structural evolution of amisk collage - james ryan

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Structural evolution of the eastern Amisk collage, Trans-Hudson Orogen, Manitoba 1

J ames J . Ryan and Paul F. Williams

Abstract : Deformation recorded in the Amisk collage in the central part of the Paleoproterozoic Flin Flon Belt(southeastern Trans-Hudson Orogen) is divided into pre-, early, late, and post-Hudsonian orogeny, distinguished bysignificant changes in metamorphic conditions and the orientation of structures. Detailed structural analysis,petrography, and high-precision geochronology, combined with previous mapping and geochemical studies, indicate astructural history spanning 180 Ma in the Amisk collage, and the database provides an excellent opportunity to studythe structural evolution of Precambrian greenstone belts. Accretion of the 1.921.88 Ga tectono-stratigraphicassemblages in the Amisk collage began prior to 1.868 Ga. The deformational history records six generations of ductilestructures (F 1F 6), followed by development of brittleductile and brittle structures (F 7), which may have continued aslate as 1.690 Ga, during exhumation of the collage. The steep, generally north-northeast macroscopic structural grain isdominated by two regional foliations (S 2 and S 5), and contrasts strongly with the less steeply inclined, eastwest grainin the adjacent Kisseynew Domain. Maximum displacements between tectono-stratigraphic assemblages occurred along

early rather than late shear zones. Vertical extension was important in post-D 1 deformations, even in the later stages.Postorogenic, low-angle extensional features that are common to many mountain belts appear to be absent, possiblyindicating that erosion was the dominant unroofing mechanism.

Rsum : La dformation enregistre dans le collage dAmisk, rgion centrale de la zone de Flin Flonpaloprotrozoque (sud-est de lorogne Trans-Hudsonien), est divise en phases orogniques hudsoniennes antrieure,prcoce, tardive et postrieure, qui se distinguent par des changements significatifs de conditions mtamorphiques etdorientation des structures. Les analyses structurales dtailles, ltude ptrographique et la gochronologie de hauteprcision, allies la cartographie gologique et aux tudes gochimiques existantes, rvlent lhistoire structurale ducollage dAmisk chelonne sur 180 Ma, et la base de donnes fournit une excellente opportunit pour tudierlvolution structurale des zones de roches vertes prcambriennes. Laccrtion des assemblages tectonostratigraphiquesdats de 1,921,88 Ga dans le collage dAmisk a dbut antrieurement 1,868 Ga. Lhistoire de la dformationdocumente six gnrations de structures ductiles (F 1F 6), suivies par le dveloppement de structures fragilesductiles etfragiles (F 7) qui sest probablement poursuivi aussi tardivement que 1,690 Ga, durant lexhumation du collage. Le grain

structural macroscopique, subvertical, orient gnralement nord-nord-est, est domin par deux foliations rgionales (S 2et S5), et il contraste fortement avec le grain inclinaison moins inclin, orient estouest, qui apparat dans leDomaine de Kisseynew adjacent. Les dplacements maximums entre les assemblages tectonostratigraphiques se sontproduits le long de zones de cisaillement prcoces plutt que tardives. Lextension verticale jouait un rle importantdans les dformations postrieures D 1, mme durant les derniers stages. Les structures dextension post-orogniques,subhorizontales, frquentes dans de nombreux domaines montagneux semblent tre absentes, ce qui indique que peut-tre lrosion fut le facteur dominant dans le mcanisme de surrection et mise laffleurement.

[Traduit par la Rdaction] R yan and Williams 273

The fabric of greenstone belts and the history of fabric de-velopment are controlled by a number of variables, perhapsthe most important being the tectonic environment in whichthe greenstone belts have formed (Condie 1981; Windley

1995). For example, greenstone belts deposited on the mar-gins of intracratonic rifts may record a fabric similar to thatof their basement if the autochthonous relationship remainsintact during collisional orogenesis (e.g., Chadwick et al.1989; Bickle et al. 1994). Greenstone belts obducted at basinmargins will likely preserve a basal dcollement, whichtends to be overprinted by upright structures during orogen-esis (e.g., Hoffman 1985; St-Onge and Lucas 1993), compli-cating the fabric. In contrast, greenstone belts that havedeveloped as intraoceanic accretionary complexes (e.g.,Hamilton 1988; Taira et al. 1992; van Staal 1994) prior tocontinentcontinent collision will likely record a greater de-gree of complexity due to the inheritance of early structures.

The Amisk collage in the Paleoproterozoic Flin Flon Belt(southeastern Trans-Hudson Orogen; Fig. 1 a ) provides anexcellent opportunity to study fabric development and the

Can. J. Earth Sci. 36: 251 273 (1999) 1999 NRC Canada

251

Received May 2, 1997. Accepted December 11, 1997.

J.J. Ryan 2 and P.F. Williams. Department of Geology,University of New Brunswick, P.O. Box 4400, Fredericton,NB E3B 5A3, Canada.1Lithoprobe Publication 903.2Corresponding author. Present address: Geological Survey of Canada, 601 Booth Street, Ottawa, ON K1A 0E8, Canada(e-mail: [email protected]).


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252 C an. J. Earth Sci. Vol. 36, 1999

Fig. 1. (a ) Overview map of the Flin Flon Belt, outlining the Morton Lake fault (MLF) as the boundary between the Amisk collageand Snow Lake segment. The age for the Little Swan Lake pluton (LSLP) is from Whalen and Hunt (1994). Abbreviations: ElbowLake shear zone (ELSZ), Iskwasum Lake shear zone (ILSZ), and Berry Creek shear zone (BCSZ). ( b) Block diagram of the southernTrans-Hudson Orogen. The Saskatchewan craton underlies the Flin Flon Belt, and is exposed in a window. The southern KisseynewDomain is a zone of imbrication between the Flin Flon Belt and the Kisseynew Domain. Subsurface geology is after Lucas et al.(1994).


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structural evolution of Precambrian greenstone belts becauseit has been well characterized by regional mapping (e.g.,Syme et al. 1993) and geochemical and geochronological in-vestigations (e.g., Whalen and Hunt 1994; Stern et al.1995 a ), and its deeper architecture has been explored byseismic surveys (Lucas et al. 1994). Despite multiple defor-mations, accretion-related structures that predate Hudsoniancontinentcontinent collision are preserved in the collage(Lucas et al. 1996; Ryan and Williams 1996 b). Decipheringthe structural history of the Amisk collage is difficult, as inmost greenstone belts, because (1) volcanic units are gener-ally discontinuous and have primary heterogeneity;

(2) marker units of well-layered rocks are absent;(3) abundant synvolcanic to postorogenic plutons have ob-scured relationships; (4) deformation is generally concen-trated in the low-grade metavolcanic rocks rather than inplutons; (5) zones of deformation are prone to reactivation,obscuring earlier foliations.

The ElbowCranberryIskwasum lakes area (Figs. 1 a , 2)is most suitable for studying the structural evolution of theAmisk collage for a number of reasons. Its fabric is repre-sentative of the collage, and it records the most completestructural history of any portion of the collage. Bedrock ex-posure was significantly improved in part of the area by a

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R yan and Williams 253

Fig. 2. Geological map of the ElbowCranberryIskwasum lakes area in the eastern Amisk collage. Modified from Ryan and Syme(1997). Geochronology sites are marked by arrows for Ansdell and Ryan (1997), and by asterisks for Whalen and Hunt (1994).Abbreviations: Anvil Lake pluton (ALP), Gants Lake batholith (GLB), Claw Bay shear zone (CBSZ), Loukes Lake shear zone (LLSZ),Grass River fault (GRF); Elbow Lake tonalite (ELT); others as in Fig. 1.


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forest fire in 1989. The area hosts three of the tectono-stratigraphic assemblages (and subordinate basaltic forma-tions) that constitute the Amisk collage, and their boundingstructures (Stern et al. 1995 a , 1995 b; Syme 1995; Lucas etal. 1996; Ryan and Williams 1996 a ). Intrusive rocks varyfrom synvolcanic mafic intrusions, to syntectonic to late tec-tonic calc-alkaline successor arc plutons (Stern and Lucas

1994; Whalen and Hunt 1994; Lucas et al. 1996; Ansdelland Ryan 1997), and aid in bracketing the age of structures.We distinguished generations of structures primarily by

overprinting relationships between folds, fabrics, and shearzones. Where direct overprinting relationships are absent orcontinuous outcrop is lacking, generations of structures wereidentified by style, orientation pattern, and their timing rela-tive to metamorphic mineral growth or intrusion of igneousbodies. Even where different generations of foliations areparallel, they can be distinguished using the above criteria,though caution is required (see also, Williams 1985). Corre-lating generations of structures with episodes of deformationis problematic because multiple generations of structures candevelop locally in a single progressive deformation episode

(Hobbs et al. 1976; Tobisch and Paterson 1988), and thesame structural features need not be developed everywhereduring a heterogeneous deformation. Generations of struc-tures (F n, Sn, etc.) are grouped into episodes of deformation(D i) where distinct temporal or kinematic frameworks couldbe identified.

