further 64 300 000 t contained in 43 subeconomic deposits (Syme and Bailes, 1993; Syme et aI., in press). Recent work has fundamentally altered earlier perceptions of the evolution of the Flin Flon Belt and the setting of VMS deposits in it, indicating that economic deposits are only found in juvenile-arc rocks (Syme and Bailes, 1993). As a group, VMS deposits in the Flin Flon Belt 1) occur in tholeiitic and calc­ alkaline suites dominated by basalt and basaltic andesite; 2) are stratigraphically associated with isotopi­ cally primitive (positive initial ENd) rhyolite, commonly the most primitive rock in the sequence; 3) occur at major stratigraphic and compositional 'breaks', recognized by contrasting major-element, trace-element, and isotopic characteristics ofthe underlying and overlying mafic rocks; 4) are commonly underlain by vol­ caniclastic rocks; and 5) commonly have discordant footwall chloritic alteration zones (Syme and Bailes, 1993). Where the stratigraphic context of the VMS deposits in the Flin Flon arc assemblage is preserved

30 through overprinting deformation, magmatism, and sedimentation, it can be demonstrated that some were temporally associated with arc-rifting processes (Syme et aI., in press; Bailes and Galley, in press). Critical observations include evidence for extensional faulting, erosion and development of unconformities; extru­ sion of MORB-like basalts and associated rhyolites; and development of depositional basins with thick sequences of shoshonitic turbidites (Syme et ai. in press).

It is important to note that the large-scale tectonic interleaving and juxtaposition we observe between tec­ tonostratigraphic assemblages are reproduced at a more detailed (camp) scale. For example, within a 20 km radius of Flin Flon, 14 VMS deposits occur in a number of tectonically juxtaposed arc slivers, separated by major accretion-related shear zones, slivers of ocean-floor basalts, and slivers of successor-basin sedimen­ tary deposits. As a result, VMS-hosting stratigraphy usually cannot be correlated between deposits. Detailed mapping, geochemistry, and geochronology are required to define the various tectonostratigraphic components and their bounding structures. However, the first-order association between juvenile-arc rocks and VMS deposits provides a powerful screen to focus exploration in the Flin Flon Belt, given the contrast­ ing lithologic, stratigraphic, and geochemical characteristics of the two dominant assemblages (Syme and Bailes, 1993; Syme et ai. in press).

Volcanic-hosted massive-sulphide deposits at Snow Lake can be subdivided into Cu-rich, Zn-rich, and Cu­ Zn-Au types. Copper-rich deposits, mainly at Anderson and Stall lakes, occur in a flow-dominated, bimodal (basalt-rhyolite) sequence composed mainly of primitive-arc tholeiite. Zinc-rich types (e.g. Chisel Lake) occur in a volcaniclastic-dominated, relatively more evolved sequence. A recently discovered Au-rich Cu-Zn deposit at Photo Lake (Bailes and Simms, 1994) also occurs in the more evolved arc sequence, but within a rhyolite-dominated section. As at Flin Flon, stratigraphic and geochemical evidence suggests that VMS deposition occurred during a period of arc extension and rifting and is associated with the most primi­ tive initial ENd values in the associated stratigraphic sequences (Stem et aI., 1992; Syme et aI., 1996, Bailes and Galley, 1996; Bailes and Galley, in press).

Gold deposits

Gold mineralization in the NATMAP Shield Margin area (Fig. 1 b, 5) can be subdivided into two main types: late (post-peak metamorphism), mesothermal, vein-type deposits associated with shear zones, and early (synvolcanic) epigenetic deposits. The most common type is quartz-vein deposits intimately associated with brittle-ductile D3-D4 shear zones (e.g. New Britannia, Herb Lake camp, Tartan Lake, Rio; Fedorowich et aI., 1991; Galley et aI., 1986, 1989; Ansdell and Kyser, 1992; Gale, 1997, Schledewitz, 1997). Textural relationships in quartz-carbonate-albite-chlorite-muscovite-pyrite-arsenopyrite alteration envelopes and Ar-Ar data in gold deposits in the central and western Flin Flon Belt indicate that mineralization occurred after peak regional metamorphism at about 1.790 to 1.760 Ga. The only clearly identified example of earlier epigenetic gold mineralization is the Laurel Lake Au-Ag deposit (Ansdell and Kyser, 1991), which consists of quartz- muscovite-carbonate-pyrite-galena-sphalerite-tennantite-electrum veins surrounded by a zone of K-metasomatism and hosted by 1.887 Ga felsic volcanic rocks. This deposit predates regional metamorphism and deformation, and may have been similar to Au-bearing volcanic-associated epithermal­

31 exhalative systems. Galley et ai. (1986, 1989) considered gold deposits in eastern Flin Flon Belt to be postmetamorphic, but preliminary investigations of deposits north of Snow Lake by Gale (1997) and Schledewitz (1997) suggest at least some of this gold mineralization is premetamorphic.

