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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
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A.H. Bailes1, K.E. Ashtod, and F.M. Haidl 3
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Manitoba Geological Services Branch
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3 Saskatchewan Geological Survey
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Table of Contents 1
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Introduction
Comments on map presentation
Trans-Hudson Orogen overview
Precambrian geology
Introduction
Metamorphism
Superior Boundary Zone (unit A)
Paleoproterozoic and Archean rocks in the Pelican Window (units A2, A3, A4)
1.92-1.87 Ga volcanic, intrusive, and sedimentary rocks (units J, F, U, E)
Juvenile arc assemblages (unit J)
Flin Flon arc assemblage
Snow Lake arc assemblage
Fourmile Island arc assemblage
West Amisk arc assemblage
Birch Lake arc assemblage
Hanson Lake arc assemblage
Kisseynew Domain 'south flank'
Ocean-floor assemblages (unit F)
MORB-like basalts (units Fl, F2, F4, FS)
Synvolcanic mafic to ultramafic complexes (unit F6)
Ocean-island basalts (unit F3c)
Ocean-plateau basalts (units F3a, b, d)
Unknown geochemical affinity (unit U)
Isotopically evolved Proterozoic and Archean rocks
Archean crustal fragments (unit AI)
Evolved arc (unit E)
Tectonic setting of arc and ocean-floor assemblages
Neodymium-isotopic and geochronological evidence
1.88-1.83 Ga intrusive rocks (unit P)
Geochemistry and source constraints
Rocks of uncertain age (unit W)
Sub-Phanerozoic Precambrian geology
1.87-1.83 Ga sedimentary, volcanic and intrusive rocks (units S, M, B)
Schist-Wekusko assemblage (unit S)
Missi Group (unit M)
Burntwood Group (unit B)
Late intrusive rocks (unit L)
Mineral deposits
Volcanic-hosted massive-sulphide deposits
Gold deposits
Tectonic evolution
D 1: Accretion
D2 : Postaccretion magmatism, sedimentation, and deformation
D3: Collision
D4-DS: Postcollision
Phanerozoic bedrock geology
Paleozoic geology
Mesozoic geology
Quaternary geology
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 of federal, 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. % Si0 2), basaltic andesite (52-57 wt. % Si0 2), andesite (57-63 wt. % SiO), dacite (63-70 wt. %
Si0 2) and rhyolite (70-80 wt. % Si0 2). 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 of geochemical 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 of key elements of the 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 of North 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 of these
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 of the Cree Lake Zone now included in Hearne Prov­
ince (Hoffman, 1989; Lewry and Stauffer, 1990).
,
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Phanerozoic
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Sedimentary rocks
Paleoproterozoic
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Sedimentary rocks
Trans-Hudson Orogen
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plutonic rocks
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collisional sedimentary
and plutonic rocks
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L...-_ _ _ _ _ _ _ _ _ _ _ _ _.......;;...;.._ _ _ _ _.........:._ _ _.........
Faults ----------
Archean cratons/
Pikwitonei Granulite Belt
Archean (exposed in internal domains)
LlTHOPROBE seismic reflection lines
Figure 2. Map of the 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 of voluminous 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
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3D Model of the Flin Flon Belt (Manitoba & Saskatchewan)
(NATMAP Shield Margin Project, LITHOPROBE Trans-Hudson Orogen Transect)
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Digital elevation models of the Quaternary, Phanerozoic and Precambrian
surfaces (>2000 drillhole intersections)
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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).
signature, age, and inferred plate-tectonic
setting (see below; Lucas et aI., 1996). The
1.92-1.87 Ga tectonostratigraphic assem­
blages recognized during the Shield Margin
Project include those with juvenile arc,
juvenile ocean-floor (including MORB-like
basalt, ocean-plateau basalt and ocean­
island basalt) and evolved arc affinities
(Fig. 5; Syme and Bailes, 1993; Stem et aI.,
1995a, b; David and Syme, 1994; Lucas et
aI., 1996). This has important economic
implications as not all of the tectonostra­
tigraphic assemblages are equally endowed
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.
Conventional stratigraphy of
the Flin Flon Belt
Missi GP.
-1.845 Ga
Amisk GP.
-1 .9 Ga
'. _, ;;Ol~a~ic'}~
"." .; rocks :' '
sandstone~
conglomerate
New tectonostratigraphy of
the Flin Flon Belt
Amisk
Collage
Metamorphism
1
Figure 4: Upper: Cartoon showing the conventional stratigra­
phy of the Flin Flon Belt (Bruce, 1918; Harrison, 1951).
