Metallogeny of the Paleoproterozoic Flin Flon Belt, Manitoba and Saskatchewan

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METALLOGENY OF THE PALEOPROTEROZOIC FLIN FLON BELT,

MANITOBA AND SASKATCHEWAN

ALAN G. GALLEY1, R IC SYME2, AND ALAN H. BAILES2

1Geological Survey of Canada, 601 Booth Street, Ottawa, Ontario K1A 0E8 2 Manitoba Geological Survey, 360-1395 Ellice Avenue, Winnipeg, Manitoba R3G 3P2

Corresponding author’s email:[email protected]

Abstract

For the past 75 years, the Paleoproterozoic Flin Flon Belt has been a significant contributor of base metals and Auto the Canadian economy. It is host to 27 significant volcanogenic massive sulphide deposits that originally containedover 154 Mt of ore, including the 62 Mt Flin Flon deposit. It is also host to several past-producing orogenic Au deposits,including the second >1 million ounce producer (New Britannia) within a Canadian Paleoproteroic terrane, the first beingthe Flin Flon VMS deposit. The Flin Flon Belt is part of the dominantly juvenile core of the Trans-Hudson Orogen, andconsists of a series of 1.92 to 1.87 Ga accreted oceanic arcs and intervening back-arc basins, all intruded by successor arc plutons. The resulting collage has undergone several phases of deformation associated with subduction-related colli-sion, oblique compression, and uplift during the closing of the original ocean basin. During the formation of its variousdomains and succeeding tectono-metamorphic history, the Flin Flon Belt was host to three main metallogenic events. Thefirst is characterized by the formation of several types of VMS deposits within the various oceanic rifted arc terranes, withmafic-bimodal being the dominant VMS type. A period of successor arc and magmatism included development of maficgabbroic intrusions with Ni-Cu-PGE and PGE potential, plus subeconomic Cu-Mo porphyry-type mineralization. Crustalthickening during continued subduction-related collision was accompanied by formation of orogenic, shear-hosted Au

mineralization.The extension of the Flin Flon arc and ocean basin collage under the Phanerozoic cover rocks leaves much of this

Paleoproteroic juvenile terrane open to further discovery of VMS and orogenic Au deposits. This will require a better understanding of the sub-Phanerozoic geology through high-resolution geophysics, a better understanding of deformationhistory and resulting structural architecture, geochemical and mineralogical zoning of large-scale VMS hydrothermalsystems, and geochemical remote sensing.

Résumé

Au cours des 75 dernières années, la ceinture de Flin Flon du Paléoprotérozoïque a fourni une contribution importanteen métaux communs et en or à l’économie canadienne. Elle recèle 27 importants gisements de sulfures massifs volcano-gènes qui renfermaient à l’origine plus de 154 Mt de minerai, dont le gisement de Flin Flon de 62 Mt. On y trouve égale-ment plusieurs gisements de Au orogéniques dont l’exploitation a cessé, incluant le deuxième plus important producteur de plus d’un million d’onces (New Britannia) dans un terrane paléoprotérozoïque au Canada, le premier ayant été legisement de SMV de Flin Flon. La ceinture de Flin Flon fait partie du cœur principalement juvénile de l’orogène trans-hudsonien et se compose d’une série d’arcs océaniques et de bassins d’arrière arc intercalaires accrétés datant de 1,92 à1,87 Ga et tous pénétrés par des plutons d’arcs successeurs. Le collage qui en résulte a subi plusieurs phases de déforma-tion associées à la subduction liée à la collision, à une compression oblique et à un soulèvement pendant la fermeture du bassin océanique d’origine. Pendant la formation de ses divers domaines et l’histoire tectonométamorphique qui a suivi,la ceinture de Flin Flon a été le siège de trois principaux événements métallogéniques. Le premier est caractérisé par laformation de plusieurs types de gîtes de SMV à l’intérieur des divers terranes océaniques d’arcs en distension, les gîtesmis en place dans des assemblages bimodaux à dominante mafique étant la variété dominante des gîtes de type SMV. Une période de formation d’arcs successeurs et de magmatisme a permis la mise en place d’intrusions mafiques (gabbroïques) présentant un potentiel pour les minéralisations de Ni-Cu-ÉGP et de ÉGP, ainsi que de minéralisations subéconomiquesde Cu-Mo de type porphyrique. Un épaississement crustal pendant la subduction liée à la collision qui s’est poursuivies’est accompagné de la formation de minéralisations orogéniques de Au dans des cisaillements.

Le prolongement du collage d’arcs et de bassins océaniques de Flin Flon sous les roches de couverture du Phanérozoï-que fait en sorte qu’une grande partie de ce terrane juvénile du Paléoprotérozoïque pourrait receler d’autres gîtes de SMVet de Au orogéniques. L’exploration à la recherche de ces gîtes exigera une meilleure compréhension de la géologie sub- phanérozoïque par l’application de méthodes géophysiques haute résolution, une meilleure compréhension de l’histoirede la déformation et de l’architecture tectonique résultante, la zonalité géochimique et minéralogique des grands réseauxhydrothermaux à l’origine des gîtes de SMV ainsi que la télédétection géochimique.

Introduction and Mining History

The Paleoproterozoic Flin Flon Belt of central Manitobaand Saskatchewan (Fig. 1) contains both syngenetic, poly-metallic base metal and orogenic Au deposits, but it is bestknown for its two volcanogenic massive sulphide (VMS)

camps: Flin Flon and Snow Lake (Fig. 2). Tom Creighton andthe Mosher and Dion brothers first discovered sulphide min-eralization along the Manitoba-Saskatchewan border in theFlin Flon area in 1914. This mineralization was the surfaceexpression of the 62 Mt Flin Flon VMS deposit. The depositlay dormant until 1930 due to the lack of economic methods

Galley, A.G., Syme, E.C., and Bailes, A.H., 2007, Metallogeny of the Paleoproterozoic Flin Flon Belt, Manitoba and Saskatchewan, in Goodfellow, W.D., ed.,Mineral Deposits of Canada: A Synthesis of Major Deposit Types, District Metallogeny, the Evolution of Geological Provinces, and Exploration Methods:Geological Association of Canada, Mineral Deposits Division, Special Publication No. 5, p. 509-531.



A.G. Galley, R. Syme and A.H. Bailes

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Prospecting from the 1920s through to the 1950s utilizedthe extensive system of interconnected lakes along the GrassyRiver system to push east and west along the length of the exposed Flin Flon greenstone belt. This resulted in the discovery

of a number of VMS deposits, including the Sherritt-Gordon(1928), Don Jon (1930), and North Star (1949). With the intro-duction of ground and airborne electromagnetic surveys in theearly 1950s, Hudson Bay Exploration and Development Ltd.(HBED) discovered the majority of the VMS deposits foundfrom 1952 until the present. This includes the seven VMS deposits developed in the Snow Lake area, most of which werenoted on a 1955 geology map by Russell (1957) as sulphideoccurrences prior to their discovery.

for separating Zn from sphalerite. The Mandy VMS depositon the shore of nearby Schist Lake was the first to produceore in the region (1916) from a high-grade chalcopyrite lens.The development of an electrolytic method to produce Zn,

coupled with the ability to produce inexpensive hydroelec-tric power from the nearby Saskatchewan River, resulted inthe construction of a smelter complex and opening of theFlin Flon deposit. The town that grew up on the Manitobaside of the border was named “Flin Flon” by prospectorsin 1915, after the character Josiah Flintabbatey Flonatin inthe dime novel “The Sunless City.” The twin town on theSaskatchewan side of the border was named Creighton after Tom Creighton, co-discoverer of the Flin Flon deposit.

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FIGURE 1. Location map after Hoffman (1989) illustrating the position of the Flin Flon Belt in relation to the Precambrian geology of North America(A) and central Canada (B). The latter illustrates the components of the dominantly juvenile core to the Trans-Hudson Orogen (THO) in relation tothe bordering Archean terranes, middle Proterozoic Athabasca Basin, and overlying Phanerozoic strata of the western Canadian sedimentary basin.FF=Flin Flon Belt; GD=Glennie Domain; HLB=Hanson Lake Block; LL-LR =Lynn Lake-LaRonge Belt; KD=Kisseynew Domain; RD=RottenstoneDomain; RL=Rusty Lake Belt; TH=Thompson Belt; TB=Tabbernor fault zone; WB=Wathaman-Chipewayan Batholith; WD=Wollaston Domain.Modified from Lucas et al. (1996), from the original by Hoffman (1989). Dashed box represents area shown in more detail in Figure 2.



Metallogeny of the Paleoproterozoic Flin Flon Belt, Manitoba and Saskatchewan

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Deposit Discovery Opened Closed Tonnes** Au (g/t) Ag (g/t) Cu (%) Zn (%)

Birch Lake Back-arc Assemblage

Konuto 1994 1998 2005 1,948,500 2.07 8.64 4.13 1.41

Centennial 1969 1975 1986 1,473,767 1.58 20.23 1.41 2.48

Flexar 1952 1969 1972 305,939 1.30 6.51 3.76 0.50

Birch Lake 1950 1957 1960 272,898 0.10 4.11 6.21 Tr

Flin Flon Arc Assemblage

Flin Flom 1915 1930 1992 62,520,625 2.64 41.49 2.17 4.13

Trout Lake 1976 1982 21,269,476 0.99 16.05 1.60 4.93

Triple 7 1993 1998 14,495,288 2.12 31.08 2.64 4.98

Callinan 1984 1989 8,395,667 2.05 25.55 1.29 4.02

Schist lake 1947 1950 1976 1,846,656 1.30 37.03 4.30 7.27

West Arm 1973 1977 1997 1,394,149 1.58 17.49 3.21 1.48

Coronation 1953 1960 1965 1,281,726 2.06 5.14 4.25 0.20

White Lake 1963 1972 1982 849,784 0.72 27.09 1.98 4.64

Cuprus 1941 1948 1954 462,096 1.30 28.8 3.25 6.40

Birch lake 1950 1957 1960 272,898 0.10 4.11 6.21 Tr

North Star 1950 1955 1958 241,692 0.34 8.57 6.11 Tr

Mandy 1915 1917 1944 102,252 3.05 55.54 5.63 13.69

Don Jon 1950 1955 1957 79,329 0.96 15.09 3.09 Tr

Four Mile Arc Assemblage

Spruce Point 1973 1981 1992 1,865,095 1.68 19.54 2.06 2.40

Dickstone 1935 1971 1975 775,504 0.41 12.00 2.46 3.12

Snow Lake Arc Assemblage

Chisel Lake 1956 1960 1994 7,153,536 1.77 44.77 0.54 10.60

Stall lake 1956 1964 1994 6,381,129 1.41 12.34 4.41 0.50

Chisel North 1985 2000 2,847,194 0.40 21.54 0.15 9.36

Osborne 1953 1968 1983 2,807,471 0.27 4.11 3.14 1.50

Anderson 1963 1970 1988 2,513,290 0.62 7.54 3.40 0.10

Rod 1970 1984 1991 735,219 1.71 16.11 6.63 2.90

Photo Lake 1994 1995 1998 689,885 4.87 29.49 4.58 6.35

Lost & Ghost 1956 1972 1988 581,437 1.20 39.09 1.34 8.60

TABLE 1. Grade and tonnages, and discovery and production dates, Flin Flon deposits*.

