UBC Theses and Dissertations

UBC Theses Logo

UBC Theses and Dissertations

Structure and metamorphism at the western margin of the Omineca belt near Boss mountain, east central… Fillipone, Jeffrey Alan 1985

Your browser doesn't seem to have a PDF viewer, please download the PDF to view this item.

Item Metadata


831-UBC_1985_A6_7 F54.pdf [ 14.13MB ]
JSON: 831-1.0052798.json
JSON-LD: 831-1.0052798-ld.json
RDF/XML (Pretty): 831-1.0052798-rdf.xml
RDF/JSON: 831-1.0052798-rdf.json
Turtle: 831-1.0052798-turtle.txt
N-Triples: 831-1.0052798-rdf-ntriples.txt
Original Record: 831-1.0052798-source.json
Full Text

Full Text

STRUCTURE AND METAMORPHISM AT THE WESTERN MARGIN OF THE OMINECA BELT NEAR BOSS MOUNTAIN, EAST CENTRAL BRITISH COLUMBIA by JEFFREY A. FILLIPONE B.S., STATE UNIVERSITY OF NEW YORK AT ALBANY, 1982 A THESIS SUBMITTED IN PARTIAL FULFILMENT OF THE REQUIREMENTS FOR THE DEGREE OF MASTER OF SCIENCE in THE FACULTY OF GRADUATE STUDIES DEPARTMENT OF GEOLOGICAL SCIENCES We accept this thesis as conforming to the required standard THE UNIVERSITY OF BRITISH COLUMBIA JULY, 1985 ® JEFFREY A. FILLIPONE, 1985 In presenting this thesis in partial fulfilment of the requirements for an advanced degree at the The University of British Columbia, I agree that the Library shall make it freely available for reference and study. I further agree that permission for extensive copying of this thesis for scholarly purposes may be granted by the Head of my Department or by his or her representatives. It is understood that copying or publication of this thesis for financial gain shall not be allowed without my written permission. GEOLOGICAL SCIENCES The University of British Columbia 2075 Wesbrook Place Vancouver, Canada V6T 1W5 Date: JULY. 1985 Abstract Rocks of the Hadrynian and Early Paleozoic (?) Snowshoe Group comprise the core of the Boss Mountain area at the western margin of the Omineca Belt near Crooked Lake. Structurally overlying these are rocks of the Intermontane Belt: the Permian Slide Mountain Group (Antler Formation), Triassic fine grained sediments (unnamed), and Jurassic volcanic rocks (Takla Group). In the Snowshoe Group, a large, lensoid intrusion of coarse grained granitic rock (Boss Mountain gneiss) was emplaced during the mid-Paleozoic, and later deformed and metamorphosed with the enclosing metasediments. The rocks of the Snowshoe Group act as basement to the overlying Late Paleozoic/Early Mesozoic cover rocks. Within the basement, four phases of regionally significant deformation have been recognized, and are manifest as fold generations designated F l through F4. Earliest structures, F l , in the Snowshoe Group are isoclinal folds, accompanied by a transposed foliation of regional extent, which are overprinted by penetrative deformation related to easterly verging F2 nappe structures. The F3 folds are upright or inclined to the northeast, and give a consistent southwesterly sense of vergence. These folds are responsible for the regional map pattern, and have folded both the basement and cover into an antiformal culmination in the Boss Mountain area. Fourth phase structures refold the other features, but do not appreciably affect the F3 geometry. In the cover sequences, the first phase of deformation is equivalent to the second phase within the basement During the Phase 2 deformational episode the cover rocks were emplaced over rocks of the Snowshoe Group. West-dipping imbricate faults characterize the western margin of the area, where basement rocks contain fault-bounded slivers of the cover, and the tectonic contact between basement and cover rocks is marked by a zone of mylonitization. Similarly, the F2 and F3 folding ii phases in the cover are equivalent to the F3 and F4 structures in the basement, respectively, but are only weakly developed in the cover. An early, enigmatic metamorphic event accompanied Phase 1 deformation in rocks of the Snowshoe Group. Field relations suggest that this was probably coeval with the mid-Paleozoic emplacement of the Boss Mountain gneiss. Metamorphism during the Jurassic was synchronous with F2 deformation in rocks of the Snowshoe Group, and resulted in Barrovian type mineral assemblages ranging from the biotite through sillimanite zones. The metamorphic grade increases from west to east; with only low grade metamorphism of the cover rocks in the study area. Phase 2 structures in the Snowshoe Group were overprinted by the peak of this metamorphic event, as indicated by staurolite through sillimanite zone assemblages. The Boss Mountain area is structurally correlative with rocks of the Shuswap Complex. These rocks appear to comprise a portion of the continental margin sedimentary wedge, which was overridden by an allochthonous terrane accreted to the western margin of North America in post-Early Jurassic times. iii A C K N O W L E D G E M E N T S Field and laboratory support were provided by NSERC grant A67-2134 to J. V. Ross, whose interest and enthusiasm were responsible for the success of this project. Discussions with fellow students have led to an improvement in my thinking and perception during the course of this work. John Montgomery, Darren Elsby, Clark Isachsen, Ian Moffat, Dave Parkinson and Jim Logan have influenced ideas developed during this study. I would like to thank Drs. H. J. Greenwood and R. L. Armstrong for critically reviewing the thesis manuscript, and offering useful suggestions at various stages in its development. J. K. Mortensen contributed his time and enthusiasm in dating samples from the area, and discussing the more obscure aspects of terrane philosophy. Special thanks to J. K. Russell for carefully editing part of the preliminary manuscript J. Knight assisted in the collection of analytical data for the geothermometry study. Steve Garwin enthusiastically lent his expertise in metamorphic petrology and geothermometry, and developed a computer program for processing temperature data. During the 1983 field season, I was privileged to have the assistance of Mary Anne Bloodgood, whose support and continuing encouragement are deeply appreciated. Any errors or flaws in logic remain the sole responsibility of the author. iv Frontispiece View north from the Boss Mountain area, with Crooked Lake and the Interior Plateau distance. Table of Contents I. Introduction 1 II. Stratigraphy 7 1. Eastern succession 10 2. ML Beisig succession 13 3. Boss Mountain succession 17 4. Western succession 20 5. Permian Slide Mountain Group 23 6. Triassic phyllite 24 7. Stratigraphic correlations - discussion 25 III. Structure - Description of Phases 27 1. Basement - Snowshoe Group 27 2. Cover - Slide Mountain Group and Triassic phyllite 52 3. Discussion 55 IV. Metamorphism and textural description 57 1. M l - Early Metamorphism 57 2. M2 - Jurassic Metamorphism 58 3. M2 Metamorphic reactions - metapelites 70 4. Calc-silicate reactions 81 5. Discussion 85 V. Structural synthesis 92 1. Basement - Phase 1 92 2. Basement - Phase 2 / Cover - Phase 1 94 3. Basement - Phase 3 110 VI. Tectonic interpretation - discussion 120 REFERENCES 130 APPENDIX A - METAMORPHIC MINERALS 137 APPENDIX B - ELECTRON MICROPROBE ANALYSES 140 vi DR=0 LIST OF FIGURES F i g u r e Page 1. T e c t o n i c map o f t h e Canadian C o r d i l l e r a 2 2. G e o l o g i c map o f t h e Crooked Lake a r e a 4 3. R e g i o n a l g e o l o g y o f t h e w e s t e r n m a r g i n o f t h e Omineca B e l t near Quesnel Lake 5 4. S t r u c t u r a l s u c c e s s i o n - Snowshoe Group 8 5. S t r u c t u r a l domains i n t h e Boss Mountain a r e a 28 6. F l i n t r a f o l i a l f o l d s - basement 30 7. M e s o s c o p i c F l f o l d - basement 31 8. F2 f o l d s i n t h e Snowshoe Group and t ' v s . o<. t h i c k n e s s v a r i a t i o n p l o t 33 9. E q u a l a r e a p r o j e c t i o n o f F2 a x i a l p l a n e s 37 10. F2 f o l d s i n g r a n i t i c p e g m a t i t e 39 11. Phase 2 f a u l t s l i n e d by g r a p h i t i c p h y l l i t e 40 12. C r o s s - s e c t i o n o f phase 2 f a u l t s i n t h e w e s t e r n p a r t o f t h e map a r e a 42 13. F3 f o l d d a t a - basement, and m e s o s c o p i c F3 f o l d s 46 14. F4 f o l d d a t a - basement, and m e s o s c o p i c F4 f o l d s 50 15. F l f o l d d a t a - c o v e r 53 16. C r o s s - s e c t i o n o f t h e I n t e r m o n t a n e - Omineca B e l t boundary i n t h e Boss M o u n t a i n a r e a 56 17. M2 metamorphic zones 59 18. SI f o l i a t i o n 61 19. Garnet i n p e l i t i c s c h i s t 61 20. S y n - k i n e m a t i c F2 g a r n e t i n p e l i t i c s c h i s t 63 21. L a t e s y n - k i n e m a t i c F2 s t a u r o l i t e 63 22. P o s t - k i n e m a t i c F2 k y a n i t e 64 23. F i b r o l i t e c r e n u l a t e d by F3 65 24. Syn t o p o s t - k i n e m a t i c F3 g a r n e t and mica 67 25. Metamorphic m i n e r a l / d e f o r m a t i o n a l t i m i n g c h a r t 69 26. M2 g a r n e t 72 27. M2 s t a u r o l i t e 72 28. M2 k y a n i t e 75 29. M2 s i l l i m a n i t e 76 30. Garnet - b i o t i t e geothermometry 79 31. C a l c - s i l i c a t e - k y a n i t e zone 82 32. S c a p o l i t e i n s i l i c e o u s m a r b l e 83 33. M i n e r a l r e a c t i o n s , P-T r e l a t i o n s , and b a t h o z o n e s 86 33. AFM diagrams o f m i n e r a l assemblages i n p e l i t i c s c h i s t s 87 34. F l h y d r a u l i c f r a c t u r e s 91 35. F l - F2 s u p e r p o s i t i o n o f f o l d s 93 36. Deformed F2 augen - basement 95 37. I n c l u s i o n s i n metamorphic p o r p h y r o b l a s t s 98 38. Deformed i n t e r n a l f a b r i c i n g a r n e t p o r p h y r o b l a s t s 99 39. F2 c r e n u l a t i o n c l e a v a g e f o r m a t i o n 100 40. K i n e m a t i c model f o r F2 c l e a v a g e f o r m a t i o n 102 41. P r e s s u r e shadows i n m e t a - v o l c a n i c s c h i s t - A n t l e r F o r m a t i o n 104 42. Phase 2 f a u l t d a t a i n g r a p h i t i c p h y l l i t e 107 43. Phase 2 d u c t i l e shear zones i n q u a r t z o f e l d s p a t h i c and 108 Boss Mountain G n e i s s 109 v i i . 44. F3 s l ip directions - basement 111 45. F2 - F3 fold interference c lass i f icat ion 113 46. Interference patterns - basement 114 47. Pulled apart garnet crenulated by F3 119 48. Fl and F2 quartz segregations in Antler Formation meta-volcanics 124 49. , Plate tectonic model for the evolution of the Crooked Lake area 128 50. Geologic map of the Boss Mountain area separate 51. Structural map of the Boss Mountain area ^ SmCc/*"^ \ s e P a r a t e 52. Structural cross-section : X - X' \\ \ ' ) separate 53. Structural cross-section : Y - Y' C^Uch^ < s e p a r a t e 54. Generalized cross-section through the Boss / Mountain area : A - B separate v i i i . LIST OF TABLES Page Table 1 M2 metamorphic mineral assemblages i n the Snowshoe Group 71 Table 2 Garnet - b i o t i t e geothermometry temperatures 90 Table 3 Interference pattern c l a s s i f i c a t i o n r e s u l t s 117 Table 4 Metamorphic mineral abbreviations 137 Table 5 Metamorphic mineral formulae 138 Table 6 M2 plagioclase compositions 139 Table 7 Microprobe standards 141 Table 8 Sample locations 142 Table 9 Garnet - b i o t i t e geothermometry r e s u l t s 143 i x . I. INTRODUCTION The Canadian Cordillera is divisible into five major tectonic belts (Fig. 1) consisting of distinct assemblages or terranes of different affinities (Wheeler and Gabrielse 1972; Monger and Price 1979). Extensive deformation and metamorphism of Jurassic age in the Omineca Belt, were followed by detachment and eastward transport of the sedimentary cover in the Rocky Mountain Belt. Terranes, at one time outboard of the western margin of North America, were thrust eastwards over the continental margin during the Jurassic. Basement rocks related to North America - the Omineca Belt, were subjected to deformation initiated by the accretion of predominantly volcanic and sedimentary sequences (Brown 1981). Rocks of the Quesnellia Terrane comprising Upper Paleozoic (?) and Mesozoic eugeoclinal units, were emplaced, causing the formation of extensive mylonite zones and folding within both the accreted packages and rocks related to North America (Monger et al. 1982). Location The Boss Mountain area is located along the south shore of Crooked Lake, and extends about 5 km south of Boss Mountain (Fig. 2). The eastern boundary is the Clearwater River Valley, and in the west the area is bordered by Bosk Lake. Boss Mountain is a deeply dissected series of peaks, which physiographically lie at the western margin of the Quesnel Highlands. The area is a western continuation of the Cariboo Mountains south of Quesnel Lake, B.C.. Access is fairly good on the numerous improved dirt roads eastwards from Williams Lake, B.C.. The two main routes that exist are from 100 Mile House via Hendrix Lake, or from 150 Mile House via Black Creek and Miocene. In terms of regional geology, the area comprises high grade metasedimentary rocks and deformed granitic intrusions. The sediments are correlated with the Hadrynian to Early Paleozoic(?) Snowshoe Group of the Omineca Belt In the western part of the area, sediments and volcanics of Permian(?) and Early Mesozoic age represent units 1 Figure 1 - Generalized tectonic map of the Canadian Cordillera showing the major structural divisions and location of the study area, (after Wheeler and Gabrielse 1972). 3 of the Intermontane Belt or Quesnellia Terrane (Monger et al. 1982) in tectonic contact with the underlying Proterozoic and younger rocks (Fig. 2). During the 1983 field season an area of approximately 60 km 2 was mapped, from the Intermontane/Omineca Belt boundary, to east of Boss Mountain. Summary of Previous Work Figure 3 is a generalized geologic map of the Crooked Lake area, showing the distribution of major rock types in relation to the Boss Mountain area. A number of workers have contributed to the stratigraphic terminology applied to the widely distributed rocks of Proterozoic and Early Paleozoic age in the western Cariboo Mountains. The rocks east of the Intermontane/Omineca Belt contact have been assigned to the Snowshoe Formation (Cairnes 1939; Holland 1954; Sutherland Brown 1963). Holland (1954) described the Snowshoe Formation as an upper unit within the Precambrian Cariboo Group. Sutherland Brown (1963) interprets rocks of the Snowshoe Formation, as being the uppermost sequence of Hadrynian metasediments seen below the Late Paleozoic Slide Mountain Group. This interpretation is consistent with the succession observed at Boss Mountain in terms of structural position. Studies conducted by the Geological Survey of Canada have also attempted to make regional stratigraphic correlations of these rocks (Campbell et al. 1973; Struik 1981, 1982, 1984). Campbell and others (1973) view is that the sequences of rocks in the Cariboos from Wells, B.C. to south of Crooked Lake, are a western finer grained facies of the "gritty" units of the Kaza Group to the east Campbell's study was carried out as part of a reconnaissance project conducted from the early 1960's to about 1970. A number of unpublished theses have also attempted in part, to clarify the structural and stratigraphic divisions in areas nearby Boss Mountain. Campbell (1971), and Campbell and Campbell (1970), discuss the structure and local geology in the Crooked Lake area. However, the tectonic framework and polydeformed nature of these rocks remained unclear. 4 EARLY JURASSIC LATE TRIASSIC LATE PALEOZOIC MIDDLE t.-_--_-_-PALEOZOIC SLIDE MTN. GROUP ANTLER FM. Matavolcanie phyMta and aoMat Qranttlo gnalaa Shaar zona . . . , Fold axial traca Antlfonn Synform PROTEROZOIC E. PALEOZOIC SNOWSHOE GROUP QuartzofaMapathte aehlat and gnalaa, quartztta Figure 2 - Geologic map of the Crooked Lake area, showing the generalized structure and lithologic units defined in the area (modified from Campbell 1978). Inset shows the location of the map-area. Tb - Tertiary basalt TB - Takomkane batholith RB - Raft batholith Kg -Cretaceous granite Jg - Jurassic granite Js - Jurassic slate Tr/Jr - Triassic & Jurassic volcanics, includes Takla Group UTr - Triassic sediments and volcanics SM - Slide Mtn. Group Pec - Cache Creek Group QLG - Quesnel Lake Gneiss H-Pz - Hadrynian and Paleozoic metasediments HK - Kaza Group HS - Snowshoe Group 123°Oo' 63* 3 120*00' -T -63 #30' William* Lake -f 52*00' N ' 123* 00 Figure 3 - Simplified regional geologic map of the Intermontane/Omineca Belt boundary £% w T ^ ^ distTihution o f Hadrynian and possible Early Paleozoic rocks , 5 otTtml (SsSC MaSpTo5A)G r O U P' " * ( H K ) • G l 0 U P - M ° d i f i e d ^ T J P p e r « « in is 6 Struik (1981, 1982) has demonstrated that the Snowshoe Formation as described in the main ranges of the Cariboo Mountains does not readily correlate with that in the west as described by Holland (1954). The Cariboo Group in the two locations is dissimilar in many of its formations, and a major thrust may separate the two. This has led to a revision of the Snowshoe Formation, now renamed Snowshoe Group (Struik 1984). The purpose of this study is to present a more detailed stratigraphic and structural interpretation, and to elucidate the relationship of the structural/metamorphic succession to the tectonic evolution of this zone. Structural mapping and petrographic analysis form the basis for the following descriptions and their interpretation. II. STRATIGRAPHY Two distinct stratigraphic assemblages exist in the Boss Mountain area. Hadrynian to Early Paleozoic(?) schists and gneisses of the Snowshoe Group comprise basement to the overlying slate, phyllite, and volcanic rocks of the Permian to Early Jurassic cover sequence. Hadrynian to Early Paleozoic (?) Snowshoe Group Stratigraphic subdivision of the metasedimentary rocks in the Boss Mountain area is hampered due to intense metamorphism and deformation, which has homogenized contacts and obliterated primary sedimentary features. The method of stratigraphic analysis involved the correlation of similar rock types within each structural succession. Four distinct structural successions have been defined within the map area (Fig. 4). These are referred to as the Eastern (E), Beisig (B), Boss Mountain (BM) and Western (W) structural successions, from east to west respectively. The rock units in each of the successions appear to represent a sequence of variable but distinctive lithologies. Some of the units have been correlated between successions where structural control was good, or compositional similarities were sufficiently established. Boss Mountain gneiss Quartzofeldspathic gneiss underlies much of the central part of the map area (Fig. 4). The gneiss consists of quartz + K-feldspar + plagioclase + muscovite + biotite. Textural variants observed in the area include: granoblastic granitic gneiss, coarse quartz-feldspar augen gneiss and strongly foliated, finely laminated gneiss. The name Boss Mountain gneiss was applied by Campbell (1971), and has been used to describe the granitic, sill-like bodies present within Snowshoe Group metasediments in the map area. 7 12 0* 2 8' 62* 12'—|— 82* 10' Tr-Triassic phyllite SM-Slide Mountain Group BMQ-Boss Mountain Gneiss Hs-Snowshoe Group marbi* unit* Boundary between • structural successions *• 1 Km \^late fault \ Imbricate faults \ Figure 4 - a) Distribution of rock types in the Boss Mountain area, showing major lithologic units, b) The structural successions and relationships between individual units are indicated schematically for the four successions defined in the area. 9 Western Succession Mt. Beisig Succession Eastern Succession (4B) 10 1. EASTERN SUCCESSION This is the lowest structural level recognized in the area, and consists of feldspathic schists, micaceous quartzites and lesser pelitic schists. Local, thin calc-silicate horizons are present, but discontinuous. Coarse grained quartz veins and granitoid pegmatites are commonplace, with porphyroblasts of staurolite, kyanite, and sillimanite abundant in the more pelitic zones. Thin amphibolite and equigranular siliceous gneiss appear as members of limited extent. Upwards in the succession, injections of quartz-K-feldspar-muscovite-garnet pegmatite are abundant These intrusives locally crosscut compositional layering, and are commonly folded. Unit IA approximate thickness = 150 m At the base of the succession exists a series of quartzose metasediments, generally lacking distinctive carbonate units. They are, for the most part, coarse grained quartzites with very little pelitic component, often finely laminated with an average layer thickness of 0.5 - 3 mm. Individual layers are composed of equigranular to elongate quartz and feldspar, separated by thin (1 mm to nearly 1 m) pelitic layers of coarse biotite and muscovite. The quartz layers commonly consist of thin bands of fine, granular material'. Plagioclase, minor K-feldspar, and both biotite and muscovite comprise large grains within much of the unit, where coarse micas outline the compositional layering. Porpyroblasts of garnet are concentrated within the thin pelitic interlayers. Several coarse grained, greenish-white mottled calc-silicate horizons are present as thin bands less than 0.5 m thick at the base of the observed section. Calcite, pyroxene and epidote are evident in outcrops of this rock type. Generally, thickness of individual compositional layers increases upward through the section. Unit IB approximate thickness = 20 - 30 m The upper portions of Unit IA are marked by an increase in thick grey to tan quartzite in gradational contact with Unit IB. Unit IB is dominantly homogeneous 11 quartzite, and is distinguished from Unit IA by the lack of pelitic interlayers. Fine lamination planes are still often closely spaced within layers of a specific rock type, but compositional layering ranges from 2 - 20 cm for individual "beds." Carbonate horizons are not observed in this unit Unit IC approximate thickness = 2 - 5 m In sharp contact above the quartzite of Unit IB is a distinctive very coarse grained, well foliated biotite schist and amphibolite unit Coarse grains of dark brown biotite, minor muscovite, quartz, and iron oxides comprise this rock type. Garnet is rare. Dark brown to rusty weathering black schist averages 30 or 40 cm, up to 1 m in thickness, and is interlayered with strongly laminated amphibolite. Colour laminations in the amphibolite due to thin quartz/plagioclase layers, are common as 1 - 3 mm compositional bands. Interdigitation of the two units is typical, and contacts with adjacent units are sharp. Where the amphibolite is in contact with the underlying quartzites, the biotite schist occurs as a 1 - 2 m interlayer between thick layers (1 -3 m) of amphibolite. The rarity of amphibolite within the Snowshoe Group in the Boss Mountain area renders it as a prominent marker horizon. Unlike amphibolite units associated with marble in other portions of the area, Unit