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Structure and metamorphism of Penfold Creek area, near Quesnel Lake, central British Columbia Fletcher, Christopher John Nield 1972

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STRUCTURE AMD METAMORPHISM OF PENFOLD CREEK AREA, NEAR QUESNEL LAKE, CENTRAL BRITISH COLUMBIA by CHRISTOPHER JOHN NIELD FLETCHER B . S c , U n i v e r s i t y of S t . Andrews, 1966 M.Sc. Queen's U n i v e r s i t y , 1968 A THESIS SUBMITTED IN PARTIAL FULFILMENT OF THE REQUIREMENTS FOR THE DEGREE OF DOCTOR OF PHILOSOPHY i n the Department o f Geology We accept t h i s t h e s i s as conforming to the r e q u i r e d s tandard THE UNIVERSITY OF BRITISH COLUMBIA June , 1972 In presenting t h i s thesis i n p a r t i a l f u l f i l m e n t of the requirements for an advanced degree at the University of B r i t i s h Columbia, I agree that the Library s h a l l make i t f r e e l y available for reference and study. I further agree that permission for extensive copying of t h i s thesis for scholarly purposes may be granted by the Head of my Department or by h i s representatives. It i s understood that copying or publication of t h i s thesis for f i n a n c i a l gain s h a l l not be allowed without my written permission. Department The University of B r i t i s h Columbia Vancouver 8, Canada ABSTRACT The Quesnel Lake area l ies within the Omineca Crystal l ine Belt , and is underlain by the northern extremity of the Shuswap Metamorphic Complex. Closely spaced and steeply dipping isograds mark the margins of the metamorphic be l t . In the Penfold Creek area only one and a half miles separate the b iot i te and s i l l imanite isograds. Related to this sharp increase in metamorphic grade there is a marked change in the fold style from similar folds, showing a strong axial-plane cleavage, in the chlorite zone to tight refolded isocl ines in the s i l l imanite zone. Three periods of deformation and two periods of prograde metamorphism have been recognized, with the f i r s t metamorphic period being associated with Phase 2 deformation and the second being post Phase 2. Mineralogical changes in p e l i t i c and calcareous assemblages suggest that the increase in metamorphic grade was a function of both temperature and composition of the f lu id phase, and that total pressure remained relat ively constant. Compositional variations of major minerals in the p e l i t i c rocks are consistent with a model of increasing P u with increasing temperature, moreover they suggest that in the highest grades of meta-morphism P ^ = P T o t a l . An i soc l ina l l y folded granit ic gneiss crops out along the shores of Quesnel Lake. It i s believed to represent a metamorphosed granite s i l l which had been intruded into the surrounding Proterozoic sediments. i i TABLE OF CONTENTS INTRODUCTION 1 O u t l i n e of the geology of B r i t i s h Columbia . 1 O u t l i n e of the geology o f Quesnel Lake a r e a . 3 F i e l d s t u d i e s . 10 PART I. STRUCTURE AND METAMORPHISM OF PENFOLD CREEK AREA 11 1. RELATIONSHIPS BETWEEN DEFORMATION AND METAMORPHISM 11 S t r a t i g r a p h y . 11 S t r u c t u r e . 13 Metamorphic framework. 18 Petrography and t i m i n g o f minera l growth. 19 c h l o r i t e 1? b i o t i t e 19 muscov i te 20 garnet 20 s t a u r o l i t e 23 k y a n i t e , 24 s i l l i m a n i t e 24 i l m e n i t e . 2 4 R e l a t i o n s h i p s between f o l d s t y l e and metamorphic grade. 27 2 . MINERALOGY OF THE PELITIC ROCKS 28 Methods of s t u d y . 28 M i n e r a l assemblages. 3 1 . M i n e r a l o g y . 34 b i o t i t e 34 muscov i te j 39 c h l o r i t e 44 garnet 46 s t a u r o l i t e 53 p l a g i o c l a s e 57 i l m e n i t e 59 i i i 3 . ATTAINMENT OF EQUILIBRIUM 59 4 . REACTIONS OCCURRING IN THE PELITIC ROCKS 62 L i n e a r r e g r e s s i o n t e c h n i q u e . 63 E r r o r p r o p a g a t i o n . 64 B i o t i t e i s o g r a d . 65 Garnet i s o g r a d . ; 66 S t a u r o l i t e / k y a n i t e i s o g r a d 68 S i l l i m a n i t e i s o g r a d s . 69 5 . PETROLOGY OF THE CALC-SILICATE BEARING ASSEMBLAGES 74 6. CONTROLS OF METAMORPHISM 82 PART I I . THE QUESNEL LAKE GNEISS I n t r o d u c t i o n . S t r u c t u r e . Metarnorphism. P e t r o l o g y . P e t r o c h e m i s t r y . Age and o r i g i n of the g n e i s s . 