* Data from HudBay Minerals unpublished data.** Tonnages and grades are not National Instrument 43-101 compliant.

To date, 27 deposits have been mined in the Flin FlonBelt (Table 1). Many of these deposits are polymetallic andcontain significant contents of precious metals. The FlinFlon VMS deposit alone produced over 120 tonnes Au. Themost recent discoveries include Callinan (1984), Photo Lake(1994), Konuto (1994), and Triple 7 (1993; Table 1). As of

2007, there are five operating base metal mines in the FlinFlon Belt, with HudBay Minerals Ltd. pursuing an aggres-sive exploration program investigating some of the numer-ous untested SPECTRUM® airborne electromagnetic anom-alies.

At approximately the same time that prospectors were dis-covering polymetallic sulphide occurrences in the western part of the Flin Flon Belt, others were discovering Au show-ings 200 km to the east at Wekusko Lake (Fig. 2). The firstAu discoveries were made in 1914 by M.J. Hackett and R.Woosey on the east shore of Wekusko Lake (formerly HerbLake; Stockwell, 1937). By 1918, the small mining town of

Herb Lake had sprung up around several small arsenopyrite- bearing shear-hosted lode Au vein deposits. Less than 10,000ounces of Au was produced from this camp, the largest beingthe Rex-Laguna, which produced 7,122 ounces between 1918and 1925 (Table 2). In 1926, Chris Parres discovered aurif-erous quartz-carbonate-arsenopyrite veins outcropping along

the north shore of Snow Lake at what was to become the Nor Acme Au deposit. Mining began at Nor Acme in 1949 andcontinued until its closure in 1956. In 1999 it re-opened as the New Britannia mine, and became the first and only million-ounce primary Au producer in the Flin Flon Belt. The 1980ssaw production from several other vein-hosted Au deposits,none of which remain open today (Table 2). These include thePuffy Lake, Tartan Lake, and Rio deposits.

Other minor deposit types represented in the Flin Flon Beltinclude the Dion Lake porphyry Cu-Mo mine southwest of Flin Flon. There is also potential for Ni-Cu-PGE mineraliza-tion in mafic layered intrusion and gabbroic intrusions, as




513

well as Ta mineralization in pegmatites.Geological Setting of the Flin Flon Belt

The Paleoproterozoic Flin Flon Belt is approximately 200km in strike length and has an exposed width of up to 70km. The Flin Flon Belt is overlain to the south by Paleozoicsandstone, limestone, and dolomite of the Western CanadaSedimentary Basin, and is bordered to the north by high-grade paragneiss and granitoid rocks of the Kisseynew Domain(Fig. 1).

The Flin Flon Belt is interpreted to be an accreted assemblage of oceanic to continental margin arc terranes, inter-spersed with oceanic basins representing back-arc, fore-arc,and oceanic settings (Lucas et al., 1996; Syme et al., 1996).

It is part of the Reindeer zone, a largely juvenile portion of the Trans-Hudson Orogen separating the Archean Superior and Hearne provinces (Fig. 1). Recent tracer isotope stud-ies have confirmed the presence of >3.0 Ga Archean crust(the Sask Craton) below parts of the Trans-Hudson (Lucaset al., 1996). The Shield Margin National Mapping Program(NATMAP; Lucas et al., 1996) traced the Flin Flon Belt as-semblages below the Phanerozoic to the south and recognizedhighly metamorphosed and deformed Flin Flon volcanic andsedimentary formations (Zwanzig, 1990, 1999) within theKisseynew Domain to the north. To the east, the Flin FlonBelt is separated from the Paleoproterozoic Thompson NickelBelt by Kisseynew Domain rocks. To the west, the Flin FlonBelt is terminated against the Tabernor fault zone (Fig. 1).

The recent Geological Survey of Canada (GSC)-Manitoba-Saskatchewan NATMAP Shield Margin Project andLITHOPROBE Trans-Hudson Orogen transect built on anextensive existing geological database that led to a much im- proved understanding of the components and evolution of thesoutheastern Reindeer zone, including the Flin Flon Belt (e.g.,Lucas et al., 1996). These investigations have shown that, ata crustal scale, the Flin Flon “greenstone” belt is only oneof three components in a northeast-dipping stack, juxtaposedduring 1.84 to 1.80 Ga collisional deformation (Lucas et al.,1999):

At the lowest structural level (exposed in the PelicanWindow, Fig. 2, within the Hanson Lake Block): meta- plutonic rocks and paragneisses (3.20–2.40 Ga) of the“Sask craton” (Corrigan et al., 2007).

At intermediate structural levels: Flin Flon Belt (nowdefined to include the Attitti Block and Paleoproterozoicrocks in the Hanson Lake Block) and Glennie Domain,shown in Figures 1 and 2, (together comprising the “FlinFlon-Glennie Complex” [FFGC]; Lucas et al., 1996).

At the highest structural levels: marine turbidites(Burntwood Group; 1.85–1.84 Ga) and distal faciesof alluvial-fluvial sandstones (Missi Group) in theKisseynew Domain (Fig. 1).

Geological Evolution of the Flin Flon Belt

The stratigraphy of the Flin Flon Belt has been previ-ously subdivided into two major groups, the Amisk Groupmetavolcanic rocks and Missi Group continental metasedi-mentary rocks (Bruce, 1918; Harrison, 1951). The Flin FlonBelt is now recognized as consisting of several 1.9 to 1.88Ga terranes comprised of four main tectono-stratigraphicassemblages that represent both juvenile and continentallyunderlying oceanic segments of a Paleoproterozoic ocean basin that were accreted during formation of the Trans-Hudson orogen (Syme, 1990; Syme and Bailes, 1993; Sternet al., 1995a,b; Lucas et al., 1999; Figs. 1, 2). The orogenwas formed by oblique collision between the Superior andHearne Archean terranes, and the resulting collage is separ-ated into a number of assemblages distinguished by uniquetectono-stratigraphy, and dismembered by fault systems thatwere originally thrust surfaces (Syme, 1995, Lucas et al.,1996). Each tectonostratigraphic assemblage is a distinct package of rocks in terms of its stratigraphy, geochemistry,isotopic signature, age, and inferred plate tectonic setting(see below; Syme and Bailes, 1993; Lucas et al., 1996;Corrigan et al., 2007).

The tectonostratigraphic assemblages were juxtaposed in

1.

2.

3.

Deposit Discovery Opened ClosedTonnesMilled

Oz AuProduced

Production& Reserves*

Grade(g/t) Source

New Britannia 1926 1949 2004 1,323,000 4,994,871 3.9 Northern Miner, 2004

Laguna (Rex) 1914 1918 1940 58,962 unknown unknown Richardson & Ostry, 1996

Gurney 1919 1937 1939 94,650 28,045 unknown unknown Richardson & Ostry, 1996

Tartan Lake 1931 1987 1988 592 18,510 375,000 6.27 Richardson & Ostry, 1996

Ferro-Rainbow 1923 1932 1933 1,006 730 66,000 13 Richardson & Ostry, 1996Moose Horn(Ballast)

1914 1917 1918 42 194 unknown unknown Richardson & Ostry, 1996

Bingo 1915 1926 - - 128 unknown unknown Richardson & Ostry, 1996

Gold Hill 1934 1935 1963 103 147 unknown unknown Richardson & Ostry, 1996

Century 1919 1941 1942 - 59.6 272,000 12 Richardson & Ostry, 1996

Apex 1918 - - - - 758,000 3.63 Richardson & Ostry, 1996

Newcor - - 1947 - - 43,000 10.56 Coombe Geoconsultants Ltd., 1984

Rio - 1984 1985 - - 124,000 12.54 Coombe Geoconsultants Ltd., 1984

Henney-Maloney - - 1941 - - 15,000 14.85 Coombe Geoconsultants Ltd., 1984

TABLE 2. Historical Statistics for Flin Flon Belt Gold Mines.

* Not instrument 43-101 compliant.




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of the Kisseynew turbidite basin (now part of the KisseynewDomain, Figs. 1, 2) was synchronous with continental sedi-mentation in the Flin Flon Belt (Ansdell, 1993; David et al.,1993, 1996; Machado and Zwanzig, 1995; Connors, 1996;Connors et al., 1999).

The transition from Kisseynew basin extension to col-lisional collapse occurred rapidly at about 1.840 Ga, al-though sedimentation and magmatism continued throughto ~1.830 Ga (Ansdell and Norman, 1995; Machado andZwanzig, 1995; David et al., 1996; Connors et al., 1999). TheKisseynew Domain was thrust over the Amisk collage alongthe southernflank of the Kisseynew Domain (Harrison, 1951;Zwanzig, 1990; Lucas et al., 1994; Connors, 1996; Connorset al., 1999; Zwanzig, 1999). Following collisional thicken-ing and peak metamorphism at 1.83 to 1.80 Ga, the Flin FlonBelt experienced protracted intracontinental deformation toca. 1.69 Ga (Lucas et al., 1996; Stern et al., 1999).