IC always occurs with biotite schist, often to the exclusion of any calc-silicate bearing lithologies. Unit 2 approximate thickness = 120 - 150 m Immediately overlying the schist and amphibolite marker Unit IC, is garnetiferous pelitic schist of Unit 2. These fine grained rocks locally contain thin quartzites 3 - 10 cm thick, and very abundant coarse grained quartz/feldspar segregations. The coarse grained veins are nearly parallel to the main compositional layering, but are occasionally more irregular and crosscut the foliation. They are often very knotted looking, and range in thickness from a few cm to 5 or 10 cm maximum, and exhibit variable stages of distortion by folding (see structure section Phase 1). The pelitic schist is very dark grey on fresh surfaces, brownish red 12 weathering, and frequently contains abundant granitic pegmatite (quartz + muscovite + perthite + garnet). Garnet is ubiquitous, and consists of rounded to strongly oblate porphyroblasts averaging from a few mm to commonly 5 mm in diameter. It is concentrated within pelitic interlayers (2 - 25 cm thick) which alternate with light grey quartzite layers. Unit 3 approximate thickness = 3 - 5 m In abrupt contact with Unit 2 is a thin, but very easily distinguished quartzofeldspathic gneiss. Very thin seams of pelitic schist are intercalated with the gneiss. The gneiss marks the general contact between the underlying Unit 2, and the next highest unit - Unit 4. Its thickness varies from 2 m to about 5 m. In outcrop, the gneiss is very resistant, light coloured, and light grey to white weathering. On fresh surfaces, it is usually medium grained, light grey granular quartz and plagioclase, which is thickly layered on the scale of tens of cm up to a metre. Both the upper and lower contacts are well exposed, and are parallel to the compositional layering. Unit 4 approximate thickness = 200 m maximum Near the top of the Eastern Succession is a very repetitively layered sequence. It varies upwards from medium grained quartzite and schist in thin 5 - 20 cm interlayers, with local amphibolite layers 10 - 15 cm thick, through a well laminated to more massive grey quartzite and lesser pelitic schist in the upper portions. At least 100 to 150 m of the dark grey, brown-red weathering, rhythmically layered material is present before any lateral breaks consisting of pelite with subordinate quartzite are encountered. Granitic pegmatite is also abundant within this unit, where it has been injected to considerable thickness; in places up to 35 or 40 m thick, roughly parallel to the compositional layering. Unit 5 approximate thickness = 50 m The top of the Eastern Succession, which extends to the southern foot of ML Beisig, is another mixed pelitic schist and quartzite sequence with occasional gneissic 13 intervals. Muscovite schist is a subordinate, but important member. Tan to brownish weathering pelitic schist, granular schist, and whitish quartzite comprise the remainder of the unit, and appear to be limited in extent when traced to the north. Thickness of individual layers is not a firm basis for recognizing these members as the internal makeup of the unit is quite variable. The white quartzite is easily separable, and occurs as layers 1 - 5 cm, within 1 - 2 m thick layers of pelitic schist Granitic pegmatite Large tabular and irregular bodies of granitic pegmatite are characteristic of the eastern succession. Outcrop relations with the country rock reveal the crosscutting nature of the pegmatite with respect to the compositional layering. Smaller dykes and sill-like bodies 1 - 2 m thickness are common. Chilled margins were not observed in any of the outcrops studied. Mineralogy is easily determined in hand specimen, and is, in decreasing order of abundance: K-feldspar (incl. perthite), quartz, muscovite, plagioclase, biotite and garnet Graphic intergrowth between K-feldspar and quartz is common in some locales, while other exposures reveal a more granular intergrowth of the two phases. Muscovite and biotite often occur together; biotite always as the subordinate phase. Muscovite is present mostly as large platy crystals 2 - 5 mm across, or as rare prismatic grains. Tiny red garnets up to 1 mm across are frequently present as an accessory phase. 2. MT. BEISIG SUCCESSION The basis for defining the M t Beisig Succession is the occurrence of units not seen below in the Eastern Succession, the presence of a high angle fault between the two, and a complete lack of granitic pegmatites in the M t Beisig Succession. An abundance of carbonate units further distinguishes the M t Beisig Succession from neighbouring successions. 14 Unit 6A approximate thickness = 250 m The lowermost unit seen in the M L Beisig Succession is equivalent to Unit 5 of the Eastern Succession. Only a small portion of this unit is seen below the south flank of ML Beisig, however, possibly correlative rocks exist between ML Beisig and the next prominent peak to the west (elev. 7715 ft, Fig. 4). Unit 6A is easily recognized by its largely pelitic composition, and includes a locally thick marble member. The unit also commonly contains large idioblastic crystals of kyanite and staurolite. The fine to coarse grained pelitic schist is frequently silvery blue weathering, owing to the high volume of muscovite and aluminosilicate minerals. The schist is locally interrupted by the presence of several cm to 10 cm thick medium grained grey quartzite layers. Further north in the col between peak 7715 and the ridge adjoining ML Beisig, thickly layered quartzite makes up a large part of this uniL These layers are often 10 cm to 1 m or more in thickness and are light grey weathering, composed of round granular quartz not exceeding 1 mm in diameter. Unit 6B approximate thickness = 350 m In marked contrast to the pelitic schist of Unit 6A, the rocks of Unit 6B are very feldspathic pelites. The most striking effect in outcrop is the coarse weathering, chalky feldspar and staurolite. Dark red porphyroblasts of staurolite 1 to 5 mm long and abundant garnet dominate the coarse metamorphic grains, as opposed to the blue kyanite blades of Unit 6A. In the rare fresh specimen, the rocks are grey to dark grey, and weather readily to light grey or orange-red. Two distinct marble members are present in conjunction with a dense amphibolite marker layer. These marble layers probably represent one individual layer which has been structurally repeated. In the region of the south flank of ML Beisig, some of the marble and amphibolite have been metamorphosed to an extremely mottled looking calc-silicate. On the ridge adjoining ML Beisig to the west, the amphibolite and marble may be observed essentially intacL The two marble layers 15 flank the discontinuous amphibolite, and vary in thickness: the upper marble averaging 2 - 3 m, and the lower marble unit ranging from 1 m, diminishing to zero. The composition of the marble is dominantly rhombs of calcite (± dolomite) 0.5 - 3 mm across, plus a great abundance of dark brown biotite which defines the compositional layering. Quartz is also important, and occurs as layers, and to a lesser extent as veins. Tan weathering layers are common throughout the marble, and are interspersed with clots of pyrite and magnetite(?). The marble is particularly susceptible to ductile deformation, and locally displays intricate fold geometries. The remainder of Unit 6B is medium to coarse grained, quartz-rich schist, which is terminated abruptly by the Boss Mountain gneiss in the vicinity of M t Beisig, and continues on the north side of Boss Mountain. Boss Mountain Gneiss maximum thickness in the M t Beisig section = 305 m; average thickness = 210 m The Boss Mountain gneiss covers much of the west flank of M t Beisig. Contact relations between the gneiss and the metasediments show an obvious discordance, but some degree of structural conformity is maintained. Contacts are frequently covered, but where observable, they have been intensely deformed and are sub-parallel to the compositional layering in the metasediments. The obviously disparate compositions of the metasediments, versus the strongly foliated granitoid, provides a basis for defining the contact between the two rock types. Complications arise due to the extreme degree of foliation development within both the sediments and the intrusive. Additionally, the gradation from igneous to metamorphic textures from the internal portions to the margins of the intrusive are very pronounced. The border (or marginal) zone of the granitic gneiss is fine to medium grained, and very felsic. Much of the gneiss is mylonitic, with a very strong, nearly parallel laminated foliation. Very little interdigitation of the main body of gneiss with the surrounding metasediments has occurred at this location. The central portion of the 16 intrusion is a coarse grained, quartz-feldspar megacryst augen gneiss, which grades upward into a more thinly foliated, siliceous, almost schistose material. An upward compositional and textural variation that is analogous to that seen at the lower contact zone, probably indicates the upper portion of the granitic body has been barely exhumed. Thus, the maximum original thickness of the Boss Mountain gneiss in the vicinity of Mt. Beisig, was probably no more than 400 - 500 m. The augen gneiss is exceptionally distinctive in outcrop, with light coloured polycrystalline augen of quartz and feldspar in a matrix of coarse crystalline quartz, K-feldspar + plagioclase, and abundant biotite and muscovite. The more schistose evenly textured marginal variants are brownish weathering and look nearly "grit-like" in outcrop. They are however, clearly equivalents of the augen material exposed in the core of the intrusion. Unit 7 approximate thickness = 200 m The top of the. ML Beisig Succession is a locally feldspathic, medium to coarse grained quartzite and metapelite uniL The outcrop characteristics allow easy identification of this unit by the red-orange Fe staining after weathering, and by its massive appearance. Compositionally, the dominant constituents are grey, very granular quartzite with lesser granular schisL In some locations, kyanite is ubiquitous within coarse grained quartz-muscovite schist and quartzite. Quartzite in this unit is a mixture of two varieties; medium grained grey homogeneous quartzite, consisting of quartz, plagioclase, minor biotite and muscovite. The second type is quartzofeldspathic, very micaceous quartzite that varies from schistose to strongly gneissic in texture. Quartzofeldspathic units are extremely dense, finely laminated greyish to blue-grey and invariably weather tan and very orange-red due to the presence of Fe oxides. Coarse flakes of white mica and biotite 1 - 2 mm long comprise up to roughly 15 to 20% of a typical specimen, in contrast to the few percent seen in the more homogeneous quartzite. The quartzofeldspathic type generally appears structurally above the grey 17 quartzites. Quartz veins are locally abundant, but are volumetrically less important than in many other units in the area. 3. R O S S M O U N T A I N S U C C E S S I O N Stratigraphy on the slopes of Boss Mountain and the adjacent ridges to the north and west, represent a different variation of Snowshoe Group metasediments, and their relationship to the intrusive rocks in the area. Compositionally, the succession is equivalent to parts of the M L Beisig sequence, and shares some gross characteristics in types of lithologic units. However, the internal variations are greater, and an overall more well layered nature is observed. Large thicknesses of rhythmically layered Snowshoe metasediments occur on the north slope of Boss Mountain, and on the adjacent plateau to the north. Colour laminations within the interbedded quartzite and pelite units are common in this succession. Varying tones of brown, orange, tan and grey are repeated over several hundred metres of section. Contact areas between augen gneiss and sedimentary rocks show a strong mixing, and intense leucocratic dyke injection related to the intrusion of the Boss Mountain gneiss protolith. Unit 8 approximate thickness = 80 - 100 m The base of Unit 8 is similar to the highest unit described in the ML Beisig Succession - Unit 7. Above this, is buff to tan weathering marble, and amphibolite/chlorite schist with minor feldspathic schisL A prominent sequence of metasediments consisting of quartzite, quartzo-feldspathic schist and pelitic schist, is encountered above, and is in turn abruptly and discordantly in contact with Boss Mountain gneiss. Structurally above the thick marble (20 - 60 m observed) exposed at the base of the north face and upper south flank of Boss Mountain, are very rythmically layered quartzite with lesser feldspathic and pelitic schists. An extreme variation exists, but the unit is dominated by 50% to 80% quartzite, with well 18 developed fine parallel compositional banding of several mm spacing within layers, giving a finely "bedded" appearance. The layers of banded quartzite are restricted in thickness from 6 - 2 0 cm, and are exposed in excess of 60 - 70 m of uninterrupted, but tightly folded section. Thin interlayers of dark grey biotitic schist, usually less than 5 - 10 cm thick, disrupt the otherwise homogeneous medium grained quartzites. Micaceous quartzite with excessive amounts of very fine garnet, and discontinuous amphibole-bearing rocks, are locally present within the pelitic zones. Fine granular garnet may comprise up to 10% of the rock, with garnet grains only about 0.1 - 0.5 mm in diameter (Fig. 24a). Unit 9 approximate thickness = 20 - 30 m In sharp contact above Unit 8 is a series of mica schists with lesser amounts of micaceous quartzites, coarse calcite marble and calc-silicate. The mica schists are dominated by reddish brown weathering, muscovite -biotite schist and less regularly interlayered micaceous quartzite. The intermittent quartzites comprise up to 30% of the outcrop, and rarely exceed 1 m in thickness, with the majority only 20 cm or less, similar to those in Unit 8. Unit 10 approximate thickness = 40 - 50 m A marked contrast in composition and texture occurs in the unit at the contact between the Boss Mountain gneiss and Snowshoe Group rocks north of Boss Mountain. The compositional changes begin above Unit 9. Silvery mica schists are transformed into chalky looking feldspathic gneiss. The gneiss is relatively schistose, but gives the overwhelming appearance of a strongly foliated igneous rock. Fine scale layering separates whitish layers of quartz and plagioclase, from dark mafic rich layers, 1 - 3 mm thick, containing biotite, garnet, kyanite and minor amphibole (hornblende). Granular, euhedral apatite is concentrated along some of the compositional layers, and locally comprises a large proportion of some of these layers. Leucocratic dykes composed mainly of quartz and plagioclase mark the boundary between the pelitic 19 metasediments, and coarse grained augen gneiss of the Boss Mountain intrusives. The dykes are typically 5 - 10 cm in width, and are intensely deformed by zones of concentrated ductile deformation. Boss Mountain Gneiss approximate thickness = 350 - 400 m Contact relationships clearly indicate that truncation of Snowshoe Group stratigraphy by the granitic gneiss has occurred, and that intrusive activity has produced a narrow zone of migmatites immediately adjacent to the Boss Mountain gneiss at this location. Disharmonic folding accompanied by quartzofeldspathic dyking of the country rocks characterizes the 10 - 15 m wide contact zone between gneiss and metasediments. Medium to coarse grained quartz, plagioclase, muscovite, biotite and K-feldspar comprise the granitic rocks that show strongly interfolded contacts with the pelitic metasediments. Rootless folded, white coloured dykes comprise up to 10% or more of the immediate zone of contact between the two groups of rocks. In contrast with the enclosing metasediments, are schistose rocks that appear to be strongly altered metasediments or possibly a more pelitic border phase of the Boss Mountain gneiss. Structurally above the granite/metasediment contact zone, the granitic gneiss coarsens away from the contact, and develops large quartz and feldspar augen in a strongly deformed matrix of the same plus biotite and muscovite. The light coloured dykes persist within the gneiss for tens of metres from the contact zone, where they are both parallel to, and crosscut,the compositional layering within the gneiss. A recent U - P b zircon date from a sample collected at this locality has yielded a minimum age of approximately 320 Ma for the Boss Mountain gneiss (J. K. Mortensen personal communication 1985). The relationship of the gneiss to the Snowshoe metasediments, and the dykes to both of the other rock types suggests the following sequence of events : Snowshoe Group rocks were intruded by the granitic rocks dyking within both granite and metasediments, and across the contact 20 subsequent deformation by successive folding and faulting affects the entire package. 4. WESTERN SUCCESSION Boss Mountain Gneiss - approximate thickness = 500 m A large exposure of variable-texture granitic gneiss caps much of the group of peaks west of Boss Mountain to the east slope of peak 7698. The prevailing variations within the gneiss includes a decrease in grain size toward the margins of the body, accompanied by a complement of leucocratic, fine grained quartzose dykes. The dykes occur as crosscutting features within the gneiss, and also appear external to the gneiss as concordant sills and layers within the metasedimentary country rocks. The variations at the margins of the granitic body are best exposed on the east face of peak 7874, and at the westernmost extremity of the Boss Mountain gneiss. The contact relations between the granitic gneiss and the country rocks are rarely simple in this succession. An extremely mixed and lithologically heterogeneous zone characterizes the contact west of Boss Mountain. Amphibolite and feldspathic units occur below the summit of peak 7874, where they are intercalated within the quartzite and mica schists at the margin of the Boss Mountain gneiss. Unit 11 approximate thickness = 1750 m Pelitic schist, lesser quartzite and carbonate dominates the succession structurally above and west of Boss Mountain. Massive, to locally well-layered grey quartzite and mica schist are present immediately above the strongly injected granitic gneiss/metasediment contact, where homogeneous quartzite commonly exceeding 1 m in thickness is interlayered with subordinate pelitic schist. A local compositional variation within the pelitic schist occurs where quartz-biotite-muscovite-garnet schist and quartzose leucogneiss produce an alternating pelitic/felsic zone. The gneiss is strongly foliated but granular, white and made up of quartz, fine grained muscovite and very 21 fine grained biotite, in sharp contact with the coarse grained pelite. The gneiss layers average 10 cm to nearly 1 m in thickness. In places the gneiss is a more biotite-feldspar rich augen textured rock, and appears to be a variant of the white felsite. These felsic rocks are interpreted to be aplitic dykes derived from the Boss Mountain gneiss. Approximately 20 - 30 m of this rock type is exposed at the western margin of the granitic gneiss. Coarse-grained pelitic schist, micaceous quartzite and marble dominate the remainder of the lower portion of Unit 11. The schists are characterized by their well developed anastamosing foliation of predominantly coarse muscovite, large garnet porphyroblasts, and reddish brown weathering hues. Closely spaced schistosity planes enclose large ( 2 - 5 mm) garnet porphyroblasts. Staurolite grains are often less conspicuous, but in places are visible as corroded crystals intergrown with garnet Disruption of the local succession by faulting nearly parallel to lithological boundaries is evident Where intact the pelitic schists are overlain by a very fissile amphibolite/chlorite schist which appears to be similar to the amphibolite present in the M t Beisig Succession. The sequence of amphibolite and several (1 - 2 m) coarse marble layers is recognized at the western margin of the granitic gneiss body. These marble layers are of the same general composition as those seen in the east; medium to very dark grey, and composed of calcite, white mica, biotite (+ quartz). The marble often hosts a zone of calc-silicates, as metamorphism has caused reactions between the carbonate and adjacent pelitic rocks (see metamorphism section). Structurally upwards in Unit 11, the marble, amphibolite and schist are overlain by slivers of black graphitic phyllite and schist interpreted to represent fault-bounded slices. The faults are responsible in part for the repetition of the schist, marble and amphibolite. Adjacent to the last fault to the west, a different lithology occurs, comprising coarse grained garnetiferous schist and fine grained mylonitic quartzite. Tiny K-feldspar augen are observable in the quartzite, with fine grained muscovite and 22 biotite giving the rocks a lustrous appearance. Above this interval occurs more coarse grained pelitic schist, lesser thinly laminated quartzite and another group of thin (1 -3 m), tan weathering micaceous marbles. Unique to this stratigraphic interval, whose thickness measures 300 - 400 m thick, are several 10 - 15 m wide zones which appear to have been silicified. The top of the succession is dominated by thick pelitic schist, minor marble layers and thin micaceous quartzite, similar to some of the lower portions of the unit Unit 12 approximate thickness = 1200 m The western margin of the area adjacent to the Intermontane Belt contact contains a thick succession of marble, amphibolite, laminated quartzite, granular quartzitic schist and thinly layered augen gneiss. The break between the thick sequence of Unit 11 and the beginning of Unit 12 is defined by a thick (15 - 75 m) tan micaceous marble. The rock is a coarse- to medium-grained, grey calcite marble. Its estimated maximum thickness is greater than 60 m, but is inferred to be as great as 100 m, although this is indeterminate as the base is not exposed. The evidence which suggests that these rocks are very different from previously described units, is the presence of dark grey quartzose schist and gneiss, and distinctive bands of what appear to be originally igneous "quartz-eye" augen gneiss. These "orthogneiss" layers show well preserved igneous textures, and extremely sharp contacts with the pelitic sediments. They have not been recognized elsewhere in the Boss Mountain area, but do exist in the marginal sequence of pelite, carbonates, and thin quartzofeldspathic gneiss within Snowshoe Group rocks immediately below the same tectonic contact in the M L Perseus area (D.C. Elsby - personal comm. 1984). A nearly identical sequence of rocks has been documented in the same structural position on the west flank of Eureka Peak above Crooked Lake (M. A. Bloodgood, personal comm. 1984). 