Six generations of ductile deformation structures and ageneration of late brittleductile to brittle features were dis-tinguished in the eastern Amisk collage. Its north-northeast-trending structural grain in dominated by two regionalfoliations (S 2 and S 5). The regional structural grain has anapparent southward continuation below the Paleozoic coverrocks (Leclair et al. 1997). An internally consistent model ispresented for the development of fabric and structural evolu-tion of the eastern Amisk collage. It is demonstrated that, al-though the collage was structurally modified during finalassembly of lithotectonic domains in the southeastern Trans-Hudson Orogen, the bulk of its structural architecture pre-dates the Hudsonian orogeny sensu stricto. Becausegreenstone belts generally form the oldest portions of theinternal zones in Precambrian orogens, distinguishing de-formation that predates collisional orogenesis is crucial for acomplete understanding of how and when the orogen wasassembled. If interpretations about orogenesis are to bedrawn from structural grain as imaged by means other thatdirect observation (e.g., linear anomalies on geophysicalmaps; reflectors in seismic profiles), it is important to under-stand that some of the structural grain may predate the orog-eny. Both statements are also true of Paleozoic and youngermountain chains.

Hudsonian orogeny was used by Stockwell (1961) to referto metamorphism, plutonism, and associated deformation at1.901.70 Ga in the Churchill province, as determined byKAr dating. The Trans-Hudson Orogen (Fig. 1), as definedby Hoffman (1981), was believed to represent a collision be-tween the Archean Superior and Hearne cratons (Gibb and

Walcott 1971). However, recent geochronological,lithogeochemical, and structural constraints indicate that cer-tain lithotectonic elements in the orogen record tectonismthat predates continentcontinent collision, and significantstructural modification postdates the main collision (Sternand Lucas 1994; Stern et al. 1995 a , 1995b; Lucas et al.1996; Ansdell and Ryan 1997). The structural analysis pre-

sented in this paper uses the orientation of structures to dis-tinguish four main tectonic shortening polarities, which arereferred to as: pre -, early , late, and post -Hudsonian defor-mation (Table 1).

The Flin Flon Belt constitutes the middle portion of a tri-partite, north- to northeast-dipping, crustal-scale thrust pilethat is the southernmost exposed portion of the Trans-Hudson Orogen (Figs. 1 a , 1b). The Archean Saskatchewancraton (3.002.45 Ga) forms the lower structural slice belowthe Flin Flon Belt, separated by an ~2 km wide ductile shearzone called the Pelican dcollement (Lewry et al. 1990,1994; Lucas et al. 1994). To the north, the Kisseynew Do-main forms the upper structural slice (Zwanzig 1990; Nor-man et al. 1995) and is separated from the Flin Flon Belt bya broad zone of imbrication referred to as the southernKisseynew Domain (Fig. 1 b). The Flin Flon Belt is subdi-vided into the Amisk collage (Lucas et al. 1996) and SnowLake segment (Lucas et al. 1999), separated by the MortonLake fault (Fig. 1 a ).

The Amisk collage comprises a variety of distinct 1.921.88 Ga tectono-stratigraphic assemblages (Stern et al.1995 a , 1995 b) that were accreted to form a structurallythickened, intraoceanic complex early (1.881.87 Ga) in thetectonic evolution (Stern and Lucas 1994; Lucas et al. 1996;Ansdell and Ryan 1997). Voluminous calc-alkaline succes-sor arc plutonism overprinted the collage from 1.87 to 1.84Ga (Stern and Lucas 1994; Whalen and Hunt 1994; Whalen

and Stern 1996), during regional eastwest shortening of themagmatic arc (Lucas et al. 1996; Ryan and Williams 1996 b;Ansdell and Ryan 1997). This early portion of the tectonichistory (1.881.84 Ga) is referred to here as pre-Hudsoniandeformation. The thickened magmatic arc, which probablyemerged as a stable microcontinent (Lucas et al. 1996), wasunconformably overlain by the 1.851.83 Ga alluvialfluvialMissi group (Bailes and Syme 1989; Stauffer 1990; Ansdell1993). In the Kisseynew Domain to the north (Fig. 1 a ), thedominant lithology comprises the Burntwood group, whichconstitutes thick accumulations of immature, volcaniclasticmarine turbidites (Bailes 1980), deposited between 1.86 and1.84 Ga (Gordon et al. 1990; David et al. 1996) in a basin(s)on the northeast margin of the microcontinent. During the

early Hudsonian orogeny (1.8351.800 Ga), the Flin FlonBelt was transported southwestward over the colliding Sas-katchewan craton (Lewry et al. 1990; Ashton et al. 1996),and metaturbidites of the Kisseynew Domain underwent in-tense top-to-the-southwest-directed shear, inverting the basinand placing it structurally above the Flin Flon Belt (Zwanzig1990; Norman et al. 1995; Ansdell et al. 1995; Connors1996). The Amisk collage deformed internally along steepshear zones during overthrusting (Ryan and Williams 1995;Lucas et al. 1996). The southeastern portion of the orogenunderwent late-stage (late Hudsonian deformation) sinistraltranspression (Bleeker 1990), which postdates the peak of

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regional metamorphism in low-grade rocks, and modifiedlow-angle boundaries and metamorphic isograds (Kraus andWilliams 1998). In higher grade rocks, however, tempera-tures appear to have remained high after this deformation,because metamorphic reaction isograds transect regionalfolds (Connors 1996; Kraus and Menard 1997). Post-Hudsonian deformation from 1.775 to 1.690 Ga, during up-

lift and erosion of the Trans-Hudson Orogen (Bleeker 1990;Fedorowich et al. 1995), occurred under brittleductile tobrittle conditions.

Supracrustal rocks in the ElbowCranberryIskwasumlakes area (Figs. 1 a , 2) comprise three distinct tectono-stratigraphic assemblages: the ElbowAthapapuskow oceanfloor assemblage, the Flin Flon arc assemblage, and theocean island assemblage (Stern et al. 1995 a , 1995 b). The El-bowAthapapuskow ocean floor assemblage, which com-

prises subaqueous mafic volcanic rocks and a complex of maficultramafic rocks, makes up most of the southeasternportion of the map area (Fig. 2). On the basis of field char-acteristics and geochemistry, the assemblage is subdividedinformally into formations, composed principally of sub-aqueous basaltic flows and sparse, thin interflow sedimen-tary layers (Stern et al. 1995 a , 1995 b; Syme 1992, 1995).

The maficultramafic complex (Syme 1992), which occurssporadically throughout the eastern portion of the map area(Fig. 2), is separated by a structural break from the ClawBay formation. Based on a calculated liquid line of descent,Stern et al. (1995 b) interpreted the cumulates of the ultra-mafic complex as being related to basalts of the Claw Bayformation by crystal fractionation; however, no direct link has been established between the complex and any one of the basalt formations. Collectively, the basalts and themaficultramafic complex are interpreted as having formedin a back-arc setting, representing an ophiolite obductedonto the bordering arc rocks (Stern et al. 1995 b; Lucas et al.

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P r e - H u

d s o n

i a n

d e f o r m a t i o n

e a r l y

l a t e H

u d s o n

i a n O r o g e n y

P o s t - H u

d s o n

i a n

d e f o r m a t

i o n

Deformationepisode

Orogen-scaleshorteningdirection Structures Metamorphism

Table 1. Summary of tectonometamorphic evolution of the Flin Flon Belt and Kisseynew Domain.