The Flin Flon Accreted Terranes

The Flin Flon Belt consists of two principal segments(Amisk collage and Snow Lake arc assemblage) that were juxtaposed during southwest-verging continent–contin-ent collision between 1.84 and 1.82 Ga. To the west of theAmisk collage is a volcano-sedimentary domain of similar age (Hanson Lake arc assemblage; Fig. 2). Although nottraditionally designated as part of the Flin Flon Belt, theHanson Lake arc assemblage will be included here due tothe presence of volcano-sedimentary rock units and severalVMS occurrences and deposits of similar age (Maxeiner etal., 1993, 1999).

Although each of these three tectonic segments has a dis-tinct character, their metallogenic tenor is a product of thesame three tectonically controlled evolutionary stages of the region. The first is represented by syngenetic polymetal-lic base metal and precious metal deposits and occurrencesthat formed during a pre-accretionary stage within distinct

oceanic, suprasubduction environments. The second consistsof post-accretion intrusion-related mineralization associatedwith successor arc formation and extensional magmatism. Thethird is comprised of shear zone-related orogenic Au depositsthat formed during periods of collision, oblique compression,and crustal thickening.

Hanson Lake Block

The Hanson Lake arc assemblage structurally overliesArchean crust and contains coeval volcanic and sedimentaryassemblages that are found within the neighbouring Flin Flon

Belt. It is also host to numerous VMS deposits and show-ings. The fault-bounded area containing the Hanson Lake arcassemblage and underlying Archean crust is known as theHanson Lake Block (HLB) (Figs. 1, 2).

The HLB is composed of a highly deformed and meta-morphosed assemblage of 1.91 to 1.85 Ga volcanic and sedi-mentary rock, and 1.86 to 1.81 Ga syntectonic intrusions andmigmatitic gneisses that have been thrust over the ca. 2.5 Ga Neoarchean charnokitic and enderbitic intrusive rocks knownas the Peleican Window (surface expression of the Sask Craton;Ashton et al., 1987; Maxeiner et al., 1993, 1999; Ashton andLewry, 1994; Fig. 2). The HLB is terminated to the west bythe Tabernor fault zone and to the east by the Sturgeon-Weir fault zone, which separates the HLB from those assemblages

traditionally included within the Flin Flon Belt, and extendssouthward below the Phanerozoic cover. The principal reasonfor the original exclusion of the arc assemblages of the HLBfrom those of the Flin Flon Belt was that they were origin-ally believed to have formed upon Archean crust and were,therefore, not considered part of the Flin Flon oceanic supras-ubduction suite.

Supracrustal rocks of the HLB are dominated by metavol-canic and metasedimentary rocks. Volcanism and sedimenta-tion are coeval from 1910 to 1880 Ma, with sedimentationcontinuing to at least 1850 Ma (Maxeiner et al., 1993, 1999).

2. ACCRETIO N(1.88-1.87 Ga )

1. ARC/OCEANI CMAGMA TIS M(1.92-1.88 Ga )

3. POST -ACCRETIO NARC MAGMA TISM/SEDIMENT ATIO N(1.87-1.84 Ga )

post-accretion(successor)

arcKisseynew

back-arc basin

"evolved"arc

Snow L.ar cplateau

Archean crust

Snow L.ar c

Flin Flon juvenile

ar c

Amisk accretionar ycollage

back-ar c

NESW

?

FIGURE 3. Schematic illustrating the tectonic evolution of the Flin Flon district (modified fromLucas et al., 1996).

an accretionary complex at ca. 1.88 to 1.87Ga, probably as a result of arc–arc collision(D

1; Lucas et al., 1996; Stern et al., 1999; Fig.

3). Accretionary collage-bounding structureswere largely obliterated by subsequent de-formation and metamorphic events (D

2 –D

5),

but are inferred where juxtaposed terranes are“stitched” together by calc-alkaline plutonsrelated to a 1.866 to 1.838 Ga successor arcformation (Whalen et al., 1999). Coeval subaerial volcanism is recorded in ca. 1.87 to1.85 Ga calc-alkaline to shoshonitic volcan-iclastic sequences (Syme, 1988; Bailes andSyme, 1989; Lucas et al., 1996; Stern et al.,1996). Unroofing of the accretionary collage,development of a paleosol, and depositionof alluvial-fluvial sedimentary rocks (Missisuite; Bailes and Syme, 1989; Holland et al.,1989) occurred ca. 1.85 to 1.84 Ga (Ansdell,1993). These events were coeval with thewaning stages of post-accretion arc magma-tism (Stern and Lucas, 1994; Whalen andHunt, 1994; Lucas et al., 1996). Development




515

Volcanic strata are dominantly tholeiites (Maxeiner et al.,1993, 1999) and include pillowed basalt overlain by inter-mediate to felsic flows and volcaniclastic rocks. Also pres-ent is a large felsic hypabyssal intrusive/extrusive complex(Maxeiner et al., 1993). The volcanic assemblage is in con-tact with calc-silicate-carbonate-rich strata, silicate-facies

iron formation, and polymictic conglomerate, and overlain by psammitic greywacke and mafic wacke.The supracrustal assemblages of the HLB are intruded by

numerous synvolcanic intrusions, ranging in compositionfrom ultramafic through gabbro and quartz diorite to rhyo-litic. Large antiformal domes of migmatitic gneiss are accompanied by lit-par-lit injection into the supracrustal formations.Metamorphic grade generally increases from south to north,from upper greenschist to upper amphibolite facies, with re-gional metamorphism peaking between 1810 and 1806 Ma. Amajor folding event took place between 1860 and 1850 Ma,and was followed by 1810 to 1800 Ma continental collisionthat caused the thrusting of this terrane over Archean base-ment (Ashton and Lewry, 1994). Deformation that accompan-

ied crustal thickening and post-peak metamorphism continueduntil 1770 Ma.The HLB contains numerous metamorphosed synvolcanic

hydrothermal alteration zones and spatially and geneticallyassociated sulphide mineralization. These include felsic alum-inous schists and zones of cordierite-anthophyillite and gar-net-amphibole alteration in more mafic rocks. Two significantVMS deposits are associated with the felsic volcanic strata.The Western Nuclear Pb-Zn deposit (1967–69) produced 147Mt grading 9.99 percent Zn, 0.51 percent Cu, and 5.83 percentPb, with minor Au and Ag. The McIlvenna Bay deposit hasestimated reserves (pre-National Instrument 43-101) of 13.08Mt grading 4.95 percent Zn, 1.26 percent Cu, 0.53 g/t Au, and24.3 g/t Ag (CAMECO, 1992).

Amisk Collage

The Amisk collage is comprised of a series of fault-bound-ed tectonostratigraphic assemblages (Syme, 1995; Lucas etal., 1996; Figs. 2, 4). These are intruded by post-accretion-ary plutons and are overlain by fluvial-alluvial sediment-ary rocks of the Missi Group. The collage is bounded to thewest by the Sturgeon-Weir fault system, to the east by theMorton Lake fault zone, and to the north by the southernflank of the Kisseynew domain. It extends to the south belowthe Phanerozoic cover. The Amisk collage contains the WestAmisk, Birch Lake, Flin Flon, and Fourmile Island oceanic arcassemblages, and the Sandy Bay and Elbow-Athapapuskow back-arc basin basalt assemblages (Stern et al., 1999; Syme etal., 1999; Figs. 2, 4).

The West Amisk arc assemblage consists of island arctholeiitic basalt overlain by the West Amisk differentiatedcalc-alkalic suite (high-Al basalt to rhyolite; Figs. 2, 4). Thefelsic component of the calc-alkaline suite forms a series of rhyolite domes, associated volcaniclastic facies, and hypabys-sal intrusions that are overlain by a Hawaiian-type emergingshield volcano of basalt and andesite flows, and scoriaceousvolcaniclastic facies (Ayres et al., 1991). The arc tholeiite andcalc-alkaline suites are overlain to the north by mafic grey-wacke-argillite and minor pebble conglomerate. The West

Amisk arc assemblage contains the hybrid VMS-epithermalLaurel Lake deposit (Fig. 4) and numerous small aurifer-ous occurrences of arsenopyrite-rich quartz-carbonate veinsand vein systems (Byers and Dahlstrom, 1954; Walker andMcDougall, 1987).

The West Amisk assemblage is in fault contact to the east

with the Sandy Bay assemblage (Fig. 4), which consists of pillowed to massive basalt interpreted to represent an oceanic plateau succession that formed either in an oceanic or back-arc basin (Reilly et al., 1994; Watters et al., 1994; Ryan andWilliams, 1999; Syme et al., 1999).

The Birch Lake assemblage is flanked to the west by theMosher Lake shear zone (Fig. 4). It consists of a 2 to 8 km-thick succession of tholeiitic to boninitic eruptive flows andassociated subvolcanic sills (Watters et al., 1994; Stern et al.,1995a,b; Wyman, 2003). The volcanic succession containslittle to no volcaniclastic material and no known felsic ex-trusive strata. There are four known Cu-Zn VMS deposits,Coronation, Birch, Flexar, and Konuto (Table 1), located atthe contacts between these massive to pillowed mafic flow

units.The Flin Flon assemblage consists of a number of fault-

bounded structural blocks that have little or no stratigraph-ic correlation (Syme and Bailes, 1993; Stern et al., 1995a;Syme, 1995; Lucas et al., 1996; Fig. 4).

The Flin Flon block (Fig. 4, 5a, b) is bounded to thewest by the West Arm fault and to the east by the RossLake fault, and is host to five rhyolite-associated Zn-Cu-Au VMS deposits (Flin Flon, Triple 7, Callinan, Schist,and Mandy). The block is a primitive, bimodal oceanicarc succession dominated by a number of tholeiitic bas-alt and basaltic andesite units, including mafic flowsand volcaniclastic rocks, and minor rhyolite flows anddomes, the latter including the Flin Flon mine rhyolite(1903 +7/–5 Ma; Stern et al., 1999; Syme et al., 1999).