23 Above the contact with the marble, very granular schist consisting mainly of clear, glassy quartz, plagioclase + K - feldspar and fine grained muscovite, hosts an assortment of sub-lithologies including very dark grey, schistose micaceous quartzite and quartz-muscovite schist Immediately overlying 30 - 40 m of the schist/quartzite, are the "quartz eye" felsic gneiss layers. A notable size and frequency gradient is apparent in the quartz augen within thin repetetive compositional bands several cm to more than 10 cm thick. These bands are separated from one another by fine grained quartzose and micaceous zones. Rounded, to fairly elongate quartz phenocrysts, 2 or 3 mm by 5 mm, are scattered along the prominent foliation at the margins of individual layers, and show a gradual decrease in size toward the opposite layer margin. The layers themselves occupy an interval of approximately 25 m or more. The complete mineralogy of these layers is quartz, biotite, and minor muscovite, with less than 3-4% augen quartz within a strongly laminated quartz/mica matrix. Intermittent pelitic to quartzofeldspathic schist layers are muscovite rich, contain abundant quartz and feldspar, and often host garnet ± staurolite. A large barren quartz vein 3 - 4 m thick immediately overlies this mixed gneiss and pelite, above which no augen gneiss is observed. Pelitic schist resumes above this section, and is overlain by another 15 - 20 m thick, well layered coarse calcite marble, interpreted to be the same marble unit as at the base of this section. Above the marble, more granular pelite and compact greenish-grey quartzite occurs; another thin marble layer less than 2 m thick, and finely layered homogeneous quartzite 50 - 60 m thick, above which the Permian metavolcanics of the Antler Formation are reached. 5. PERMIAN SLIDE M O U N T A I N GROUP Antler Formation approximate thickness = 295 m The contact between the Hadrynian rocks of the Snowshoe Group with metavolcanic rocks of the Slide Mountain Group is purely tectonic and not well 24 exposed at the west margin of the Boss Mountain area. The contact has been mostly interpolated, constrained by exposures of Snowshoe metasediment and mafic schists of the Antler Formation which are observed separated from one another by a distance of 2 - 3 m. Close to the contact, the Antler volcanics are dark green, chlorite/amphibole schists with some massive zones where hornblende crystals are recognizable. Cubic pyrite crystals are locally abundant, and distributed along the major foliation planes. Near the top of the volcanic section, an important textural change takes place. A lozenge-like texture of quartz and feldspar grains surrounded by anastamosing chlorite and amphibole gives the rocks a lighter coloured augen schist appearance. Approximately 15 - 20 m of the augen schist is present, and is overlain in sharp contact, by dark green hornblende/chlorite schist identical to that observed at the lower contact with the Snowshoe Group. 6. TRIASSIC PHYLLITE approximate thickness = 800 - 1000 m The contact zone between Antler Formation volcanics and Triassic sedimentary rocks is reached at the western limit of exposure in the map area. The surface of contact is not exposed between these distinctly different units, but can be constrained within about 20 m. The slate and phyllite are dark grey to black, locally very fissile quartz- muscovite- graphite ± chlorite metapelite. Variations include a more siliceous black slate and phyllite, particularly near the lower contact Thin laminations internal to the foliation are apparently bedding/compositional layers, which upon close examination are very tightly folded about the dominant slaty cleavage/schistosity. These layers are typically less than several mm thick, and are accompanied by quartz veins 1 - 3 mm thick, which are also lightly folded about the foliation planes. Graphite, muscovite and 25 chlorite comprise the phyllosilicate minerals which are responsible for the well cleaved character of the rock. Compositional layering or bedding is obvious in some of the fine grained rocks, and may be observed as fine colour lamination or grain size variation from fine to very fine grain sizes. The top of the Triassic Phyllite sequence is inferred to be located several hundred metres east of the east shore of Bosk Lake, where Jurassic volcanic rocks are exposed. The thickness is unknown, but would appear to be nearly 1000 m from the map of Campbell (1978). 7. STRATIGRAPHIC CORRELATIONS - DISCUSSION Compilation of existing stratigraphic information from the Boss Mountain area, Eureka Peak, Mt. Perseus and the Isoceles Mountain area, has allowed a correlation of major rock units in the region to be made (Ross et al. 1985). The detailed stratigraphy within each area has shown that divisions within the Hadrynian, Permian and Triassic sequences are persistent and traceable over distances of at least several tens of kilometres. Using the Omineca Belt boundary for reference, the gross order of stratigraphy on either side of the boundary is readily observed. In the Boss Mountain area, the overall variation in the Snowshoe Group downwards from the Intermontane Belt contact, is from a predominantly pelite/carbonate facies, to a more quartzofeldspathic, granular group of rocks. Well-bedded metasediments comprise the bulk of the units at intermediate structural levels within the area (see Boss Mountain Succession), with compositional layering showing a general increase in thickness of individual layers in the eastern, and to some extent in the western portions of the area. In the McKay Creek area, a succession similar to the stratigraphy in the western part of the Boss Mountain area has been described (D.C. Elsby personal comm. 1984). The Antler Formation as seen on the west flank of ML Perseus is more 26 of a mixture of volcanics and sediments, but retains its mafic volcanic character. In the M L Perseus area and McCallum Peak area, Snowshoe Group metasediments are similarly pelitic to coarse clastic and siliceous, with abundant, thick impure carbonates. The mixed, clastic nature of rocks in the Boss Mountain area clearly indicates a diverse source terrane through time. Parent rocks were most likely crystalline and granitic rocks representing some form of continental-type crust Detrital grains, particularly zircon, sphene, and occasionally feldspar and quartz where demonstrable, indicate derivation from such a source. The original sedimentary facies have been described as a prograding platform sequence (Campbell and others, 1973), and have been equated with rocks in the main Cariboo Mountains interpreted to be turbidite fan deposits (Murphy and Rees 1983). Repetetive sequences of thin layered quartzite and pelite with a relative paucity of carbonate rocks suggests a probable flysch-type deposit with local silty limestones. An abundant feldspathic component in some units is also interpreted to represent original feldspar derived from the crystalline source. The original stratigraphic thickness of these rocks must have been large, even with a conservative estimate of 100% tectonic thickening due to successive deformations and metamorphism. Observable structural thickness of the Snowshoe Group is in excess of 5000 m in the Boss Mountain area. III. STRUCTURE - DESCRIPTION OF PHASES Four important phases of deformation are recognizable in the Boss Mountain area, and a fifth phase of a more localized nature is evident The various phases are identified and separated from one another in the field by overprinting relations of their respective S-surfaces, and by their orientation and distribution. These observations have led to the following structural progression, discussed in sequence beginning with the earliest recognizable structures in the Snowshoe Group, through the latest structures identified in the Antler Formation and Triassic/Jurassic rocks. The rocks of the Snowshoe Group are considered to be basement to the overlying cover comprising the Slide Mountain Group (Antler Formation) and the Triassic to Jurassic sediments and volcanics. These two distinct packages will often be referred to as basement and cover rocks, respectively, which will imply this sttuctural/stratigraphic division. Structural domains within the map area have been defined on the basis of the orientation and distribution of early structures. Four orientation groups comprise domains 1 through 4 (Fig. 5). Each domain has been considered separately in the analysis of data collected in the field, for F l , F2 and F3 structures. Later structures have not been analyzed in this manner since their distribution is more sporadic within the map area. 1. BASEMENT - SNOWSHOE G R O U P Phase 1 Features related to the Phase 1 deformation are the regional foliation - SI, and mineral lineation LI . Original bedding (SO) has been completely transposed within the SI surfaces, and is now recognized as a pervasive compositional layering. The combination of SO (bedding) and SI is simply referred to as SI. This earliest recognizable phase of deformation appears to be related, at least in part, to an enigmatic regional metamorphic/plutonic event In the mid-Paleozoic, granitic bodies 27 lntermontane/Omineca\ Belt contact Structural Domaln\ boundary 29 were intruded into the Snowshoe Group sediments (ie. Boss Mountain gneiss, Quesnel Lake Gneiss) and a regionally pervasive synkinematic recrystallization affected the rocks. The SI differentiated layering is composed of a combination of folded bedding, and quartz filled partings both along and across the compositional layering. These clear to more clouded, foliation-conformable quartz veins may comprise a considerable proportion of the rock. The veins are frequently folded into small scale, tight "rootless" intrafolial folds, internal to the SI foliation surfaces (Fig. 6). Dismembered bedding is seen as fine granular quartz and feldspar, with thin selveges of pelitic material, whose thickness ranges from a few mm to several metres. The SI foliation locally consists of a finely laminated mylonitic fabric. Thin bands of quartz and feldspar were produced during Phase 1 deformation, and possibly accentuated by Phase 2 deformation. The dimensions of these layers are often on the order of 5:1 in their length-width ratios. The SI foliation has not been observed as a crenulation cleavage fabric. Figure 6 shows the typical outcrop appearance of transposed FI folds. These folds are infrequently observed, but cause obvious discordant compositional layering orientations within an individual outcrop. Angles between limbs of these tight folds have been measured between 0 ° - 2 0 ° . Mesoscopic fold limbs reach lengths of more than 10 m in some exposures. The folds have profiles which show extreme thickening in their hinge regions, and a marked thinning of their limbs (Fig. 7). F I folds are commonly preserved within the more competent compositional layers as completely transposed folds. The discontinuity of F I fold structures prohibits the tracing of mesoscopic folds over any considerable distance, and renders the relationship of stratigraphy to the F I folds unclear. Vergence of these folds in the map area is indeterminate. Transposition has caused repetition of individual layers, often with no hint of fold closures. This implies a tectonic thickening of unknown, but probably large proportions. Variations in 30 F i 8 1 e , 6 r n . m i n « . f o l d s . a) u8htly folded intrafolial SI quartz segregation, b) transposed r S o n - a y s i m P U n l t E " L N e a T l y ^ l y l n g s c h i s t o s i t y i s tiie transposition 31 Figure 7 - Photograph and sketch of F l folds within quartzite and schist, Unit BM-9. 32 attitude of the SI foliation are readily observable in much of the area. In the vicinity of ML Deception and Mica Mountain, the SI foliation dips shallowly to the south and southeast (Campbell 1978). In the Boss Mountain area, this regional schistosity dips gently to the northwest or is refolded into various orientations due to several phases of later folding. The original geometry of large scale FI folds is unknown. Phase 2 Phase 2 deformation is the most significant in terms of both intensity, and development of macroscopic folds. F2 folds are widespread and well developed in all parts of the map area. They are associated with L2 mineral lineations which are always accompanied by a penetrative schistosity - S2, parallel to axial planes of F2 folds. Folds are frequently tight to nearly isoclinal, and rarely ever open in style. Occasionally, they are disharmonic folds, but the majority of F2 folds have a high amplitude to wavelength geometry, with thinned limbs and thickened hinges ("similar type" folds, Fig. 8d). F2 folds are nearly recumbent in much of the eastern part of the map area, but are inclined in the southwest portion of the area. Vergence of minor folds is variable because of later folding and faulting effects, but most mesoscopic F2 folds reveal a northeasterly sense of rotation, consistent with macroscopic F2 folds in the area (see cross-sections - in pocket). It is interpreted that much of the area is situated on one limb of a large second phase antiform, overturned to the easL The magnitude of F2 folds varies from small-scale folds having limbs several tens of cm long, up to. large second-order structures with limb lengths 1 - 2 km long. Equal area projection analysis of F2 fold elements has been done according to the structural position of observed outcrops with respect to both F2 and F3 fold structures. The structural zones are defined by the respective limb and hinge regions of macroscopic folds. 33 (b) Figure 8 - F2 folds, a) photograph of disharmonic F2 folds from the contact zone between the Boss Mountain gneiss and metasediments, b) typical F2 fold in quartzite and schist of the Snowshoe Group, Unit E-2, c) mesoscopic parasitic folds on the upper limb of northeast verging F2 antiform. d) f vs. a plot of several F2 folds from various outcrops of Snowshoe Group schist and gneiss. (8c) 35 <8d) 36 Three zones of areal distribution of F2 axial planes have been illustrated: (1) Eastern Zone - dominantly shallow northeast and west dipping F2 axial planes. (2) Boss Mountain Hinge Zone - F2 folds in the zone occupying the hinge of the F3 antiform show a distribution indicating shallow north to northeast dipping axial planes. This zone encompasses Domains 2 and 3; similar orientation of Phase 2 structures in this domain precludes the need for separate analysis. (3) Southwest Limb Zone Boss Mountain Antiform - F2 axial planes are more evenly distributed with shallow, to steep southwest dips, and occasional northerly dips. Axial planes of F2 minor folds have ben plotted, and show a girdle distribution about F3 fold axes. In the hinge region of the phase three Boss Mountain antiform, second phase folds maintain a general " M " symmetry, with shallow northwest to east dipping axial planes. This second phase fold hinge region influences the geometry of the large F3 upright antiformal structure which dominates the regional configuration. Thus an F2 nappe core coincides with the hinge region of the F3 antiform given the name "Boss Mountain anticline" by Campbell (1971). Due to the indeterminate nature of stratigraphic facing directions, the name "Boss Mountain antiform" has been adopted and henceforth used. The S2 foliation surfaces are defined by a mica schistosity and penetrative crenulation cleavage. Microfolds with compositionally differentiated quartz/feldspar and mica domains defining their axial planes are observed in both hand specimen and thin section. The crenulation fabric typically crosscuts the SI foliation at variable angles (10° - 40°), with only occasional high angle or subparallel relationships observed. Where the S2 development is most intense, it nearly obliterates the continuity of SI surfaces. L2 linear features are present as elongate mineral lineations and intersection lineations parallel to F2 fold axes. The most common type of lineation in the Boss Eastern zone Domain 1 Hinge zone Domain 2+3 Southwest limb zone Domain 4 Figure 9 - Lower hemisphere, equal area projection of poles to mesoscopic F2 axial planes (7T-S2). Data is divided into domains shown in figure 5. 38 Mountain gneiss is an intersection of the S2 schistosity and SI foliation; crenulation lineations are frequently observed within the pelitic metasediments. Mineral growths parallel to L2 consist of quartz rods up to several cm long, finer grains of biotite and muscovite, rarely the sub-parallel alignment of tabular staurolite, and needle-like grains of hornblende when present (e.g. Unit B-6B). The prominent foliation within the Boss Mountain gneiss has been strongly overprinted by F2 deformation. It is usually simple to separate S2 foliation surfaces from the SI compositional layering. The quartzofeldspathic compositional layering within the gneiss is typically crosscut at low angles by the S2 foliation. The S2 schistosity is defined by mica alignment and the preferred orientation of elongate quartz grains. The fabric is frequently mylonitic-looking and has been overprinted by later metamorphic effects. Megacrysts composed of intergrowths of K-feldspar, plagioclase and quartz may have originally been K-feldspar, but have been recrystallized into aggregates of quartz and feldspar. Granitic pegmatite is folded by F2 within Domain 1 (Fig. 10). Locally tight, mesoscopic folds of tabular pegmatite bodies 1 - 2 m thick are prominent at several locations in the Eastern structural succession. 39 Figure 10 - F2 folds in granitic pegmatite, a) Unit E-2 , and b) Unit E-4. 40 Figure 11 - Phase 2 fault zones, a) Photo of trace of fault on east side of peak elevation 7698 ft. Location of Figure b) is indicated by the arrow, b) Phyllite lined Phase 2 fault zone within Unit W - l l . Compositional layering within the phyllite truncates that in the quartzofeldspathic schist and gneiss adjacent to the zone, c) Closeup of schistosity and "shear bands" in graphitic phyllite within F2 fault Unit W - l l . View is to the southeast 41 (1 1c) sw N E Metres 2 0 0 0 -1 6 0 0 -L 4 0 0 0 Snowshoe Group pelitic schist, micaceous quartzite, marble and amphibolite Boss Mountain gneiss phyllite lined fault S1 form lines Figure 12 - Schematic coss-section showing the occurrence of Phase 2 faults in the western part of the map area. F2 folds are interpreted from SI - S2 relationships and minor fold data. F2 antiformal fold is in marble, Unit W - l l . 43 Phase 2 faulting Phase 2 faults are obvious west of the Boss Mountain gneiss, and have caused structural imbrication of the Snowshoe Group in Units W - l l and W-12 (Figs. 5, 11 and 12). The fault zones are typically less than 20 m wide, and taper downward or splay, taking up a cuspate geometry in profile. They are lined with sooty, black phyllitic rocks, rich in muscovite and graphite, with locally abundant quartz veins. Garnet has been recognized in some specimens. A well developed schistosity (S2) is present within the phyllite, about which the compositional layering has been tightly folded or transposed. The compositional layering within the phyllite appears to truncate the compositional layering within the enveloping metasedimentary schist and gneiss (figure l i b , lower right corner of photograph).. The presence of slickensides and mineral elongations nearly parallel to the dip of the foliation, and the sheared appearance of the phyllites, indicates that faulting may be responsible for the formation of these zones. The phyllites are interpreted to be fault-bounded slivers of the cover sequence (Triassic phyllite). They are more or less laterally continuous within the map area, and have been measured from 2 or 3 m, to more than 70 m in thickness (Fig. 12). They were most likely developed during the Phase 2 deformation. The best preserved large-scale F2 structures are in the vicinity of the F3 Boss Mountain antiform hinge, where easterly verging, recumbent F2 folds are outlined by structurally continuous units. The net effect of F2 folding is the further repetition of the units by tight, regional scale folds. Considering that original stratigraphy has probably been thoroughly disrupted by Phase 1 deformation, it is likely that the F2 nappe structures do not represent intact stratigraphic sequences. The presence of several large antiforms and synforms of F2 generation has been documented in the area, but are locally obscured by the lack of stratigraphic continuity. Thicknesses of folded units vary irregularly, and extreme thinning of F2 fold limbs has been observed. 