:rECTONIC EVOLUTION

0 1 : Accretion

As discussed above, the range in lithological and geochemical associations, Sm/Nd isotopic signatures, and U -Pb zircon ages has allowed subdivision ofFlin Flon Belt and Kisseynew Domain greenstones into a series of 1.92-1.87 Ga arc and ocean-floor assemblages (Fig. 1b, 5). The mechanism by which these various oceanic tectonostratigraphic assemblages were broken up and subsequently reassembled in a tectonic col­ lage clearly lies in the processes acting at consuming plate margins (e.g. Hamilton, 1979, 1988), but is unre­ solvable in detail. Similarly, little of a concrete nature can be said about the polarity of associated subduction zone(s).

These contrasting volcanic and older crustal assemblages (units J, F, and E) are separated by early high­ strain (shear) zones and are stitched by crosscutting 1.88-1.84 Ga plutons (unit P), suggesting that they were accreted to form an accretionary collage ('Amisk collage' of Lucas et aI., 1996: Fig. 5) between about 1.88 and 1.87 Ga. This initial deformation event (0[; Table 2) occurred at least 50 Ma before the start of orogen­

scale collisional deformation at about 1.840-1.830 Ga. The D j accretionary collage may have included the Flin Flon Belt (now considered to include the Attitti Block and Paleoproterozoic rocks in the Hanson Lake area), Snow Lake assemblage, and Glennie Domain (including the Scimitar Complex), and would thus be equivalent to the Flin Flon-Glennie Complex (as described previously). However, overprinting structures

and fault zones containing slices of the Burntwood Group preclude reconstruction of the D j accretionary Complex beyond the Amisk Collage (Fig. 5).

02: Postaccretion magmatism, sedimentation, and deformation

Postaccretion 1.88-1.84 Ga plutons (unit P) and volcanic rocks (in units S, M) are attributed to younger, postaccretion ('successor') arc(s) imposed on the D[ collage. These developed contemporaneously with regional steepening ofthe early collision-accretion structures (D2' Table 2; Lucas et aI., 1996; Ryan and Williams, 1996, in press). The intrusive rocks have a typical arc signature in their trace-element geochemistry (Whalen et aI., in press). As discussed above, early and middle successor-arc magmatism is likely to have had input from melting of basaltic sources (likely subducted back-arc oceanic crust) under high-pressure conditions, while late successor-arc plutons probably recycled pre-existing juvenile-arc/accretionary-complex crust.

Uplift and erosion, development of a paleosol on 1.9-1 .85 Ga volcanic and plutonic rocks, and deposition of voluminous turbidites (Burntwood Group, unit B) and continental sedimentary rocks (Missi Group, unit M) occurred ca. 1.85-1.84 Ga, coeval with the waning stages of postaccretion arc magmatism (Ansdell et aI., 1995; Machado and Zwanzig, 1995; David et aI., 1996).

32 Table 2: Deformation episodes in the NATMAP Shield Margin Project area.

Age Episode Structures Magmatism Metamorphism (Ma) Tectonic context

L1 (tectonites, None (?) ? 1880-1870 Intra-oceanic accretion D1 ~ , mylonites)