Lower: Cartoon showing the tectonostratigraphic framework
for the NATMAP Shield Margin Project area (Lucas et aI., 1997).
On the exposed Shield, peak regional meta­
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
Digel, 1992; Digel and Gordon, 1995;
Maxeiner et aI., 1995; Kraus and Menard,
1997; Menard and Gordon, in press).
A general northward-increasing gradient is
interrupted in the west by transitional upper
amphibolite- to granulite-facies assemblages,
marking a metamorphic culmination in rocks of
the Pelican Window (Ashton et aI., in press).
7
co Elbow-Athapap ocean floor assg.
20km
PRE-ACCRETION ASSEMBLAGES (1.87-1.92 Ga)
_
Juvenile-arc and undivided metavolcanic rocks
*
[ B Ocean-floor (back arc) metabasalVsynvolcanic mafic intrusive ~ Ocean-plateau metabasalt
[[[[[II] Tectonite
Ocean-island metabasalt PELICAN WINDOW GNEISSES
~ Archean charnockite
Ell! Orthogneiss and pelitic gneiss [;iii
~
VMS deposit
Au deposit
Sillimanite isograd
.c..
Iwoi3X~
c=J
L:~:~<
I
~
~
FAULT
H.V. Zwanzia 1998
SUCCESSOR-ARC and BASIN DEPOSITS
~
~
"
PHANEROZOIC
~
Missi Group (1.83-1.85 Ga)
Continental sandstone / volcanics
Burntwood Group turbidites (1.84 - 1.85 Ga)
Schist-Wekusko Suite (1.85-1.88 Ga)
FELSIC-MAFIC PLUTONS
1.76 - 1.82 Ga (Kisseynew Belt plutons)
1.83 - 1.84 Ga (late successor-arc plutons)
1.84 - 1.90 Ga (early juvenile-arc + early-middle
successor-arc plutons)
ca. 1.92 Ga
('evolved-arc' plutons)
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 of the northern limit of pre­
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 of no 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 of the 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.
Rock Type
Unit
Type
Age
(Ma)
Error
(Ma)
Method
Mineral
Reference
UTM-E
UTM-N
Zone
UNIT
No.
Jan Lake Complex
Quartzofeldspathic
gneiss
igneous
2959
±13
U/Pb
zircon
Ashton et al. (in
press)
641870
6101120
13
A2
Missi Group
Sandstone lens in
conglomerate
detrital
2721
±4
Pb-Pb
zircon
Ansdell (1993)
323173
6055822
14
MIa
Stroud Felsic
Breccia
Rhyolite breccia
inherited
2715
Pb-Pb
zircon
David et al. (1996)
427600
6074000
14
J7a
inherited
2604
Pb-Pb
zircon
Stern et al. (in
press)
Wolf Lake Turbiditic Turbiditic wacke
Wacke
detrital
2523
±12
Pb-evap
zircon
Ansdell and
Connors (1994)
663224
6053536
13'
J9b
Beaverhouse Lake
Tonalite
Tonalite
igneous
2518
+7/-4
U/Pb
zircon
David and Syme
(1994)
327453
6069129
14
Ala
Felsic Mylonite. NE
Arm Shear Zone
Granodiorite
igneous
2497
+3/-2
U/Pb
zircon
David and Syme
(1994)
322420
6061324
14
Ala
Missi Group; Flin
Sandstone (lower
Flon Basin;
sequence)
Beaverdam Member
detrital
2494
±2
U/Pb
zircon
Ansdell (1993)
702154
6072398
13'
M3a
Jan Lake Complex
Gabbro
igneous
2488
±12
U/Pb
zircon
Ashton et al. (in
press)
639800
6109140
13
Ala
Jan Lake Complex
Chamockite
igneous
2450
±8
UlPb
zircon
Ashton et al. (in
press)
636290
6104700
13
A4
Burntwood Group
Greywacke
detrital
2397
Schist Lake Road
Felsic Dyke
Aplitic dyke
inherited
2022
Jan Lake Complex
Migmatitic
paragneiss
detrital
2016
Meridian Creek
Tonalite
Tonalite
igneous
1920
Missi Group; Flin
Flon Basin
Conglomerate
(lower sequence)
detrital
Mystic Lake
Intrusive Suite
Tonalite
Mystic Lake Tonalite Tonalite
Ela
Pb-Pb
zircon
David et al. (1996)
406450
6074500
14
Bla
Pb-Pb
zircon
Stern et al. (1993)
315125
6058810
14
Pl0a
Pb-Pb
zircon
Ashton et al. (in
press)
641390
6101560
13
A2
+18/-13
U/Pb
zircon
Stern and Lucas
(1994)
309850
6067400
14
Ela
1914
±6
Pb-evap
zircon
Ansdell et al.