44 Phase 3 Macroscopic F3 fold structures are reflected in the regional map pattern. The variation in structural style in the map area is, in part, attributable to different temperatures during the waning metamorphism. Phase 3 deformation transcends any clearcut division between ductile and brittle behaviour in rocks in the area, and ranges in its manifestations from shear folding, to kink bands and fine crenulations devoid of any penetrative fabrics. The most frequently observed F3 fold features are small-scale crenulations and mesoscopic second order folds representing parasitic structures on a large regional scale antiform. The F3 Boss Mountain antiform encompasses the region between the north end of Crooked Lake, to south of M L Deception. It is generally open and uprighL with gentle northwesterly dips over much of the broad hinge zone, but with locally steep and overturned segments. Minor folds associated with the antiform give a consistent southwesterly vergence sense. The core of the structure directly underlies Boss Mountain. Average orientation of the F3 axial surface is 140°/40-50° NE, and measurements of minor fold axes and stereonet plots of L3 lineations indicate a fold axis orientation of approximately 15°/330° (see stereoneL Fig. 13a). Both the east and the west limbs of the structure achieve dips of more than 60° as adjacent downward closing fold structures are approached (ie. Eureka Synform, Elbow Lake Synform). The Boss Mountain antiform encompasses a series of minor (second order) synforms and antiforms, having wavelengths of 300 m or more, with individual limb lengths seldomly less than 150 m. These folds have relatively small amplitudes, typically on the order of several hundred metres. West of the central hinge zone of the antiform, the folded S1/S2 schistosities dip southwest at moderate to steep angles, and there is a marked decrease in frequency of minor fold structures of all generations, particularly F3. Mesoscopic F3 folds do not exceed several metres wavelength or amplitude in these localities. The style of folding from outcrop observations, inference from stereonet 45 distribution of distorted F I and F2 features, and layer thickness variation measurements indicates a modified buckle fold geometry (Class IB to IC type fold (Ramsay 1967), see Fig. 13). Outcrop scale folds of F3 generation are commonly accompanied by a crenulation cleavage - S3. These are typically more upright than the larger parasitic F3 folds. F3 crenulations are generally oriented 140°/45-65°NE, with rare dips of up to 90°. Deformation by F4 folding has distorted this value locally and more variable F3 strikes have been recorded. F3 crenulations and their attendant axial planar schistosities, are well developed in the east and central parts of the area, but are only intermittently present in the west F3 crenulations everywhere distort the L2 lineation, and are accompanied by a crenulation lineation or occasionally a subordinate mineral lineation - L3. Mineral lineations may be fine needles of biotite or elongate quartz which have grown upon the folded SI and S2 surfaces. Other metamorphic minerals such as staurolite and kyanite have not been observed to oudine this same lineation. Both of the preexisting schistosities, SI and S2, have been affected by low to high amplitude crenulations. Closely spaced (400-500M m) crenulation cleavage is associated with these, and has resulted in macroscopic cleavage spacings of 2 - 5 mm in outcrop. S3 cleavage is variably penetrative, and exists as both spaced cleavage and mica schistosities, as a function of the intensity of F3 folding. Where the F3 folds are modified buckle folds, the schistosity is, at most, an incipiently developed foliation of mica films and tabular quartz. The more intensely developed F3 folds are accompanied by a penetrative axial planar fabric. This latter type of foliation is rare. 46 solid circles - axial planes open circles - fold axes Figure 13 - F3 fold data, a) Lower hemisphere, equal area projection of F3 minor fold data plotted on a it diagram, poles to F3 axial planes and minor fold axes are shown, b) t' vs. c plot for several F3 minor folds within Snowshoe Group metasediments. c) Photograph of open, southwest verging F3 folds in Unit. BM-10 north of Boss Mountain, d) F3 antiformal fold in Boss Mountain gneiss. Western Succession. 47 3 0 6 0 9 0 pelitic schist schist and quartzite quartzite amphibolite and quartzite (13b) (13d) 49 Phase 4 Fourth phase deformational features are seen throughout the Boss Mountain area, but are best developed in association with features related to imbricate faulting of the western flank of the area. These features are, in decreasing order of structural importance and abundance: Crenulations and mesoscopic folds with axial surfaces inclined to the southwest Shear zones and attendant folds. Faults developed or reactivated during D4. Gentle open folds, to locally more sharp, angular hinged folds are common. The orientation, and timing of these folds facilitates their recognition in the field. The average F4 axial plane orientation is 150°/38° SW. L4 is variable, but plunges shallowly to the northwest or southeast (Fig. 14a). Figures 14b and 14c show typical F4 folds in outcrop, with their shallow southwest dipping axial planes, refolding the dominant foliations, SI and S2. In addition to normal crenulations and folds, localized shear zones contain discrete kinks and crenulations possibly related to faulting. The faulted zones are interpreted as the continuation of cuspate infolds of the "cover" rocks within the high grade metasedimentary "basement" rocks (see Phase 2 faulting). D2 deformation initiated these zones, and by D4 time, the faults had achieved steep dips due to F3 folding. Crenulations within the faulted rocks are clearly due to later effects of both D3 and D4 deformations. Low angle kink bands and crenulation lineations are observed internal to these zones, and kinks inclined shallowly to both the southwest and to the northeast are attributable to this episode of deformation. Foliation being crenulated is the dominant S2 schistosity. External to the narrow fault zones (2 - 50 m wide), F4 folds are generally very open warps or angular hinged kink-style folds which plunge shallowly both to the northwest and southeast It is common to find a rather well developed axial plane 50 ( 1 4 a ) Figure 14 - F4 folds, a) Lower hemisphere, equal area projection of F4 fold elements. Data are undivided with respect to structural domains defined for F2 and F3 structures, b) Photograph of F4 folds in micaceous quartzite and thin pelite of Unit W-12. Typical "Z" asymmetry is apparent in F4 folds beneath the boundary with the Slide Mountain Group (view is up-plunge to the southeast), c) Northeast verging F4 fold in Unit BM-10, refolding an east verging isoclinal F2 fold. (14c) 52 cleavage associated with these folds. Folds and crenulations are especially well developed westward beyond the margin of the Boss Mountain gneiss. Ubiquitous S4 crenulation cleavage and crenulation lineations are present in this zone. Phase 5 Crenulations croscutting all of the above features are observable within both the basement and cover rocks. These consist of a very fine crenulation about axial planes inclined 60° to 70° to the northwest Strike orientations of 040° to 060° are typical in localities from Boss Mountain to the western margin of the area. Fractures and joints parallel to this orientation are frequently seen throughout the area. 2. COVER - SLIDE M O U N T A I N GROUP A N D TRIASSIC PHYT A JTE Phase 1 Two important phases of deformation are recognized within the Antier Formation metavolcanics and the Triassic phyllites. Typical minor folds within both the metavolcanic rocks and Triassic pelitic rocks are intrafolial folds related to the SI foliation (Fig.15). Macroscopic compositional layering is difficult to recognize, with the exception of the contact between the two cover units. Mylonitic fabric is seen within the metavolcanic schists immediately below the contact with the Triassic phyllites. The fabric is lenticular, and consists of plagioclase and quartz, separated by layers of chlorite, hornblende, and epidote. Magnetite and pyrite are often present as large crystals. Feldspar phenocrysts or porphyroclasts are elongated within this foliation. Locally biotite, in addition to chlorite outlines the schistosity, and is accompanied by a well defined L I mineral lineation of hornblende crystals. Elongate quartz also grows parallel to this lineation. The Triassic slates and phyllites show similar schistosity elements of the regional SI foliation. Bedding (SO) has been tightly folded, and a well developed slaty, 53 solid circles - poles to S1 open circles - minor fold axes Figure 15 - F l fold data from the cover rocks. Lower hemisphere, equal area projection of poles to SI surfaces in the Antler Formation and Triassic phyllite. LI is curvilinear, and plunges both northwest and southeast 54 or crenulation cleavage has resulted. Linear structures within the slate and phyllite are fine crenulation lineations, or intersection lineations between compositional layering and SI. These typically plunge at shallow angles to the northwest Phase 2 Phase two features within the cover rocks are weakly developed crenulations of SO/SI surfaces, usually without an associated cleavage. Macroscopic F2 folding has developed about northeast dipping axial planes, and consist of minor warps and crenulations which do not affect the overall geometry of F l structures. Crenulation lineations - L2, are nearly parallel to the L I lineations in many instances. Angles between the two range from 10-20° as measured in the SI schistosity. L2 lineations invariably plunge to the northwest The S2 schistosity is not well exposed in the area, and was verified by examination in oriented thin sections. Phase 3 Crenulations plunging shallowly to the northwest and southeast are the result of relatively mild F3 folding. The axial plane orientation of the folds is nearly constant at 130 to 140/35° SW. This localized folding does not seem to alter the previously established structural geometry. Some brittle fracturing is also similar in orientation to the axial planes of these folds. Only several of these features were observed in the field. Phase 4 Fine crenulations with steeply northwest dipping axial planes affect the Antler Formation and Triassic phyllite. The effects of the folding are to produce small-scale interference structures with the earlier structures (Fl,2,3). Joints and minor fault scarps (?) with a similar orientation are commonplace in the areas of exposure of the cover rocks. 55 3. DISCUSSION The relationship between phases of deformation within the basement and cover has been established in the map area. Field relationships show that F2 in the basement (Snowshoe Group) is equivalent to F l in the cover sequences. This can be demonstrated by the conformable nature of the foliation across the tectonic boundary (Fig. 16). The dominant foliation within the Antler Fm and Triassic phyllite corresponds to the schistosity which is axial planar to F2 folds within the Snowshoe Group immediately below the basement-cover contact. This is also supported by the consistency of superposition of later structures between the cover and the basement The southwest verging F3 folds in the basement also affect the cover (F2), but are mildly developed in comparison. Metres 2000-1600-1000-- 4000 *- 3000 8 N o UJ -J < a. > < UJ S n o w s h o e Qroup Mart) is Augen gneiss < CC UJ 0. u % < I'l /// S l i d e M t n . Q r o u p B l a c k PhyWte / B a s e m e n t / / c o v e r b o u n d a r y phyNHe c o n t a c t Figure 16 - Detailed cTossection of the western margin of the area, illustrating the geometry of the tectonic contact between the basement (Snowshoe Group) and cover rocks (Antler Formation, Triassic phyllite). The axial plane schistosity to F2 folds in the basement is parallel to SI in the cover. IV. METAMORPHISM A N D TF.XTURAI. DESCRIPTION Two distinct phases of metamorphism have been recognized in the Boss Mountain area. The earliest recognizable metamorphism (Ml) , was of regional extent and probably coeval with Phase 1 deformation within rocks of the Snowshoe Group. The regional foliation (SI) in the basement was produced at this time. The M l metamorphism did not affect the cover rocks, hence the Late Paleozoic (Antler Formation) and Triassic/Jurassic rocks have no analogue to the SI foliation seen in the basement The second phase of metamorphism (M2) has been documented across the entire belt of rocks associated with the Shuswap Complex (Campbell 1971; Pigage 1978; Getsinger 1985). The M2 metamorphism occurred during, and outlasted the Early-Middle to Late Jurassic Phase 2 deformational event (Campbell et al. 1973, Pigage 1978; Monger et al. 1982). The metamorphism reached upper amphibolite conditions of the Barrovian facies series (Miyashiro 1961), and resulted in the formation of sillimanite, kyanite, staurolite and garnet which define metamorphic zones in the pelitic and quartzofeldspathic sequences of the Snowshoe Group (Fig. 17). 1. M l - EARLY METAMORPHISM A strong recrystallization fabric (SI), and tightly folded quartz veins are indicative of the deformation accompanying M l . The transposed quartz veins locally comprise up to 25-30% of the rock. Occasional relict fibrous quartz veins have been observed, and suggest that repeated crack-seal vein growth was operative (Ramsay 1980). The SI foliation has developed by the ubiquitous recrystallization of chlorite, biotite, plagioclase and quartz. Aluminosilicates such as staurolite, kyanite and sillimanite did not crystallize during this interval. The combination of fluid-aided deformation and the growth of micas in the absence of higher grade metamorphic minerals such as 57 58 staurolite and kyanite/sillimanite,. suggest that the M l metamorphism did not exceed greenschist grade. Plagioclase that outlines the SI foliation is albitic (An 2 5), and appears to have grown during the Phase 1 deformation (Fig. 18). Little evidence remains that would shed light on the relative intensity or variation in metamorphic conditions during the M l thermal event The intrusion of sizeable granitic bodies such as the Boss Mountain and Quesnel Lake gneisses during the mid-Paleozoic (approximately 320-340 Ma, J. K. Mortensen personal comm. 1985) may be linked to the metamorphic event. The inferred low to moderate temperatures of greenschist grade metamorphism, and the "two-mica" granite composition of the Boss Mountain gneiss (Chappell and White 1974), is compatible with a model for low grade metamorphism and subsequent intrusion of granite derived from a much greater depth. Contact metamorphic effects have not been observed in the metasedimentary rocks, which may suggest a relatively small temperature contrast existed between the two at the time of emplacement of the granite. 2. M2 - .JURASSIC MFTAMQRPHTSM The effects of the Jurassic metamorphism are the most evident in the Boss Mountain area. A readily discernable metamorphic gradient has been documented by the distribution of diagnostic mineral assemblages in the metapelitic rocks. Metamorphic grade generally increases from west to east in the area; the chlorite zone coinciding with the contact between Slide Mountain Group metavolcanic rocks and Triassic phyllite (Fig. 17). Sillimanite bearing feldspathic and pelitic schists comprise the highest grade metamorphic assemblages in rocks of the Snowshoe Group, which are confined to the southeast portion of the map area. Garnet Garnet comprises a range of crystal forms, from decomposed, resorbed crystals riddled by inclusions, to less poikiloblastic, nearly idiomorphic grains (Fig. 19). Typical N Omineca Belt contact Figure 17 - Distribution of metamorphic zones within the Boss Mountain area. The biotite-garnet zone approximately coincides with the tectonic contact between the Snowshoe Group and the Slide Mountain Group. 60 porphyroblast size ranges from 0.5 mm to more than 5 mm in diameter. Retrograde replacement of garnet by chlorite and biotite is common in samples throughout the area. Prograde metamorphism was initiated early in the Jurassic orogenic event, with garnet growing during, and outlasting the Phase 2 deformation. Textural relations reveal that the early foliation - SI, is frequently truncated by garnet porphyroblasts, and is preserved as an internal schistosity within these grains. Crystallization may have begun prior to, or early in the Phase 2 deformational episode, and terminated near the end of the deformation with the possible exception of some syn- to post-Phase 3 growth. Garnet grade rocks are best exposed near the Snowshoe Group/Slide Mountain Group contact in the west of the area - Unit W-12. A clear crosscutting relationship can be demonstrated for many of the porphyroblasts with respect to the S2 schistosity, but some examples are more equivocal, and a conclusive relationship has not been determined (Fig. 20). Magnetite, ilmenite and quartz form inclusion trails within the garnet, but continuity of these with the external foliation has not been established. Subsequent rotation and dissolution of garnet has apparently occurred in most of the samples. Garnet also exhibits concentric inclusion patterns of quartz, magnetite, ilmenite and subordinate micas and zircon. Staurolite Staurolite occurs throughout the map area, and invariably coexists with garnet, and at higher grades is present with kyanite and sillimanite in the metapelites. In hand specimen, the crystals are large, dark red to brownish in colour, and are frequently intergrown with garnet Grains up to 5 mm long are common. In thin section the crystals are straw, to deep yellow and strongly pleochroic. They display a variety of idiomorphic crystal forms and are frequently twinned. Poikiloblastic crystals are common, and contain inclusions of quartz, garnet and occasionally fine grained white mica. 61 I 2 HUB i Figure 18 - Sketch from photomicrograph of the SI foliation defined by quartz-feldspar layers, with laminations of chlorite, biotite and muscovite from sample 10, Unit E - l . Figure 19 - Sketches from photomicrographs of two representative thin sections in which garnet is present as a) resorbed poikiloblastic grains (sample 650, Unit W-12), and b) euhedral porphyroblasts relatively unaffected by later decomposition (sample 29, Unit E-5). 62 The staurolite is a syn-kinematic Phase 2 growth, that appears to postdate the height of deformation. Semi-alignment of the porphyroblasts with respect to the L2 linear structures, and deformation by F3 kinking and crenulation, demonstrates that the growth of this phase was prior to the onset of appreciable F3 folding (Fig. 21). Minor alignment of staurolite prisms sub-parallel to the L2 lineation has occurred in some examples. Kyanite Kyanite was generated in the hiatus between the F2 and F3 folding phases (Fig. 25). Nucleation of the kyanite took place within the SI compositional layering, and has overgrown the S2 schistosity. Absolute random orientation of kyanite with respect to the prominent L2 and L3 lineations is observed (Fig. 22b). Unit E-5 hosts abundant kyanite in an assemblage including staurolite and fibrolite. Large idioblastic grains, often twinned on (100), have grown in a completely random fashion across the S2 schistosity, and are extensively kinked by F3 deformation (Fig. 22). Large blades of kyanite up to 2 - 3 cm long, are present as porphyroblasts in the pelitic schist Kyanite is widespread within both pelitic schist and quartzofeldspathic rocks in the area. Inclusion of garnet crystals is common, and pseudomorphs of kyanite by muscovite have been observed. Sillimanite The maximum temperatures of metamorphism were reached during late to post-D2 times in rocks of the Snowshoe Group. The prograde metamorphism culminated with the the formation of fibrolite and partial melting of some of the metasediments in various localities including Boss Mountain (Pigage 1978; Fletcher and Greenwood 1979). The presence of voluminous granitic pegmatite is a reflection of the melting. Sillimanite is exclusively fibrolitic, and has been recorded in several paragenetic associations. 63 Figure 20 - a) Sketch from photomicrograph of late M2 garnet growth; garnet growing across quartz-feldspar domain in S2 crenulation fabric, kyanite zone, sample 53, Unit BM-8. b) Outcrop of SI surface, showing randomly oriented blades of kyanite affected by F3 crenulations. S3 1 mm Figure 21 - Sketch from photomicrograph of staurolite with pressure shadows and crenulation of the SI foliation due to F3 folding (sample 661-b, Unit W-12). 64 Figure 22 - a) Sketch from photomicrograph of kyanite crosscutting the SI and S2 fabrics and showing internal kinks and external crenulations due to F3 folding (sample 73, Unit B-7). b) Post-kinematic F2 kyanite in outcrop overprinted by F3 crenulations (note strong L3 lineation). 65 Figure 23 - Sketch from photomicrograph of fibrolite crenulated by F3 microfolds in pelitic schist, sample 29, Unit E-5. 66 Fibrolite growth postdated the critical period of deformation identified with Phase 2 structures (Fig. 23). Sillimanite is associated with aluminous pelitic schists in units E - l , E-5 and B-6A, in the core of the F3 Boss Mountain antiform south of M L Beisig (see metamorphic cross-section, in pocket). In unit E-5 , the fibrolitic rocks contain kyanite and garnet, with subordinate staurolite. Muscovite is always abundant, and is accompanied by quartz, plagioclase, biotite and oxides (incl. rutile?). The fibrolite bearing rocks of Unit E - l commonly contain garnet, but kyanite and staurolite have not been observed in any of these quartzofeldspathic schists and gneisses. The fibrolite occasionally shows crenulation effects due to F3 folding (Fig. 23), and has been locally replaced by retrograde chlorite. A late episode of garnet growth has been recognized very locally within the Snowshoe Group near Boss Mountain. F3 crenulation folding of the SI and S2 foliations has resulted in small scale crenulations, over which garnet has grown. The lack of significant distortion or flattening of the layering subsequent to the garnet growth, is suggested by the complete absence of strain shadows or rotation of the porphyroblasts. Small garnet crystals, 0.2 - 0.8 mm in diameter, show oriented inclusions of fine grained magnetite and ilmenite; the garnets reveal a crosscutting relationship to the F3 crenulations (Fig. 24). Porphyroblasts do not exceed 1 mm in diameter, and comprise nearly 10% of the sample. The extent and nature of the late garnet growth is not clear, the example cited is of very local extent, and no other examples of this type have been observed. In the west of the area, post-F3 biotite and muscovite porphyroblasts have also been observed as "static" growths across F3 microfolds (Fig. 24b). Figure 25 summarizes the relative timing of the growth of metamorphic minerals with respect to deformational episodes. 67 Figure 24 - a) Photomicrograph (crossed nichols) and sketch of late garnet growth across the hinges of F3 microfolds. b) Post-kinematic mica growing randomly across the hinges of F3 crenulations in pelitic schist, Unit W-12. crenulated muscovite porphyroblast (24b) 69 PHASE 1 PHASE 2 PHASE 3 syn post syn post syn post Chlorite Muscovite Biotite Plagioclase Garnet Staurolite Kyanite Sillimanite Figure 25 - Chart summarizing the relative timing of the growth of metamorphic minerals with repect to deformational episodes in pelitic schist and gneiss of the Snowshoe Group. 70 3. M2 METAMORPHIC REACTIONS - METAPEUTES The metamorphic mineral assemblages comprising the zones illustrated in Figure 17 were determined by field and petrographic examination of key samples from the Snowshoe Group schists and gneisses. The matrix of these rocks is dominantly quartz, muscovite, biotite and plagioclase, with porphyroblasts of garnet, staurolite and kyanite ( ± sillimanite). Mutual contact between all mineral phases has been observed, and no distinct mineral segregations are present Correlation between phases of deformation and metamorphic textures gives a reasonably precise picture of the relative timing of crystallization of these minerals. Metamorphic zones developed during the second metamorphic event (M2) are defined in the Snowshoe Group, and within a limited portion of the cover rocks including both the Slide Mountain Group and Triassic metapelites. The mineralogy of representative samples from each of the metamorphic zones are given in Table 1. Minerals in the reactions mentioned in the text are assumed to be ideal end members. Their general compositional formulae are given in Table 5. Chlorite - Biotite Zone Metavolcanics of the Antler Formation and Triassic pelitic rocks lie outside the zone in which garnet first appears (Fig. 17). Bulk rock composition alone does not appear to be responsible for the apparent increase in metamorphic grade in the vicinity of the Snowshoe Group/Slide Mountain Group contact Rocks of the Snowshoe Group are biotite-garnet grade immediately below the contact with the Slide Mountain Group in the map area. Biotite appears in the metavolcanics near the contact between the Antler Formation and Snowshoe Group metasediments. Hornblende, chlorite and clinozoisite are major phases within the Antler metavolcanics near this contact. 