The Hook Lake suite is bounded to the west by the Rosslake fault and to the east by the Inlet Arm fault (Bailesand Syme, 1989; Fig. 4). The block comprises basalt,mafic volcaniclastic rocks, and minor rhyolite and vol-canic conglomerate. It contains numerous base metalshowings but no significant deposits to date.

The Bear Lake suite consists of a four km-thick arc suc-cession bounded to the west by the Inlet Arm fault and tothe east by the Northeast Arm fault (Figs. 4, 6A), and isinterpreted to be part of a subaqueous shield volcano andassociated cauldron that is overlain by an intra-arc rift

succession and associated intrusions (Bailes and Syme,1989, Syme and Bailes, 1993; Fig. 6A). The shield vol-cano is composed of a 1 km-thick succession of massiveto pillowed flows, minor flow breccia, and several syn-volcanic gabbro/diorite sills (Fig. 6B). A cauldron fillsuccession of reworked dacitic tuff, interlayered andes-ite lapilli tuff, and finely laminated turbidite is host tothe Cuprus and White Lake Zn-Cu-Pb VMS deposits(Fig. 6A). Minor rhyolite dome construction occurred prior to the emplacement of cauldron sediment fill. Thisarc-rift succession is overlain by 1886 Ma ferrobasalt

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516

A t h

p a

a p u

s k o w L a

k e

Sask. Man.54 30' N

o

102 15' Wo

10 km

Cu

WL

VMS Deposits

Gold Deposits

Hybrid Epithermal-VMS

tholeiitic, mafic-felsic volcanics

tholeiitic basalt (age unknown)

calc-alkaline volcanics

major mafic-felsic tholeiiticintrusives

Missi Suitesandstone, conglomerate

SUCCESSOR ARCINTRUSIVE ROCKS

SUCCESSOR BASINS

JUVENILE ARC

'EVOLVED ARC'

intermediate-mafic

felsic plutonic rocks(~1.920-1.903 Ga)

granitic rocks (~2.5 Ga)

felsic

N-type MORB / E-type MORBlayered mafic-ultramafic complexOCEAN FLOOR

(BACK-ARC)

OCEANIC PLATEAU

PRE-ACCRETION ASSEMBLAGES (1.92-1.88 Ga ) POST -ACCRETIO N R OCKS (1.87-1.84 Ga)

1.90-1.88 Ga

ARCHEAN SLICES

Accretion-relatedshear zones (D1, >1.87 Ga)

Younger shear zones/faults (D2-D5)

VMS deposit

Gold deposit

Amisk Lake

TRC

F

S

W

DN

Ce

Co

BK

Fl

M

T

F

R

arc rift basalt

Schist-Wekusko suitegreywacke, mafic sills

1.90 Ga

West Amisk Assemblage

Sandy BayAssemblage

Elbow Lake-AthapapaskowAssemblag e

Birch LakeAssemblage

Flin FlonAssemblage

Flin FlonBlock

Hook LakeBlock Bear Lake

Block

Sourdough BayBlock

Scotty LakeBlock

B - Birch Lake; C - Callinan; Ce - Centennial; Co - Coronation; Cu - Cuprus; D - Don Jon; F - Flin Flon;Fl - Flexar; K - Konuto; M - Mandy; N - North Star; S - Schist Lake; T - Trout Lake; TR - Triple 7;W - West Arm; WL - White Lake

HM - Henney-Maloney; R - Rio; TL - Tartan Lake

LL - Laurel Lake

HM

LL

TL

R

M

L S Z

W A S Z

I A S Z

NAFZ

FIGURE 4. Tectonic assemblage map of the central part of the Flin Flon Belt illustrating the pre-accretion tectonostratigraphic assemblages, post-ac-cretion plutons, volcano-sedimentary basins and faults. Also shown are the present- and past-producing VMS and Au deposits. Accretion-realtedshear zones include: WASZ: West Amisk shear zone; MLSZ: Mosher Lake shear zone; IASZ: Inlet Arm shear zone; NASZ: Northeast Arm shear zone. Modified from Syme et al., 1999.




517

FIGURE 5. (A) Geology in the immediate host geology for the Flin Flon-Triple 7-Callinan VMS deposits showing their location along a stratigraphicinterval containing rhyolite flows, felsic heterolithic breccia, and finely layered felsic tuff (Stockwell, 1960; Bailes and Syme, 1989; Devine et al.,2002). (B) Schematic cross-sectional reconstruction of the host stratigraphy to the deposits illustrating the presence of a cauldron structure. Modifiedfrom Syme and Bailes (1993); Ames et al., 2002; and Devine et al., (2002).

v

>700 m7 00 m

Callinan Triple 7 Flin Flon

N

S

N S

Basalt flows andinterflow sediment

Heteroloithic mafic lapilli tuff and tuff breccia

Heteroloithic tuff breccia

Pillow fragment breccia

Rhyolite flows

Bedded felsic tuff

VMS deposits; FF = Flin Flon;C = Callinan; Tr = Triple 7

Finely layered mafic tuff

Hidden Lake & Louie Lakebasaltic andesite pillowed flows

Synvolcanic gabbro/diorite sillsand dikes

Sandstone and pebblysandstone

Monzodiorite andpyroxenite

Mine Sequence

Hanging wall Sequence

FLIN FLON ARC ASSEMBLAGE

MISSI SUITE

POST -MISSI INTRUSIONS

Footwall Sequence

Heteroloithic breccia withrhyolite clasts

Heteroloithic breccia withrhyolite clasts

ca. 5 km

A

B




518

and 1885 Ma shoshonitic scoriaceous tuff, mudstone,and chert. Although the Bear Lake basaltic andesite andoverlying the cauldron infill sequence represents riftingof a 1.9 Ga oceanic arc, the overlying ferrobasalt and sho-shonite indicate back-arc basin development (Bailes andSyme, 1989; Syme et al., 1999).

The Sourdough Bay volcanic suite is the most easterlyVMS-hosting arc suite in the Flin Flon arc assemblage,and contains the Centennial, Pine Bay, North Star, andDon Jon Cu-Zn VMS deposits (Bailes and Syme, 1989;Mitchinson, 2004; Fig. 4). It is bounded to the west bythe Centennial fault and to the east by the North Armfault (Bailes and Syme, 1989). This volcanic suite iscomposed of transitional tholeiitic to calc-alkalic basaltand basaltic andesite, along with several large rhyolitedome complexes.

The Elbow-Athapapuskow assemblage is a semi-continu-ous belt of juvenile ocean floor basalt and co-magmatic

gabbro sills and mafic-ultramafic intrusions (Syme andBailes, 1993; Stern et al., 1995b; Fig. 2). This assemblageis up to 25 km wide and over 100 km in length, and is infault contact with flanking arc assemblages. The assemblage consists principally of N-MORB and subordinateE-MORB (Syme and Bailes, 1993; Stern et al., 1995b;Syme et al., 1999). The absence of VMS deposits inthis back-arc assemblage led Syme and Bailes (1993) to postulate that the hydrothermal conditions necessary for VMS formation were not present in the Amisk collage.

The Fourmile Island assemblage is a suite of tholeiiticintermediate to felsic flows and heterolithologic volcan-iclastic rocks that host the Dickstone and Spruce PointZn-Cu VMS deposits (Table 1). This 5.5 km-thick oceanarc volcanic succession represents the eastern limit of theAmisk collage. Its eastern limit is defined by the MortonLake shear zone (Fig. 4), an early syn-accretion structure(Syme, 1990; Stern et al., 1995b; Lucas et al., 1996).

Snow Lake Area

The arc and ocean floor assemblages in the eastern part of the Flin Flon Belt are collectively suf ficiently distinct fromthe Amisk collage arc assemblages to suggest that they rep-resent remnants of unrelated arc terranes (Lucas et al., 1996;Syme et al., 1996). This eastern part of the belt is character-ized by a number of allochthons in a thrust stack that is bor-dered to the west by the Morton Lake fault zone and to theeast and north by the overthrust Kisseynew Domain (Baileset al., 1994; Syme, 1995). These allochthons are comprisedof the Snow Lake arc and the Northeast Reed and RobertsLake ocean floor assemblages (Fig. 2) that are separated bymajor bounding fault systems. The 1.89 Ga Snow Lake arc as-semblage (David et al., 1996; Bailes and Galley, 1999) is theonly one that contains significant VMS mineralization. It isexposed in a thrust stack that includes several structurally imbricated slivers of 1.84 to 1.83 Ga post-accretion sedimentarystrata of the Burntwood suite (Stern et al., 1995a; Connors et

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al., 1999). The >6 km-thick dominantly juvenile oceanic tocrustally contaminated arc succession of the Snow Lake arcassemblage consists of three conformable volcanic succes-sions that record the evolution from nascent or primitive arc(Anderson sequence), through mature arc (Chisel sequence),to rifting and opening of a back-arc basin (Snow Creek se-

quence; Fig. 7). Anderson Sequence

The lowest recognized volcanic cycle is a composition-ally bimodal primitive tholeiitic arc sequence dominated by pillowed to massive basaltic to basaltic andesite and sub-ordinate Ca-boninite flows. This mafic succession containsat least three rhyolite flow complexes consisting of massiveto lobe-hyaloclastite flows that are intruded by the 22 km-long subvolcanic Sneath Lake tonalite-trondhjemite intru-sive complex (Bailes and Galley, 1999; Fig. 7). The last rock type within the primitive arc succession is a 1 to 5 m-thick,finely layered, sulphidic, mixed argillite-epiclastic chertFoot-Mud horizon (Fig. 7).