52' F2 (Vick Lake Mafic to felsic dykes, Subgreenschist to 1880-1840 Intra-arc shortening, sheets, plutons amphibolite ('contact' uplift/erosion of Amisk D2 synform) S1/S2 (tectonites, mylonites) to regional) collage; development of Kisseynew back-arc basin S3' f3; shear bands (both None Regional peak 1840-1805 Regional collisional sinistral & dextral); high- (1840-1830 Ma metamorphism shortening and thickening via SW-vergent thrusts & folds; magmatic belt in Snow (subgreenschist to D3 angle shear zones; SW­ high-angle shear zones in vergent thrusts (e.g. Morton Lake assemblage, amphibolite facies) at 1820-1805 Ma Amsik collage Lake thrust zone, Fig. 2) south flank of the Kisseynew Domain) SW-vergenlthrusts (e.g. Pegmatites, Retrograde 1805-1770 (?) Postcollisional thrusting, Sturgeon-weirshear leucogranites folding; transpression of D4 zone, Fig. 2); F4 folds & Amisk collage kinks, high-angle shear zones Brittle to ductile shear None Retrograde 1770-1690 Postcollisional NW-SE D5 zones/faults shortening and longitudinal extension References: Ansdell and Connors (1994), Ansdell et al. (1995), Ansdell and Norman (1995), Ashton (1992), Ashton et al. (1992), Bailes and Syme (1989), David et al. (1993,1996), Digel and Gordon (1993,1995); Fedorowich et al. (1995), Gordon et al. (1990), Heaman et al. (1992,1993), Lucas et al. (1994), Reilly et al. (1993,1994), Ryan and Williams (1994,1995, 1996, in press), Stauffer (1990), Stauffer and Mukherjee (1971), Stern and Lucas (1994), Syme (1994, 1995), Thomas (1992).

Assembly of the crustal thrust stack in southeastern Reindeer Zone (D}) is interpreted to have occurred in response to collision of the 'Sask craton' with the overriding Flin Flon-Glennie Complex (D3, Table 2; Lewry et aI., 1994, 1996). Due to structural interleaving, the boundary between the Flin Flon-Glennie Com­ plex and Kisseynew Domain is best described as a structural-stratigraphic transition zone. In contrast, the boundary between the'Sask craton' and Flin Flon-Glennie Complex is a broad ductile shear zone termed the Pelican decollement zone (Ashton et aI., in press).

Two characteristics mark the history of terminal collision in Trans-Hudson Orogen: widespread felsic to mafic magmatism at 1.84-1.83 Ga, followed by a complete cessation ofmagmatism by 1.825 Ga (except for anatectic crustal melts, unit L). Scattered occurrences of felsic to mafic, generally calc-alkaline, 1.84-1.83 Ga volcanic rocks and plutons together form a broad magmatic belt along much of the south flank and core ofthe Kisseynew Domain as well as within the Glennie Domain and La Ronge belt (Bickford et aI., 1990; Gordon et aI., 1990; Ansdell et aI., in press; Ansdell and Norman, 1995). Although the 1.84-1.83 Ga magmatism has been interpreted to reflect terminal subduction of oceanic crust during closure ofthe 'Kisse­ ynew basin' (Ansdell et aI., 1995), it extends from the Superior margin across most Reindeer Zone terranes and well into the Hearne Province, and may well be related to the first stage of continental collision.

Initial thrusting was coeval with the 1.84-1.83 Ga mafic-felsic calc-alkaline magmatism and ongoing conti­ nental sedimentation (Missi Group; Connors, 1996; Connors et aI., in press); however, the direction ofinitial thrusting is uncertain (Zwanzig, 1995). Southwest-directed thrusting and folding in the Kisseynew Domain (Zwanzig and Schledewitz, 1992) and along the upper and lower boundaries of the Flin Flon-Glennie

33 Complex continued through peak regional metamorphism at 1.82-1.79 Ga (Froese and Moore, 1980; Gordon et aI., 1990; Ansdell and Norman, 1995; Parent et aI., 1995; Connors, 1996; David et aI., 1996; Kraus and Menard, 1997; Kraus and Williams, 1998, in press). Due to structural interleaving and refolding, the bound­ ary zone between the Flin Flon-Glennie Complex and Kisseynew Domain (i.e. the Kisseynew 'south flank ') is geometrically complex and locally 40 km wide (Zwanzig, 1990, 1995, 1996; Zwanzig and Schledewitz, 1992; Lucas et aI., 1994; Connors, 1996; Thomas and Tanczyk, in press).