Age(Ga)

Tectonicsetting

D1

D2

D3

D4

D5

? S : Low-angle

mylonite zones1 ? Pre-1.870 Intraoceanic accretion of

tectono-stratigraphic assemblages

EW F /S : Upright, NS

regional folds andfoliation

2 2 M : Contact

metamorphismduring plutonism

1 1.8701.845 Successor arc plutonism,regional shortening

coevalwith of magmatic arc

1.8501.835

NNESSW

NNESSW

F /S : ILSZ, which

reactivated S

mylonites

3 3

1

1.8401.805

Sedimentation of Missi andBurntwood groups

Collisional tectonics , initialcollapse of turbidite basin, andSSW transport of KD over FFB

F /S : CBSZ, and

dextral shear alongthe BCSZ

4 4

M : Peak of regional

metamorphismassociated with the

2

Hudsonian orogeny

Continued tectonic transportof KD to the SSW over the FFB

1.8301.800

NWSE F /S : NW-trending

ELSZ and regionalcrenulation cleavage

5 5 M : Chlorite

retrogression2 1.8051.770 Sinistral transpression of the

THO and intensifying of theNNE structural grain in the FFB

? F /S : EW-trending

features along thereactivated BCSZ

6 6 ?? Reactivation of the ELSZ andthe BCSZ, possibly by a changein tectonic regime outside the belt

? Brittle-ductile andbrittle features

Below blockingtemperature inAr/Ar systematics

1.7701.690 Late-stageduring

orogen-parallelmovementsuplift and erosion

Abbreviations: Berry Creek shear zone (BCSZ), Claw Bay shear zone (CBSZ), Elbow Lake shear zone (ELSZ), Flin Flon Belt (FFB),IskwasumLakeshear zone(ILSZ),KisseynewDomain(KD),Trans-HudsonOrogen(THO).


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1996). The maficultramafic complex is interpreted as layer

3 of the dismembered ophiolite (Lucas et al. 1996).Arc rocks (part of the Flin Flon assemblage) are exposedonly in the northwestern part of the map area (Figs. 1 a , 2),and are compositionally and lithologically more diverse thanthe ocean floor rocks (Syme 1991; Stern et al. 1995 a ). Theyare characterized by the association of pillowed and massivemafic flows, intermediate to felsic tuffs and breccias, andrhyolites (Syme 1991, 1992; Stern et al. 1995 a ). The oceanisland assemblage (Long Bay formation), exposed only onthe northwest side of Elbow Lake (Fig. 2), is one of the mostlithologically and geochemically distinct packages in theAmisk collage (Syme 1991; Stern et al. 1995 b ). It is a mafic

conglomerate interpreted as a product of cyclic submarine

debris flows. Clasts in the conglomerate are dominated bybasalts that appear to have been subaerially erupted, andhave geochemical signatures that are characteristic of oceanisland basalts (Stern et al. 1995 b ; J.J.Ryan, unpublisheddata). The unit forms an apron along the southeast margin of the arc assemblage and is only metres thick along the west-ern margin of the Elbow Lake shear zone (ELSZ) in thenorthern portion of the area.

Successor arc granitoids (Whalen and Hunt 1994; Morri-son and Whalen 1995) make up 6070% of the map area(Fig. 2), and grade with age from tonalite and hornblendegranodiorite (e.g., Gants Lake batholith) to granite

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Fig. 3. Form surface map for the eastern Amisk collage (see text for discussion). Domains AN are outlined. Some data adopted fromSyme and Morrison (1994).


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granodiorite (e.g., Anvil Lake pluton). U/Pb zircon and ti-tanite ages (Whalen and Hunt 1994) in the Elbow Lakerange from 1.876 to 1.826 Ga (Figs. 1 a , 2). Dykes vary incomposition and orientation, and tend to be concentrated inshear zones.

Mapping of metamorphic zones and reaction isograds isdifficult in the eastern Amisk collage because of discontinu-ous outcrop, the presence of shear zones and faults, two epi-sodes of metamorphism (M 1 and M 2), and the very nature of low- to medium-grade metamorphism in mafic rocks. Meta-morphosed pelites produce a relatively high number of min-erals (e.g., staurolite, cordierite, sillimanite) whosechemistry is diagnostic of grade (Essene 1989). Mafic rocksmetamorphosed at low to medium grade produce relativelyfew minerals (e.g., chlorite, amphibole, epidote) that havevariable chemistry, and are much less diagnostic of grade(e.g., Bgin 1992).

An episode of contact metamorphism (M 1) peaked at am-

phibolite facies in contact aureoles of the 1.8761.864 Gaplutons. Effects of M 1 are rarely observed more that 1 kmfrom the margins of plutons. M 2 is a regional metamorphismassociated with the Hudsonian orogeny sensu stricto, andpeaked at 1.8201.805 Ga (Gordon et al. 1990; David et al.1993, 1996; Ansdell and Norman 1995; Parent et al. 1995).In this study, metamorphic assemblages represent M 2, exceptin the vicinity of major plutons.

Mafic rocks contain the upper greenschist M 2 assemblageof chlorite + actinolite biotite east of the ELSZ at ElbowLake. A sharp eastward increase to epidoteamphibolite fa-cies is marked by a change in plagioclase composition fromalbite to oligoclase, coincident with the disappearance of chlorite (Liou et al. 1974; Laird and Albee 1981), and by the

presence of pargasitic hornblendes and ferrotschermakites.M2 metamorphic grade increases to amphibolite faciesnorthwards, and decreases southwards to upper greenschistat Iskwasum Lake. West of the ELSZ, M 2 metamorphicgrade varies from amphibolite facies in the north, toactinolite-bearing upper greenschist facies west of ElbowLake, to middle greenschist (chlorite magnesianchloritoid actinolite) in the area between southwest ElbowLake, First Cranberry Lake, and Iskwasum Lake. Variableamounts of hornblende (1080%) in mafic tectonites southof the ELSZ on the south side of First Cranberry Lake(Fig. 2) indicate a sharp increase to amphibolite faciessouthward toward the Berry Creek shear zone.

Retrograde M 2 assemblages are ubiquitously marked bychlorite and calcite, which are key minerals for identifyingdeformation fabrics that postdate the peak of M 2. Offset of metamorphic isograds in the eastern Amisk collage cannotreadily be used as an indicator of movement direction acrossthe retrograde shear zones, because the position of isogradsis not well known in map view, and even less so in three di-mensions.

Structural analysis is a necessary step in deciphering thehistory of deformation and the accompanying tectono-

metamorphic conditions. We divide the structures in theeastern Amisk collage into seven generations, which devel-oped during five episodes of deformation (D 1D 5; Table 1).Fourteen domains (AN; Fig. 3) are defined by the uniqueorientation of the dominant local foliation, almost all of which are near vertical.

Form surface traces for the more regionally pervasive

foliations (Fig. 3) illustrate how the structural grain is influ-enced by the orientation of intrusive bodies. Domain G en-compasses the well-layered Long Bay formation, and isdiscussed separately. The ELSZ and the Iskwasum Lakeshear zone (ILSZ) are the most prominent structures in theeastern Amisk collage (Fig. 2), and provide new insightsinto understanding the formation and modification of thecollage.

D1 deformation

First-generation structuresThe southern extension of Elbow Lake (Fig. 2) marks a

multiply reactivated high-strain corridor where foliations of different generations are parallel. The earliest fabric (S 1) oc-curs in a northsouth-trending zone of mylonitized basalt,diabase, and gabbro, and can be identified with confidenceonly where cut by a suite of pink, quartz-porphyritic tonalitedykes (Figs. 4 a , 4b). One of the dykes yielded an impreciseU/Pb zircon age of 1868 12

21+ Ma (Ansdell and Ryan 1997).

All other generations of structures overprint the suite of dykes, which we interpret as being associated with the1864 4

5+ Ma (Whalen and Hunt 1994) Elbow Lake tonalite

(Fig. 2). Although the macroscopic S 1 foliation is well pre-served, grain-scale S 1 deformation features have been oblit-erated by epidoteamphibolite facies M 2 metamorphism.

The high-strain corridor in southern Elbow Lake separatesthe McDougalls Point formation and the Claw Bay forma-tion (Ryan and Williams 1994; Syme 1995), but because thecorridor records multiple episodes of shear displacement,isolating the effect of D 1 is impossible. We interpret the gen-eral occurrence of Claw Bay basalts between rocks of themaficultramafic complex and the successor plutons as be-ing an effect of drag during emplacement of the plutons, andas strong evidence that the maficultramafic complex isstructurally above the Claw Bay formation. The structure(s)separating them predates the pluton and is thus interpretedas a first-generation structure. The structure must have beenshallow prior to pluton emplacement, and because itemplaced deeper level rocks (ultramafics) on shallower level

rocks (pillowed basalts), it was most likely a thrust. Pre-1.868 Ga S 1 mylonites are among the oldest fabrics in theFlin Flon Belt, and are interpreted here as shear zones asso-ciated with intraoceanic accretion in the Amisk collage (cf.Lucas et al. 1996). No kinematic indicators are preserved inthe S 1 mylonites to support or refute this interpretation.