Chisel SequenceThe primitive nascent arc Anderson succession is con-

formably overlain by a mature arc sequence of intercalatedflows, volcaniclastic rocks, and several discrete rhyolite flowcomplexes with associated monolithic debris flows of theChisel sequence. The mafic flows in this sequence are typ-ically more fractionated than those in the primitive arc (Sternet al, 1995a; Bailes and Galley, 1999). The volcaniclasticrocks consist of turbiditic and coarser-grained heterolitho-logic mafic, and mixed mafic and felsic debris flow deposits.The lower part of the sequence contains numerous synvol-canic intrusions, is extensively hydrothermally altered, andis capped by a series of discrete rhyolite flow complexesthat host the Chisel Lake, Lost, Ghost, Chisel North, and

Photo Lake Zn-Pb-Cu-Ag-Au-rich VMS deposits (Fig. 7).Overlying the VMS deposits are fine- to coarse-grainedheterolithologic debris flows and associated mafic effusiveflows.

Snow Creek Sequence

The debris flows of the upper part of the mature Chiselsequence arc succession are abruptly overlain by a thick suc-cession of massive to pillowed basalt flows and associated basalt sills of the Snow Creek sequence (Fig. 7). This mon-otonous flow succession contains minor inter flow hyaloclas-tite or flow breccia, indicating very rapid deposition. The N-MORB composition of these flows suggests they formed by rifting of the Snow Lake arc assemblage and develop-ment of back-arc ocean floor (Stern et al., 1995a; Bailes and

Galley, 1999). The rift succession is in thrust contact with astructurally overlying succession that includes an imbricated1.84 Ga Burntwood metagreywacke and a structural panel of arc volcanic rocks that resemble the upper part of the Chiselsequence. No VMS mineralization is known in this back-arcsequence.




519

Metallogeny of the Flin Flon Belt

The major mineralizing events recognized in the FlinFlon belt took place during the three main stages of crustaldevelopment: pre-accretion, post-accretion, and continent– continent collision. The pre-accretionary stage is represented by syngenetic base metal and Au deposits. The syn- to post-accretionary stage is characterized by several examples of intrusion-hosted base and precious metal deposits, and thecontinental collision stage by the development of orogenicAu deposits and REE-enriched pegmatites.

Pre-Accretion Mineralization

All the pre-accretion base and precious metal deposits

and occurrences occur within the oceanic arc assemblages(Syme and Bailes, 1993). These include VMS deposits andat least one hybrid VMS-epithermal deposit. The VMS min-eralization is classified by lithologic setting, which in most

cases is unique to particular stages of arc development (Barrieand Hannington, 1999; Franklin et al., 2005), and includesVMS deposits hosted within mafic-back-arc, bimodal-mafic, bimodal-felsic, pelitic-mafic, and felsic-siliciclastic. The hostlithologies comprise tectonostratigraphic assemblages thatare commonly bound unconformable footwall and hanging-wall contacts and formed during specific volcano-magmaticor sedimentary events (Franklin et al., 2005).

The Flin Flon Belt VMS deposits display a wide range of

FIGURE 6. (A) Simplified geology of the Bear Lake suite in the vicinity of the White Lake and Cuprus Cu-Zn-Pb VMS deposits. Modifed from Bailesand Syme (1989). (B) Schematic cross section of the Bear Lake cauldron. From Syme and Bailes (1993).

10

10

0 1 2 km

Benn Lake post-cauldron gabbro

Andesite tuff, lapilli tuff andinterlayered sulfidic mudstone

Felsic volcanic and volcaniclastic

Mafic flows with abundantvolcaniclastic and epiclastic

Mafic pillowed, massive and sheet flows

Mafic siliciclastic-hosted Cu-Zn-PbVMS deposits

Volcanic-hosted, contact-type PGE deposit

Mafic flows with abundant mafic sills

Pre and syn-cauldron assemblages

900m

200m

650m

3300m

1885 Ma

1886Ma

VMS

carbonate-bearing sediments

rhyolite crystal tuff tuff, mudstone, chert

SHOSHONITIC TUFF

graphitic mudstone

ARC RIFT

ARC

RIFTBASIN

FERROBASAL T

ANDESITE LAPILLI TUFFRHYOLITE

BRECCIAREWORKED FELSIC TUFF

Inlet Arm fault

Northeast Armshear zone

caldera basin

BASAL TIC ANDESITE

N S

gabbro sill

McBratney Lake

I n l e t

A r m

F a u l t

N o r t h

e s t A

r m F

a u l t

A

B

Cuprus Mine

White Lake Mine




520

metal contents and associations (Cu, Cu-Zn, Zn-Cu, and Zn-Cu-Pb VMS types; Fig. 8A), typical of the range of arc en-vironments contained within this accretionary collage. Mostof the Flin Flon VMS deposits contain between 1 and 2 g/tAu, which is in the range for most Canadian VMS deposits(Fig. 8B). Photo Lake, Coronation, Triple 7, and Flin Floncould be classified as Au-rich VMS deposits according tothe classifications of Hannington et al. (1999) and Dubé et

al. (2007). The Flin Flon Zn-Cu deposit produced over 120tonnes (almost 4 million ounces) of Au, making it the largestAu producer in the belt and the third largest VMS produ-cer of Au in Canada after the Horne and LaRonde deposits(Galley et al., 2007 and references therein)

Due to the extensive deformation history of the Flin FlonBelt, these pre-accretion syngenetic seafloor VMS depositshave undergone varying degrees of deformation, so that pri-

FIGURE 7. Generalized geology map of the Snow Lake arc assemblage, including the major synvolcanic and post-accretionary intrusions, large-scalemetamorphosed hydrothermal alteration zones, and VMS deposits (underlined) and major occurrences. A=Anderson; B=Bomber zone; C=Chisel

Lake; CN=Chisel North; G=Ghost; J=Joannie zone; LO=Lost; LD=Linda zone; M=Morgan Lake zone; P=Pot Lake zone; PH=Photo Lake; PN=Penzone; RD=Rod; RM=Ram Zone; RN=Raindrop zone; S=Stall Lake. Modified from Bailes and Galley (1999).

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ALTERED ROCKS (METAMORPHOSED AT 1.82-1.81GA)

Alteration pipes and completely altered rocks

Quartz+plagioclase-rich rocks epidote actinolite

(derived from a mafic protolith)

± ±

Chlorite+garnet+biotite-rich rocks staurolite actinolite± ±

Amphibole-rich rocks garnet±

Epidote-rich rocks

> 1.88 GA SYNVOLCANIC INTRUSIVE ROCKS

> 1.88 G A SUPRACRUSTAL ROCKS

Arc rift supracrustal rocks

(Snow Creek Sequence)

Mature arc supracrusta rocks

(Chisel Sequence)

Primitive arc supracrustal rocks

(Anderson Sequence)

ROCKS POSTDATING ALTERATION

< 1.84-1.83 Intrusive rocks

1.85-1.84 Ga Burntwood Group greywacke

SYMBOLS

Facing direction of strata

Berry Creek Fault Zone

Fault (early kinematic, late kinematic)Massive sulphide deposit (Cu-Zn, Zn-Pb-Cu)

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F o o t -

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n

Snow Lake Fault

HamLake

Lake

MorganLake

SnowLake

Anderson

Lake

WekuskoLake

L

McLeod Road Fault

54 53’ 13”o

9 9

5 2 ’ 3 0 ”

o

54 45’ 00”o

1 0 0

1 5

’ 0 0

”

o

Cook

M

P

RN

PN

B

PH

CN

C

LO

G

J

A

RM

S

RD

LD

WooseyLake




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mary features such as bedding are only rarely preserved (Fig.9A,B). In most cases, the stratigraphic horizons along whichthese deposits formed acted as the foci for strain due to the presence of extensive phyllosilicate-rich alteration and theductile nature of the sulphides, and because deposits com-monly lie along stratigraphic contacts between competentvolcanic formations. In intensely deformed cases, such as atthe Osborne Lake deposit in the Snow Lake arc, originallydiscordant footwall alteration zones are transposed semi-par-allel to stratigraphic contacts. Individual massive sulphidelenses are elongate and plunge parallel to the strong regionallineation. Other characteristics include the development of

foliation (locally strong) in both host rocks and deposits,with tectonic compositional banding (Fig. 9C), sulphide blastesis (Fig. 9D), durchbewegung (tectonic breccias; Fig.9E), sulphide cataclasis, and mylonitization (Fig. 9C), andsecondary sulphide vein formation within and around themassive sulphide lenses.

In the case of deposits now hosted within lower amphibolite grade rocks, metamorphic and structural remobil-ization of the massive sulphides in lenses can producecoarse-grained sphalerite (Fig. 9F) and finer-grained pyritesurrounded by finer-grained massive to banded sphalerite-

pyrite (Fig. 9D). Deposit wall rocks contain coarse-grainedveins of chalcopyrite-arsenopyrite or galena-sphalerite, bothwith high Au (up to 10s of ppm) and Ag (up to 100s of ppm)contents (Galley et al., 1993). Where massive sulphide bodiesare hosted in upper amphibolite grade rocks, there is also tex-tural evidence for sulphide melt development. The mobility of

base and precious metals during metamorphism and deforma-tion has radically changed the original metal zonation in theseVMS deposits.

Back-arc Ma fic VMS Deposits

Mafic-dominated tectonostratigraphic assemblages are de-fined as those that consist primarily of effusive mafic flowswith minimal inter flow volcaniclastic units and <5 volume per-cent felsic volcanic rocks (Franklin et al., 2005). Synvolcanichypabyssal mafic intrusions, in the form of sills and dikes, arecommon. These mafic volcanic successions are most commonat ocean spreading centres where they are characterized by N-MORB basalts, or in nascent, fore-, or back-arc environmentswhere the volcanic rocks can include N-MORB, low-Ti bas-alt, boninite, and arc tholeiitic basalts. Eruptive events may

be separated by a mixture of iron-rich chemical sediment andfine-grained water-lain tuff commonly referred to as exhalite(Galley and Koski, 1999, and references therein).