87 87 87 88 89 90 90 Appendix . I. Chemist ry of the metasediments - f rom the Isaac 99 Formation and Kaza Group. Appendix I I . D e s c r i p t i o n of the rock u n i t s i n P e n f o l d Creek 103 a r e a . Appendix I I I . Microprobe s t a n d a r d s . 106 Appendix IV. D i f f u s i o n i n g a r n e t s . 109 Appendix V. Thermodynamic c a l c u l a t i o n s of the s i l l i m a n i t e - 112 forming r e a c t i o n c u r v e s . Acknowledgements. 117 Refe rences . 118 P l a t e s . 124 Maps. pocket i v LIST OF FIGURES 1. Major t e c t o n i c elements of the southern Canadian C o r d i l l e r a . 2 2 . Geology of the Omineca G e a n t i c l i n e i n the reg ion of Quesnel 5 Lake. 3 . S t r u c t u r a l stereograms from P e n f o l d Creek a r e a . 14 4 . V a r i a t i o n i n t h i c k n e s s (T) of a p h y l l i t e u n i t and a q u a r t z i t e 16 u n i t measured p a r a l l e l to the f o l d a x i a l p l a n e . 5 . M i n e r a l r e l a t i o n s h i p s i n the P e n f o l d Creek s c h i s t s . 21 6. Timing of minera l growth w i t h r e s p e c t to the p e r i o d of deform- 25 a t i o n . 7. R e l a t i o n s h i p between p e r i o d of deformat ion and minera l growth. 26 8 . Fo ld s t y l e s i n the P e n f o l d Creek a r e a . 28 9 . Specimen l o c a t i o n map. 36 10. MgO content i n b i o t i t e . 36 11 . FeO content i n b i o t i t e . 38 12. T i0g contents of b i o t i t e and m u s c o v i t e . 38 1 3 . P l o t of muscov i te ana lyses i n p a r t of the SAF t r i a n g l e . 42 14 . Graph of atom% ( F e , M g ) v 1 a g a i n s t A 1 V 1 i n m u s c o v i t e . 42 15. Pa ragon i te content of m u s c o v i t e . 43 16. Phengi te content of m u s c o v i t e . 43 17. Microprobe t r a v e r s e s of s e l e c t e d g a r n e t s . 47 18. Average a n o r t h i t e content of p l a g i o c l a s e . 55 19. ZnO content o f s t a u r o l i t e . 55 20. Atomic p r o p o r t i o n of Mg a g a i n s t atomic p r o p o r t i o n o f Zn i n 56 s t a u r o l i t e . 2 1 . E f f e c t o f m u s c o v i t e - p a r a g o n i t e s o l i d s o l u t i o n on the s i l l i m - 70 a n i t e - f o r m i n g r e a c t i o n s . 22 . C a l c - s i 1 i c a t e assemblages i n P e n f o l d Creek a r e a . 76 2 3 . T-X diagram f o r the c a l c - s i l i c a t e r e a c t i o n s r e l e v a n t to 77 P e n f o l d Creek assemblages. V 24. The e f f e c t o f Fe s o l i d s o l u t i o n on the c a l c - s i l i c a t e r e a c t i o n s . 79 2 5 . The e f f e c t o f f l u i d p ressure v a r i a t i o n on the c a l c - s i l i c a t e 81 assemblages. 26 . The e f f e c t o f f l u i d compos i t ion v a r i a t i o n on the c a l c - s i l i c a t e 81 assemblages. 27 . A ' s e l f f l u s h i n g 1 system i n the p e l i t i c and c a l c a r e o u s r o c k s . 85 28. T r i a n g u l a r p l o t of metasediment and gne iss a n a l y s e s . 92 2 9 . T r i a n g u l a r p l o t of metasediment and gne iss a n a l y s e s . 93 30. Two p o s s i b l e c r o s s - s e c t i o n s o f the Quesnel Lake G n e i s s . 94 3 1 . T r i a n g u l a r p l o t s of metasediment ana lyses from the Isaac 102 Format ion and Kaza Group. 32. A d i f f u s i o n curve from a microprobe t r a v e r s e of a g a r n e t . 110 v i LIST OF TABLES 1. D e s c r i p t i o n o f rock types i n P e n f o l d Creek a r e a . 12 2 . Modal ana lyses o f the ana lysed p e l i t i c r o c k s . 33 3 . Chemical ana lyses o f b i o t i t e . 35 4 . Chemical ana l yses of m u s c o v i t e . 40 5 . Chemical ana lyses of c h l o r i t e . 45 6 . Chemical ana lyses of g a r n e t . 50 7. Chemical ana lyses of s t a u r o l i t e . 54 8 . Chemical ana lyses o f p l a g i o c l a s e . 58 9 . Chemical ana lyses of i l m e n i t e . 60 10. Regress ion equat ions f o r the minera l assemblages from 65 P e n f o l d Creek a r e a . 11 . Chemical and modal ana lyses of the Quesnel Lake G n e i s s . 91 12. Chemical ana lyses o f the s c h i s t s from the Quesnel Lake 100 a r e a . v i i LIST OF PLATES 1. Phase 1 f o l d s i n in terbedded p h y l l i t e s and q u a r t z i t e s 124 from the b i o t i t e zone. 2 . Fanned a x i a l - p l a n e c leavage i n a Phase 1 f o l d . 