Oblique collision with Superior craton occurred at about 1.81 Ga (D4 , Table 2) and led to postcollisional sin­ istral transpression of eastern Trans-Hudson Orogen (Bleeker, 1990), with wrench faulting and refolding of the thrust stack producing considerable structural relief (Lewry et aI., 1990). Continued convergence of Superior craton after 1.80 Ga resulted in phases of upright north-south- and later northeast-southwest­ trending folds throughout the Reindeer Zone internides (Ds' Table 2). Intrusion oflate to post-tectonic leu­ cogranites and pegmatites (ca.1.78 Ga, Table 1; Bickford et aI., 1990), generated in significant part by melt­ ing of the underthrust 'Sask craton' basement (Bickford et aI., 1992; Mock et aI., 1993), was followed by postcollisional uplift, cooling and progressive isotopic closure (e.g. Ar-Ar) by approximately 1.7 Ga (Fedor­ owich et aI., 1995). Orogen-parallel upper-crustal escape tectonics may have occurred above lower-crustal detachments (Hajnal et aI., 1996), related either to collisional indentation by Superior craton or possibly to rotation ofHearne Province. The Namew Gneiss Complex, mapped beneath the Phanerozoic cover south of Flin Flon (Leclair et aI., 1997), is interpreted to represent the middle crust of the Flin Flon-Glennie Complex juxtaposed against the exposed upper crustal levels by D4-Ds wrench faulting.

PHANEROZOIC BEDROCK GEOLOGY

The regional compilation map ofPhanerozoic bedrock geology is based on data from 1) 33 mineral exploration drill cores in Saskatchewan, 20 of which are from the Cumberland Delta, Cumberland Lake, and N amew Lake areas (examined 1991 to 1993); 2) 4 drill cores from a kimberlite exploration program in the Cumberland Delta area (examined in 1995); 3) drill cores from stratigraphic holes in Manitoba drilled in 1993 and earlier; 4) drill-core descriptions in assessment files; 5) limited field mapping on the shores of Namew Lake and Athapapuskow Lake; and 6) existing maps (Kupsch, 1952; Manitoba Energy and Mines, 1993a, 1993b).

Present-day distribution of Paleozoic and Mesozoic strata can be attributed to minor depositional thinning from south to north, to major erosional truncation genetically associated with several unconformities, and to post-Silurian flexing probably caused by Phanerozoic reactivation of the Tabbernor Fault Zone in Saskatchewan and the Churchill Superior Boundary Zone in Manitoba.

Paleozoic geology

Precambrian rocks in the southern NATMAP Shield Margin map area are unconformably overlain by the clastic Winnipeg Formation deposited during the Middle to early Late Ordovician (Fig. 8). In the southern half of the study area, the Winnipeg Formation comprises a basal sandstone unit characterized by fair to

34 good sorting of fine to coarse quartz grains, and an upper unit domi­ Manitoba Saskatchewan nated by argillaceous siltstone/sandstone, commonly burrowed, with U) o::J interbeds of sandstone and shale" In the northern half of the study area, u.. co the upper unit is not present and the commonly friable basal sandstone ue u unit is directly overlain by an arenaceous dolostone at the base of the 'iii coU) overlying Red River Formation (Baillie, 1952; Byers, 1957; Padgham, ..,~ 1968; Gent, 1993). Most Winnipeg Formation exposures in the study area occur close to the shield edge beneath a cap of Red River dolostone; these have not been shown on the map. A small surface exposure is mapped along Highway 39 south ofReed Lake (Fig. 1a). c co c The dolostones of the Red River, Stony Mountain, Stonewall, and o > Interlake formations are the dominant rock types in the map area. oQ) These carbonate strata, characterized by cyclic sedimentation, were deposited in shallow warm seas that covered most of the North American craton during much of the Late Ordovician and Early Silu­ rian. Correlation of strata in the study area with those in the Hudson Bay Basin indicates continuity of deposition between the two areas

Upper Interlake (Norford et al., 1994).

C "C East Arm CO :::J "C Q) Atikameg Q) Three shallowing- and brining-upward cycles that are well defined in :::J (f) .::£ .::£ CO CO Lower the Red River Formation in southern Manitoba and southern (f) ~ -;:: Moose -;:: Q) Q) Lake Q) Interlake Saskatchewan (Kendall, 1976; Norford et al., 1994; Haidl et al., 1997) ~ c c 0 - - .....J Fisher also appear to be preserved in the Shield Margin area. The lower Red Branch River, a sparsely fossilferous, burrow-mottled dolomudstone/wacke­ Stonewall Stonewall ------stone unit, represents the lower part of the initial cycle. The upper Red T·marker - T-marker River Formation, in this area composed of interbeds of dolomudstone, C Stony Stony CO '(3 Mountain Mountain argillaceous dolomudstone, and minor dolomitic shale, encompasses C "S; the upper portion of the first Red River cycle and most likely includes CO "(3 0 Fort "C '- Garry '- Herald the second and third cycles. 'S; ~ Q) Q) 0 0 > > "C ~ a: Selkirk a: ~ Q) "C "C The Stony Mountain Formation consists of thick-bedded (>0.5 m), 0 a.. Q) Q) a.. a: ? Cat Head a: Yeoman sparsely fossiliferous, burrow-mottled, commonly nodular, dolomud­ :::J stone which grades into more thinly bedded to laminated dolomud­ Winnipeg Winnipeg stone in the upper half of the formation. Flat-lying 'table top' outcrops characterize this formation.