Foliations that predate the S 2 regional foliation inintervolcanic sedimentary layers on western Elbow Lake andcentral Iskwasum Lake are assigned to S 1. These fabrics arepreserved in microlithons of the S 2 crenulation cleavage, andtheir enveloping surface is parallel or at a low angle to bed-ding.

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D2 deformation

Second-generation structuresThe first regional foliation (S 2) in the map area is also

the most widespread (Fig. 3). It varies in intensity fromweak flattening of primary features (e.g., west side of ElbowLake) to tectonite-gneissic layering (Fig. 4 c) adjacent to thelarger plutons (especially the Gants Lake batholith). Itchanges in orientation across the area, and is consistantlyoverprinted by the peak M 2 mineral assemblage (see below).S2 is most pervasive along the eastern side of the ELSZ fromnortheast Elbow Lake (Fig. 3) to the southern limit of theexposed Amisk collage. It is moderately to weakly devel-

oped along the Cranberry lakes, and strongly developed

south of First Cranberry Lake. It is generally a differentiatedcrenulation cleavage where it overprints finely spaced bed-ding laminae or S 1 at high angles. In plutonic rocks, it is de-fined by inequant quartz and feldspar, or by a gneissiclayering.

Map-scale trends of S 2 are largely controlled by the shapeof the older plutons, many of which are slightly elongate ina north-northeast direction, and S 2 is strongly developed ontheir east and west margins. North and south margins of theplutons appear to form strain shadows in which S 2 is at ahigh angle to the contacts. Early phases of the plutons, andlate-stage dykes of varying composition (tonalite, diabase,

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Fig. 4. (a ) S1 mylonite south of Elbow Lake, derived from gabbrodiabasebasalt. Mafic xenoliths in the tonalite dyke also contain S 1.Pencil, at arrow, is 14 cm long. ( b) Closer view of dyke contact in ( a ), shows thin ultramylonite layers in a gabbroic host, truncatedby tonalite that contains a weak S 2 defined by inequant quartz grains, which is similar in orientation to S 1. (c) Typical S 2 layeredtectonite on the eastern side of Elbow Lake. Hammer handle, 35 cm long, points toward 150. ( d ) Hinge area of a large S-asymmetricF2 fold in laminated mafic sandstone of the Long Bay formation. Hammer handle, 35 cm long, points toward 040. ( e) Intrafolial S-asymmetric drag fold (at arrow) in south-southwest-trending S 3 layering on Leaping Moose Island, where the Claw Bay shear zoneintersects S 3. Z-asymmetric F 4 folds indicate dextral shear along the Claw Bay shear zone. Pencil, 14 cm long, points toward 330,

parallel to the weak S 4.


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quartzfeldspar porphyry) within the plutons, all record S 2.This fact, combined with the strain shadows on the northand south ends of plutons, leads us to believe that the en-hanced strain at the margins of the plutons is due topostemplacement tectonic deformation rather than forcibleemplacement deformation (cf. Paterson et al. 1989).

Orientation of S 2: The average trend of S 2 in domains A, B,and C varies between northwest and northeast (Figs. 5 a5c).In domain D, S 2 has been strongly transposed by the ELSZ.

In domain E, S 2 trends generally north-northeast (Fig. 5 d ),in contrast to the north-northwest trend of bedding that itoverprints (Fig. 5 e), and both show the effect of F 5 folding.In domain F, S 2 trends predominantly northwest near theELSZ, where most measurements were recorded, but locallytrends northeast on the short limb of F 5 macroscopic S folds(Fig. 5 f ).

The trend of S 2 varies between east-northeast and north-northeast in domains K, L, and M (Figs. 5 g5 i). In the southpart of First Cranberry Lake (domain N), S 2 is defined by an

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Poles to Sdomain J

2

N = 11

Poles to Sdomain I

2

F = 138/814

N = 138

Poles to Sdomain H

2

N = 14

Poles to Sdomain L

2

N = 81

Poles to Sdomain M

2

N = 22

Poles to Sdomain N

2

F = 095/876

N = 43

Poles to Sdomain K

2

N = 12

Poles to Sdomain F

2

N = 112

Poles to Sdomain E

0

N = 67

Poles to Sdomain E

2

F = 120/835

N = 21

Poles to Sdomain C

2

N = 99

Poles to Sdomain B

2

N = 56N = 211

Poles to Sdomain A

2

N

NN N

NNN

NNNN

N N

a) b) c)

d) e) f) g)

h) i) j) k)

l) m)

Fig. 5. Equal-area lower hemisphere projections of S 2 foliations, where N is the number of data points.


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amphibolite facies tectonic layering (locally gneissic). Here,S2 is locally folded by centimetre-scale Z folds, which ap-pear to have developed in the hornblende stability field, andmay be F 3 structures regionally. S 2 in domain N is deflectedtoward an eastwest orientation between the ELSZ to thenorth and the Berry Creek shear zone to the south. The in-tensity of deflection increases southward, toward the Berry

Creek shear zone. The macroscopic fold in S 2 that plunges87 towards 095 (Fig. 5 j) is interpreted as F 6, because it isparallel to the average F 6 fold axis measured in these rocks(see later).

In domain H, north of Iskwasum Lake, S 2 trends predomi-nantly northwest (Fig. 5 k ) except where affected by F 5 folds.The orientation of S 2 in domain I strongly reflects the influ-ence of F 4 and F5 folding (Fig. 3). The great circle girdledistribution of S 2 (Fig. 5 l), plunging 81 towards 138, coin-cides most closely with measured F 4 fold hinges. In domainJ, S2 is deflected into an eastwest orientation (Fig. 5 m), dueto a large-scale F 4 fold and outcrop-scale F 6 folds. The mostintense S 2 fabrics in the entire map area are associated withthe north-northeast-trending Loukes Lake shear zone, whichlies within the Gants Lake batholith (Morrison and Whalen1995).

Regional significance and timing of S 2: The best-preservedmacroscopic F 2 folds are in the Long Bay formation and pil-lowed flow sequences in the arc assemblage of northwest El-bow Lake. F 2 folds are tight and have upright axial planes.F2 axes are predominantly steeply plunging, parallel to theF5 folds. However, a location in the Long Bay formation ex-hibits large-scale (tens of metres), subhorizontal, upright F 2folds (Fig. 4 d ). Dome and basin interference patterns be-tween shallow F 2 folds and steep F 5 folds (both have uprightaxial planes) are present at three locations around Elbow

Lake. We conclude that F 2 folds were originally shallowlyplunging, and F 5 horizontal shortening at a high angle to F 2axial planes resulted in steep F 5 folds and steepening of F 2hinges. For this reason, we interpret D 2 as being the defor-mation episode responsible for steepening the strata (and S 1shear zones), and development of a regional foliation andprominent structural grain. Subsequent deformation has lo-cally intensified the grain (Ryan and Williams 1996 b). Lucaset al. (1996) made similar conclusions for F 2 folds in theFlin Flon area, based on a shallow enveloping surface.

Timing of S 2 development can be bracketed by the1868 1221

+

Ma tonalite dykes (Ansdell and Ryan 1997) that cut S 1 andcontain S 2 (Fig. 4 b), and by the Little Swan Lake pluton(Fig. 1 a ), which truncates S 2 features and yielded an U/Pb

titanite age of 1826 5 Ma (Whalen and Hunt 1994).Ansdell and Ryan (1997) scrutinized the effects of early D 2shear in southern Elbow Lake by dating two grey tonalitedykes that are parallel to S 2. The first dyke is intenselystrained and is intruded by the second that contains a weak S2 fabric. High-precision U/Pb zircon ages of 1866 2

3+ Ma for

the first dyke and 1866 12

+ Ma for the second dyke (Ansdell

and Ryan 1997) indicate broadly synkinematic intrusion. In-tensity of S 2 in the plutons decreases with age (Figs. 1 a , 2),such that it is strong in the ~1.8761.864 Ga plutons, weak to moderate in the ~1.845 Ga plutons, and absent in the~1.830 Ga plutons (e.g., Little Swan Lake and Anvil Lake

plutons). This indicates that D 2 deformation was waning by1.845 Ga. S 2 is interpreted as a regional foliation from ReedLake (Fig. 1 a ) in the east (Syme et al. 1995) to Flin Flon inthe west (Lucas et al. 1996), because pervasive foliations inthose areas are similar in orientation, style, and timing to S 2in this study.