71 Sample 29 23 71 60 21 661 115 644 126 c 2 o, a a> o 667 = o to o CO 2 669 CO a 3 o O 9 O (0 * o c CO Metamorphic Rock zone unit k y - s l E-4 k y - s l B-6B ky BM-9 ky B-6A ky B-6B g t - s t W-12 g t - s t W - l l gt W-12 gt W-12 Phases present q z - m s - b i - p l - g t - s t - k y - s l - m t . q z - m s - b i - p l - g t - k y - s l - m t q z - m s - c h - b i - p l - g t - k y - m t q z - m s - c h - b i - p l - g t - k y - m t - i m q z - m s - b i - p l - g t - s t - k y - m t q z - m s - c h - b i - g t - s t - p y - m t q z - m s - b i - p l - g t - s t - m t q z - m s - c h - b i - p l - g t - p y q z - m s - c h - b i - p l - g t - m t b i A n t l e r q z - c h - b i - c c - p l - h b - e p - p y F o r m a t i o n c h T r i a s s i c qz-ms-ch-gp-py P h y l l i t e Table 1 - Mineral phases present in representative samples from each of the M2 metamorphic zones. In addition to the phases shown, many of the samples contain varying amounts of apatite, zircon, sphene, and tourmaline. 72 I 1 mm Figure 26 - Sketch from photomicrograph of sample 647, showing fractured garnet partially replaced by quartz and a mantling of retrograde chlorite. 1 mm Figure 27 - Sketch from photomicrograph of sample 21; garnet being consumed by staurolite reaction (3). Fragments of originally continuous garnet (note foliation internal to garnet being truncated) enclosed by staurolite with quartz filling the embayments in the porphyroblasL 73 Biotite - Garnet Zone Typical assemblages in rocks of the garnet zone consist of: quartz + plagioclase + K-feldspar + muscovite + chlorite + biotite + garnet + magnetite. Additional phases include ilmenite, tourmaline, sphene, zircon, hematite and apatite. Graphite is uncommon within most of the units except fault-related lithologies (Fig. 11). Chlorite has broken down in the presence of muscovite to form almandine-rich garnet. The reaction: (1) 3 chlorite + muse + 3 quartz = 4 almandine + biotite + 12 H 2 0 is potentially applicable to the assemblage and texture in sample 647 (Fig. 26). In this example, chlorite occurs as two distinct forms. Matrix chlorite is mostly lath-like, green pleochroic grains, showing blue interference colours. This is interpreted as prograde chlorite formed during M2. Retrograde chlorite and biotite intergrowths (on 001) are common, and slender grains of ilmenite are typically intergrown with these. Finer grained, less euhedral chlorite is associated with the garnet porphyroblasts as retrograde alteration. Garnets are resorbed to some extent and have retrograded to chlorite, quartz and biotite. Biotite and muscovite make up 30-40% of some samples in this zone. Prograde chlorite is abundant, and comprises up to 3 to 4% of the rock. 74 Garnet - Staurolite Zone Coexisting garnet and staurolite have been mapped in units of the Western Succession, where their coexistence is interpreted as representing the "staurolite in" isograd (Winkler 1979). The feldspathic schists of units B-6 and B-7 exhibit modally abundant staurolite, as opposed to the poorly developed crystals seen in units W - l l and W-12, which contain less than 1% plagioclase. The matrix of the latter units is composed mainly of quartz, micas, and some K-feldspar. The staurolite has crystallized from the groundmass, and is rarely seen in contact, with garnet Staurolite crystals are sub-idiomorphic to xenomorphic, and extremely poikiloblastic. The staurolite isograd is First encountered in the upper part of Unit W-12 approximately 600 m east of the contact between the Snowshoe Group and the Slide Mountain Group (Fig. 17). The textural relations preserved in specimens 115 and 450 suggest the following reaction may be responsible for the first appearance of staurolite. (2) chlorite + 3 muscovite = staurolite + 3 biotite + 7 quartz + 14 H 2 0 (Hoscheck 1969) The staurolite-in reaction defines an isograd in the compositionally similar rocks in the western part of the map area (Units W - l l and W-12). Reaction (2) is potentially applicable to the rocks units in the Western Succession, where chlorite coexists with staurolite, and garnet appears to be unaffected by reaction. The reaction involving the breakdown of garnet in the generation of staurolite is shown in Figure 27. (3) 4 chlorite + 9 muscovite + 5 almandine = 2 staurolite + 9 biotite + 12 quartz + 12 H 2 0 (sample 21 - Unit B-6B, sample 73 - Unit B-7) The reaction involving garnet is inferred from textural relationships observed in samples from the staurolite-kyanite zone. 75 1 mm Figure 28 - Kyanite formed by the breakdown of garnet (reaction 4). Fine grained biotite and pools of quartz have crystallized adjacent to the garnet 76 i 0.2 mm | (a) (b) Figure 29 - Sketches from photomicrographs of sample 29, a) muscovite replacing kyanite as reaction rims in quartz, and b) fibrolite intergrowth with biotite. These two textures combine to produce the reaction : kyanite = sillimanite (reaction 5 + reaction 6). 77 Temperatures calculated from garnet-biotite geothermometry (Table 2), indicate minimum metamorphic temperatures of 485° to 515 ° C in the staurolite zone for sample 115. Staurolite - Kyanite Zone The growth of kyanite has taken place at the expense of muscovite, biotite and possibly staurolite. Sample 21 contains several examples of incipientiy developed kyanite in a dominant! y staurolite-rich feldspathic schist Fine grained biotite is present as post-kinematic (F2) crystals growing in the vicinity of the kyanite. (4) muscovite + almandine = 2 kyanite + biotite + quartz Reaction 4 is inferred from textures seen in sample 21 from the staurolite-kyanite zone. Late syn-F2, to post-F2 garnet has reacted to form post-F2 kyanite (Fig. 28). The grain size of kyanite ranges from large idioblastic crystals (Unit B-7), to small grains such as in Unit B-6B. Kyanite - Sillimanite Zone Fibrolite has been generated by the breakdown of muscovite, biotite and possibly by staurolite to a lesser extent Units E-5 and B-6A comprise aluminous schists which host considerable amounts of fibrolite, in an assemblage including both kyanite and staurolite. In the feldspathic schists of Unit B-6B, only small amounts of sillimanite have been produced. Muscovite pseudomorphs of kyanite, and biotite overgrowths of garnet have been recognized in the sillimanite zone. The muscovite pseudomorphs are represented by the hypothetical reaction (5) (Fig. 29a). (5) 3 kyanite + 3 quartz + 2 K + + 3 H 2 0 = 2 muscovite + 2 H + (6) 2 muscovite + 2 H* = 3 sillimanite + 3 quartz + 2 K + + 3 H 2 0 (Carmichael 1969) Needles of fibrolite are observed to project up to 0.5 mm into aggregates of clear quartz from the clusters of muscovite and biotite. Where the sillimanite has 78 grown in this manner, combination of reaction (5) with reaction (6) would complete the transition kyanite=sillimanite (Carmichael 1969). Fibrolite and biotite have developed as intergrowths of felted, fine grained material separate from the kyanite (Fig. 29b). The biotite-Fibrolite intergrowths are extremely common in samples from the kyanite-sillimanite and sillimanite zones (samples 23, 29, 10 and Fig. 17). Reactions where biotite breaks down to form sillimanite have also been described in the same fashion as above by Yardley (1977). The local migration of components leads to the net reaction of the kyanite to sillimanite transition. The frequent intergrowths of biotite and Fibrolite in the presence of quartz in the kyanite-sillimanite zone suggests that biotite is also a product of the sillimanite formation (Carmichael 1969). In summary, sillimanite is formed mainly by the reaction of muscovite and biotite. Fine, recrystallized white mica is present with kyanite, and needles of fibrolite have grown along the (001) cleavage traces of large idiomorphic muscovite, and as felted masses intergrown with biotite. This is probably a result of kyanite reacting to form sillimanite. Small remnants of staurolite have been recognized in what appears to be staurolite pseudomorphed by muscovite, with intergrowths of biotite. Within these patches of mica, there is an abundance of fibrolite. The staurolite may have began breaking down in the kyanite zone, and has in rare instances been nearly completely replaced by white mica. The growth of Fibrolite in the absence of kyanite is observed in Unit E - l (sample 10). The assemblage kyanite + sillimanite, and staurolite + kyanite + sillimanite first appears in Unit E -5 of the Eastern succession. This may mark the presence of an additional metamorphic zone characterized by the assemblage : quartz + plagioclase + muscovite + biotite + garnet + sillimanite. PRESSURE 5 kbar A A 700 A • _ • o A <P • o 111 OC • D 600 • < cr LU Sample Zone Q_sill A - 10 LU • - 21 s t - k y • _ _ 1 . 500 • • • - 53 s t - k y • - 115 g t - s t i . i i i i i 1 1 1 _ • 0.10 0.11 0.12 0.13 0.14 0.15 0.16 0.17 0.18 Mg/Fe GARNET Figure 30 - Temperature calculations from garnet-biotite geothermometry in the Boss Mountain area. Temperatures were calculated using the model of Newton and Haselton (1981). -4 80 Conditions of metamorphism - metapelites The minimum metamorphic temperatures during progressive (M2) metamorphism were estimated using coexisting biotite and garnet compositions and inferred mineral reactions. The calibration of the temperature dependence of Mg-Fe partitioning between garnet and biotite by Ferry and Spear (1978), leads to an expression for the equilibrium distribution constant for Fe-Mg. The derivation considers the partitioning of Fe and Mg between garnet and biotite, assuming ideal mixing between the two phases. Pigage and Greenwood (1982), have demonstrated that the effects of other components, specifically Ca and Mn, will cause errors in the temperature calculation. Garnet crystals from the area are compositionally zoned, with an increase in Mg at the rims, and a greater concentration of Ca and Mn in the cores. This provides evidence that garnet grew during prograde metamorphism, and did not reequilibrate throughout the event (Tracy et al. 1976). This seems likely since some temperature calculations have yielded lower than average or higher than average temperatures for rocks whose assemblages would suggest otherwise (Spear and Selverstone 1983). A modified equation by Newton and Haselton (1981), which considers the effects of the Ca content of garnet, has been used in the garnet-biotite temperature calculations. Samples from four different localities have yielded temperatures ranging from 500 ° C in the staurolite zone, to more than 700 ° C in the sillimanite zone (Fig. 30). The mineral assemblages in the kyanite-sillimanite, and sillimanite zones are characteristic of Bathozone 4 - 5 of Carmichael (1978; see Fig. 33). This gives an approximate pressure range of 5 - 6 kb, which is compatible with metamorphic pressures for assemblages of this type from the region (Pigage 1978, Getsinger 1985). 81 4. C A L C - S I L I C A T E REACTIONS Reaction zones between carbonates and quartzo-feldspathic schists and gneisses have been recognized throughout the area, and are best developed in the high grade metamorphic zones. Coarse calcite marble with varying amounts of quartz, plagioclase (Anjo to An 6 0 ) , K-feldspar, micas and opaques (including Fe-sulphides), have marginal zones composed of clinozoisite (± epidote), chlorite, Ca-amphibole and diopside where they contact the pelitic rocks. Marble units are rarely greater than 3 or 4 m thick, and are accompanied by calc-silicates usually less than 1 m thick. The width and extent of the reaction zones is probably a function of the original thickness of the interlayers of carbonate and pelite, "mixing" of the contact due to prior deformation and mass transfer between the carbonates and pelites (Thompson 1975). Clinozoisite or epidote bearing assemblages characterize the calc-silicate rocks in many locations. Garnet is rare, and occurs where calcite comprises only a small percentage of the rock. Scapolite is characteristic of the highest grade rocks in the kyanite-sillimanite zone. Textural evidence from a sample of siliceous carbonate within the kyanite zone (Fig. 31), suggests that clinozoisite grew after plagioclase (Ferry 1976), which then subsequently broke down to calcic amphibole (hornblende?), and eventually clinopyroxene as the metamorphism progressed. Tourmaline has also grown as porphyroblasts, with some large zoned grains 1 - 1.5 mm long overgrowing the clinozoisite. Diopside occurs as large idioblastic to sub-idioblastic grains within layers of clear, glassy quartz. These grains are typically 1 - 2 mm long. Granular aggregates of diopside (Fig. 31) have crystallized at the interface between layers of quartz and calc-silicate minerals, probably due to local mobility of the fluid phases at sites of reaction (ie. qtz-calcite veins). These quartz (+ carbonate) veins represent the same type of SI features as illustrated in Figure 6a and Figure 34. The veins seem to behave as reaction "barriers", especially for the diopside, which has been observed 82 I 1 mm Figure 31 - Sketch from photomicrograph of calc-silicate from Unit B-6B, comprising clinozoisite + plagioclase + hornblende + diopside in addition to quartz and minor calcite. 83 Figure 32 - Scapolite bearing marble, sample 24, Unit B-6B. a) Schematic A b - A n - C c plot for the inferred scapolite producing reaction. The estimated composition of the scapolite is Me*,, the plagioclase is approximately A n 5 J . The sample is from the sillimanite zone. Unit B-6B. b) Sketch from photomicrograph of sample 24, showing the porphyroblastic scapolite in quartz-calcite layers which also include biotite, white mica and minor diopside. 84 concentrated at the margins of these 1 - 5 mm thick veins, and also completely enclosed within them. Mobility of the reactants at this interface may have facilitated the crystallization of diopside during vein growth. Fine grained carbonate appears to have been partially consumed by reaction, and comprises only 5-10% maximum of this type of calc-silicate rock. Within, and adjacent to the kyanite-sillimanite zone, marble units have been metamorphosed under very high temperatures and pressures. Assemblages within the marble consist of : quartz + calcite ( +dolomite?) + biotite + white mica + K-feldspar + plagioclase + diopside + scapolite. Minor phases include apatite, sphene, chlorite, tourmaline and magnetite: The scapolite occurs as sizeable porphyroblasts, 1 - 2 mm across, which have been susequently corroded and commonly replaced by quartz and biotite (Fig. 32b). Scapolite bearing rocks are relatively free of epidote group minerals, do not contain amphibole, and have only small amounts of feldspar (less than 1% plagioclase + K-feldspar). Composition of the scapolite in sample 24 has been estimated by birefringence and refractive index (Deer, Howie and Zussmann 1977). The composition is between Me 8 0 and M e 9 0 ; close to the calcium end-member meionite (Me). Plagioclase is approximately A n 5 5 , from optical determinations on several grains in sample 24. From the plagioclase and scapolite compositions within the high grade micaceous marbles, a schematic compositional triangle was constructed to illustrate the reaction assemblage for: calcite + plagioclase = scapolite The Ca end-member reaction is : (7) 3 anorthite + calcite = meionite (Orville 1975; Goldsmith and Newton 1977). Reaction (7) defines an isograd (Ferry 1976), which, when used in conjunction with the garnet-biotite temperature provides additional constraints on the P -T conditions of the M2 metamorphism. 85 Pigage (1985 in preparation) has compiled compositional and experimental data on coexisting scapolite and plagioclase from various authors, and from the Azure Lake, B.C. area (Pigage 1978). Plagioclase composition within the scapolite bearing marble (sample 24, approximately An 5 5 ) indicates a maximum metamorphic temperature of about 685 ° C to 700 ° C for the sillimanite zone in the Boss Mountain area (Pigage 1985). This temperature should be interpreted as a crude estimate, since conclusive compositional information is lacking, and the effects- of additional components such as Cl are poorly understood (Orville 1975). 5. DISCUSSION Metamorphic mineral zones have been defined in the Boss Mountain area, where metapelites and quartzo-feldspathic schists have undergone progressive Barrovian-type metamorphism. The development of mineral parageneses through increasing pressure and temperature, can be described by reactions crystallizing garnet through sillimanite. Conditions of the Jurassic metamorphism (M2), have been evaluated using three lines of evidence. 1. By metamorphic mineral assemblages characteristic of the observed metamorphic zones. 2. Garnet-biotite geothermometry from the different metamorphic zones. 3. Structural correlations between folds, foliations, and the relative timing of metamorphic mineral growth. Metamorphic isograds Isograds have been identified within the area by the first appearance of garnet, staurolite, kyanite and the transition from kyanite to sillimanite. Garnet is first generated in the metapelites by reaction (1). Chlorite is consumed in the prograde 86 500 600 700 800 TEMPERATURE (°C) (a) Figure 33 - a) Inferred mineral reactions in the Boss Mountain area. The aluminum silicate diagram is that of Holdaway (1971). The bathozones of Carmichael (1978) are shown, with the approximate position of rocks from the sillimanite zone in the Boss Mountain area constrained by temperatures calculated for rocks of the kyanite-sillimanite zone, b) Schematic illustration of A F M assemblages in pelitic rocks of the Snowshoe Group. Staurolite zone Kyanite zone b i - b i o t i t e g t - g a r n e t s t - s t a u r o l i t e Sillimanite zone (33b) 88 reaction. Tne "staurolite in" isograd has been modelled by reactions (2) and (3), indicating metamorphic conditions of approximately 5 - 6 kb, and 560 ° - 600 ° C (Hoscheck 1969; Richardson 1968) prevailed in the western part of the area during the peak of M2 metamorphism (Fig. 17). Reaction (2) was deduced from assemblages and textures in rocks of the garnet-staurolite zone. The reaction: kyanite = sillimanite, is responsible for the presence of fibrolite in metapelites in the eastern map area. This transition is interpreted to be the resultant of several coupled reactions, including the breakdown of kyanite and muscovite. Conditions in this zone are estimated at 5 - 6 kb pressure, and approximately 680 ° - 690 ° C (Holdaway 1971; Thompson 1976). Reaction (7) : plagioclase + calcite = scapolite, from impure marble in the kyanite/sillimanite zone, suggests a temperature of about 700 0 C (Pigage 1985). An upper limit on the temperature, is the coexistence of staurolite and sillimanite, which constrains the temperature as being below 700° C (Pigage and Greenwood 1982; Archibald et al. 1983; see Fig. 33). Evidence for this reaction is lacking, however, and it remains unclear whether the "staurolite-out" reaction has occurred. Fibrolite has not been observed growing within staurolite crystals, but it may be possible that staurolite has broken down in the kyanite-sillimanite zone. These isograds reflect contrasting conditions of metamorphism within rocks of similar composition, in stratigraphically distinct units. The reactions are inferred, however, and conclusive evidence from assemblages and textures is lacking. Garnet - biotite geothermometry Garnet-biotite thermometry has yielded temperatures from the garnet-staurolite, staurolite-kyanite and kyanite-sillimanite zones. Table 2 shows the results from calculations employing the model of Newton and Haselton (1981). Average values for Mg/Fe ratios of garnet and biotite are given. Temperatures were calculated at both 5 and 6 kb, but have only been reported at 5 kb since they do not vary by more than 10-15 °C/kb . 90 Table 2 Sample Metamorphic zona 10 SI 2 1 Ky-SI 53 Ky 1 15 St T CO 689 665 6 29 505 Mineral composition data for garnet and biotite is presented in Appendix B, as are operating conditions used in the electron microprobe analysis. Problems with local disequilibrium and machine drift may be responsible for inconsistencies. Zoning of garnet was considered in the geothermometry study, and care was taken to calculate temperatures using only matrix biotite in contact with garnet rims (ie. those that appeared to be in textural equilibrium). The assumption is that these represent the greatest degree of chemical equilibrium (Spear and Selverstone 1983). Structural and metamorphic correlations The relationship between microstructures and metamorphic textures has received a gTeat deal of attention recently (Rosenfeld 1970; Vernon 1977; Vernon and Flood 1979; Bell and Rubenach 1983). The development of syn-metamorphic Phase 2 structures, including differentiated crenulation cleavage, metamorphic layering and shear fabrics, are suggestive conditions of medium to high grade metamorphism (Etheridge et al. 1983, Marlow and Etheridge 1976). 91 Quartz ribbons, and dissolution of porphyroblast margins, particularly garnet, substantiate this (Etheridge et al. 1983). Dislocation creep and grain boundary diffusion are mechanically compatible with the rheology of Snowshoe Group rocks during Phase 2 deformation under P -T conditions of the M2 metamorphic regime. V. STRUCTURAL SYNTHESIS An analysis of the structural relationships between successive deformational -phases has been made in the Boss Mountain area. Mesoscopic overprinting relationships from field observations were sufficient to establish a structural chronology, which is supported by microstructural evidence. 