(1992)
697745
6076389
13"
MIa
igneous
1906
±2
U/Pb
zircon
Heaman et al.
(1992)
309165
6060575
14
Ela
Wo~ Lake Turbiditic Turbiditic wacke
Wacke
detrital
1906
±20
Pb-evap
zircon
Ansdell and
Connors (1994)
663224
6053536
13'
J9b
Athapapuskow
Diabase
Diabase
igneous
1904
±4
U/Pb
zircon
Stern et al.
(1995b)
323550
6047350
14
F5a
Mine Rhyolite
Rhyolite
igneous
1903
+7/-5
U/Pb
zircon
Stern et al. (in
press)
314096
6073283
14
J4a
West Arm Tonalite
Tonalite
igneous
1903
+6/-4
U/Pb
zircon
Stern et al. (1993)
318378
6050352
14
Ela
Claw Lake Gabbro
Gabbro pegmatite
dyke
igneous
1901
+6/-5
U/Pb
zircon
Stern et al.
(1995b)
385286
6073221
14
F6a
Herblet Lake Gneiss Melanocratic gneiss
Dome
igneous
1901
±4
U/Pb
zircon
David et al. (1996)
445550
6102385
14
J14
Welsh Lake
Assemblage
Greywacke
detrital
1894
±3
U/Pb
zircon
Heaman et al.
(1993)
665131
6074326
13
J9b
Stroud Felsic
Breccia
Rhyolite breccia
igneous
1892
±3
UlPb
zircon
Machado and
David (1992)
429129
6073267
14
J7a
Herblet Lake Gneiss Granodiorite
Dome
igneous
1890
+8/-6
U/Pb
zircon
Gordon et al.
(1990)
443643
6089923
14
J14b
±2
Richard Lake Pluton Tonalite
igneous
1889
+8/-6
U/Pb
zircon
Bailes et al. (1990)
426880
6081044
14
J12c
Bakers Narrows
Rhyolite Dyke
Porphyritic rhyolite
igneous
1888
+11/-9
U/Pb
zircon
Stern et al. (in
press)
330454
6067770
14
JI3c
Neagle Lake
Rhyolite
Rhyolite (?)
igneous
1888
±3
U/Pb
zircon
Heaman etal.
(1993)
666713
6071812
13
J7b
igneous
1887
±3
U/Pb
zircon
Heaman etal.
676450
6065750
13
J4a
Laurel Lake Rhyolite Rhyolite
• approximate location
.
•
I
Table 1. (cont.)
I
Rock Type
Unit
Type
Age
(Ma)
Error
(Ma)
Method
Mineral
igneous
1886
±1
u/Pb
zircon
UTM·E
UTM·N
Zone
UNIT
No.
Stern et al. (in
press)
317250
6076000
14
J12c
Reference
,
(1993)
Cliff Lake Pluton
Quartz diorite
Sneath Lake Pluton Tonalite
igneous
1886
+17/-9
U/Pb
zircon
Bailes et al. (1990)
428914
6069820
14
J12c
Two Portage
Rhyolite crystal tuff
Rhyolite Crystal Tuff
igneous
1886
±2
U/Pb
zircon
Gordon et al.
(1990)
325860
6067880
14
J2a
Vick Lake Tuff
igneous
1885
±3
Pb-Pb
zircon
Stern et al. (1993)
302670
6068240
14
J8a
Herblet Lake Gneiss Melanocratic gneiss
Dome
igneous
1884
±6
U/Pb
zircon
David et al. (1996)
445550
6102385
14
J14
Mystic Lake
Intrusive Suite
Tonalite-quartz
diorite
igneous
1883
±4
U/ Pb
zircon
Heaman et al.
(1992)
309110
6060240
14
Ela
West Amisk
Andesitic Tuff
Andesitic tuff
igneous
1882
±3
U/Pb
zircon
Stern and Lucas
(1994)
672055
6057596
13
J6a
Mikanagan Lake Sill Gabbro pegmatite
igneous
1881
+3/-2
U/Pb
titanite
Stern et al. (in
press)
327627
6071249
14
Jl0
Gants Lake Batholith Granodiorite
igneous
1876
+7/-6
U/Pb
zircon
Whalen and Hunt
(1994)
385740
6081260
14
P7a
Puella Bay Dacite
Dacite
igneous
1876
±2
U/Pb
zircon
Ansdell et al. (in
press)
450960
6064440
14
S3b
Hanson Lake
Rhyolite
Rhyolite
igneous
1875
±1
U/Pb
zircon
Heaman etal.