D3 deformationThird-generation structures

The ILSZ is the most extensive S 3 structure in the maparea (Fig. 2). Despite being reactivated by the S 5 ELSZalong its northern segment, the ILSZ is preserved trendingsouth-southwest (Fig. 6 a ) along the western margin of theElbow Lake tonalite, and southeast through Iskwasum Lake.The northernmost occurrence of S 3 is in the Leaping MooseIsland area (Fig. 2), where it trends south-southwest and hasintrafolial S-asymmetric drag folds (Fig. 4 e) with verticalaxes (Fig. 6 b). S folds are consistent with S 3 sinistral offsetsin dykes in the southern Elbow Lake area. S 3 trends south-east in domains H and I (Figs. 6 c, 6d ), where it is folded,

and contains a steep L 3 stretching lineation (Figs. 6 e, 6 f ).The ILSZ sweeps into an eastwest orientation near theBerry Creek shear zone to the south (Fig. 3).

S3 can be distinguished from other fabrics by its uniquerelationship to regional metamorphism. S 1 and S 2 foliationsare overgrown by M 2 minerals, whereas S 4- and S 5-relatedmylonites postdate M 2 and tend to cleave well because theyare defined by retrograde chlorite. S 3 mylonites are finegrained and are defined by the peak mineral assemblage.Epidoteamphibolite facies S 3 mylonites in southern ElbowLake exhibit dynamic recrystallization of amphiboles(Fig. 7 a ), which were not coarsened by metamorphism sub-sequent to the deformation. Matrix quartz in greenschist fa-cies S 3 mylonites at Iskwasum Lake is well recovered but

still fine grained (


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same quadrant as the macroscopic fold axes determinedfrom the distribution of S 2 and S3 in domain I (Figs. 5 l, 6d ).S4b axial planar crenulation cleavage is vertical and trends115 (Fig. 6 k ). In an outcrop outside the ILSZ in the south-east part of Iskwasum Lake, spectacular overprinting amongS1, S2, S4, and S 5 in the hinge area of a large F 4b fold(Figs. 11 a 11 d in Ryan and Williams 1996 a ) mimics themap-scale overprinting. The large F 4b fold apparently devel-oped in response to dextral shear along the eastwest-trending Berry Creek shear zone (Figs. 1, 2), which placedthe originally northwest-trending ILSZ in the shorteningfield.

D4 deformation

Fifth-generation structuresS5 is the second regional foliation in the eastern Amisk

collage. In contrast to S 2, however, S 5 has a consistent trendbetween 010 and 050, and comprises either ultramylonitezones or penetrative crenulation cleavage in their wall rocks.Similar to S 2, the orientation of S 5 is strongly influenced bythe proximity to large plutons. The ELSZ is the most region-

ally significant D 4 structure in the map area (Figs. 1, 2), butit does not form a continuous strand between Elbow Lakeand Third Cranberry Lake, as discussed below. It varies inwidth from 2500 m in central Elbow Lake to tens of metreselsewhere, possibly largely controlled by rheology, or the ef-fects of local thickening and thinning by late-stage faults.

Orientation of fifth generation structures: In domain D,where the ELSZ has been mapped in most detail, S 5 trendsnorth-northeast, and is rarely more than 10 from vertical

(Fig. 8 a ). L5 stretching lineations are generally steep(Fig. 8 b). Only lineations that are demonstrably due tostretching (stretched mineral grains and pebbles) are in-cluded (Fig. 8 b); lineations that may or may not involvestretching (intersection lineations, preferred mineral orienta-tion) are excluded. Scattered F 5 intrafolial fold axes are dis-persed in the S 5 plane (Fig. 8 c).

The trend of the S 5 axial plane fabric in domains A, B, E,and F varies between 010 and 035 (Figs. 8 d , 8e, 8g, 8h).F5 folds of S 2 and S4 foliations in domain A plunge steeplynorth and south (Fig. 8 f ). In domains K, L, and M, S 5 in theELSZ trends between 040 and 050 (Figs. 8 i, 8k , 8n). In

1999 NRC Canada


N = 21N = 27 N = 14

N = 7N = 212N = 14N = 7

F = 163/824

N = 65

F = 195/825

N = 45N = 17N = 33

L lineations

domain A4

Poles to S

domain I4bF axes

domain I4b

F axesdomain A

4aPoles to Sdomain A

4aL lineationsdomain I

3L lineationsdomain H

3

Poles to Sdomain I

3Poles to Sdomain H

3F axesdomain D

3Poles to Sdomain D

3

N

NN NN

NNN

N NN

a) b) c) d)

e) f) g) h)

i) j) k)

Fig. 6. Equal-area lower hemisphere projections of S 3, L3, F3, S4, and F 4, where N is the number of data points.


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Fig. 7. Photomicrographs. ( a ) Epidoteamphibolite facies S 3 mylonite from south of Elbow Lake. Amphiboles have an average grainsize of 35 m, of which most grains exhibit polygonization and subgrain development, with local strain-free new grains (by arrow).The texture resulted from dynamic recrystallization of amphibole during F 3 shear. Cross-polarized light, field of view is 1.3 mm.(b) Hornblende (Hbl) crystal, which overgrew S 2, was boudinaged parallel to S 5 crenulations. Hbl in the boudin neck has been alteredto chlorite (chl) and has carbonate (carb) infill. Plane-polarized light, field of view is 6.5 mm. ( c) S5 crenulation cleavage developedacross an annealed S 3 mylonite, shown in plane-polarized light (left) and cross-polarized light (right). S 5 crenulation septa lack quartzand have concentrated opaque minerals. Field of view is 1.3 mm. ( d ) S/C fabrics (top to the left) in a sheared tonalite from south of

Elbow Lake. Ultramylonitic gouge, defining one of the C planes, is itself folded with S asymmetry (at arrow). Cross-polarized light,field of view is 6.5 mm. ( e) Quartz mylonite with well-developed S/C relationships. The C plane is defined by layers of quartz grainswith similar crystallographic orientations (light and dark bands). The S fabric is oblique 25 clockwise to the C plane, indicating aneast-side-up sense of shear. Average grain size ~90 m. Cross-polarized light, field of view is 6.5 mm.


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contrast to the vertical linear features at Elbow Lake, L 5 andF5 are more shallowly plunging in the southwestern domains(Figs. 8 j, 8l, 8m, 8o, 8 p).

S5 trends between 040 and 045 in domains H and I(Figs. 9 a , 9c), defining the axial plane crenulation cleavageto outcrop- and map-scale F 5 folds of the ILSZ (Fig. 3). F 5hinges in domain H plunge steeply south (Fig. 9 b), consis-tent with the macroscopic fold indicated by the orientationof S3 (Fig. 6 c). F5 hinges in domain I plunge steeply north-

east (Fig. 9 d ), inconsistent with the folds in S 2 and S 3(Figs. 5 l, 6d ), indicating that F 4 had a more significant effecton S2 and S3 in domain I than did F 5.

Characteristics, timing, and shear sense of the ELSZ: S5 ev-erywhere overprints the peak M 2 mineral assemblage. Inrocks that contain M 2 hornblende (Fig. 7 b), S5 is generallydefined by aligned chlorite and fine-grained biotite, indicat-ing that S 5 developed in the chlorite + biotite stability field.

1999 NRC Canada


N = 161

N = 21N = 29N = 111N = 12

N = 9N = 67N = 7N = 18

N = 115N = 139N = 39N = 26

N = 22N = 215N = 751

L lineationsdomain M

5 F axesdomain M

5Poles to Sdomain M

5F axesdomain L

5

L lineationsdomain L

5Poles to Sdomain K

5 L lineationsdomain K

5 Poles to Sdomain L

5

Poles to Sdomain F

5Poles to Sdomain E

5F axesdomain A

5Poles to Sdomain B

5

Poles to S

domain A

5

F axes (sz)domain D

5L lineations (sz)domain D

5

Poles to S (sz)

domain D

5

N

NN NN

NNN

N

N

N

N

N

N

N

N

a) b) c) d)

e) f) g) h)

i)

m)

j)

n)

k)

o)

l)

p)

Fig. 8. Equal-area lower hemisphere projections of fifth-generation structures from the Elbow Lake area, where N equals number of data points.


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In the Iskwasum Lake area, S 5 generally occurs as a weaklydifferentiated crenulation cleavage overprinting S 2 and S 3. Inthe microlithons (Fig. 7 c), quartz exhibits a relatively fine-grained polygonal foam texture, characteristic of the S 3 foli-ation. Quartz is completely dissolved from the S 5 septa, con-centrating fine-grained opaque minerals, and appears to haveleft the system. This observation is common to differentiatedS5 crenulation cleavage throughout the map area. Quartz in

the microlithons exhibits undulose extinction, but no high-strain features, probably indicating strong partitioning of strain into the septa during S 5 development.