1. BASEMENT - PHASE 1 The deformation and metamorphism responsible for the regionally pervasive SI foliation, is interpreted to have resulted from both homogeneous and inhomogeneous local strain regimes. Compositionally differentiated layering related to bedding, has developed largely by a fluid-aided mode of deformation. This is supported by the high proportion of veins and segregations which crosscut the compositional layering and are folded about SI /LI . At low grade metamorphic conditions, critical values of differential stress probably lead to the development of differentiated layering by hydraulic fracturing and pressure solution (Etheridge 1984; Norris and Henley 1976). Solution transfer initiated at low temperatures is also likely to cause significant strain during the dewatering stages of deformation (Durney 1972; Williams 1972). The regionally extensive foliation associated with the Phase 1 deformation is exceptionally well preserved in much of the Boss Mountain area, and suggests that a widespread, nearly homogeneous bulk strain has affected the rocks (Talbot and Hobbs 1968). The parting along, and at low angles across bedding surfaces, may have been initiated in the early stages of FI folding and cleavage formation (Fig. 34). The strain would be the result of a combination of crystal deformation processes, and dilation of the preexisting bedding plane anisotropy. Figure 34 illustrates a possible sequence by which Phase 1 deformation initiated and progressed, and wherein the formation of vein arrays, cleavage and folding was accomplished. Internal crenulations are not associated with the SI foliation, and may 92 93 Figure 34 - Sequential development of hydraulic fractures, and the eventual development of SI-conformable quartz veins during Phase 1 deformatioa The three stages include : a) tensile fracturing, preferably along the bedding anisotropy, or at small angles across it, b) dewatering and cleavage development, and c) distortion and eventual transposition of bedding and veins by continued folding. 94 indicate a lack of this type of intense microfolding during Phase 1 deformation. If a relationship exists between deformation and intrusive activity, it may be possible that the F l foliation and folding occurred during the emplacement of the mid-Paleozoic granitic gneiss. The F l folding may have not reached the stage of development characterized by the F2 folding. Crenulation cleavage would indicate a more advanced stage of fold-related deformation had been achieved. Greenschist grade metamorphic conditions have been inferred by the texture and crystallization of the SI foliation constituents, dominated by quartz, plagioclase, chlorite and micas. Under such metamorphic conditions, macroscopic tensile fracturing of rocks "bleeding off" metamorphic fluids would be an expected mode of failure. This would serve to lower the pore pressure of fluids, and allow deposition of any transported solute in areas of lower effective stress (Elliot 1973; Etheridge et al. 1984). 2. BASEMENT - PHASE 2 / COVER - PHASE 1 Relationship between F l and F2 folds Superposition of mesoscopic F2 folds upon F l is occasionally observable in some outcrop exposures (Fig. 35). The occurrence of such structures is recognized in all parts of the map area within rocks of the Snowshoe Group, but are absent in the Boss Mountain gneiss. This is probably due to the ductility contrast which must have existed between the two rock types during F2 folding. Internal deformation of the gneiss does not appear to be associated with the same style of fold structures seen within the metasediments. Deformation affecting both the basement and the cover has been studied in an attempt to relate these features to a tentative structural history. F2 fold data has been plotted on equal area projections to illustrate the geometry of F2 planar and linear structures. F2 microfolds and schistosity were examined in thin section by viewing them in sections cut normal to the SI compositional layering containing L2 (L2 - basement, 95 (b) Figure 35 - Refolded FI folds, a) approximate coaxial refolding of FI isoclinal fold in marble of Unit B-6B, by tight F2 fold, showing "S" symmetry typical of F2 folds east of ML Beisig (view is down plunge to the north), b) FI isocline (below pencil) within transposed layering of quartzofeldspathic gneiss, Unit BM-10, being refolded by tight F2 minor folds. The F2 overprinting typically tightens the FI fold limbs and obliterates the original FI geometry. 96 L I - cover). This was taken to be the approximate orientation of the X Y plane of the Phase 2 Finite strain ellipse (Darot and Boudier 1975; Williams 1976). Other sections were cut normal to the compositional layering and L2 lineation (approximate orientation of X Z plane). Samples from the Boss Mountain gneiss, Snowshoe Group schist and gneiss, and Antler Formation metavolcanics were studied. Elongation and occasional rotation of grains and syn-F2 pressure shadows is evident in all specimens when viewed in sections containing L2. Quartz-feldspar porphyroclasts in samples of Boss Mountain gneiss indicate that these augen behaved more rigidly than the ductile micaceous matrix (Fig. 36). Fine grained augen within mylonitic quartzite in Unit W - l l , and in thin sheets of "quartz eye" augen gneiss, reveal similar .elongation of porphyroclasts and their syn-F2 pressure shadows. Considering that S2 is a plane of compressive strain, the elongation fabric is a result of incremental stretching parallel to L2 during this deformation (ie. plane strain with a considerable constriction component). Al l deformed objects, such as elongate grains and pressure shadows, are contained within the SI schistosity. Syn-kinematic dynamic recrystallization of porphyroclasts is evident in samples of the Boss Mountain gneiss, and some mylonitic quartzite. Apparently, these did not simply "passively" rotate during the ductile deformation, but were also involved in grain size reduction (Berthe et al. 1979; Simpson and Schmid 1983). Relationships between F2 foliation and porphyroblast growth Microscopic textures observed in the schists and gneisses in the Boss Mountain area have facilitated determination of the relative liming of metamorphic mineral growth. The presence of oriented inclusions within large metamorphic grains often indicates the timing of their growth with respect to foliations. A relationship between the porphyroblasts and the S2 schistosity allows a further constraining within the F2 folding episode. The various stages of S2 cleavage development during progressive deformation are preserved as internal fabrics within many of the M2 porphyrobasts 97 (Fig. 37). The SI foliation has not been observed in association with any type of crenulation fabric, which aids in the determination of S2 schistosity/porphyroblast liming relations. This also gives some insight as to the nature of progressive foliation development in some of the metasedimentary rocks. Classes of inclusions within porphyroblasts Three fundamental types of inclusions within porphyroblasts have been recognized (Fig. 37). These have been chosen due to the frequency of their occurrence throughout the area. 1. ' Matrix inclusions 2. Oxides and opaque inclusions 3. Other porphyroblasts as inclusions (1) Matrix inclusions These are the most common, and are composed mainly of quartz, feldspar, micas and other accessories such as sphene and zircon. They are present within porphyroblasts of both garnet and staurolite, and frequently define an internal planar fabric. These inclusion "trails" may be planar, or exhibit varying degrees of distortion by pre or syn-metamorphic deformation. It has been postulated that syn-metamorphic sigmoidal inclusion trails may represent the progressive overgrowth of matrix material in pressure shadows adjacent to the growing crystal (Elliot 1972; Rosenfeld 1972; Vernon 1977). Inclusions which are much finer grained than their host matrix, preserve the pre-annealing texture of the groundmass (Fig. 38). Inclusions represent either SI or an intermediate stage of S2. S2 crenulation cleavage is well developed in the matrix as a differeniated crenulation schistosity. There are also many examples in which garnet crystals have enclosed an earlier crenulated foliation (Fig. 37a). This represents the advanced stages of S2 crenulation cleavage formation, and indicates that these garnets are late F2 (Fig. 37b). 98 Figure 36 - L2 elongation fabrics, a) Sketch of L2 elongation lineation as indicated by asymmetric augen of recrystallized quartz and feldspar within Boss Mountain gneiss, sample 108. b) Photograph of Kspar augen within mylonitic quartzite, Unit W - l l , showing similar elongation parallel to L2. 99 Figure 37 - Sketches from representative thin sections illustrating the three types of inclusions common in porphyroblasts in the area, a) Inclusions of Fine grained quartz defining an internal crenulation fabric, sample 21, Unit B-6B. b) Elongate magnetite grains within M2 garnet, sample 649, Unit W-12. c) Staurolite porphyroblast overgrowing garnet in sample 21, Unit B-6B. 100 Figure 38 - Garnet porphyroblast showing discordant internal and external schistosities, sample 115, Unit W - l l . Matrix consists of coarse grained quartz, feldspar and micas. 101 (2) Oxides and opaque inclusions These include magnetite and ilmenite, plus graphite dust and stringers (typically in biotite) which also show preferred orientation. Pyrite is occasionally present, but is most common in rocks of the cover sequences, which typically have poorly developed porphyroblasts, and are not emphasized in this section. Opaque minerals are rare as foliate grains within the matrix of rocks in the area. The presence of closely spaced, aligned, opaque inclusions within garnet and occasionally staurolite, is evidence that these grains predate the redistribution or consumption which has eliminated them from the host rock matrix. F3 microtextures do not involve the preferred orientation of opaque minerals in most instances, thus implying that these grains represent either the SI or early S2 foliations (Fig. 37b). (3) Overgrowths of early porphyroblasts Mineral reactions causing the overgrowth of porphyroblasts by others are developed in response to increasing temperature. Particularly good examples occur in the staurolite-kyanite zone (Fig. 37c), where early M2 garnets have in some instances been completely enclosed by large sub-idiomorphic staurolite crystals. Staurolite, and occasionally kyanite crystals frequently envelope xenoblastic garnet grains. In some examples, a chemical reaction relationship has been deduced, others do not seem to have developed by any direct replacement Prograde overgrowths are a fairly straightforward means of assigning a relative mineral growth sequence, and relating this to the fabrics present Both staurolite and kyanite have randomly overgrown the S2 fabric, and therefore postdate i t Many of these crystals contain coarse grained matrix inclusions (Type 1), besides the frequent garnet inclusions. Sillimanite similarly grows within and across a coarse grained, foliated matrix (SI and S2 schistosities). Polygonized quartz which has recrystallized from numerous quartz subgrains is common. These relationships are suggestive of a post-F2 annealing prior to the peak of M2 metamorphism. F2 folding had, in fact ceased by this time. 102 Figure 39 - Four stages of progressive development of the S2 crenulation cleavage have been recognized in the area. Stage a) incipient mica fabric typical of the kyanite-sillimanite zone, b) open crenulations with sparse mica growth parallel to F2 crenulation axial planes, c) well developed penetrative crenulation cleavage accompanied by recrystallization of quartz-feldspar and mica domains, and d) complete differentiation of the S2 crenulation fabric with rare preservation of crenulation hinges (i.e. "differentiated crenulation cleavage", after Williams 1972). 103 Coarsening of matrix phases has occurred in samples of all metamorphic grades due to prograde recrystallization associated with the M2 metamorphism. Therefore, many of the M2 metamorphic minerals in which inclusions are finer grained than the recrystallized matrix, were likely inherited into the porphyroblast at an earlier stage in the S2 formation. Phase 2 - coaxial and non-coaxial strain Heterogeneous strain is indicated by the variable development of F2 folds and schistosity throughout the area. A general increase in penetrative mica growth and closer cleavage spacing of the S2 schistosity is observed from east to west in the area. This trend runs counter to the increase in metamorphic grade. Locally, the development of the S2 schistosity appears to intensify with decreasing metamorphic grade. Four stages of progressive development of the S2 crenulation foliation have been recognized in the area (Fig. 39). The stages of F2 crenulation development may be modelled by considering the Phase 2 strain ellipsoid in relation to the stages of cleavage formation illustrated in Figure 40. This model for crenulation cleavage formation requires a small amount of incremental shortening before flattening will dominate the strain configuration. Deformation will then proceed, with the crenulation cleavage (S2) nearly approximating the X Y plane of the incremental strain ellipse as long as the preexisting anisotropy undergoing shortening (in this case SI) remains in the field of incremental shortening (Gray and Durney 1979). From the lack of asymmetric pressure shadow mineral growths, it is inferred that incremental flattening strain was important, at least in the later stages of F2 folding. Locally in the Boss Mountain area, Phase 2 bulk strain may be the result of considerable coaxial increments. This is true especially for those rocks proximal to the tectonic contact between basement and cover. Phase 3 deformation, by contrast, appears •to be compatible with a simple shear model. Figure 40 - L^agrarnrnatic representation of the incremental strain ellipsoid during S2 crenulation formation (after Gray and Durney 1979). S2 approximates the X Y plane of the ellipsoid at an early stage in the cleavage development 105 Recent models for deformation in other portions of the Shuswap and Monashee Complexes have invoked a simple shear strain mechanism, based upon mylonitic fabrics, and rotational structures (Brown and Murphy 1982; Rees and Ferri 1983). The Phase 2 deformation in the Boss Mountain area appears to have been associated with locally high flattening strains, where inhomogeneous strain distribution has affected the cover rocks adjacent to the tectonic contact, and has also acted upon rocks in the basement toward the western margin of the map area. Snowshoe Group rocks further to the east may have suffered a Phase 2 deformational history more equivalent to simple shear strain. "Flattening fabrics" are a phenomena recognized mainly in the west, and seem to indicate a contrasting strain history across the area. Porphyroclasts and pressure shadows - Antler Formation The earliest recognizable foliation (SI) in cover rocks of the Antler Formation is equivalent to the second schistosity present within the Snowshoe Group and in the Boss Mountain gneiss (S2). This provides, the opportunity to compare and relate structures within the basement to those in the cover. Pyrite, K-feldspar and plagioclase crystals show extensive pressure shadows related to Phase 1 cover deformation. Euhedral pyrite crystals exhibit well-developed pressure shadows of quartz and chlorite (Fig. 41). The pressure shadow minerals are fibrous growths which grew syn-kinematically during FI deformation. The pressure shadows can be used as incremental strain indicators if they have not been distorted or extremely recrystallized during later events (Durney and Ramsay 1972). Samples from the Antler Formation have pressure shadows developed adjacent to euhedral and subhedral crystals of pyrite. The matrix of the rocks is fine-grained quartz, plagioclase and needles of hornblende. The pressure shadow samples have been sectioned normal to the compositional layering; both parallel and perpendicular to the LI lineation. 106 Pressure shadows from the specimens have planar Fibres or lateral boundaries, indicating coaxial deformation during incremental and progressive straining (Durney and Ramsay 1972; Gray and Durney 1979; Ramsay and Huber 1983). The pressure shadows have grown antitaxially, approximately parallel to the principal extension direction l + eu Pressure shadow growths reach lengths of up to 5 mm, adjacent to crystals 2 - 3 mm across. The pyrite grains range from less euhedral types, 0.1 - 0.5 mm, to euhedral crystals 1 - 2 mm (Fig. 41). The spacing of the SI foliation in the specimens is always less than the smallest pyrite grains examined. The majority of pressure shadows in the cover rocks show very little apparent rotation in sections containing LI , as do most examples from the basement rocks. Phase 2 faulting Imbrication of the western margin of the area presumably related to the bounding fault between Antler and Snowshoe rocks, is shown in Figure 10. The bifurcating nature of the faults has resulted in packages of rock that isolate and repeat adjacent stratigraphy in rocks of the Snowshoe Group. Within the fault-bounded packages, there frequently occurs a sooty, graphite-muscovite-quartz phyllite or schist, of proposed Triassic age. S2 cleavage is recognized as being axial planar to tightly transposed folds within these zones (Fig. 42). Deformed linear features within the phyllite have been folded by F3 crenulations and kink-style folds, indicating that these lineations predate Phase 3 deformation. Slickenside and mineral lineations at high angles to L2 fold axes, may indicate a general direction of Phase 2 movement of the faults. Fibrous chlorite and quartz Filled seams parallel to the S2 layering suggest that dilation of S2 has allowed veins to permeate the schistosity. This must have occured subsequent to the M2 metamorphism since garnet porphyroblasts show effects of the chloritization associated with the veins. Fine quartz veins crosscutting SI and folded by S2 are common, and are similar to those in some examples of the Triassic phyllite. These veins have been 107 Figure 41 - Photomicrograph and sketch of elongate pressure shadows composed of quartz and chlorite adjacent to pyrite crystals in a sample of metavolcanic schist from the Antler Formation (sample 665) on the western flank of the Boss Mountain area. SI is the prominent foliation. 108 Figure 42 - Lower hemisphere, equal area projection of lineations and foliation in graphitic phyllite from several locations within Phase 2 fault zones. Figure 43 - Phase 2 ductile shear zones in basement rocks, a) sheared limb of F2 folds in Unit W-12, adjacent to the western contact of the Boss Mountain gneiss, b) deformed mafic xenoliths within Boss Mountain augen gneiss, adjacent to the contact with Unit W - l l . no folded by F2, and probably represent fractures produced by a similar process as shown in Figure 32. Metamorphism producing garnet and staurolite clearly overprints the faulted rocks associated with the imbricate shear zones. Ductile fault zones within the Snowshoe Group schist and gneiss are present adjacent to the imbricate faults and phyllite zones. Extreme thinning of F2 fold limbs in the vicinity of the shear zones has caused attenuation of the fold limbs (Fig. 43a). Deformed mafic xenoliths (biotite + plagioclase + quartz) within the Boss Mountain gneiss, are also evidence that localized ductile strain has affected the rocks during Phase 2. The xenoliths have been elongated parallel to the F2 axial direction (L2), and rotated parallel to S2 (Fig. 43b). 3. BASEMENT - PHASE 3 Kinematics of F3 folding From the fold classification of Ramsay (1967), the F3 folds generally approximate class IC or rarely class 2 folds (Fig. 13b). The folds are flattened flexural folds (Ramsay 1962b); the amount of flattening is heterogeneous throughout the area. Figure 44 represents the domainal analysis of deformed F2 structures within the area from various positions on the F3 Boss Mountain antiform. Contoured plots of L2 minor fold axes reveal a slip direction - a3 , associated with the F3 folding (Ramsay 1962a). The general implication is that the movement responsible for the F3 folding is at a small angle to the F3 fold axis. The movement or slip during the F3 folding episode has produced relatively low amplitude, open macroscopic folds in the Boss Mountain area. This is also evidenced in stereonet plots of deformed SI surfaces, which show an even distribution in relation to F3 structures (Fig. 13). Relationship between F2 and F3 folds Interference between F2 and F3 folds has typically resulted in a type 1 to type 2 interference pattern (Ramsay 1967). The relative scale of the fold sets is critical F i g u r e 44 L2 l i n e a t i o n s c o n t o u r e d a t i n t e r v a l s o f : Domain 1 - 2%, 4%, 8%, 12% p e r 1% a r e a T o t a l = 54 p o i n t s Domain 2 - 1.5%, 4%, 7%, 10% p e r 1% a r e a T o t a l = 80 p o i n t s Domain 3 - 1.5%, 5%, 8%, 12.5% p e r 1% a r e a T o t a l = 73 p o i n t s Domain 4 - 2%, 4%, 8% p e r 1% a r e a T o t a l = 52 p o i n t s Figure 44 - F3 slip directions, a) lower hemisphere, equal area plots of deformed L2 lineations from the four structural domains, b) F3 fold axes and F3 axial planes for each of the sub-domains are shown, with the variation in slip direction for each of these. 112 (44b) 112a F i g u r e 4 5 - T e r m i n o l o g y a n d r e l a t i o n s h i p s f o r f o l d i n t e r f e r e n c e c l a s s i f i c a t i o n . a = s h e a r d i r e c t i o n w i t h i n t h e a x i a l p l a n e o f s u p e r p o s e d f o l d s b = d i r e c t i o n n o r m a l t o a w i t h i n t h e a x i a l p l a n e c = p o l e s t o a x i a l p l a n e s f = d i r e c t i o n o f f o l d a x e s d = n o r m a l t o f o l d a x e s w i t h i n t h e i r a x i a l p l a n e s a = a n g l e b e t w e e n b 3 a n d f 2 /3 = a n g l e b e t w e e n a 3 a n d c 2 7 = a n g l e b e t w e e n c 3 a n d f 2 AP 2 AP 3 = F 2 a x i a l p l a n e = F 3 a x i a l p l a n e Figure 45 - Angular relationships used in the fold interference classification (after Ramsay 1967, and Thiessen and Means, 1980). 114 0 30 60 90 9 0-4 1 L_ 1 1 1_ feo-p I 30-Type 1 / Type 1 - 2 Type 2 H «< T3 © CO Type 2 Crescent Figure 46 - a) Interference partem types after Ramsay (1967), and their relative fields of distribution (after Thiessen and Means 1980). b) Transitional dome and basin (Type 1 - Type 2) interference pattern between F2 and F3, north side of Boss Mountain, illustrating the curvilinear nature of the 12 fold axis. 115 in making this evaluation, and although these patterns are commonly observed at individual outcrops, the larger extent of the interference may remain obscure unless outcrops throughout the area are examined. A method for qualitatively determining the type of interference between superposed fold sets produced by imposing a shear direction of a later phase upon preexisting folds, has been described by Ramsay (1967). More recently, a modified method has been described by Thiessen and Means (1980), and has been used here to verify the assertions made from field observations. Figure 45 illustrates the angular relationships and terminology used in this determination. Measurements from individual locations were chosen by structural domain for consistency with the determined a3 slip direction from that region. A systematic east-west variation is noted, from a type 1 (dome and basin) to a transitional type 2, only locally bordering type 3. The results of this classification are presented in table 3. The angular relationship between the a3 movement direction and the S2 surfaces is partially responsible for the variation in amplitude and scale of the F3 folds. The a3 slip direction is relatively constant throughout the map area, variations in the attitude of S2 (and to a lesser extent, SI) has resulted in the interference patterns noted in Table 3. Low amplitude folds characterize F3 in the Boss Mountain area. This contrasts with areas nearby, which were affected by the same deformational history. Northeast of Crooked Lake, rocks of the Snowshoe Group are deformed by high amplitude F3 folds due to a high angle relationship between the a3 slip direction and the preexisting layering (Elsby 1985; Montgomery 1985). Lower resolved shear strain parallel to S3 surfaces is a consequence of the small angle between a3 and the preexisting layering in the Boss Mountain area. Larger values of constrictional strain due to the high resolved longitudinal strain parallel to the F3 axial direction would be expected (Ramsay 1967). 