(1993)
639500
6061000
13
J4a
Jungle Lake Pluton
Tonalite
igneous
1874
±2
u/Pb
zircon
Machado et al. (in
press)
374075
6116600
14
P6a
Finger Peninsula
Sheared Granite
Granite
igneous
1873
±1.6
U/Pb
zircon
Heaman et al.
(1994)
638130
6066280
13
J12a
Ragged Lake Pluton Granite gneiss
igneous
1873
±4
U/Pb
zircon
Hunt and Zwanzig
(1990)
374303
6100053
14
P9a
Hanson Lake Mine
Road Granite
Granite
igneous
1870
+7/-5
U/Pb
zircon
Heaman etal.
(1994)
635740
6063500
13
J12a
Schist Lake Road
Felsic Dyke
Aplitic dyke
igneous
1869
±1
Pb-Pb
zircon
Stern et al. (1993)
315125
6058810
14
Pl0a
Scheiders Bay
Rhyolite crystal tuff
Rhyolite Crystal Tuff
detrital
1867
±17
U/Pb
zircon
Stern et al. (in
press)
327406
6059436
14
SIc
Annabel Lake Pluton Granodiorite
igneous
1866
±3
U/ Pb
zircon
Stern and Lucas
(1994)
311450
6074850
14
P7a
East Elbow Tonalite Quartz megacrystic
Stock
tonalite
igneous
1864
±3
U/Pb
zircon
Whalen and Hunt
(1994)
382630
6081518
14
P6a
Elbow Lake Pluton
Tonalite
igneous
1864
+5/-4
U/Pb
zircon
Whalen and Hunt
(1994)
381225
6074488
14
P6a
Missi Group
Cross-bedded
sandstone
detrital
1864
±7
Pb-Pb
zircon
Ansdell et al. (in
press)
West Arm
Sandstone
Turbiditic sandstone
detrrtal
1863
±3
Pb-Pb
zircon
Stern et al. (in
press)
320782
6071821
14
SIb
Pluton intruding
Hanson Lake
metasediments
Quartz-feldspar
porphyry
igneous
1861
+4/-3
U/Pb
zircon
Heaman et al.
(1997)
639327
6065954
13
Pl0b
Missi Group; Flin
Flon Basin
Conglomerate
(lower sequence)
detrital
1861
±17
Pb-evap
zircon
Ansdell et al.
(1992)
697745
6076389
13"
MIa
Puffy Lake Tonalite
Tonalite
igneous
1860
±2
U/Pb
zircon
Machado et al. (in
press)
370150
6099500
14
P7e
igneous
1859
+14/-7
U/Pb
zircon
Hunt and Zwanzig
(1993)
391400
6114890
14
P8c
Shoshonitic tuff
Batty Lake Intrusive Tonalite gneiss
Complex
" approximate location
M3a
\
Table 1. (cont.)
Rock Type
Unit
Type
Age
(Ma)
Error
(Ma)
Method
Mineral
Reference
UTM-E
UTM-N
Zone
442400
6075300
14
Burntwood Group
Greywacke
detrital
1859
±7
Pb-Pb
zircon
David et al. (1996)
Missi Group
Cross-bedded
sandstone
detrital
1859
±8
Pb-Pb
zircon
Ansdell et al. (in
press)
UNIT
No.
Bla
M3a
Neso Lake Pluton
Quartz diorite
igneous
1858
±3
UlPb
zircon
Syme et al. (1991)
332482
6059654
14
P5a
Schist Lake
Conglomerate
Trachyandesite
igneous
1858
±4
U/Pb
zircon
Stern et al. (in
press)
320374
6058788
14
Sla
Schist Lake
Conglomerate
Qz-fp phyric rhyolite
cobble
igneous
1858
±1
UlPb
zircon
Stern et al. (in
press)
320782
6058583
14
Sla
Missi Group; Flin
Sandstone (lower
Flon Basin;
sequence)
Beaverdam Member
detrital
1857
±2
U/Pb
zircon
Ansdell (1993)
702154
6072398
13'
M3a
East Wekusko
Rhyolite
igneous
1856
±1
UlPb
zircon
Ansdell et al. (in
press)
449180
6069140
14
S3a
Kaminis Lake Pluton Granodiorite
igneous
1856
±2
U/Pb
zircon
Stern and Lucas
(1994)
313750
6059750
14
P7a
Mari Lake Pluton
Biotite granodiorite
igneous
1855
+4/-3
U/Pb
zircon
Heaman etal.