Margins of the ELSZ, which are not commonly exposed,have high-strain gradients (Fig. 10 b), and S 5 has the sameorientation in the wall rocks and in the shear zone. Evenwhere S 5 is subparallel to earlier fabrics, it is readily distin-guished because it developed under retrograde metamorphicconditions. Where S 5 is the first foliation in fine-grainedchloritic wall rocks, it generally forms a slaty cleavage, orit may be defined by flattened primary features (e.g., pil-lows). S 5 shear zones in mafic rocks are generally character-

ized by chloritic phyllonites, whereas in granitoids, they arecharacterized by a finely spaced domainal cleavage, givingthe rocks an almost shredded appearance (Fig. 10 a ). Cleav-age surfaces accommodated slip, and are coated with finechlorite sericite. Local narrow S 5 ultramylonite shearzones (120 cm) in the granitoid rocks have an average ma-trix grain size of less than 5 m. Intrafolial carbonate andiron oxide staining are ubiquitous in all rocks containing S 5

fabrics.Timing of S 5 development is poorly constrained because itis not cut by dated igneous bodies or minerals. S 5 postdatesthe peak of regional metamorphism (1.8201.805 Ga;Gordon et al. 1990; David et al. 1993, 1996; Ansdell andNorman 1995; Parent et al. 1995) because it overprints thepeak M 2 mineral assemblage. Regional considerations mustbe made to bracket the lower limit of F 5 development. S 5 inthe Snow Lake segment appears axial planar to regional-scale (30 km) F 5 folds (e.g., Threehouse synform; F 3 of Kraus and Williams 1998) that control the structural grainnortheast of Reed Lake (Fig. 1). F 5 folds in the Snow Lake

1999 NRC Canada


N = 7 N = 30 N = 5

N = 24

N = 6 N = 7

N = 23

N = 15N = 14

N = 27

N = 30

S planesdomain J

6b

Poles to dextralshear planesdomain J

mean dextralshear plane

F axesdomain J

6bF axesdomain D

6a

Poles to Sdomain N

6b

Poles to Sdomain D

6a

F axesdomain I

5

F axesdomain N

6b

Poles to Sdomain I

5Poles to Sdomain H

5 F axesdomain H

5

N

NN N

N

N

N

N

NN

N

mean of Sin domain D

5

a) b) c) d)

i)

e)

j)

f) g) h)

k)

Fig. 9. Equal-area lower hemisphere projections of fifth-, sixth-, and seventh-generation structures, where N equals number of datapoints. Foliations in ( h) are plotted as great circles. In ( k ), the mean of the dextral shears (118/90) is plotted as a great circle (dashedline). The mean of S 5 in the ELSZ in domain D (147/85) is also plotted as a great circle (solid line).


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segment appear to predate tighter, shorter wavelength, re-gional-scale folds in the Thompson nickel belt (F 3 structuresof Bleeker 1990). Timing of F 3 development in the Thomp-son nickel belt is well constrained by a suite of pegmatitedykes associated with retrograde metamorphism. The oldestdykes were mylonitized during F 3, intermediate-aged dykesare lesser deformed, and the younger ones crosscut the F 3

folds at high angles and are themselves only weakly de-formed (Bleeker and Macek 1996). These three differentages of pegmatites yielded high-precision U/Pb zircon agesof ca. 1.786, 1.771, and 1.770 Ga respectively (W. Bleeker,unpublished data). If regional correlation of structures is cor-rect and the deformation is not too diachronous, S 5 in theAmisk collage and Snow Lake segment must predate the1.7861.770 Ga folding in the Thompson nickel belt (Ta-ble 1).

L5 lineations, which are best defined by stretched clastsand phenocrysts, have aspect ratios of the order of 20:3:1 inmost outcrops. Although they represent the accumulated fi-nite strain of multiple deformations, their vertical orientationis associated with F 5 deformation. Reliable kinematic indica-

tors (macroscopic and microscopic) in the ELSZ demon-strate sinistral shear for S 5 (Figs 7 d , 10c). Along the entireELSZ, west-side-down shear sense indicators are seen inthin sections cut parallel to the vertical stretching lineation(Fig. 7 e), but tend to be more rare, and developed at asmaller scale than the more prominent sinistral kinematic in-dicators. A similar retrograde mineral assemblage (chlorite +muscovite + calcite + quartz) occurs in boudin necks andpressure shadows associated with the vertically plunging L 5stretching lineation, and in asymmetric pressure shadows as-sociated with both sinistral and west-side-down kinematicindicators. We interpret these similarities as indicating thatthe lineations and pressure shadows are related in time. Im-plications of vertical stretching lineations, and both sinistraland dip-slip kinematic indicators are addressed in the Dis-cussion.

Domain G (Long Bay area): Domain G in the Long Bayarea is treated separately because it is the only domain withgood primary layering, and five generations of ductile struc-tures (locally in the same outcrop). Bedding trends eastwest(Fig. 11 a ), parallel to the gross trend of the formation, dueto a map-scale F 5 fold. S 1 is defined by flattened clasts, andtrends around a fold plunging steeply to 114 (Fig. 11 b).Changes in bedding orientation delineate 100 m scale F 2folds. Well-developed S 2 axial plane foliation is strongly re-oriented by later folds plunging steeply to 110 (Fig. 11 c).F3 folds of S 0, S1, and S 2 plunge moderately to steeply east

(Fig. 11 d ), coaxial with F 2 folds (Fig. 11 e). A weak axialplane S 3 cleavage trends broadly northwest (Fig. 11 f ).Foliations that trend eastwest and overprint S 3 are assignedto S4 (Fig. 11 g). Steeply plunging F 5 folds (Fig. 11 h), with awell-developed north-northeast-trending axial planecrenulation cleavage (Fig. 11 i) overprint all earlierfoliations.

Sixth-generation structuresF6 structures occur only in the southern part of Elbow

Lake (domain D) and the southern extremity of the map areaalong the Berry Creek shear zone (domains J and N). Direct

correlation is impossible and respective structures are thusdesignated F 6a (domain D) and F 6b (domains J and N). In thesouthern part of domain D, F 6a S folds occur within the S 5ELSZ and plunge at moderate to shallow angles toward 035(Fig. 9 e). An axial plane crenulation cleavage (S 6a) dipssteeply northwest (Fig. 9 f ), slightly oblique to S 5.

In domain J, F 6b structures comprise a set of isoclinal

folds that overprint S 2 tectonic layering that was rotated intoan eastwest orientation on the south limb of the large F 4bfold (Fig. 3). The F 6b folds vary in scale from 1 to 50 m,have predominantly S asymmetry (Fig. 10 d ), and plungesteeply east-northeast (Fig. 9 g). Strong S 6b axial planecrenulation cleavage (best developed in the chlorite-rich lay-ers) trends 080 (Fig. 9 h), parallel to the Berry Creek shearzone (Fig. 3). F 6b folds, which can be confused with F 4bfolds, are not overprinted by S 5, in contrast to pre-S 5 struc-tures. We interpret F 6b folds as a discrete set of structures,associated with the reactivation of the Berry Creek shearzone by sinistral shear.

In domain N, 10100 m scale, S-asymmetric F 6b folds oc-cur in S 2 mafic tectonic layering. S and Z folds both occur at

1 m scale. F 6b folds plunge steeply east (Fig. 9 i), parallel tothe macroscopic fold determined from the S 2 orientationdata (Fig. 5 j). The S 6b axial plane fabric trends 080(Fig. 9 j), parallel to the Berry Creek shear zone. The degreeto which S 2 is reoriented eastwest in domain N increasessouthward toward the Berry Creek shear zone, and the S 6bfabric changes in character southward. In the northern partof domain N, where amphiboles have altered to chlorite, S 6bgenerally forms a penetrative crenulation cleavage. In thesouth, where the mafic rocks are composed of up to 80%amphibole, S 6b is defined by millimetre-scale chevron folds,with no penetrative cleavage.

The timing of F 6 structures is poorly constrained. Theypostdate F 5 structures, and predate brittleductile deforma-tion (Table 1). Retrograde chlorite associated with S 6 fabricsindicates that S 6 developed under metamorphic conditionssimilar to those under F 5, and hence shortly thereafter.