In the Flin Flon Belt, the Sandy Bay and Birch Lake as-semblages of the Amisk collage and certain parts of theElbow-Athapapuskow back-arc and ocean floor assemblageare dominated by basalt flows (Figs. 2, 4). Of these, only theBirch Lake assemblage contains significant VMS mineral-ization (the Birch, Flexar, Coronation, and Konuto deposits;Table 1; Fig. 4). The deposits are localized at the contacts be-tween either mafic effusive flows or between mafic flows andoverlying flow breccias and mafic tuff, and consist of mul-tiple massive sulphide lenses and footwall vein stockworksenclosed in strongly chloritized and silicified mafic rock. The

sulphide mineralogy is dominated by chalcopyrite, pyrrhotite,and pyrite with subordinate magnetite and sphalerite (Byers etal., 1965; Tourigny et al., 2002). The Konuto deposit is composed principally of sulphide remobilized into syntectonicquartz and carbonate veins containing high concentrations of both Cu and Au (Tourigny et al., 2002).

Bimodal-Ma fic VMS Deposits

Mafic-bimodal volcanic successions are dominated bymafic flows and associated autoclastic units and subordinatediscrete felsic flow centres (Franklin et al., 2005). The rock assemblages hosting VMS deposits may contain significantvolumes of mafic volcaniclastic strata and subordinate felsicautoclastic debris associated with dome complexes. Exhalitehorizons are associated with the VMS mineralization and can be laterally extensive. Synvolcanic intrusions are commonas dike swarms that are spatially associated with centres of seafloor hydrothermal activity and as larger quartz diorite-tonalite-trondhjemite sill-like intrusive complexes (Galley,2003; Franklin et al., 2005). High abundances of footwallvolcaniclastics, isolated felsic flow complexes, and discretedike swarms are all characteristic of extensional graben or cauldron features that commonly host the bimodal-mafic-typeVMS deposits.

Bimodal-mafic VMS deposits, the principal type of VMS

FIGURE 8. (A) Tonnage and grade of Flin Flon VMS deposits plotted ona Cu-Zn-Pb ternary plot for all Canadian VMS deposits as calculated byGalley et al. (2007). Data from Table 1. (B) Binary plot of Au contentvs. tonnage for the Flin Flon VMS deposits. Although four of the deposits could be defined as Au-rich, almost all have Au contents that fallwithin the normal range of 1–2 g/t Au (Franklin et al., 2005).

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deposit in the Flin Flon Belt, tend to form in clusters, a featuretypical of mafic-bimodal VMS deposits. The pyrite-sphaler-ite-pyrrhotite-chalcopyrite massive sulphide lenses that areunderlain by chlorite-sericite-quartz-rich alteration zonescontaining pyrrhotite-chalcopyrite-pyrite vein stockworks

(Franklin et al., 2005; Galley et al., 2007). The Flin Flon(62.5 Mt), Callinan (8.4 Mt), and newly developed Triple 7(14.2 Mt) deposits occur in a single felsic volcanic-domin-ated interval (Table 1; Fig. 5A) within a 4 km-thick tholeiitic basalt-dominated sequence (Bailes and Syme, 1989;

FIGURE 9 Photograph of representative hydrothermal facies. (A) Sulphide layers within the argillite unit hosting the West Arm VMS deposit; (B) Ar-

gillite-chert layers hosting the White Lake deposit are replaced by pyrite, pyrrhotite, and sphalerite; (C) Strong mylonitization and cataclasis of pyrite-sphalerite ore from the Mandy deposit; (D) Well-developed sucrose texture due to the development of metamorphic blastesis in sphalerite- pyrite ore from the Chisel VMS deposit. (E) Well-developed durchbewegung texture in pyrite-chalcopyrite-pyrrhotite ore from the Callinan deposit.Clasts include bedded chert, altered rhyolite, and quartz vein material; (F) Coarse-grained development of sphalerite and pyrite blastesis in the Ghostdeposit. Note large, twinned sphalerite crystal (arrow).

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Thomas, 1992; Syme and Bailes, 1993). Mafic breccias infilla cauldron structure that is at least 4 km in diameter (Symeand Bailes, 1993; Devine et al., 2002) and is truncated to thenorth by a fault contact with overlying fluvial sedimentaryrocks of the Missi Group. The three orebodies lie along a 3km strike length containing several felsic flows, domes, and

associated heterolithologic debrisfl

ows. The felsic volcaniccentres are underlain by a thick sequence of mafic debrisflows that are interlayered with mafic effusive flows. TheVMS deposits and associated felsic volcanic strata are over-lain by a sequence of shallowly emplaced basaltic andesite peperitic sills in finely layered mafic tuff. These are in turnoverlain by a thick sequence of pillowed mafic flows.

The Mandy and Schist Lake deposits (Fig. 4) are locatedto the southeast of and have a footwall stratigraphy similar to the Flin Flon-Callinan-Triple 7 deposit cluster, but their stratigraphic hanging wall is faulted away (Simard, 2006).Other examples of bimodal-mafic-hosted VMS deposits in-clude the Dickstone and Spruce Lake deposits in the FourmileIsland arc assemblage (Fig. 2), and the Anderson-Stall-Rod-

Linda cluster of VMS deposits in the Snow Lake arc assemblage (Fig. 7). The latter are hosted within a tholeiitic fel-sic rhyolite complex near the top of a >3 km succession of subaqueous low-Ti tholeiitic basalt to basaltic-andesite flowsthat includes a unit of high-Ca boninite (Stern et al., 1995a).High-Ca boninites are relatively rare in modern settings,having only been identified within the northern end of theTongan forearc (Falloon et al., 1992) and in the upper PillowLavas of the Cretaceous Troodos ophiolite (Duncan andGreen, 1987). On the basis of the high-Ca boninite and low-Ti basalt association, Stern et al. (1995a) suggested that the primitive bimodal volcanic phase occurred in a nascent arctectonic setting. This primitive arc succession is intruded bya 15 km-long quartz diorite-tonalite-trondhjemite intrusivecomplex (Fig. 7). The deposits are spatially associated with

a sulphidic exhalite known as the Foot-Mud horizon that ex-tends for 15 km (Bailes and Galley, 1999). The Anderson,Stall, Rod, and Linda deposits are all localized along dif-ferent internal flow contacts with an 8 km-long felsic domecomplex that is intruded by the latest phase of the SneathLake subvolcanic intrusive complex.

Bimodal-Felsic VMS Deposits

VMS-hosting bimodal felsic assemblages are definedas those that include equal proportions of mafic and felsicvolcanic strata, with a large component of both mafic andfelsic volcaniclastic rocks typical of more mature island-arcassemblages (Franklin et al., 2005). The VMS deposits arespatially associated with felsic flow complexes underlain by

differentiated tholeiitic to transitional tholeiite-calc alkaline basalt-andesite-dacite assemblages. The deposits are Zn-Pb-Cu-Ag-Au-rich and characterized by sphalerite-pyrite-ga-lena-rich massive sulphides and a pyrrhotite-chalcopyrite-sphalerite-rich footwall vein and stockwork complex. Theformation of seafloor carbonate appears to pre-date the gen-eration of the massive sulphides, which partially replaces thecarbonate, forming skarnoid alteration assemblages (Galleyet al., 1993). The presence of carbonate is a characteristicof this deposit type in the Paleoproterozoic in several ter-

ranes worldwide (Galley, 1996). Footwall alteration gener-ally includes extensive zones comprising dominant sericite-quartz-pyrite and more restricted chlorite-quartz-pyrite. This bimodal-felsic-type of VMS deposits occurs within the Chiselmature arc sequence of the Snow Lake arc assemblage (Fig.7), and is represented by the Chisel Lake, Chisel North, Lost,

and Ghost deposits (Table 1).In the Hanson Lake Block (Fig. 2), a fractionated arc as-semblage of tholeiitic oceanic arc rocks contains a large felsiccomplex consisting of dacite and rhyolite flows, pyroclastic,and autoclastic flows. The eruptive complex is intruded byquartz-feldspar porphyry and overlain by mixed volcaniclas-tic units followed by well-bedded greywacke. Mixed oxide-silicate-sulphide facies iron formation is also present in placesnear the base of the greywacke succession. Massive sulphidemineralization in located within and on top of the felsic se-quence and includes both Pb-Zn-Cu (Western Nuclear) andCu-Zn (McIlvenna Bay) massive sulphide deposits (Maxeiner et al., 1993).

Hybrid Epithermal VMS Deposits

There are several examples in the literature of a sub-set of bimodal-felsic VMS deposit types that share some of the char-acteristics of epithermal deposits (Hannington et al., 1999;Galley et al., 2007). These deposits are commonly associatedwith rhyolite dome complexes, have high precious metal con-tents, and are dominated by aluminous proximal alteration as-semblages, such as quartz-sericite-pyrite. They are postulatedto have formed in shallow-water environments in which thehydrothermal fluids boiled in the subsurface (Sillitoe et al.,1996; Hannington et al., 1999). This resulted in the oxida-tion and acidification of the rocks and formation of aluminousalteration mineral assemblages. The deposits are commonly pyrite-rich, with abundant tennantite-tetrahedrite, and galena,and have high As, Sb, and Hg, along with Au and Ag. There

may also be a magmatic volatile component involved in theformation of the deposit as a result of degassing of an under-lying, shallow felsic intrusion (Sillitoe et al., 1996).