(1992)
308025
6096090
14
P8b
Kakinagimak Lake
Leucotonalite
Biotite leucotonalite
igneous
1852
+6/-4
U/Pb
zircon
Heaman et al.
(1993)
669565
6104335
13
P8b
Reynard Lake
Pluton
Porphyritic
granodiorite
igneous
1850
±3
U/Pb
zircon
Stern et al. (1993)
689775
6062125
13
P7a
Wekach Lake Pluton Gabbro
igneous
1850
±2
UlPb
zircon
Heaman etal.
(1992)
311400
6063775
14
P2a
Missi Island
Trondhjemite
Trondhjemite
igneous
1848
±11
Pb-evap
zircon
Ansdell and Kyser
(1991)
678000
6064250
13
P8a
Puffy Lake
Metadacite
Dacite
igneous
1848
±4
UlPb
zircon
Machado et al. (in
press)
372150
6100100
14
M5d
Lynx Lake Pluton
GranodiOrite
igneous
1847
±4
U/Pb
zircon
Gordon et al.
(1990)
333749
6052482
14
P7a
Syn-Meridian-West
Arm Shear Zone
Dyke
Plagioclase­
porphyry dyke
igneous
1847
±2
U/Pb
zircon
Stern and Lucas
(1994)
309400
6069012
14
Pl0b
MissiGroup
Pebbly sandstone
detrital
1847
±2
Pb-Pb
zircon
Ansdell (1993)
317480
6073800
14
M2a
igneous
1846
+14/-6
Pb/Pb
titanite
Ashton (1992)
667191
6073948
13
P6c
detrital
1846
±6
Pb-Pb
zircon
Ansdell (1993)
323173
6055822
14
Mia
Big Rat Lake Pluton Granodiorite
igneous
1845
±3
Pb-Pb
zircon
Whalen and Hunt
(1994)
377307
6077068
14
P7a
Hanson Lake Pluton Granodiorite
igneous
1844
±2
UlPb
zircon
Heaman et al.
(1993)
642921
6064730
13
P7a
Hanson Lake Pluton Gabbro
igneous
1843
±2
UlPb
zircon
Heaman etal.
(1994)
640833
6060206
13
P2a
Pelican
Decollement Zone
Quartz monzodiorite
igneous
1843
±2
U/ Pb
zircon
Ashton et al. (in
press)
652450
6086575
13
P9c
Boundary Intrusion
Metagabbro
igneous
1842
±3
U/Pb
zircon
Heaman etal.
(1992)
315770
6065500
14
P2g
Burntwood Group
Psammite
detrital
1842
±2
U/Pb
zircon
Machado et al. (in
press)
390650
6116675
14
Bib
Post-Meridian-West M icrocline-porphyry
Arm Shear Zone
dyke
Dyke
igneous
1839
±3
U/Pb
zircon
Stern and Lucas
(1994)
309475
6068919
14
Pl0b
Boot Lake Intrusion
igneous
1838
±2
U/Pb
zircon
Heaman et al.
(1992)
315770
6065500
14
P4c
Rhyolite
Neagle Lake Pluton Granodiorite to
quartz monzodiorite
Missi Group
Sandstone lens in
conglomerate
Monzodiorite
, approximate location
.
Table 1. (cont.)
Unit
Rock Type
Type
Age
(Ma)
Error
(Ma)
Method
Mineral
Reference
UTM-E
UTM-N
Zone
UNIT
No.
Phantom Lake
Granodiorite Dyke
Granodiorite
igneous
1838
±2
U/Pb
zircon
Heaman et al.
(1992)
315770
6065500
14
P7a
Tramping Lake
Pluton
Granite
igneous
1837
+8/-6
U/Pb
zircon
David et al. (1996)
436260
6072890
14
P9a
Missi Group
Metasandstone
detrital
1837
±4
Pb-Pb
zircon
Machado et al. (in
press)
J
Mlb
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
press)
452840
6071083
14
M5a I
Wekusko Granite
Granite
igneous
1834
+8/-6
U/Pb
zircon
Gordon et al.
(1990)
438981
6068838
14
P9a
Herb Lake Felsic
Volcanic
Rhyolite
igneous
1833
+6/-2
U/Pb
zircon
Ansdell et al. (in
press)
450860
6071110
14
M5a
Nelson Bay Gneiss
Dome
Granodiorite gneiss
igneous
1832
±2
U/Pb
zircon
David et al. (1996)
413226
6085967
14
P7e
Rex Lake Plutonic
Complex
Granodiorite
igneous
1832
+4/-3
UlPb
zircon
Gordon et al.