D5 deformation

Seventh-generation structuresA variety of late brittleductile and brittle features are

grouped as structures of the seventh generation; however,they may represent multiple generations of structures devel-oped during a protracted D 5 episode. Within the multiply re-activated corridor in the southern Elbow Lake area (Fig. 2),a northwest-trending, 1 km or more, broad zone of dextral,brittleductile shear bands overprints the ELSZ and the S 1

S2S 3 composite foliation. The shear bands are manifest asnarrow ductile zones in chloritic rocks, and narrow discretefaults in more competent rocks. They are vertical and trend045 (Fig. 9 k ), with an average obliquity to S 5 of about 25.Apparent dextral displacements greater than 1 m occuracross zones of less than 30 cm width (Fig. 10 e), and thesum of their displacements may be large. The shear bandsdo not penetrate the Elbow Lake tonalite, rather they mergeinto the ELSZ at the eastern margin of the tonalite. It is pos-sible that the zone of shear bands caused the apparentdextral step of the ELSZ in this region; however, no discretedisplacement zone has been mapped. Alternatively, an un-

1999 NRC Canada



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recognized brittle fault, with a large dip-slip component,may have offset the ELSZ to its present configuration. Lack of outcrop makes this conjectural.

The Grass River fault (Syme 1991) is a discrete brecciatedzone forming a prominent lineament from the west side of Elbow Lake, southward through Grass River to Third Cran-berry Lake, then southwestward along the western shore of

Second Cranberry Lake. It is narrow (12 m wide) for mostof its exposed length, but is 200 m wide in the segment be-tween Elbow Lake and Third Cranberry Lake (Fig. 2). It hasapproximately equal proportions of clasts and carbonate ma-trix, and clasts are typically lensoid in cross section(Fig. 10 f ) with a long vertical dimension. Much of the ma-trix carbonate is coarse grained and unstrained; however,

1999 NRC Canada


Fig. 10. (a ) S5 shear fabric in a granitoid, characterized by a finely spaced anastomosing domainal cleavage. Faint sinistral C shearbands are oblique counterclockwise to S 5. Pencil, 14 cm long, points toward 200. ( b) Western margin of the ELSZ on Elbow Lake(looking toward 020) showing a narrow strain gradient from pillow basalts to chloritic phyllonite is about 10 cm wide. Pencil (rightof arrow), 14 cm long, lies along the contact zone. ( c) Vertically plunging F 5 S folds, generated where the ELSZ overprints S 4 of theClaw Bay shear zone north of Leaping Moose Island (Fig. 2). Pencil, 14 cm long, is parallel to the axial plane and the marker pen liesparallel to long limbs of F 5 folds. ( d ) Tight to isoclinal F 6b S folds overprint S 2 tectonic layering, which was deflected into an eastwest orientation during F 4b deformation south of Iskwasum Lake (see text). Hammer, 35 cm long, points towards 085. ( e) Late dextral

brittleductile shear bands across an S 1S 2S 3 composite foliation in a mixture of highly strained gabbro, basalt, and tonalite. Hammerhandle, 35 cm long, points toward 045. ( f ) Brecciated basalt with abundant carbonate matrix in the Grass River fault, southwest of Elbow Lake. The 14 cm long pencil (at arrow) points 010, parallel to the fault. The mylonitic band in carbonate (top right of photo)illustrates the brittleductile nature of this deformation. ( g) Closer view of the breccia clasts in ( f ) shows subhorizontal carbonateveins, indicating vertical extension during brittleductile conditions.

N = 75 N = 21 N = 97

F = 114/713F = 108/823

N = 13 N = 14N = 14

N = 11 N = 81N = 8

Poles to Sdomain G

4 F axesdomain G

5 Poles to Sdomain G

5

Poles to Sdomain G

3F axesdomain G

2F axesdomain G

3

Poles to S

domain G2Poles to S

domain G0 Poles to S

domain G1

N

NN N

NN

NNN

a) b) c)

d) e) f)

g) h) i)

Fig. 11. Equal-area lower hemisphere projections of all structures in domain G, where N is the number of data points.


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narrow mylonite zones do occur locally in carbonate-richrock (Fig. 10 f ), illustrating the brittleductile nature of de-formation in the Grass River fault. A chlorite lineation onthe surface of the clasts appears to be due to the intersectionof the preexisting cleavage in the clast with the surface, andis not a stretching lineation. The long dimension in the brec-cia clasts is probably more a function of anastomosing shear

surfaces in map view than significant ductile vertical stretch.Subhorizontal carbonate-filled fractures that cut across thebreccia clasts (Fig. 10 g) indicate that vertical extension con-tinued to play a role during the later stage deformation. Lo-cal slivers of preexisting chloritic phyllonite between thebreccia and the undeformed wall rocks are similar in charac-ter to rocks in the ELSZ. In our opinion, the Grass Riverfault reactivated an S 5 shear zone, which was possibly anoffset portion of the ELSZ. This might explain the greaterwidth of the Grass River fault in that region.

Conjugate brittleductile kink bands locally overprint theELSZ for most of its length. Brittle faults trending north,north-northeast, and north-northwest, which form prominentlineaments on aerial photographs, are the youngest struc-

tures. Large faults are rarely exposed, whereas associatedminor faults typically form narrow (0.515 cm) cataclasticzones. Movement direction and age of the faults are notknown.

Hunt and Roddick (1992) presented KAr and 40Ar/ 39Armineral ages for a suite of hornblende and biotite separatesfrom granitoid plutons at Elbow Lake. These include a KArhornblende age of 1757 21 Ma; KAr biotite ages of 1768 22, 1763 14, 1758 17 Ma; and an 40Ar/ 39Ar (to-tal fusion) biotite age of 1766 12 Ma (Hunt and Roddick 1992). This cluster of ages at ~1.760 Ga could reflect ametamorphic pulse (M 3), which reset the K and Ar systemat-ics; however, no textural evidence supports an M 3 episode.More likely, the ~1.760 Ga ages represent the time at whichthese minerals cooled from peak conditions of M 2, throughtheir blocking temperatures, during regional exhumation.The temperatures at which intracrystalline diffusion is im-peded in minerals such as biotite may broadly coincide withthe onset of brittleductile conditions. Fedorowich et al.(1995) determined an 40Ar/ 39Ar age of 1691 6 Ma frompristine potassic feldspar within a vein associated with a latefault in the Flin Flon. They interpreted the vein as havingbeen emplaced at a temperature close to the blocking tem-perature of Ar in feldspar, and thus approximating the age of movement along the fault. This fault represents the youngestdeformation so far recognized in the Flin Flon Belt.

Despite discontinuous outcrop, extensive plutonism, mul-tiple reactivation of high-strain zones, and a lack of goodmarker units of well-layered rocks, the fabric developmenthistory of the eastern Amisk collage, spanning more than180 Ma, has been deciphered. Deformation in the ductileflow regime initiated prior to 1868 12

21+ Ma, and continued af-

ter ~1.800 Ga. Brittleductile deformation probably initiatedafter 1.760 Ga, and brittle deformation may have continueduntil about 1.690 Ga. We present a tectonic model in Fig. 12illustrating the ductile portion of the deformation history,making reasonable assumptions about the tectonic shorten-

ing direction for generations of structures. Because kine-matic information is not preserved along S 1 structures, theF1 shortening direction is unknown. Although the S 2 struc-tural grain in the eastern Amisk collage trends north-northeast, it is northsouth in portions of the Flin Flon Beltwith fewer plutons, indicating significant eastwest shorten-ing of the Amisk collage during D 2. The shortening direction

within the Flin Flon Belt during F 3 and F4 (D3) is not wellconstrained, because deformation was restricted to shearzones without development of regional foliations. We tenta-tively correlate F 3 and F4 in the Flin Flon Belt with F 1 andF2 in the Kisseynew Domain (Zwanzig and Schledewitz1992), based on the relative timing of deformation and M 2metamorphism. F 1 and F 2 structures there are believed tohave developed during south-southwest shortening (earlyHudsonian deformation) between the Flin Flon Belt and theKisseynew Domain (Zwanzig and Schledewitz 1992;Ansdell et al. 1995; Norman et al. 1995; Connors 1996),placing metaturbidites of the Kisseynew Domain at a higherstructural level than the Flin Flon Belt. The north-northeasttrend of the S 5 regional foliation indicates significant west-

northwest shortening of the Flin Flon Belt during lateHudsonian (D 4) deformation. The F 5 block (Fig. 12) illus-trates how S 5 intensified preexisting upright fabrics, andhow the upright F 5 folds affected the shallow eastern bound-ary between the Flin Flon Belt and the Kisseynew Domain.The tectonic shortening direction during F 6 is not well con-strained. Regionally, F 6 folds may correlate with F 4 folds inthe southern Kisseynew Domain (Kraus and Williams 1998;Connors 1996), which trend eastwest, and indicate northsouth shortening during late D 4.