The Laurel Lake Au-base metal deposit hosted by the WestAmisk arc assemblage is perhaps the only known exampleof a hybrid epithermal-VMS hydrothermal system in the FlinFlon Belt (Fig. 4; Walker and McDougall, 1987; Ansdell andKyser, 1991). The Laurel Lake deposit occurs within a ma-ture oceanic arc, fractionated basalt-dacite-rhyolite sequenceof rocks. The Au-Ag mineralization is hosted within a cluster of hydrothermally altered porphyritic rhyolite and rhyodaciteextrusive complexes, and associated hypabyssal stocks anddikes. Although the host volcanic suite is believed to be approximately 1.88 Ga, the nearby hydrothermally altered syn-

to post-accretionary Missi Island quartz diorite, trondhjemite,and quartz monzonite intrusive suite is considerably younger (1860–1837 Ma; Ansdell and Kyser 1991). The deposit con-sists of a number of anastomosing quartz-sulphide veinscontaining quartz, muscovite, pyrite, tennanite, chalcopyrite,sphalerite, galena, electrum, and carbonate within an altera-tion halo of predominantly sericite-pyrite. Gold and Ag arecontained within the electrum and tennanite. The veins aredeformed and the alteration assemblage is metamorphosedand crosscut by D

2deformation features. The deposit forms




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the core of a broader sericite-pyrite-rich alteration zone thatcontains several other auriferous quartz-sulphide stockwork systems (Byers and Dahlstrom, 1954). The mineralizingfluids were saline (<10.3 wt. % NaCl equivalent) and CO

2-

bearing, and are believed to represent dominantly seawater that was modified at temperatures of around 300˚C (Ansdell

and Kyser, 1991). Pelitic-Ma fic VMS Deposits

Pelite-mafic VMS deposits are a category that includesseafloor deposits that formed in volcanic basins or rifts partlyinfilled by sediment. This can include discrete cauldrons or larger back-arc rifts (Franklin et al., 2005). The infill sedi-ment are commonly a mixture of pelite and turbidite, whichare usually intruded by mafic sills. There is no spatially as-sociated effusive mafic or felsic volcanic strata. The depositsconsist typically of laterally extensive interlayered sulphideand clastic sediment, and include the Cuprus and White LakeZn-Cu-Pb-Ag deposits. These deposits are hosted within theshaley laminated turbidites in a cauldron infill sequencefloored by reworked dacitic tuff and are overlain by andes-

itic lapilli tuff (Bailes and Syme, 1989; Fig. 6). The depositsare composed of both massive sulphide and finely beddedto interbedded sulphide-shale. The sulphides are composedof pyrite, sphalerite, pyrrhotite, chalcopyrite, and galena andare underlain by chlorite-quartz-rich hydrothermal alteration pipes and vein stockworks.

The West Arm Cu-Zn deposit (Fig. 4) is a stratiform, mas-sive to finely bedded ore lens hosted by a sequence of graph-itic argillite, pyritic siltstone, and banded chert-hematite ironformation (Syme, 1987). This deposit is underlain by vari-ably silicified and carbonatized pillowed basalt, and related pillow fragment breccia.

VMS-related Hydrothermal Alteration

One of the distinctive features of VMS deposits is thescale and composition of accompanying hydrothermalalteration (Galley, 1993; Franklin et al., 2005; Galley et al.,2007). VMS deposits form on or near the seafloor from thefocused discharge of hot, saline, low-pH hydrothermal flu-ids (Galley, 1993 and references therein). These ore-formingfluids are generated by the interaction of circulating seawater and sub-seafloor strata at various temperatures, with the heatcommonly supplied by shallowly emplaced subvolcanicintrusions (Galley, 2003). The result of this seawater-rock interaction at temperatures that range from approximately140°C to >400°C is the formation of alteration zones of varying composition and size (Galley, 1993 and referencestherein; Alt, 1995 and references therein). These zones mayrange in size from several tens of kilometres to several tens of metres, depending on whether they represent widespread dif-fuse recharge or discharge of hydrothermal fluid, or focuseddischarge. Each alteration facies has a particular mineralogyand chemistry depending on rock composition, fluid composition, and temperature of rock-fluid reaction. Because of their close genetic and spatial relationship to VMS deposits,alteration zones provide an effective vector in locating mas-sive sulphide mineralization.

In the Flin Flon Belt, seafloor hydrothermal alteration

zones is commonly accentuated through mineralogical chan-ges that accompany regional metamorphism. Where the ori-ginal alteration is characterized by hydrated mineral assemblages, the secondary metamorphic mineral assemblages can be quite distinctive. Large-scale distal alteration can take theform of epidotization, silicification, or Mg-Fe metasoma-

tism. At Snow Lake, the top of the primitive arc Andersonsequence is affected by a zone of silicification and bleachingover 15 km long (Figs. 10; 11B). Extensive zones of Fe-Mgmetasomatism focused about the rhyolite flow complexes arecharacterized by a chlorite-aluminosilicate-rich assemblage.In the Bear Lake Block (Fig. 4), there is a zone of epidotealteration >10 km in strike length and up to 4 km thick (Bailesand Syme, 1989; Fig. 11A). Zones of silicification (Fig. 11C),Fe-Mg metasomatism, and garnet-amphibole alteration (Fig.11D) underlie the Chisel-Lost-Ghost VMS deposits withinthe Chisel sequence (Skirrow and Franklin, 1994; Bailes andGalley, 1999; Figs. 6, 10). Subvolcanic intrusion in both of these arc sequences are also affected by hydrothermal altera-tion (Fig. 11E,F).

Post-accretion MineralizationThe pre-accretion 1.92 to 1.88 Ga oceanic volcanic arcs and

intervening oceanic basin successions were intruded by 1.87to 1.84 Ga post-accretion granitoid intrusions associated withsuccessor arc magmatism (Lucas et al., 1996). Of these, sev-eral contain occurrences of Cu-Mo mineralization (Baldwin,1980; Galley and Franklin, 1987). The 1.84 Ga Phantom-BootLake intrusive suite is spatially associated with numerous oc-currences of Au, Cu, Mo, and W (Galley and Franklin, 1987).The intrusive complex contains a series of sills, stocks, anddike swarms of diorite, hornblende porphyritic quartz monzo-diorite, monzodiorite, quartz monzonite, and microcline porphyritic granite. Mineralization is present as: 1. chalcopyrite,scheelite molybdenum and Au in quartz breccias associated

with hematite-quartz-ferrodolomite-epidote alteration, and2. Au-scheelite-chalcopyrite mineralization in fracture-con-trolled, sericite-quartz-carbonate-pyrite alteration zones.Mineralized vein and breccia-vein systems are hosted either by the various phases of the intrusive complex or by the sur-rounding volcanic strata where they are intruded by feldspar porphyritic granite dikes.

Also formed during this tectonic stage in the evolution of the Flin Flon Belt are Ni-Cu-PGE-rich ultramafic to differen-tiated gabbroic intrusions that are pre-Missi Group sediment-ation but post-date arc accretion and compression. Although poorly constrained, they appear to have been emplaced about45 million years after arc formation between 1.84 and 1.845Ga (Bailes and Theyer, 2006). These intrusions are commonly

small (<1000 m long and 500 m wide) and occur within sev-eral tectonic assemblages of the Flin Flon Belt. These includethe Namew Lake intrusion with associated a Ni-Cu-rich pyr-rhotite-pentlandite-violerite massive sulphide lens (Menardet al., 1996) and the Wonderland gabbro suite (Oliva et al.,2002; Bailes and Theyer, 2006). The Namew Lake intrusionis an 1847±6 Ma komatiitic amphibolite that intruded into anupper amphibolite gneiss terrane, believed to be equivalentto Flin Flon Belt strata, underlying Phanerozoic cover rocks.The McBratney Lake sulphide showing is spatially associatedwith a small gabbro body that is part of the Wonderland in-




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trusive suite dated at 1839±3.9 Ma (Bailes and Theyer, 2006).This PGE-Au-rich showing consists of veins containing chal-copyrite, Ni-pyrrhotite, pentlandite, millerite, and magnetite, plus several PGE-bearing Te-As-Sb minerals (Oliva et al.,2002).

Collisional Tectonic Mineralization

The Flin Flon Belt is host to numerous shear zone-hostedorogenic Au deposits and occurrences that have been ex- ploited at various times over the last 100 years. Only sev-

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en deposits have recorded production, of which only the New Britannia mine has produced over 30 tonnes (1 mil-lion ounces) of Au (Table 2). This compares to the Flin FlonVMS deposit, which alone has produced over 120 tonnes of

Au (Table 1; Fig. 8B).Orogenic Au mineralization in the Flin Flon Belt is local-

ized within ductile oblique-slip, high-angle shear zones andsouthwest-verging thrusts that developed during D

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FIGURE 11. Photographs of representative alteration types. (A) Pillowed basalt flows strongly altered to an epidote-quartz-plagioclase assemblagewithin a zone of extensive hydrothermal alteration below the Cuprus and White Lake VMS deposits in the Bear Lake Block; (B) Silicified marginson pillows from the basal pillowed basalt formation of the Bear Lake Block. (C) Close-up of the complex alteration of a clast of feldspar-phyric basalt, now represented by the metamorphic assemblage of plagiocalse-actinolite-epidote-garnet. The mafic breccia is part of the silicified EdwardsFormation, Chisel sequence, in the deep footwall to the Chisel, Lost, and Ghost Zn-Pb-Cu VMS deposits, Snow Lake; (D) Intense quartz-feldspar-amphibole-garnet alteration at the contact between finely layered mafic volcaniclastic and overlying mafic volcanic breccia of the Edwards LakeFormation. These altered strata lie approximately 1,200 m below the Chisel-Lost-Ghost VMS deposit horizon and are interpreted to represent the

metamorphosed equivalent of a sub-seafloor high-temperature reaction zone (Bailes and Galley, 1999); (E) Columnar-jointed facies in a trondjemite phase of the Sneath Lake subvolcanic intrusion, Anderson sequence (Fig. 6). Fracture surfaces are defined by a quartz-feldspar-staurolite meta-morphic assemblage typical of felsic igneous rocks affected by extensive sub-seafloor Si-Al-Fe-enriched alteration, and then metamorphosed to biotite-almandine facies; (F) Fracture pattern within the margin of the Sneath Lake syn-volcanic intrusion. This trondhjemite phase of the RichardLake subvolcanic intrusion, Chisel sequence (Fig. 6), was hydrofractured by circulating hydrothermal fluids. Subsequent regional metamorphismformed a (Mg-chlorite)-(Mg-biotite)-staurolite-quartz-rich mineral assemblage. Modified from Bailes and Galley (1999).