116 (46b) 117 Outcrop Number Type of i n t e r f e r e n c e p a t t e r n 1 4 10 12 22 47 49 60 39° 23° 35° 60° 19° 0° 32° 77° 78° 88° 74° 74° 90° 85° 57° 77° 80° 68° 77° 83° 62° 1-2 1 0- 1 1- 2 1-2 0 1 93 1 07 1 10 1 28 1 57 160 86° 18° 35° 20° 14° 18° 82° 77° 77° 88° 80° 84° 56° 66° 70° 84° 77° 82° 1-2 1-2 1-2 0-3 1 (1-2) 1 (1-2) 263 309 312 390 386 336 27° 65° 38° 44° 6° 16° 75° 64° 68° 82° 89° 75° 80° 70° 7 9 ° 85° 77° 73° 1-2 (3) 2 1-2 3 1 1-2 497 549 647 485 620 645 51° 57° 68° 19° 23° 37° 63° 70° 71 0 87° 61 ° 78° 79° 72° 20° 89° 83° 68° 2 1-2 1- 2 0 2- 3 1-2 Table 3 - Results of the fold interference pattern classification. 118 Phase 3 deformation - non-coaxial strain Post-metamorphic (M2) structures have produced folding of both SI and S2 surfaces and pulling apart of M2 metamorphic mineral phases. Deformation of garnet, staurolite and kyanite are recognized in a geometrical relationship to F3 minor fold structures. Porphyroblasts on the limbs of F3 microfolds have been pulled apart and the resulting opening filled with quartz. Garnets containing a deformed internal schistosity are common (Fig. 37a), and therefore must be syn-F2 growths, which locally show the texture described above. Pulled-apart garnet textures which postdate Phase 2 deformation are common, and appear to be synchronous with F3 folding (Fig. 47). Progressive deformation appears to have caused local extension of the S1/S2 foliation(s) related to F3 folding. The initiation of a more brittle type of behaviour and S3 pressure solution cleavage in some samples from Unit W-12, may indicate that a brittle-ductile transition was occurring at this time. The frequent occurrence of F3 brittle folds and kink bands support this interpretation. Uplift and unroofing during Phase 3 deformation may be responsible for the change in deformational style. 119 Figure 47 - Quartz filled pull-apart in garnet, sample 115, Unit W - l l . Polygonized quartz filling and fine grained inclusions within garnet affected by the extension suggest syn-F3 deformation. VI. TECTONIC INTERPRETATION - DISCUSSION The Boss Mountain area comprises the flank of the westernmost anticlinorium of rocks related to the Shuswap Complex (Campbell et al. 1973; Okulitch 1984). The rocks have undergone two phases of metamorphism and at least four episodes of deformation by folding and faulting. The metamorphic imprint left by the post-Early Jurassic thermal event can be ostensibly related to the metamorphic zones defined within rocks of the basement and its allochthonous cover. This indicates that the metamorphism culminated subsequent to the emplacement of the accreted terrane represented by the Antler Formation and Triassic/Jurassic rocks. The boundary between the two also seems to coincide with a rapid decrease in metamorphic intensity. Thus a possibility exists that this structural zone may have undergone a more complex and diachronous sequence of events than can be constrained by the present evidence. A systematic variation in structural style with structural level in the rocks related to the Cordilleran metamorphic core have been described by previous workers (Campbell 1970, 1973). The Boss Mountain area shows a similar transition over a distance of 10 - 15 km, from the high grade metasediments of the Snowshoe Group to the tectonically overlying greenschist facies cover rocks. The structural sequence observed in the Hadrynian and Early Paleozoic basement rocks is dominated by a second phase folding and faulting event, modified by subsequent folding during progressive deformation. In the cover rocks, the correlative deformation is evidenced by deformation of a more discrete nature, with an apparent increase in the intensity of foliation and shear fabrics. Explanations regarding the nature of the margin between the Hadrynian basement and the Late Paleozoic/Early Mesozoic cover range from a deformed unconformity (Campbell 1971) to a shear zone of concentrated mylonitic deformation (Rees 1981), Results from the present work in the Boss Mountain area clearly demonstrate that the strain associated with the emplacement of the cover rocks has been accomodated inhomogeneously. Zones of high 120 121 strain are not unique to the tectonic boundary, and appear to be distributed within the basement over a broad zone (Fig. 12), and probably well into the upper levels of the cover sequences (Montgomery 1978; Bloodgood 1985). The presence of major Phase 1 - cover / Phase 2 - basement shear zones has been documented as structural imbrication over a distance of 2 - 3 km on the western flank of the Boss Mountain antiform as shear zones within the Boss Mountain gneiss, and as mylonite zones within the metavolcanic Antler Formation rocks. Strain and related movements have occurred both inboard of, and external to the actual contact between the basement and cover, and cumulatively account for the deformation associated with the emplacement of the allochthonous cover. Zones of mylonitization cannot be equated solely with the contact between the terranes. Localized zones of movement in the basement rocks typically parallel the SI foliation trend (Figs. 10 and 40). These almost invariably coincide with broad zones (5 - 50 m wide) of moderate to intense hydrothermal alteration where an apparent fluid "flushing" has altered the mineral assemblages resulting in chloritization of garnet and widespread growth of hematite due to oxidation. In the vicinity of Elbow Lake, mylonitized Triassic sediments and volcanics are exposed. Zones of concentrated strain associated with emplacement of the cover rocks are distributed both above and below the fundamental contact (Ross et al. 1985). Recent models for the evolution of the area describe the geologic setting in terms of a collision zone between an accreted terrane (Quesnel Terrane or Quesnellia) and the western margin of North America (Monger et al. 1982; Brown 1981). The accretion of arc-related assemblages to the highly deformed Proterozoic (Windermere) and Paleozoic platform sediments of the Cariboo Mountains is interpreted to have occurred in pre-mid-Jurassic times (Ross et al. 1985). The timing of individual events in the tectonic evolution of the region is speculative, and only crudely bracketed by isotopic ages of plutonic rocks in the area (Pigage 1977, 1978; Armstrong 1982). 122 Additionally, the uncertain stratigraphic relations between similar rocks in the Kaza and Snowshoe Groups, which comprise the western flank of the Shuswap metamorphic core makes structural correlation difficult (Struik 1982; Pigage 1978). A significant difference between the two groups of ancient platform sediments is the lack of post-kinematic plutonic rocks in the Snowshoe Group, with respect to Kaza Group rocks which host numerous granitic intrusions of this type (Pigage 1977, 1978; Campbell 1978). Metamorphism and timing of convergence - The western margin of the Omineca Belt represents the stratigraphic and metamorphic equivalents of rocks east of the metamorphic core zone (Ghent et al. 1981; Price et al. 1985). Above the tectonic boundary, the adjacent cover rocks have been affected by the same regional metamorphism, but at lower pressure and temperature (Campbell 1978, Ross et al. 1985). Geothermometry from the Snowshoe Group, and the distribution of mineral assemblages in both the basement and the cover rocks, indicates the presence of a metamorphic and structural culmination in the Boss Mountain area. The implications of the rapidly decreasing temperature effects between the basement and the cover suggests that a "heat sink" effect may have operated during the metamorphism which affected both the basement and the cover. Dissipation of heat into the overthrust package of sediment and volcanics from the deeper basement material, could lead to the apparent rapid temperature "dropoff' observed between the two. The transition from staurolite grade rocks a mere 600 m below the tectonic boundary, to chlorite-muscovite bearing assemblages in the adjacent volcanic and sedimentary rocks, seems to indicate that metamorphic isograds were compressed, and nearly paralleled the margin. Thick, vitreous quartz veins are present between the the Snowshoe Group and Antler Formation at the western edge of the map area, where they parallel the boundary and the SI foliation in the cover (Fig. 16). These veins and the presence of fractures and veins at higher levels within the cover (Fig. 48), are further evidence that fluids being produced during metamorphism were escaping and depositing material within extensional 123 fractures in the cover. Fluid flowing nearly parallel to the gravity gradient may have participated in the buoying of slices of the cover during progressive deformation, thus aiding in translation of the overthrust package. High temperatures and pressures in the eastern part of the map area are responsible for the kyanite-sillimanite zone rocks observed in the lower portion of the Snowshoe Group. Widespread fibrolite in the Boss Mountain area (samples 23, 29, 10), is clearly a product of the breakdown of biotite during the transition from kyanite to sillimanite (see metamorphism section). Tectonic model - F2 basement structures appear to be a result of the convergence between the Quesnel Terrane and rocks of the North American continental margin during Early to Middle Jurassic time. The F2 geometry was established before the peak of M2 metamorphism in the basement rocks. F3 has refolded both the F2 folds and metamorphic isograds, as cohesion between the basement and cover was achieved by this time. Simple shear deformation during tectonic emplacement of the accreted terrane was apparently accomodated at various structural levels in a number of ways within the basement and cover. The variation in structural style is directly related to the viscosity contrast between the two terranes throughout the progressive deformation. the features which evolved in response to inhomogeneous Phase 2 strain include: Shear folding and ductile deformation of Snowshoe Group rocks at structural levels up to 2 - 4 km below the basement/cover contact Faulting and mylonitization in discrete zones 1 - 2 km below the tectonic contact, and faulting within the cover rocks especially near major lithologic boundaries (i.e. Triassic sediments/Jurassic volcanics). Deformation within the cover sequences by pressure solution and folding, and the 124 Figure 48 - Intrafolial folds in quartz veins parallel to the SI foliation, and thick, vitreous quartz vein in the ruptured hinge of a minor F2 antiformal fold in Antler Formation metavolcanics, near Bassett Creek, north of Boss Mountain. 125 formation of the prominent slaty cleavage - SI. The convergence between the Quesnel Terrane and the western margin of North America thus involved internal deformation within two bodies of contrasting rheology. The immediate zone of contact is characterized by tightly compressed folds and shear fabric within the cover, and well developed folds and schistosity within the basement Thrusting probably initiated along bedding planes in the cover during the west over east translation of the allochthonous terrane, resulting in the telescoping of this terrane over the basement Cumulative displacement on cover and basement faults, and shortening by folding and cleavage formation in both blocks contributed to the bulk strain related to the convergence event Folding in the cover alone may have involved as much as 100% shortening by pressure solution effects and transposition of bedding. The evolution of the convergent zone by simple shear deformation may be interpreted as being initiated upon a shallow west-dipping thrust surface during Early to Middle Jurassic times (Ross et al. 1985). The convergence may have been facilitated by a regional topographic and structural slope resulting from the marginal basin filled by Triassic and Jurassic rocks offshore to the west Cover strata whose compositional layering was probably at low angles to this thrust zone would have undergone translation with little initial folding (Ramsay 1983). The angular discordance between the overthrust beds and the previously deformed basement rocks would have tended to fold the metasediments by varying amounts with continued convergence. The prevailing distribution of F2 structures in the Boss Mountain area is suggestive of large strains adjacent to the tectonic boundary, and decreasing strain toward the centre of the area. A similar strain gradient is seen upwards in the cover rocks present within the, adjacent Eureka synform (Bloodgood 1985). Plate tectonic interpretation - Descriptions of the composite deformed belt of Proterozoic (Windermere) and Paleozoic rocks comprising the Selkirk Mountains and 126 their northern equivalents in the Cariboo Mountains have been presented in several major works (Bally et al. 1966; Campbell et al. 1973). More detailed work has been published sporadically, describing the rocks related to the contact zone between the Mesozoic sequences and the highly deformed basement complex west of the Rocky Mountain Trench (Campbell 1971; Montgomery 1978; Rees 1981; Struik 1981, 1982). This study and several related projects (Elsby 1985; Montgomery 1985; and Bloodgood 1985) have been conducted in the area in an effort to define the detailed structural and stratigraphic details of the western margin of the Omineca core and its adjacent cover. Although precise stratigraphic facing directions are not everywhere known, the detailed results have led to a proposed tectonic model consistent with the information from the Crooked Lake area. In the Crooked Lake area, convergence has taken place between an island arc (Takla Group) and associated marginal basin (Triassic phyllite) and the western margin of North America. It is unclear, however, if the rocks comprising the Snowshoe Group formed a contiguous part of the pre-Mesozoic continental margin, or whether they themselves were structurally decoupled from the craton to the east (Brown 1981, Struik 1985). A proposed Upper Paleozoic metamorphism and deformation has affected the Snowshoe Group rocks in the area, and appears to be related to the intrusion of large granitic bodies during Devonian to Early Carboniferous time. This phase of deformation within the Hadrynian rocks in the Crooked Lake area is apparently "absent" in equivalent age rocks to the north and east (Struik 1982, Rees 1981). It may thus be further suggested that the Snowshoe Group is parautochthonous with respect to sequences related to the North American craton. Struik (1984, 1985) has proposed that a low angle east-dipping thrust emplaced Kaza Group rocks (Cariboo Terrane) over Snowshoe Group rocks (Barkerville Terrane) in the vicinity of Quesnel Lake prior to the major Mesozoic deformation. Evidence for such a structure is not present in the Boss Mountain area. 127 The relationship of the Crooked Lake area to similar metamorphic complexes to the north and south is complicated by ambiguities in stratigraphic and structural correlations. There is little doubt that elsewhere the relationships between rocks of the Shuswap Complex and their overlying cover are suggestive of similar tectonic histories in different localities (Ross 1968; Ross and Christie 1979; Ross 1981; Read and Brown 1981; Okulitch 1984). Early recumbent folds characterize the western margin of the Shuswap complex as far south as the •southern Okanagan Valley (Ross and Christie 1979), and on the western edge of the Monashee Complex (Read and Brown 1981). Later folding and faulting of the metamorphic rocks signalled the collision of outboard terranes and the deformation of the accretionary margin. Amalgamation of smaller terranes including the Quesnel Terrane of the Intermontane Belt, occurred prior to their collision with the North American margin (eg. Terrane I of Monger et al. 1982). By the Late Triassic, "Quesnellia" may have been positioned between the North American margin and older terranes, including the Permo-Triassic Cache Creek Group to the west Concomitant translation and compression of these blocks with the craton as a backstop must have occurred after the beginning of the Jurassic (Ross 1981; Ross et al. 1985). Subsequent tectonic burial resulted in increasing metamorphism of the basement rocks, tapering off into greenschist facies within the Triassic and Jurassic rocks as seen in the Eureka synform above Crooked Lake (Ross et al. 1985; M . Bloodgood personal comm.). Actinolite and prehnite have been generated in the upper sequences of the Jurassic Takla-equivalent volcanics. The metamorphism was synchronous with F l deformation in the cover rocks, and apparently outlasted F l event The evidence in the Crooked Lake area overwhelmingly leads to the conclusion that eastward translation of the allochthonous arc terrane took place over a west-dipping subduction zone during Early to Mid-Jurassic time (Fig. 49a). With progressive deformation, the convergent margin underwent folding (F3-basement) 128 L. Triassic - E. Jurassic (?) Quesnellia Late Jurassic (?) Phase 2 Jura -Cretaceous (?) Figure 49 - Plate tectonic model for the evolution of the Intermontane/Omineca Belt boundary near Crooked Lake (modified from Dewey 1976). 129 producing southwesterly verging folds and tightening the earlier convergence-related structures (F2-basement). Steepening of F2 structures by F3 folding and rotation, and flattening caused by additional motion along convergence related faults, may have caused decoupling at higher structural levels within the cover rocks (Fig. 49b). A strong viscosity contrast between the progressively more "active" basement during the latter part of the Phase 3 deformation, seems to have caused further straining adjacent to the basement-cover contact This resulted in the flattening of Phase 2 cuspate fold/fault structures, which continue as narrow ductile shear zones in the basement Renewed underthrusting to the east may be the cause of the southwesterly verging folds which affect the basement and cover couple (Fig. 49c). REFERENCES Archibald, D.A., Glover, J.K., Price, R.A., Farrar, E. and Carmichael, D .M. 1983. Geochronology and tectonic implications of magmatism and metamorphism southern Kootenay Arc and neighbouring regions, southeastern British Columbia. Part I : Jurassic to mid-Cretaceous. Canadian Journal of Earth Sciences, 20, pp. 1891-1913. Armstrong, R.L. 1982. Cordilleran metamorphic core complexes from Arizona to southern Canada. Annual Review of Earth and Planetary Sciences, 10, pp. 129-154. Bally, A.W., Gordy, P.L. and Stewart, G.A. 1966. Structure, seismic data, and orogenic evolution of the southern Canadian Rocky Mountains. Bulletin of Canadian Petroleum Geology. 14, pp. 337-381. Bell, T.H. and Rubenach, M.J. 1983. Sequential porphyroblast growth and crenulation cleavage development during polyphase deformation. Tectonophysics, 92, pp. 171-194. Bence, A.E. and Albee, A.L. 1968. Empirical correction factors for the electron microanalysis of silicates and oxides. Journal of Geology, 76, 382-403. Berthe, D., Chokroune, P. and Jegouzo, P. 1979. Orthogneiss, mylonite and non-coaxial deformation of granites: the example of the South Amorican Shear Zone. Journal of Structural Geology, 1, 31-42. Bloodgood, M.A. 1985. Structure and stratigraphy of the Eureka Peak area, Cariboo Mountains, British Columbia. Geological Society of America, Abstracts with programs, 17, number 6, Cordilleran Section Meeting, Vancouver, B.C., May 1985. Brown, R.L. 1981. Metamorphic complex of SE Canadian Cordillera and relationship to foreland thrusting. In Thrust and nappe tectonics. The Geological Society of London, pp. 463-473. Brown, R.L. and Murphy, D.C. 1982. Kinematic interpretation of mylonitic rocks in part of the Columbia River fault zone, Shuswap Terrane, British Columbia. Canadian Journal of Earth Sciences, 19, pp. 456-465. Brown, R.L. and Read, P.B. 1983. Shuswap Terrane of British Columbia: a Mesozoic "core complex". Geology, 11, pp. 164-168. Cairnes, C E . 1939. The Shuswap rocks of southern British Columbia. Proceedings, 6th Pacific Science Congress, vol. 1, pp. 259-272. Campbell, K.V. 1971. Metamorphic petrology and structural geology of the Crooked Lake area, Cariboo Mountains, British Columbia. Ph.D. thesis, University of Washington, Seattle, Washington, 192 pages. Campbell, K.V. and Campbell, R.B. 1970. Quesnel Lake map area, British Columbia. In report of activities, part A. Geological Survey of Canada, Paper 70-1, pp. 32-35. Campbell, R.B. 1970. Structural and metamorphic transitions from infrastructure to 130 131 suprastructure, Cariboo Mountains, British Columbia. Geological Association of Canada, Special Paper No. 6, pp. 67-72. Campbell, R.B. 1973. Structural cross-section and tectonic model of the southeastern Canadian Cordillera. Canadian Journal of Earth Sciences, 10, pp. 1607-1620. Campbell, R.B. 1978. Quesnel Lake (93A) map area, British Columbia. Geological Survey of Canada, Open File map 574. Campbell, R.B., Mountjoy, E.W. and Young, F.G. 1973. Geology of the McBride map area, British Columbia (93H). Geological Survey of Canada, Paper 72-35, 104 pages. Carmichael, D .M. 1969. On the mechanism of prograde metamorphic reactions in quartz-bearing pelitic rocks. Contributions to Mineralogy and Petrology, 20, pp. 244-267. Carmichael, D .M. 1978. Metamorphic bathozones and bathograds: a measure of the depth of postmetamorphic uplift and erosion on the regional scale. American Journal of Science, 278, pp. 769-797. Chappell, B.W. and White, A.J.R. 1974. Two contrasting granite types. Pacific Geology, 8, 173-174. Coney, P.J., Jones, D.L. and Monger, J.W.H. 1980. Cordilleran suspect terranes. Nature, 288, pp. 329-333. Crittenden, M.D.Jr., Coney, P.J. and Davis, G.H., eds. 1980. Cordilieran metamorphic core complexes. Geological Society of America Memoir 153, 490 pages. Darot, M . and Boudier, F. 1975. Mineral lineations in deformed peridotites, kinematic meaning. Petrologie, 3, pp. 225-236. Deer, W.A., Howie, R.A. and Zussman, J. 1977. An introduction to the rock forming minerals. Longman Group Ltd. 528 pages. Dewey, J.F. 1976. Ophiolite obduction. Tectonophysics, 31, pp. 93-120. Durney, D.W. 1972. Solution-transfer, an important geological deformation mechanism. Nature, 235, pp. 315-317. Durney, D.W. 1976. Pressure solution and crystallization deformation. Philosophical Transactions of the Royal Society of London, 283, pp. 229-240. Durney, D . M . and Ramsay, J.G. 1972. .Incremental strains measured by syntectonic crystal growths. In DeJong, K.A., and Scholten, R., eds., Gravity and tectonics, John Wiley and Sons, New York, pp. 67-96. Elliot, D. 1972. Deformation paths in structural geology. Geological Society of America Bulletin. 