(1990)
450113
6076491
14
P7a
Ham Lake Pluton
Granodiorite
igneous
1830
+27/-19
U/Pb
zircon
Gordon etal.
(1990)
417475
6083464
14
P7a
Mirond Lake
Enderbite
Homogeneous calc­
alkaline enderbite
igneous
1830
±1
U/Pb
zircon
Ashton et al. (in
press)
645450
6103970
13
P5d
Touchbourne
Intrusive Suite
Enderbite
igneous
1830
+11 /-5
U/Pb
zircon
Gordon et al.
(1990)
414000
6130900
14
P5d
Little Swan Lake
Pluton
Granodiorite
igneous
1828
±6
Pb-Pb
titanite
Whalen and Hunt
(1994)
338380
6073200
14
P9b
Missi Group
Quartz porphyry 1
pink gneiSS
igneous
1826
+11/-5
U/Pb
zircon
Hunt and
Schledewitz (1992)
353834
6097861
14
M6b
Anderson Rhyolite
Rhyolite
metamorphic
1812
±15
U/ Pb
titanite
Hanson Lake
Psammopelitic
Schist
Psammopelitic
schist
metamorphic
1808
±2.6
U/Pb
monazite
Pelican
Decollement Zone
Quartz monzodiorite
metamorphic
1808
±3
U/Pb
Attitti Lake Felsic
Volcanic
Felsic gneiss
(volcanic ?)
metamorphic
1807
+3/-2
metamorphic
1807
igneous
Herblet Lake Gneiss Tonalite gneiss
Dome
David et al. (1996)
438778
6078845
14
J4a
Heaman etal.
(1994)
639450
6066280
13
J9b
titanite
Ashton et al. (in
press)
652450
6086575
13
P9c
U/Pb
zircon
Heaman et al
(1992)
664360
6104205
13
Ulb
±3
U/Pb
zircon
David et al. (1996)
446780
6102510
14
J14a
1806
±2
UlPb
zircon
Ashton (1992)
658530
6095230
13
L4
metamorphic
1804
+5/-4
U/Pb
titanite
Heaman et al.
(1994)
634400
6064980
13
J4a
±2
U/Pb
monazite
Ansdell and
Nonman (1993)
337814
6097681
14
Ll
Pb-Pb
zircon
Heaman etal.
(1991)
308910
6059950
14
L1
Hunt and Zwanzig
(1993)
388430
6116480
14
l1
Pre-Sturgeon-weir
Shear Zone
Pegmatite
Granitoid pegmatite
Hanson Lake Mine
Road Rhyodacite
Rhyodacite
Pegmatitic Granite
Granite
igneous
1799
Mystic Lake
Intrusive Suite
Porphyritic rhyolite
dyke
igneous
1797
Pegmatite Dyke
Pegmatite
igneous
1796
±1
Pb-Pb
monazite
Kakinagimak Lake
Leucotonalite
Biotite leucotonalite
metamorphic
1789
±3
U/Pb
titanite
Heaman et al.
(1993)
669565
6104335
13
P8b
Jan Lake Granite
Suite
Aplitic-pegmatitic
granite
igneous
1773
±9
UlPb
zircon
Bickford et al.
(1987)
633396
6086198
13
L3a
Post Sturgeon-weir
Shear Zone
Pegmatite
Granitoid pegmatite
metamorphic
1767
±1
U/ Pb
monazite
Ashton (1992)
653620
6091600
13
L4
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
20-50 km across (Hanson Lake, West Amisk, Birch
Lake, Flin Flon, Fourmile Island, and Snow Lake
assemblages; Fig. 1b, 5). These are separated by
major faults or intervening ocean-floor rocks (unit
F), Burntwood Group turbidites (unit B), plutons
(unit P), or a combination ofthese rock types. The arc
assemblages are internally complex, comprising
numerous fault-bounded and folded volcanic suites
(e.g. Bailes and Syme, 1989), rendering correlation
of volcanic stratigraphy within and between the
assemblages nearly impossible. It is unclear whether
the segments represent the fragmented parts of a for­
merly single are, or were generated in completely
different arcs (e.g. Syme et aI., 1995; Lucas et aI.,
1996).