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O c e a n f l o o r ( b a c k a r c ) m e t a b a s a l t / s y n v o l c a n i c m

a f i c i n t r u s i v e

O c e a n p l a t e a

u m e t a b a s a l t

O c e a n i s l a n

d m e t a b a s a l t

T e c t o n i t e

F E L S I C - M A F I C P L U T O N S

1 . 7 6 - 1 . 8 2 G a

( K i s s e y n e w B e l t p l u t o n s )

1 . 8 3 - 1 . 8 4 G a

( l a t e s u c c e s s o r a r c p l u t o n s )

1 . 8 4 - 1 . 9 0 G a

( e a r l y j u v e n i l e a r c +

e a r l y - m i d d l e

s u c c e s s o r a r c p l u t o

n s )

c a

. 1 . 9 2 G a

( ' e v o l v e d a r c ' p l u t o n s )

M i s s

i G r o u p ( 1 . 8 3 - 1 . 8 5 G a )

C o n t i n e n t a l s a n d s t o n e / v o l c a n i c s

B u r n

t w o o d G r o u p t u r b i d i t e s ( 1 . 8 4 - 1 . 8 5

G a )

S c h i s t - W e k u s k o S u i t e ( 1 . 8 5 - 1 . 8 8 G a )

S U C C E S S O R

A R C

a n d B A S I N

D E P O

S I T S

F F

2 0 k m

P H

A N E R O Z O I C

C O V E R

N

G

N B

G

C C

T TR R

HM H M

F F

W e k u s k o

R R L

P P

MB M B

a

a

a

a

a

a

a

a

F I G U R E

1 2 . D i s t r i b u t i o n o f A u

o c c u r r e n c e s ( y e l l o w d o t s ) a n d p a s t p r o d u c e r s ( r e d d o t s ) i n t h e F l i n F l o n B e l t . T h e A u m i n e r a l i z a t i o n t e n d s t o f o r m a s c l u s t e r s o f o c c u r r e n c e s a l o n g

m a j o r f a u l t s y s t e m s .

G o l d m i n e s : C = C e n t u r y ; F = F e r r o ; G = G u r n e y ; H M = H e n n e y - M a l o n e y ; M B = M o o s e h o r n - B a l l a s t ; N B = N e w B r i t a n n i a ; P = P u f f y L

a k e ; R = R i o ; R L = R e x - L a g u n a ; T = T a r t a n L a k e . F

r o m R i c h a r d s o n a n d

O s t r y ( 1 9 9 6 ) .




528

1900 1880 1860 1840 1820 1800 1780 1760 1740 1720 1700

VOLCANISM Amisk Group (1)

Laurel Lake area (4) Ar-Ar

Tartan Lake Ar-Ar (5)

Rio Rb-Sr (4)

Neagle Lake (2)

Burial

Peak regionalmetamorphism

Low-grade

Cooling

High-grade (1)

Peak thermalconditions inlow-gradeareas (2)

P1

P2

P3

P4

P5

Uplift

Phantom Lake (2)

Ma

PLUTONISM

DEFORMATION

METAMORPHISM

MESOTHERMAL GOLDMINERALISATION

MOLASSESEDIMENTAION

Cliff Lake (1)

Missi Formation, Flin Flon Basin (3)

Annabel Lake/Reynard Lake/Missi Island (2)

FIGURE 13. Summary of tectonic evolution and major deposit-generating events, Flin Flon Belt. Modified from Ansdell and Kyser (1992).

metamorphic conditions between 1840 and 1805 Ma (Ansdelland Kyser, 1991; Fedorowich et al., 1995). These D3 shearsformed as a result of regional collisional shortening, crustalthickening, and uplift (Ansdell and Kyser, 1991; Fedorowichet al., 1995; Lucas et al., 1996). Many of these shear zoneswere then reactivated under post-peak metamorphic brittle-ductile conditions during the D

4, dominantly strike-slip fault-

ing. D4 brittle-ductile shearing and faulting was accompanied

by intense refolding of the D3shear foliations, with steep plun-

ges on the fold axes (Fedorowich et al., 1995). Most of theshear zone-hosted Au occurrences and deposits form in clus-ters about these reactivated fault systems. These include theWekusko Lake, Snow Lake, Northstar, Elbow Lake, PhantomLake, East, and West Amisk Lake clusters (Fig. 12).

Major Au deposits and occurrences are localized at litho-logic contacts that are characterized by a large competencycontrast. In the case of the Tartan Lake and Rio deposits (Figs.4, 12), Au is concentrated along volcanic-intrusive contactswhere there is a dilation or jog in the host shear zones, or anapophysis of volcanic or sedimentary rock in the intrusion(Pearson et al., 1986; Fedorowich et al., 1995). Several of theWekusko Lake deposits lie along the contact between MissiGroup sedimentary strata and shallowly emplaced quartz porphyry sills (Stockwell, 1937). At the New Britannia de-

posit the ore lenses lie along a folded basalt-rhyolite contactnear the apex of a fold (Galley et al., 1988). In the case of the Ferro Mine of the Wekusko Lake camp, the deposit ishosted by a shear zone occupying the axis of a F

2isoclinal

fold (Stockwell, 1937).The principal auriferous vein systems are oriented paral-

lel to the steeply dipping C-fabric within the shear zones,and have a crack-and-seal texture with thin discontinuous bands commonly represented by oriented and sheared frag-ments of mineralized wall rock (Ansdell and Kyser, 1991).Secondary mineralized veins formed as steeply dippingextensional features crosscutting the shear-parallel veins.Auriferous zones within the veins commonly occur near theintersection of anastomosing shears with moderate to steep plunges parallel to the intersection lineation (Fedorowichet al., 1995). The auriferous vein events are commonly pre-dated by the formation of quartz-tourmaline-carbonateveins and overprinted by shallowly dipping quartz veins be-lieved to be associated with development of D

5brittle fault

systems.Gold mineralization formed over an extended period of

post-peak metamorphic cooling between 1790 Ma and 1740Ma, as indicated by the Ar-Ar and Rb-Sr dating of musco-vite and tourmaline from auriferous Au veins (Ansdell and




529

Kyser, 1991; Fedorowich et al., 1995). The O and H isotopesystematics of hydrothermal minerals associated with Aumineralization are compatible with the formation of aurifer-ous fluids by devolatilization reactions during prograde meta-morphism (Ansdell and Kyser, 1992). This is supported byH

2O-CO

2-NaCl fluid inclusion compositions and pressure-

temperature calculations indicating temperatures and pres-sures of vein quartz formation from 250˚ to 420˚C and 1.2 to2.4 kbars (4.3–8.6 km), respectively. Orogenic Au mineraliza-tion began to form in the Flin Flon Belt at approximately 1805Ma and continued for several tens of millions of years duringregional metamorphic cooling with accompanying transitionfrom ductile to brittle deformation.

Summary

There are three major metallogenetic periods in the FlinFlon Belt. The period of pre-accretion mineralization includ-ed formation of VMS deposits within a variety of submarineoceanic arc and extensional arc settings formed between 1.91and 1.88 Ga (Fig. 13). The differing character and distributionof the Flin Flon Belt VMS deposits can be related to different

extensional arc settings.The arc and intra-arc basins were compressed together

to form a series of accreted arc assemblages between 1.88and 1.87 Ga. This post-accretionary period (1.87–1.84 Ga)was characterized by the emplacement of various types of Tonalite-Trondjhemite-Granodiorite (TTG) granitoids associ-ated with successor arc magmatism and tholeiitic to alkalicintermediate to mafic igneous rocks related to faulting and ex-tension. Non-economic Cu-Mo mineralization is recognizedin some of the former, whereas the latter includes intrusionsand related Ni-Cu and PGE mineralization.

The following period of terrane collision and crustal thick-ening between 1.84 and 1.80 Ga was accompanied by the for-mation of fault- and shear-related orogenic Au deposits. Thedeformation and metamorphic history of these dominantlyvein-hosted Au systems indicate formation during faultingrelated to peak regional metamorphic conditions, with con-tinued vein deformation and Au deposition during post-peak cooling conditions.

Exploration Guidelines

Paleoproterozoic volcano-sedimentary terranes are mostcommonly explored for VMS deposits. Next to the Devonian,the Paleoproterozoic is the most prolific era for the genera-tion of VMS deposits (Franklin et al., 2005). As with mostwidely prospected VMS terranes, there is little known of thegeology of the Flin Flon Belt below the surface. Furthermore,much of the southern part of the Flin Flon Belt is overlain

by Phanerozoic sedimentary rocks of varying thickness,which further complicates evaluating the VMS potential of the belt. The most successful method to date for discoveringVMS deposits in the Belt has been airborne electromagneticsurveys. An increased knowledge of the tectonostratigraphyhas allowed companies to use lithogeochemistry to focus on bimodal primitive rifted arc sequences, which host most of the VMS deposits. These tectonostratigraphic domains arenow being traced northwards into high-grade metamorphicterranes originally believed to be part of the sediment-richKisseynew Domain. The presence of amphibolite regional

velop these deposits is the same for most terranes, regardlessof age. These include subsidiary faults and shears parallelto major regional lineaments, volcanic-sedimentary reversefault contacts, competency contrast linked with jogs and off-sets in fault/shear zones, and chemical traps. There is some potential for hybrid epithermal-VMS Au deposits, but thevolumes of the required shallow-water to subaerial terranesare very limited.

Acknowledgements

This summary of the present knowledge of the tectonicand metallogenic evolution of the Flin Flon Belt is the prod-uct of the work of many different geologists, whose contribu-tions we hope are appropriately referenced in the text. Thissummary also benefited from conversations with D. Ames,

K. Ansdell, C. Beaumont-Smith, C. Devine, M. Fedikow,H. Gibson, K. Gilmore, D. Zeihlke, H. Zwanzig and manyothers. We very much appreciated the thorough and thought-ful reviews of the draft manuscript by J. Peter, S. Piercey,and W. Goodfellow. Their comments did much to improvethe clarity of the manuscript.

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Metallogeny of the Paleoproterozoic Flin Flon Belt, Manitoba and Saskatchewan

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