The change in tectonic shortening direction from eastwest during D 2, to south-southwest during F 3 and F4, towest-northwest during F 5, delineate pre-Hudsonian deforma-tion, and early and late Hudsonian orogeny sensu stricto.The change in tectonic regime marked by the late Hudsoniandeformation likely represents the influence of the Superiorcraton colliding from the southeast (cf. Connors 1996). Post-Hudsonian deformation under brittleductile and brittle con-ditions probably represents plate movement readjustmentsduring final convergence of the Superior and Hearne cratons.

The fabric and tectono-metamorphic history of theKisseynew Domain differs significantly from that of the FlinFlon Belt in that (1) layering and tectono-stratigraphic unitswere shallowly north dipping during regional metamorphism(Zwanzig and Schledewitz 1992; Norman et al. 1995;Connors 1996); (2) the thermal gradient was relatively high(high temperature, low to moderate pressure) during regionalmetamorphism (Gordon 1989; Norman et al. 1995; Krausand Menard 1997). Based on high thermal gradients and thelack of evidence for upright structures prior to metamor-phism in some Proterozoic belts, like the Kisseynew Do-main, Norman et al. (1995) questioned whether modern-dayplate tectonic models apply to the Proterozoic. In the FlinFlon Belt, however, layering was steep during regional meta-morphism (Ryan and Williams 1996 b), which appears tohave peaked at moderate pressure and temperature (Digeland Gordon 1993). Apparently abnormal tectono-metamorphic conditions documented in the Kisseynew Do-main are best considered in terms of the local tectonic set-ting, rather than general tectonic processes.

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Fig. 12. Schematic block diagrams depicting the F 1 to F5 structural evolution of the eastern Amisk collage. F 3 marks the onset of south-southwest-directed overthrusting of the Flin Flon Belt (FFB) by the Kisseynew Domain (KD), and the Saskatchewan craton bythe FFB. The Iskwasum Lake shear zone (ILSZ) was active at that time. Deformation within the FFB was along steep structures. F 5sinistral transpression of the belt, which developed the Elbow Lake shear zone (ELSZ) and a regional crenulation cleavage, had asignificant effect on the map pattern of the Snow Lake segment, and its shallow boundary with the KD. Abbreviations: ocean floor(OF).


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The map distribution of the McDougalls Point formationhas implications for the amount of displacement on shearzones of different generations. This formation occurs on thewest side of the ELSZ at Elbow Lake, and on its east side inthe Cranberry Lakes region (Fig. 2). It occurs predominantlyon the west side of the ILSZ, except for a thin sliver that lies just to the east. The sliver is actually separated from the

Claw Bay formation by a narrow unexposed structure, whichmay be interpreted as part of an S 1 shear zone. We concludethat the McDougalls Point formation was juxtaposed withthe Claw Bay formation, and possibly the maficultramaficcomplex was juxtaposed with the Claw Bay formation, alongS1 shear zones that were partially reactivated as the ILSZduring regional metamorphism. The direction from whichthe maficultramafic complex was structurally emplaced re-mains unresolved, although a definitive link to one of the ba-salt formations through isotopic study might elucidate thisquestion. The ELSZ reactivated part of the ILSZ in the El-bow Lake area, whereas a new shear zone developed throughthe Cranberry lakes region. The S 5 shear zones are impres-sive in extent; however, the small amount of offset of the

McDougalls Point formation indicates that they do not havelarge displacement.

The origin of F 6a folds in southern domain D, the onlyshallowly plunging late folds in the area, is enigmatic. Theymay represent post-S 5, east-side-up rotation of the shortlimbs of F 6a folds during reactivation of the ELSZ, consis-tent with higher grade rocks occurring on the east side of theELSZ. In such a deformation, however, the stretchinglineation on the limbs would remain steep, even though thefold axes are shallow. This is not the case. The stretchinglineation is parallel to the F 6a fold axes, and at some loca-tions, a transition from steep, S-asymmetric F 5 folds to shal-low F 6a folds occurs in the same outcrop. We interpret theF6a folds as having developed in a steep orientation, and ro-tated into a shallow orientation with progressive deforma-tion. The ELSZ narrows significantly in the area between theElbow Lake tonalite to the east, the Big Rat Lake pluton tothe west, and a large sheet of tonalite between the two(Fig. 2). The deformation path in this segment of the shearzone would have changed significantly due to the alteredboundary conditions imposed by the plutons, because a simi-lar amount of deformation had to be accommodated in thenarrow zone compared with where it is over 2 km wide incentral Elbow Lake. Variation in the orientation of linearfeatures (F 6a axes and L 5 lineations) probably reflects heter-ogeneity in the local extension direction and local shearstrain rate in the abruptly narrowed segment.

Within the ELSZ, the combination of vertical stretchinglineations, sinistral transcurrent, and east-side-up dip-slip ki-nematic indicators inicates deformation with triclinic sym-metry, which is best resolved in terms of a sinistral obliquetranspressive shear model (Ryan and Williams 1994; Lin etal. 1998). The displacement vector in the ELSZ, which isnot tracked by any strain features (e.g., lineations), mustplunge shallowly south. The transpression episode probablypredates formation of the shallow F 6a folds.

The sharp southward increase in metamorphic grade in thesouthern First Cranberry Lake area can be explained by acomponent of south-side-up late-stage dip-slip movement.Leclair et al. (1997) determined an increase in grade to up-

per amphibolite facies only a few kilometres to the south,and concluded that the Berry Creek shear zone experiencedlate-stage wrench faulting. However, the offset of metamor-phic isograds is not an indicator of shear sense, because onlythe displacement of a line across a structure will provide ashear sense.

(1) The steep, north-northeast fabric of the Flin Flon Belt,which contrasts that of the adjacent Kisseynew Domain, de-veloped primarily during F 2 upright folding of accreted as-semblages and earlier shear zones between 1.864 and 1.845Ga, predating the Hudsonian orogeny sensu stricto. UprightF5 folding intensified the fabric, and locally altered its north-northeast trend.

(2) Vertical extension was important in post-D 1 deforma-tion, even in late-stage structures such as the Grass Riverfault. Postorogenic, low-angle extensional features that arecommon to mountain belts like the North American Cordil-lera (Wernicke 1981) and the Pyrenees (Vissers 1992) arenot prevalent in the Flin Flon Belt, possibly indicating thaterosion was the dominant unroofing mechanism.

(3) The F 5 ELSZ transpression episode reactivated shearzones that were in a favorable orientation (e.g., southern El-bow Lake), and macroscopically folded those that were not(e.g., at Iskwasum Lake). The ELSZ has triclinic symmetry;its stretching lineation is at a high angle to the sinistral sheardirection, which had an east-side-up component.

(4) Maximum displacements between tectono-stratigraphicassemblages occurred on early shear zones. The map distri-bution of the maficultramafic complex probably indicatesthat it was emplaced along an S 1 shear zone, which was re-activated during F 3 deformation.

(5) The Berry Creek shear zone strongly deflected thenorth-northeast structural grain into an eastwest orientationduring F 4 dextral shear. Sinistral reactivation of the BerryCreek shear zone developed F 6b S folds along its margin.Brittle faults further reactivated the zone.

(6) Seven generations of structures in the eastern Amisk collage indicate marked changes in tectonic shortening di-rections across the Flin Flon Belt during a structural evolu-tion exceeding 180 Ma. These changes are consistent withthose of other domains in the southeastern Trans-HudsonOrogen, and can be divided into pre-, early, late, and post-Hudsonian deformation.

Capable field assistance was provided by SandraMacDougall, Natasha Connell, and Scott Gilliss. Previousmapping, field trips, and input by Ric Syme have contributedgreatly to this work. Logistical support was provided by theGeological Survey of Canada. Manitoba Energy and Minesis thanked for further contributions to logistics, and many of its staff members are thanked for geological discussions.Steve Lucas initiated the project and has provided sound, en-thusiastic discussion throughout. Earlier reviews by PeterStringer and journal reviews by Peter Cawood and WouterBleeker significantly improved the manuscript. Dave Pirie,Ancel Murphy, and Calvin Nash are thanked for their expe-

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dient, high-quality thin section preparation. Bob McCullochprinted the photographs. Funding was mainly from a NaturalSciences and Engineering Research Council of Canada re-search grant and a one year Lithoprobe grant to P.F.W.

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structural evolution of amisk collage - james ryan

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