+4 FUN
FLON
MANTLE
---------~~~77~!r
+2
Missi Group
(fluvial sediments)
CHUR
o~-------------------------+~~~
Postaccretion Rocks
(Successor Arc & Basins)
-2
(Amisk Collage)
"C
Z
c::
..2
'iii
Pre-Accretion
Assemblages
Evolved Arc -4
0.
(Mystic Lake Assem blage) w
-6
Sask Craton
-8
Postcollisional
-12
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
ronment (Bailes and Syme, 1989). Basalt and
basaltic andesite flows dominate the assemblage, and
belong mainly to tholeiitic (unit J1a) and calc­
alkaline (unit J1 b) suites (Stem et aI., 1995a, b).
Locally, the calc-alkaline rocks are stratigraphically
associated with terrigenous turbidites (unit J9),
16 1770
1790
1810
1830
1850
1870
1890
1910
Age (Ma)
Figure 6: CNd VS. time plot for units from the Flin Flon Belt.
Neodymium-isotopic data from Stern et al. (1992, 1993,
1995a, b, unpub. data), Whalen et al. (in press), Bickford
et al. (1992) and Mock et al. (1993) .
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 of bubble-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 of either 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 of sheeted 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 of their 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 Al 20 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 of the
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 of the 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 of metres 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 of the 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 of these 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 of arc 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 Ti0 2
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) of the 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 of detrital 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 of continental crust locally forming the basement to arcs (e.g.
Mystic Lake assemblage and possibly part of the 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 of the 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. % Si0 2 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 of adjacent 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 of up 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 of the 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 of mUltiple intrusive phases (Leclair et al., 1997). Uranium-lead zircon dating of drill-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
103 0
102 0
lor
102
lor
1000
99 0
55 0 540
0
1000
54·
990
99°
1030
55 0 103 0
102 0
101 0
1000 54"
99 0
55 '
54'
26
Figure 7: Upper: Shaded relief representation of
the total field aeromagnetic data for the NATMAP
Shield Margin Project area (cf. Broome and Viljoen,
in press).
Middle: Bouger anomaly gravity map of the
NATMAP Shield Margin Project area (cf. Broome
and Viljoen, in press).
Lower: Major lithotectonic domains and regional
tectonic framework of the exposed and buried Flin
Flon Belt (Leclair et al. 1997). The Hanson Lake
Block, eastern Glennie Domain, Tabbernor Fault
Zone, Kisseynew Domain, and the West Amisk,
Elbow-Athapap and Snow Lake assemblages
extend southward into the subsurface. Abbreviations
are: BCF - Berry Creek Fault; CBF - Crowduck
Bay Fault; ELSZ - Elbow Lake Shear Zone; NLS­
Namew Lake Structure; SASZ - South
Athapapuskow Shear Zone; SLF - Suggi Lake
Fault; SRSZ - Spruce Rapids Shear Zone;
SWSZ - Sturgeon-weir Shear Zone.
oflinear north-northwest-striking structures, possi­
bly representing fractures or faults. Seven widely
spaced drillholes in the Cormorant Batholith
magnetic-gravity low have all intersected unde­
formed monzogranite, two of which have yielded
U-Pb zircon ages of 1.830 Ga (Leclair et aI., 1997).
1.87-1.83 Ga sedimentary, volcanic
and intrusive rocks (units S, M, B)
Volcanic, volcaniclastic, and sedimentary rocks
that are younger than the 1.92-1. 87 Ga tectonostra­
tigraphic assemblages (Fig. 5) have been docu­
mented across the central part of the Flin Flon belt,
and are termed 'successor' basin deposits, in paral­
lel with terminology for the 1.88-1.83 Ga mag­
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 of this 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. 1b, 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.
:r
ECTONIC 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 of Flin 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 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 accretionary
Complex beyond the Amisk Collage (Fig. 5).
j
j
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.
Episode
D1 D2
Structures
Magmatism
~ , L1 (tectonites,
mylonites) None (?)
52' F2 (Vick Lake
synform) S1/S2
(tectonites, mylonites)
Mafic to felsic dykes,
sheets, plutons
None
S3'
f3; shear bands (both
Metamorphism
Age
(Ma)
1880-1870
Intra-oceanic accretion
Subgreenschist to
amphibolite ('contact'
to regional)
1880-1840
Intra-arc shortening , uplift/erosion of Amisk collage; development of Kisseynew back-arc basin Regional peak
metamorphism
(subgreenschist to
amphibolite facies)
at 1820-1805 Ma
1840-1805
Regional collisional
shortening and thickening via
SW-vergent thrusts & folds ;
high-angle shear zones in
Amsik collage
?