.!!!" .5 E u'" Precambrian

Stratigraphic nomenclature Figure 8: Correlation chart for Phanerozoic bedrock of Phanerozoic bedrock, units in the NATMAP Shield Margin Project area. NATMAP Shield Margin area

35 The Stonewall Formation is a sparsely fossiliferous dolostone with several thin argillaceous marker beds, including one designated as the 'T marker' . The marker beds appear to reflect the regressive phase of deposi­ tional cycles. Conodont data provided by GSC Calgary indicate that the Ordovician-Silurian boundary is located within the upper 3 m ofthe Stonewall Formation, above the T marker, at a Stonewall outlier between Cormorant and Little Cormorant lakes, Manitoba (McCabe, 1986, 1988a, b; Nowlan, 1989; Bezys, 1991). The boundary occurs at a similar stratigraphic position in the Interlake area of Manitoba and in a well near Esterhazyin southern Saskatchewan (Nowlan, 1995, 1996; Haidl, 1991, 1992).

The Interlake Formation consists of dolostone ranging from variably fossiliferous and stromatolitic to pre­ dominantly finely crystalline, and dense to sublithographic (Baillie, 1951; Stearn, 1956; Bezys, 1989, 1991; Haidl, 1992). Several thin, sandy, argillaceous marker beds are present. Glacio-eustatic sea level changes during Interlake deposition may have been responsible for the transgressive-regressive cycles observed in these rocks (Johnson and Lescinsky, 1986). In the eastern part of the area, the Interlake can be subdivided into the East Arm, Atikameg, Moose Lake, and Fisher Branch formations and, therefore, has Group status. This stratigraphic nomenclature is modified from that of Stearn, 1956 (Bezys, 1989, 1991; Norford et aI., 1994).

Middle Devonian sediments (Winnipegosis and Ashern formations) were probabl y deposited over the entire study area, but are preserved only in the southwest. The Ashern Formation, which unconformably overlies the Interlake, is composed of argillaceous dolostone and dolomitic mudstone, and the Winnipegosis Forma­ tion of reefal and inter-reefal carbonate strata.

Mesozoic geology

Clastic sediments tentatively correlated with the Success Formation (Jura-Cretaceous) in southern Saskatchewan are preserved in paleovalleys eroded into Ordovician carbonates in the western portion of the study area. At Pinechannel Mossy River Kimberlite 67B, 6-19-60-7W2 in the Stonewall outcrop belt, Red River dolostone is unconformably overlain by 25 metres ofSuccess mudstone and sandstone (Gilboy, 1995). Core descriptions in the assessment files from an area south of Twigge Lake describe anomalous thinning of carbonates and associated thickening of 'overburden' . This relationship suggests that a Success-filled paleo­ valley may also be preserved in this area.

QUATERNARY GEOLOGY

Surficial geology mapping is a major component of the Quaternary geology studies initiated in 1991 as part of the NATMAP Shield Margin Project in central Manitoba and Saskatchewan. The surficial geology maps were compiled at the scale of 1: 100 000, based on extensive fieldwork, air-photo interpretation, and LANDSAT TM images. To date, 10 maps have been compiled digitally and released as GSC color Open File maps (Campbell and Henderson, 1996; Campbell et aI., 1997; McMartin, 1993, 1994, 1997a, b, c, 1998; McMartin and Boucher, 1995; McMartin et aI., 1995). The remainder of the maps will be published in 1998 (Campbell et aI., in press).