83, pp. 2621-2638. Elliot, D. 1973. Diffusion flow laws in metamorphic rocks. Geological Society of America Bulletin, 84, pp. 2645-2664. Etheridge, M.A. 1983. Differential stress magnitudes during regional deformation and 132 metamorphism. Geology, 11, 231-234. Etheridge, M.A., Wall, V.J., Cox, S.F. and Vernon, R.H. 1984. High fluid pressures duringregional metamorphism and deformation: implications for mass transport and deformation mechanisms. Journal of Geophysical Research, 89, pp. 4344-4358. Etheridge, M.A., Wall, V.J. and Vernon, R.H. 1983. The role.of the fluid phase during regional deformation and metamorphism. Journal of Metamorphic Geology, 1, pp. 205-226. Ferry, J.M. 1976. Metamorphism of calcareous sediments in the Waterville-Vassalboro area, south-central Maine: mineral reactions and graphical analysis. American Journal of Science, 276, pp. 841-882. Ferry, J.M. and Spear, F.S. 1978. Experimental calibration of the partitioning of Fe and Mg between biotite and garnet Contributions to Mineralogy and Petrology, 66, 113-117. Fillipone, Jeffrey A. 1985. Polyphase deformation at the Intermontane - Omineca tectonic boundary, southern Cariboo Mountains, British Columbia. Geological Society of America, Abstracts with Programs, 17, Number 6, Cordilleran Section Meeting, Vancouver, B.C., May 1985. Fletcher, C.J.N., and Greenwood, H.J. 1979. Metamorphism and structure of Penfold Creek area, near Quesnel Lake, British Columbia. Journal of Petrology, 20, part 4, pp. 743-794. Froese, E. and Gasparrini, E. 1975. Metamorphic zones in the Snow Lake area, Manitoba. Canadian Mineralogist 13, pp. 162-167. Getsinger, J.S. 1985. Geology of the Three Ladies Mountain/Mount Stevenson area, Quesnel Highland British Columbia. Ph.D. thesis, University of British Columbia, Vancouver, British Columbia. Gray, D.R. and Durney, D.W. 1979. Investigations on the mechanical significance of crenulation cleavage. Tectonophysics, 58, pp. 35-79. Ghent E.D., Simony, P.S. and Raeside, R.P. 1981. Metamorphism and its relation to structure within the core zone west of the Southern Rocky Mountains. Field guides to geology and mineral deposits, Calgary 81, Geological Association of Canada, Mineralogical Association of Canada, Canadian Geophysical Union. Goldsmith, J.R. and Newton, R.C. 1977. Scapolite-plagioclase stability relations at high pressures and temperatures in the system NaAlSiaOs-CaAliSijOs-CaCOj-CaSO*. American Mineralogist 62, pp. 1063-1081. Hobbs, B.E., Means, W.D. and Williams, P.F. 1976. An outline of structural geology. John Wiley and Sons, 571 pages. Holdaway, M.J. 1971. Stability of andalusite and the aluminum silicate phase diagram. American Journal of Science, 271, pp. 97-131. Holland, S.S. 1954. Geology of the Yanks Peak- Roundtop Mountain area, Cariboo District, British Columbia. British Columbia Department of Mines Bulletin, 34, 90 pages. 133 Hoscheck, G. 1969. The stability of staurolite and chloritoid and their significance in metamorphism of pelitic rocks. Contributions to Mineralogy and Petrology, 22, pp. 208-232. Marlow, P.C. and Etheridge, M.A. 1977. Development of a layered crenulation cleavage in mica schists of the Kanmantoo Group near Macclesfield, Australia. Geological Society of America Bulletin, 88, pp. 873-882. Miyashiro, A. 1961. Evolution of metamorphic belts. Journal of Petrology, 2, pp. 277-311. Monger, J.W.H. and Price, R.A. 1979. Geodynamic evolution of the Canadian Cordillera progress and problems. Canadian Journal of Earth Sciences, 16, pp. 770-791. Monger, J.W.H., Price, R.A. and Tempelman- Kluit, D. 1982. Tectonic accretion and the origin of two major metamorphic and plutonic welts in the Canadian Cordillera. Geology, 10, pp. 70-75. Montgomery, J.R. 1985. Structural relations of the southern Quesnel Lake Gneiss, Isoceles Mountain area, central British Columbia. M.Sc. thesis, University of British Columbia, Vancouver, British Columbia. Montgomery, S.L. 1978. Structural and metamorphic history of the Dunford Lake map area, Cariboo Mountains, British Columbia. M.S. thesis, Cornell University, Ithaca, New York. Murphy, D.C. and Rees, C.J. 1983. Structural transition and stratigraphy in the Cariboo Mountains, British Columbia. In Current Research, Part A., Geological Survey of Canada, Paper 83-IA, pp. 245-252. Newton, R.C. and Haselton, H.T. 1981. Thermodynamics of the garnet-plagioclase-Al2SiOj-quartz geobarometer: In Thermodynamics in minerals and melts, eds. R.C. Newton, A. Navrotsky and B.J. Wood, Springer Verlag, New York. Norris, R.J. and Henley, R.W. 1976. Dewatering of a metamorphic pile. Geology, 4, pp. 333-336. Okulitch, A.V. 1984. The role of the Shuswap Metamorphic Complex in Cordilleran tectonism. Canadian Journal of Earth Sciences, 21, pp. 1171-1193. Orville, P.M. 1975. Stability of scapolite in the system A b - A n - N a C l - C a C 0 3 . Geochemica et Cosmochemica Acta, 39, pp. 1091-1105. Pigage, L.C. 1977. Rb-Sr dates for granodiorite intrusions of the northeast margin of the Shuswap Metamorphic Complex, Cariboo Mountains, British Columbia. Canadian Journal of Earth Sciences, 14, pp. 1690-1695. Pigage, L.C. 1978. Metamorphism and deformation on the northeast margin of the -Shuswap Metamorphic Complex, Azure Lake, British Columbia. Ph.D. thesis, University of British Columbia, Vancouver, British Columbia, 289 pages. Pigage, L.C. and Greenwood, H.J. 1982. Internally consistent estimates of pressure and temperature: the staurolite problem. American Journal of Science, 282, pp. 943-969. 134 Price, R.A., and Monger, J.W.H. and Roddick, J.A. 1985. Cordilleran cross-section; Calgary to Vancouver. In Field guides to geology and mineral deposits in the Southern Canadian Cordillera. Geological Society of America Cordilleran Section Meeting, Vancouver, British Columbia, May 1985. Ramsay, J.G. 1962a. The deformation of earlier linear structures in areas of repeated folding. Journal of Geology, 68, 75-93. Ramsay, J.G. 1962b. The geometry and mechanics of "similar" type folds. Journal of Geology, 70, pp. 309-327. Ramsay, J.G. 1967. Folding and fracturing of rocks. McGraw-Hill , New York, N.Y., 568 pages. Ramsay, J.G. 1980. The crack-seal mechanism of rock deformation. Nature, 284, pp. 135-139. Ramsay, J.G., Casey, M . and Kligfield, R. 1983. Role of shear in the development of the Helvetic fold-thrust belt of Switzerland. Geology, 11, pp. 439-442. Ramsay, J.G. and Huber, M.I. 1983. The techniques of modern structural geology, volume 1,: strain analysis. Academic Press, New York, N.Y., 307 pages. Read, P.B. and Brown, R.L. 1981. Columbia River fault zone: southeast margin of the Shuswap and Monashee Complexes, southern British Columbia. Canadian Journal of Earth Sciences, 18, pp. 1127-1145. Rees, C.J. 1981. Western margin of the Omineca Belt at Quesnel Lake, British Columbia. In Current Research, Part A., Geological Survey of Canada, Paper 81-IA, pp. 223-226. Rees, C.J. and Ferri, F. 1983. A kinematic study of mylonitic rocks in the Omineca-Intermontane Belt tectonic boundary in east-central British Columbia. In Current Research, Part B, Geological Survey of Canada, Paper 83-IB, pp. 121-125. Richardson, S.W. 1968. Staurolite stability in part of the system F e - A l - S i - O - H . Journal of Petrology, 9, pp. 467-488. Rosenfeld, J.L. 1970. Rotated garnets in metamorphic rocks. Geological Society of America Special Paper, 129, 105 pages. Ross, J.V. 1968. Structural relations at the eastern margin of the Shuswap Complex, near Revelstoke, southeastern British Columbia. Canadian Journal of Earth Sciences, 5, pp. 831-849. Ross, J.V. 1981. A geodynamic model for some structures within and adjacent to the Okanagan Valley, southern British Columbia. Canadian Journal of Earth Sciences, 18, pp. 1581-1598. Ross, J.V. and Christie, J.S. 1979. Early recumbent folding in some westernmost exposures of the Shuswap Complex, southern Okanagan, British Columbia. Canadian Journal of Earth Sciences, 16, pp. 877-894. Ross, J.V., Fillipone, J.A., Montgomery, J.R., Elsby, D.C. and Bloodgood, M.A. 1985. 135 Geometry of a convergent zone, central British Columbia, Canada. In: N.L. Carter and S. Uyeda (Editor), Collision Tectonics: Deformation of Continental Lithosphere. Tectonophysics, 119: 000-013. Simpson, C. and Schmid, S.M. 1983. An evaluation of criteria to deduce the sense of movement in sheared rocks. Geological Society of America Bulletin, 94, pp. 1281-1288. Spear, F.S. and Selverstone, J. 1983. Quantitative P - T paths from zoned minerals: theory and tectonic applications. Contributions to Mineralogy and Petrology, 83, pp. 348-357. Struik, L.C. 1981. Snowshoe Formation, central British Columbia. In Current Research, Part A., Geological Survey of Canada, Paper 81-IA, pp. 213-216. Struik, L.C. 1982. Snowshoe Formation (1982), central British Columbia. In Current Research, Part B, Geological Survey of Canada, Paper 82-IB, pp. 117-124. Struik, L.C. 1984. Geology of Quesnel Lake and part of Mitchell Lake, British Columbia. Geological Survey of Canada, Open File Map, 962. Struik, L.C. 1985. Pre-Cretaceous terranes and their thrust and strike slip contacts, Prince George east half and McBride west half, British Columbia. In Current Research, Part A, Geological Survey of Canada, Paper 85-1 A. Sutherland Brown, A. 1963. Geology of the Cariboo River area, British Columbia. British Columbia Department of Mines, Bulletin 47. Talbot, J.L. and Hobbs, B.E. 1968. The relationship of metamorphic differentiation to other structural features at three localities. Journal of Geology, 76, pp. 581-587. Thiessen, R.L. and Means, W.D. 1980. Classification of fold interference patterns: a reexamination. Journal of Structural Geology, 2, pp. 311-316. Thompson, A.B. 1975. Mineral reactions in calc-mica schist from Gassetts, Vermont, U.S.A. Contributions to Mineralogy and Petrology, 53, pp. 105-127. Thompson, A.B. 1976a. Mineral reactions in pelitic rocks: I. Prediction of P - T - X (Mg/Fe) phase relations. American Journal of Science, 276, pp. 401-424. Thompson, A.B. 1976b. Mineral reactions in pelitic rocks: II. Calculation of some P - T - X (Mg/Fe) phase relations. American Journal of Science, 276, pp. 425-454. Tipper, H.W., Woodsworth, G.J. and Gabrielse, H . 1978. Tectonic assemblage map of the Canadian Cordillera and adjacent parts of the United States of America. Geological Survey of Canada, Map 1505A. Tracy, R.V. 1976. High grade metamorphic reactions and partial melting in pelitic schist, west central Massachusetts. American Journal of Science, 278, pp. 150-178. Tracy R.V., Robinson, P., and Thompson, A.B. 1976. Garnet composition and zoning in the determination of temperature and pressure of metamorphism, central Massachusetts. American Journal of Science, 61, pp. 762-775. Vernon, R.H. 1977. Relationships between microstructures and metamorphic assemblages. 136 Tectonophysics, 39, pp. 439-452. Vemon, R.H. and Flood, R.H. 1979. Microstructural evidence of time relationships between metamorphism and deformation in the metasedimentary sequence of the Northern Hill End Trough, New South Wales, Australia. Tectonophysics, 58, pp. 127-137. Wheeler, J.O. and Gabrielse, H. 1972. The Cordilleran structural province. Geological Association of Canada, Special Paper 11, pp. 1-81. Williams, P.F. 1972. Development of metamorphic layering and cleavage in low grade metamorphic rocks at Bermagui, Australia. American Journal of Science, 272, pp. 1-47. Williams, P.F. 1976. Relationships between axial plane foliations and strain. Tectonophysics, 30, pp. 181-196. Winkler, H.J.F. 1979. Petrogenesis of metamorphic rocks. Springer Verlag, New York, N.Y., fifth edition, 348 pages. ' Yardley, B.W.D. 1977. The nature and significance of the mechanism of sillimanite growth in the Connemara schists, Ireland. Contributions to Mineralogy and Petrology, 65, 53-58. APPENDIX A - METAMORPHIC MINERALS TABLE 4 - Mineral Abbreviations Abbreviation Mineral name Am almandine Ap apatite Bi biotite Cc calcite Ch chlorite Di diopside Ep epidote Gt garnet Gp graphite Gr grossular Hm hematite Hb hornblende Im ilmenite Ks K-feldspar Ky kyanite Ms muscovite Mt magnetite PI plagioclase Py pyrite Pr pyrope Qz quartz Ru rutile Sc scapolite SI sillimanite Sp sphene St staurolite To tourmaline Zr zircon Zo zoisite 137 T A B L E 5 - Mineral Formulae Mineral name quartz albite anorthite muscovite chlorite biotite garnet staurolite kyanite sillimanite calcite zoisite scapolite General formula Si0 2 NaAlSi 3 0 8 CaAlSi 2 0 8 KAl 3 Si 3 O 1 0 (OH) 2 (Fe,Mg) 5Al 2Si 3O 1 0(OH) 8 K(FeJ^g) s AlSiiO 1 0 (OH) 2 (Fe,Mg,Mn,Ca)3 A l , Si 3 012 (Fe,Mg) 4 Al 1 8 Si 8 0 4 4 (OH) 4 A l 2 S i 0 5 AljSiOs CaC0 3 Ca 2 Al 3 Si 3 0 1 2 (OH) 3 CaAl 2 Si 2 0 8 • CaC0 3 139 TABLE 6 - Plagioclase Compositions Sample Number Metamorphic Zone Rock Unit Composition 10 sillimanite E-1B AII40 ~ AJI50 29 sillimanite E-5 A n 3 2 21 kyanite B-6B A n 2 4 - A n 2 8 31 kyanite B-6B A n 4 0 73 kyanite B-7 A n 3 S 71 kyanite BM-10 A n 4 0 93 staurolite BM-8 A O u 115 staurolite W - l l A n 3 6 647 staurolite W-12 A n 3 8 APPENDIX B - E L E C T R O N MICROPROBE ANALYSES Microprobe analyses of garnet and biotite were collected utilizing the three channel, automated A R L SEMQ at the University of British Columbia. Four specimens of pelitic schists from the different mineral zones in the area were analyzed using mineral standards from the University of British Columbia collection. Spot analyses were done on several crystals in each sample, and both rims and cores of garnets were analyzed for the presence of chemical zonation. Instrument operating conditions were identical for each run performed. accelerating potential - 15 kV specimen current - 40 nanoamps electron beam diameter - 10 to 12 microns counting time - 10 to 20 seconds All analyses were corrected for background and dead time using the Bence -Albee correction method (Bence and Albee 1968), with the alpha factors of Albee and Ray (1970). Considerable problems with machine drift were encountered, and poor analyses were rerun or discarded. 140 T A B L E 7 - Standards for garnet and biotite analyses Element Siandaid I^ocalitv Si wollastonite A l andalusite Ti rutile Fe fayalite Mn pyroxmangite Mg forsterite Ca wollastonite Na albite K orthoclase F F - phlogopite Ba benitoite garnet - Si, A l , Mg pyrope 21 26 13 250 245 22 21 20 96 24 35 235 New York Brazil synthetic synthetic Japan synthetic New York Oregon New York synthetic California New Zealand 142 TABLE 8 - Sample Locations Sample Lat(N) Long. (W) Rock type 115 52° 10 24 120° 38 00 gt-st pelitic schist 10 52° 09 50 120° 29 30 gt-sl pelitic schist 21 52° 09 32 120° 32 30 gt-st-ky pelitic schist 53 52° 10 00 120° 34 00 gt pelitic schist TABLE 9 - Garnet-Biotite Geothermometry (5 kb) Sample Xsha Xm Xgr Xss Xsh Xann T °c 10 . 0.78 0.13 0.06 0.04 0.42 0.58 689 21 0.79 0.13 0.04 0.04 0.44 0.57 665 53 0.79 0.11 0.07 0.03 0.41 0.59 629 115 0.83 0.09 0.07 0.00 0.47 0.47 505 144 Sample 10 Garnet 3-A 3-B 3-B2 S i 0 2 36 .76 37. 36 37.62 T i 0 2 0 .02 - 0.02 A1 20 3 21 .07 20. 81 20.59 FeO 35 .43 35. 40 35.20 MnO 1 .06 1 . 07 2.17 MgO 3 .49 3. 44 3.03 CaO 2 .18 2. 1 4 2.06 Total 100 .00 100. 23 100.7 Si 2 .96 3. 00 3.02 Ti - - -Al 2 .00 1 . 97 1 .95 Fe 2 .39 2. 38 2.36 Mn 0 .07 0. 07 0.15 Mg 0 .42 0. 41 0.36 Ca 0 .19 0. 18 0.18 Total 8 .04 8. 02 8.01 3-B6 3-B7 3-B8 4-A 4-C 37. 77 36. 28 37. 20 37. 1 1 36.76 0. 10 0. 03 0. 01 - -20. 97 20. 35 20. 56 21 .01 20.42 35. 40 35. 57 35. 53 35.68 35. 18 1 . 96 2. 02 1 . 16 2.21 2.81 3. 01 3. 13 3. 22 2.89 2.88 1 . 99 1 . 60 2. 10 1 .63 1 .40 101 .2 98. 98 99. 77 100.53 99.45 3. 01 2. 97 3. 01 2.99 3.00 1 . 97 1 . 97 1 . 96 1 .99 1 .96 2. 36 2. 44 2. 40 2.40 2.40 0. 13 0. 14 0. 08 0.15 0.19 0. 36 0. 38 0. 39 0.35 0.35 0. 17 0. 1 4 0. 18 0.14 0.12 8. 00 8. 04 8. 01 8.02 8.02 145 Sample 21 G a r n e t 5 - A 5 - B 6 - A 7 - A 8 - A 9 - A S i 0 2 3 8 , . 0 0 3 7 . 2 4 3 6 . 3 6 3 7 , . 3 8 3 6 . 7 7 . 3 6 . 6 2 T i 0 2 0 , . 0 4 0 . 14 0 . 0 1 0 , . 0 4 - 0 . 0 4 A l 2 0 3 2 0 , . 4 4 2 0 . 18 2 0 . 3 3 2 0 , . 9 0 21 . 21 1 9 . 3 2 F e O 3 5 , . 6 9 3 5 . 0 6 3 6 . 2 2 3 5 , . 0 5 3 5 . 4 4 3 4 . 4 8 MnO 1 . . 3 7 2 . 17 1 . 7 9 1 . . 6 7 1 . 9 6 1 . 8 3 MgO 3 . . 3 7 3 . 16 3 . 2 4 3 , . 4 6 3 . 2 4 3 . 0 4 C a O 1 . . 3 3 1 . 4 3 1 . 2 3 1 . . 7 0 1 . 2 8 3 . 12 T o t a l 1 0 0 , . 2 6 9 9 . 3 9 9 9 .18 1 0 0 , . 2 2 9 9 . 91 9 8 . 4 4 S i 3 - , 0 5 3 . 0 2 2 . 9 8 3 . . 0 0 2 . 9 7 3 . 0 2 T i - - - - - _ A l 1 . , 9 3 1 . 9 3 1 . 9 6 1 . . 9 8 2 . 0 2 1 . 8 8 F e 2 . , 3 9 2 . 3 8 2 . 4 8 2 , . 3 5 2 . 4 0 2 . 3 8 M n 0 . , 0 9 0 . 15 0 . 1 2 0 , . 1 1 0 . 13 0 . 13 M g 0 . . 4 0 0 . 3 8 0 . 4 0 0 , . 4 1 0 . 3 9 0 . 3 7 C a 0 . . 1 2 0 . 13 0 . 1 1 0 , . 1 5 0 . 1 1 0 . 2 8 T o t a l 7 . . 9 9 8 . 0 0 8 . 0 4 8 , . 0 1 8 . 0 2 8 . 0 4 146 Sample 53 Garnet 1 1-A 1 2 13 1 4 1 5-A 1 5-B 1 2-B S i 0 2 36.93 36.97 37. 12 37.59 38.40 37.33 36.43 T i 0 2 0.04 0.03 - 0.03 0.01 0.03 0.33 A 1 2 0 3 21 .45 19.22 21 . 50 21 .31 20.71 20.65 21 .23 FeO 35.73 35.94 35. 40 35.48 35.89 35.06 29. 13 MnO 1 .24 0.99 2. 36 1 .08 1 .65 1.16 4.38 MgO 2.84 2.73 2. 58 2.63 2.68 2.62 1 .40 CaO 2.39 2.55 1 . 80 2.87 1 .97 2.71 6.87 T o t a l 100.62 98.44 100. 76 100.99 101.3 99.57 99.76 S i 2.96 3.05 2. 98 3.00 3.05 3.02 2.95 T i - - - - - - -A l 2.03 1 .87 2. 03 2.00 1 .94 1 .97 2.02 Fe 2.40 2.48 2. 37 2.37 2.39 2.37 1 .97 Mn 0.08 0.07 0. 1 6 0.07 0.11 0.08 0.30 Mg 0.34 0.34 0. 31 0.31 0.32 0.32 0.17 Ca 0.21 0.23 0. 16 0.25 0.17 0.24 0.60 T o t a l 8.02 8.02 8. 01 8.00 7.98 7.99 8.02 •147 Sample 115 Garnet 1 -A 1-B 1-C S i 0 2 37. 07 34.59 37.32 T i 0 2 0. 05 0.06 0.01 A 1 2 0 3 20. 45 21 .05 20.14 FeO 37. 20 37.27 36.85 MnO 0. 10 0.19 0.38 MgO 2. 39 2.34 1 .87 CaO 2. 56 2.64 3.34 T o t a l 99. 82 98. 1 4 99.91 S i 3. 01 2.88 3.03 T i - - -A l 1 . 96 2.07 1 .93 Fe 2. 53 2.60 2.50 Mn - 0.01 0.03 Mg 0. 29 0.29 0.23 Ca 0. 22 0.24 0.29 T o t a l 8. 01 8.08 8.01 2-A 2-B 2. 5-A 2.5-B 36.76 36.82 37.25 37.03 0.08 0.06 0.04 0.09 20.38 20.70 20.33 20. 1 5 36.21 37.46 37.56 36.73 0.35 0.19 0.22 0.99 1 .99 2.36 2.36 1.31 3.36 2.62 2.50 3.93 99.13 100.22 100.27 100.24 3.01 2.98 3.01 3.01 1 .96 1 .98 1 . 94 1 .93 2.48 2.54 2.54 2.50 0.02 0.01 0.02 0.07 0.24 0.29 0.29 0.16 0.29 0.23 0.22 0.34 8.01 8.03 8.01 8.02 148 Sample 10 B i o t i t e 5 5-B 5 - c 6 7 8 S i 0 2 36. 23 36. 1 1 35. 73 36. 50 34. 89 35 .31 A 1 2 0 3 19. 15 19.80 19. 42 19. 39 19. 84 20 .00 T i 0 2 1 . 76 1 .83 2. 19 1 . 93 2. 65 1 .56 FeO 21 . 74 21 .36 21 . 1 1 22. 1 1 21 . 55 21 .91 MnO 0. 15 0.13 0. 14 0. 13 0. 13 0 .13 MgO 8. 89 8.64 8. 97 8. 92 8. 90 9 .31 BaO 0. 17 0.17 0. 12 0. 12 0. 08 0 .08 Na 20 0. 27 0.34 0. 29 0. 26 0. 29 0 .24 K 20 8. 64 8.62 8. 73 8. 72 8. 75 8 .40 CaO - - - 0. 02 0. 01 0 .01 F 0. 24 0.11 0. 36 0. 45 0. 20 0 .48 T o t a l 97. 24 97.25 97. 08 97. 54 97. 29 97 .45 Si 5. 46 5.43 5. 35 5. 43 5. 25 5 .39 A l 3. 48 3.58 3. 43 3. 40 3. 52 3 .50 T i 0. 19 0.20 0. 25 0. 22 0. 30 0 .18 Fe 2. 67 2.61 2. 77 2. 75 2. 71 2 .72 Mn 0. 02 0.02 0. 02 0. 02 0. 02 0 .02 Mg 1 . 94 1 .88 2. 00 1 . 98 2. 00 2 .06 Ba - - - - - -Na 0. 08 0.10 0. 08 0. 08 0. 09 0 .07 K 1 . 62 1 .61 1 . 67 1 . 66 1 . 68 1 .59 Ca - - - - - -F 0. 1 1 0.11 O. 1 7 0. 21 0. 10 0 .23 T o t a l 15. 57 15.546 15. 74 15. 73 15. 67 15 .95 149 Sample 21 B i o t i t e 9 10 S i 0 2 37.16 36.90 A 1 2 0 3 19.02 18.98 T i 0 2 1.99 2.07 FeO 21 .00 21 . 12 MnO 0.06 0.06 MgO 8.69 9.22 BaO 0.12 0.11 Na 20 0.40 0.35 K 20 8.46 8.53 CaO 0.04 0.01 F 0.24 0.16 T o t a l 97.18 97.52 Si 5.55 5.49 A l 3.35 3.33 T i 0.22 0.23 Fe 2.62 2.63 Mn -Mg 1.93 2.05 Ba - -Na 0.12 0.10 K 1.61 1.62 Ca F 0.12 0.08 T o t a l 15.53 15.55 11 12 13 36.23 36.69 35.69 18.54 18.60 19.16 1.94 1.95 1.85 21.00 21.67 21.08 0.07 0.08 0.07 8.64 9.28 9.99 0.05 0.14 0.02 0.40 0.40 0.42 8.61 8.69 8.59 0.01 0.45 0.36 0.30 95.92 97.86 97.19 5.51 5.42 5.36 3.33 3.41 3.39 0.22 0.22 0.21 2.67 2.68 2.65 1.96 2.04 2.35 0.12 0.12 0.09 1.67 1.64 1.62 0.21 0.17 0.14 15.71 15.71 15.77 1 5 0 Sample 53 B i o t i t e 16 16-B 17 18 19 20 21 22 S i 0 2 36.33 34.46 37. 18 35.85 35.29 36. 19 35.46 37.75 A 1 2 0 3 18.98 19.53 24. 35 19.54 19.78 20.50 19.33 22.1 1 T i 0 2 1 .93 1 .77 1 . 48 2.42 2.23 2.32 2.38 2.11 FeO 21 .95 21 .55 16. 74 21 .53 21 .28 20.68 21 .42 18.34 MnO 0.11 0.10 0. 05 0.11 0.08 0.07 0.08 0.08 MgO 8.65 8.85 6. 92 8.11 8.38 7.91 8.78 7.18 BaO 0.01 0.07 0. 17 0.11 0.14 0.14 0.14 -Na 20 0.19 0.16 0. 39 0.21 0.22 0.31 0.18 0.41 K 20 8.63 8.90 9. 12 8.95 8.85 8.66 8.95 8.92 CaO 0.04 0.02 0. 01 0.05 - 0.01 - -F 0.24 0.45 0. 37 0.41 0.24 0.45 0.57 0.21 T o t a l 97.08 95.87 96. 83 97.30 96.49 97.24 97.30 97. 1 2 Si 5.47 5.29 5. 52 5.40 5.35 5.41 5.36 5.54 A l 3.37 3.53 4. 1 5 3.47 3.53 3.61 3.44 3.82 T i 0.22 0.20 0. 16 0.27 0.25 0.26 0.27 0.23 Fe 2.76 2.77 2. 22 2.71 2.70 2.59 2.71 2.25 Mn 0.01 0.01 - 0.01 0.01 - 0.01 0.01 Mg 1 .94 2.03 1 . 49 1 .82 1 .89 1 .76 1 .98 1 .57 Ba - - - - - - - -Na 0.06 0.05 0. 1 1 0.06 0.07 0.09 0.05 0.12 K 1 .66 1 .74 1 . 68 1 .72 1 .71 1 .65 1 .72 1 .67 Ca - - - - - - - -F 0.12 0.22 0. 1 7 0.20 0.24 0.21 0.27 0.21 T o t a l 15.61 15.85 15. 32 15.68 15.64 15.61 15.82 15.31 1-51 Sample 115 B i o t i t e S i 0 2 A 1 2 0 3 T i 0 2 FeO MnO MgO BaO Na 20 K 20 CaO F T o t a l 1 36.00 19.81 1 .54 20.64 10.63 0.11 0.33 8.57 0.02 0.04 97.68 4 36. 10 19.84 1 .36 20.75 0.02 1 0.24 0.07 0.38 8.57 0.56 97.89 3 35.90 19.96 1 .52 20.76 0.02 10.28 0.07 0.31 8.67 0.16 97.66 S i 5.34 5.45 5.42 A l 3.46 3.43 3.45 T i 0.17 0.15 0.17 Fe 2.56 2.55 2.55 Mn -Mg 2.35 2.24 2.25 Ba -Na 0.09 0.11 0.09 K 1 .62 1.61 1.62 Ca -F 0.02 0.26 0.07 T o t a l 15.63 15.80 15.62 


Citation Scheme:


Citations by CSL (citeproc-js)

Usage Statistics



Customize your widget with the following options, then copy and paste the code below into the HTML of your page to embed this item in your website.
                            <div id="ubcOpenCollectionsWidgetDisplay">
                            <script id="ubcOpenCollectionsWidget"
                            async >
IIIF logo Our image viewer uses the IIIF 2.0 standard. To load this item in other compatible viewers, use this url:


Related Items