D3
sinistral & dextral) ; highangle shear zones; SW­
vergent thrusts (e.g. Morton
Lake thrust zone, Fig. 2)
(1840-1830 Ma
magmatic belt in Snow
Lake assemblage,
south flank of the
Kisseynew Domain)
Pegmatites,
leucogranites
Retrograde
1805-1770 (?)
D4
SW-vergenlthrusts (e.g.
Sturgeon-weirshear
zone, Fig. 2); F4 folds &
kinks, high-angle shear
zones
None
Retrograde
1770-1690
D5
Brittle to ductile shear
zones/faults
Tectonic context
Postcollisional thrusting, folding; transpression of Amisk collage Postcollisional NW-SE
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 of magmatism 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 of initial
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 (D 4 , 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 (D s' 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 of Hearne 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
Manitoba
Saskatchewan
U)
..
::J o
u
co
e
u
u
'iii
U)
co ..,~
c
co
c
o
>
Q)
o
Upper Interlake
C
CO
:::J Q) Q)
Atikameg .::£
:::J (f) .::£
CO
CO
~
Moose
-;::
(f) Q) -;::
Q)
Q)
-
~ c
0
.....J
Lake
Stonewall
T·marker
C
0
Stonewall
-
----T-marker
Stony
Mountain
"S;
CO 0
"(3 "C
"C
~
-
Stony
Mountain
CO
'(3
'S;
Lower
Interlake
c
Fisher
Branch
-----
C
Fort
Garry
'Q)
~
0 a:>
~
The dolostones of the Red River, Stony Mountain, Stonewall, and
Interlake formations are the dominant rock types in the map area.
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
(Norford et al., 1994).
East Arm
"C "C
good sorting of fine to coarse quartz grains, and an upper unit domi­
nated by argillaceous siltstone/sandstone, commonly burrowed, with
interbeds of sandstone and shale" In the northern half of the study area,
the upper unit is not present and the commonly friable basal sandstone
unit is directly overlain by an arenaceous dolostone at the base of the
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 of Reed Lake (Fig. 1a).
>
Selkirk
Q) "C
Q)
0 a..
a.. a:
Herald
'-
Q)
Three shallowing- and brining-upward cycles that are well defined in
the Red River Formation in southern Manitoba and southern
Saskatchewan (Kendall, 1976; Norford et al., 1994; Haidl et al., 1997)
also appear to be preserved in the Shield Margin area. The lower Red
River, a sparsely fossilferous, burrow-mottled dolomudstone/wacke­
stone unit, represents the lower part of the initial cycle. The upper Red
River Formation, in this area composed of interbeds of dolomudstone,
argillaceous dolomudstone, and minor dolomitic shale, encompasses
the upper portion of the first Red River cycle and most likely includes
the second and third cycles.
a:
"C
Q)
?
Cat Head
:::J
Winnipeg
a:
Yeoman
Winnipeg
The Stony Mountain Formation consists of thick-bedded (>0.5 m),
sparsely fossiliferous, burrow-mottled, commonly nodular, dolomud­
stone which grades into more thinly bedded to laminated dolomud­
stone in the upper half of the formation. Flat-lying 'table top' outcrops
characterize this formation.
"
.!!!
.5
E
u'"
Precambrian
Stratigraphic nomenclature
of Phanerozoic bedrock,
NATMAP Shield Margin area
Figure 8: Correlation chart for Phanerozoic bedrock
units in the NATMAP Shield Margin Project 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 of the 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 of Success 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).
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37
Ansdell, K.M., Kyser, K., Stauffer, M., and Edwards, G.
1992:
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1995:
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1992:
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1995:
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1994:
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38
1971:
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1985:
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1996:
10
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1978:
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Bailes, A.H. and Simms, D.
1994:
Implications of an unconformity at the base of the Threehouse Formation, Snow Lake (NTS 63
K/16); Manitoba Energy and Mines, Minerals Division, Report of Activities, 1994, p. 85-88 .
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1990:
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Baillie, A.D.
1951:
Silurian geology of the Interlake area, Manitoba; Manitoba. Department of Mines and Natural
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1989:
Stratigraphic and industrial minerals core hole program; in Manitoba Energy and Mines,
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39
Bickford, M.E., Collerson, K. D., Lewry, J. F., Van Schmus, W.R., and Chiarenzelli, J. R.
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10
press: Applications of digital methodology to the NA TMAP Shield Margin Project; Canadian Journal
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10
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40
The Nistowiak and Guncoat gneisses: implications for the tectonics of the Glennie and La Ronge
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