36 The surficial geology component of the NATMAP Shield Margin Project has generated a comprehensive Quaternary knowledge base through 1) regional mapping of surficial geology, 2) drift-prospecting studies as an aid to mineral exploration for gold, base metals, and diamonds, 3) reconstruction of glacial and deglacial history, 4) development of a surficial materials database for land-use planning, and 5) environmental studies on distribution and stability of base metals in soils, with a focus on smelter emissions around Flin Flon (Henderson, 1995a, b; McMartin, 1996; McMartin and Pringle, 1994; McMartin et aI., 1996). A fundamen­ tal advance has been the recognition of at least nine distinct ice-flow events across the project area during Quaternary glaciations, which has profound implications for drift prospecting using indicator minerals and/or till geochemistry. Studies on the distribution and speciation of smelter-related metals in the surficial environment have helped to assess the relative contribution of the metals from natural and anthropogenic sources (Henderson and McMartin, 1995; Henderson et aI., in press; McMartin et aI., in press).

REFERENCES

Ansdell, K.M. 1993: U -Pb zircon constraints on the timing and provenance of fluvial sedimentary rocks in the Flin Flon and Athapapuskow basins, Flin Flon Domain, Trans-Hudson Orogen, Manitoba and Saskatchewan; in Radiogenic Age and Isotopic Studies: Report 7; Geological Survey of Canada, Paper 93-2, p. 49-57.

Ansdell, K.M. and Connors, K.A. 1994: Geochronology of continental sedimentary and volcanic rocks in the Flin Flon Domain; in LITHOPROBE: Trans-Hudson Orogen Transect Workshop, LITHOPROBE Report 38, p.205-213.

Ansdell, K.M. and Kyser, K. 1991: The geochemistry and fluid history of the Proterozoic Laurel Lake Au-Ag deposit, Flin Flon greenstone belt; Canadian Journal of Earth Sciences, v. 28, p. 155-171.

1992: Mesothermal gold mineralization in a Proterozoic greenstone belt: Western Flin Flon Domain, Saskatchewan, Canada; Economic Geology, v. 87, p. 1496-1524.

Ansdell, K.M. and Norman, A. R. 1993: The timing of magmatism, deformation and metamorphism along the southern flank of the Kisseynew Domain; Trans-Hudson Transect, LITHOPROBE Report No. 34, p. 219-225.

1995: U-Pb geochronology and tectonic development of the southern flank of the Kisseynew Domain, Trans-Hudson Orogen, Canada; Precambrian Research, v. 72, p. 147-167.

Ansdell, K.M. and Stern, R.A. 1997: SHRIMP geochronology in the Trans-Hudson Orogen: detrital zircons in turbiditic and fluvial sedimentary rocks; in LITHOPROBE: Trans-Hudson Orogen Transect Workshop , LITHOPROBE Report 62, p. 119-127.

Ansdell, K.M., Connors, K. A., Stern, R. A., and Lucas, S.B., in press: Coeval sedimentation, magmatism, and fold-thrust belt development in the Trans-Hudson Orogen: geochronological evidence from the Wekusko Lake area, Manitoba, Canada; Canadian Journal of Earth Sciences.

37 Ansdell, K.M., Kyser, K., Stauffer, M., and Edwards, G. 1992: Age and source of detrital zircons from the Missi Group: a Proterozoic molasse deposit, Trans-Hudson Orogen, Canada; Canadian Journal of Earth Sciences, v. 29, p. 2583-2594.

Ansdell, K.M., Lucas, S.B., Connors, K. A., and Stern, R. A. 1995: Kisseynew metasedimentary gneiss belt, Trans-Hudson Orogen (Canada); back-arc origin and collisional inversion; Geology, v. 23, p. 1039-1043.

Ashton, K.E. 1992: Geology of the Snake Rapids area: update; in Summary of Investigations 1992, Saskatchewan Geological Survey, Saskatchewan Energy and Mines, Miscellaneous Report 92-4, p. 97-113.

Ashton, K.E. and Balzer, S. 1995: Wildnest-Tabbernor Transect: Pelican Lake-Tabbernor Fault area (part of 63M-3); in Summary of Investigations 1995, Saskatchewan Geological Survey, Saskatchewan Energy and Mines, Miscellaneous Report 95-4, p. 13-22.

Ashton, K.E. and Digel, S. 1992: Metamorphic pressure-temperature results from the Attitti Block; in Summary of Investigations 1992, Saskatchewan Geological Survey, Saskatchewan Energy and Mines, Miscellaneous Report 92-4, p. 114-116.

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48 McMartin,1. (cont.) 1997c: Surficial geology, Reed Lake area, Manitoba and Saskatchewan (NTS 63 K/9, K/l 0); Geological Survey of Canada, Open File 3406, scale 1: 100 000.

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