124 3 . I s o c l i n a l Phase 1 f o l d s from s i l l i m a n i t e Zone 1. 124 4 . Refo lded f o l d s from s i l l i m a n i t e Zone 2 . 124 5 . I s o c l i n a l Phase 1 f o l d s from s i l l i m a n i t e Zone 3 . 124 6 . 'F low f o l d s ' from s i l l i m a n i t e Zone 3 . 124 7 . Photomicrograph showing two pe r iods o f garnet growth. 125 8 . Photomicrograph of the a l t e r a t i o n of s t a u r o l i t e to 125 muscov i te and s i l l i m a n i t e . 9 . Photomicrograph o f the a l t e r a t i o n of k y a n i t e to c r y s t a l l i n e 125 s i l l i m a n i t e . 10. Photomicrograph of the a l t e r a t i o n of garnet to muscov i te 125 and s i l l i m a n i t e . 1 1 . Photomicrograph of the t e x t u r a l r e l a t i o n s h i p s between 125 second g e n e r a t i o n g a r n e t , s t a u r o l i t e , and k y a n i t e . 12 . Photomicrograph of the ' b l o c k y ' t e x t u r e to second g e n e r a t i o n 125 g a r n e t . 13 . Photomicrograph o f t e x t u r a l r e l a t i o n s h i p s i n a t y p i c a l 126 s i l l i m a n i t e Zone 3 g n e i s s . 14. S i l l i m a n i t e augen from the breakdown of s t a u r o l i t e , 126 15. General view of P e n f o l d Creek a r e a . 126 16. Low grade augen gne iss from the Quesnel Lake G n e i s s . 126 17. M a f i c i n c l u s i o n s i n the Quesnel Lake G n e i s s . 126 18. High grade f o l i a t e d gne iss from the Quesnel Lake G n e i s s . 126 1 INTRODUCTION Outline of the Geology of B r i t i s h Columbia The Canadian C o r d i l l e r a (Roddick et a l . 1967; Wheeler 1970; Douglas et a l . 1970) may be divided into two geosynclinal belts (Fig. 1) ; a narrow miogeosynclinal belt to the east, referred to as the Marginal Zone (Rocky Mountain Thrust B e l t ) , and a wide eugeosynclinal belt to the west. The Marginal Zone i s characterized by numerous south-west dipping thrust faults and i s bounded to the east by the Interior Platform. The eugeosynclinal belt contains two major c r y s t a l l i n e b e l t s ; the Coast Geanticline consisting mainly of plutonic rocks, and the Omineca Geanticline (Eastern Core Zone) consisting mainly of metamorphic rocks. Separating these two geanticlines and to the west of the Coast Geanticline l i e two troughs: the Interior Trough (Intermontane Zone) and the Insular Trough, respectively. They both consist of marine sedimentary and volcanic rocks, and are r e l a t i v e l y unmetamorphosed. The Interior Trough i s divided into the Nechako and Quesnel Troughs by the Pinchi Geanticline. The geanticlines were repeatedly the si t e s of intense deformation and regional metamorphism, i n contrast to the intervening troughs which were, i n general, affected by a single deformation. 2 FIG. I. MAJOR TECTONIC ELEMENTS OF THE SOUTHERN CANADIAN CORDILLERA (after Wheeler 1970). Outline of the Geology of Quesnel Lake Area 3 The late Proterozoic Kaza Group (10,000'+) i s the oldest unit exposed (Sutherland Brown 1957), consisting mainly of feldspathic quartzites, g r i t s , p h y l l i t e s and minor thin limestones. If, constitutes the major part of the Omineca Geanticline i n the area, and was presumably deposited along the western margin of the craton. In the immediate v i c i n i t y of Quesnel Lake rocks of the Kaza Group contain a body of gr a n i t i c gneiss, the age and o r i g i n of which i s discussed i n Part I I of this thesis. The Cariboo Group (Holland 1954; Sutherland Brown 1957, 1963; R.B. Campbell 1968) i s a conformable sequence, ranging i n age from late Proterozoic to uppermost Lower Cambrian, and resting conformably on sediments of the Kaza Group. The lowest member i s the Isaac Formation (2,000'-5,000') which i s composed of s l a t e s , p h y l l i t e s , quartzites, g r i t s and limestones. I t i s dominantly argillaceous i n comparison to the arenaceous sediments of the underlying Kaza Group. The Cunningham Formation (1,200') overlies the Isaac Formation and consists of grey limestone and minor p h y l l i t e and a r g i l l i t e . Both the Isaac and Cunningham Formations are regarded as l a t e s t Precambrian (R.B. Campbell 1970), and are overlain by the Yankee Belle Formation (500'-900'), which i s an assemblage of alternating beds of p h y l l i t e s , quartzites, and minor limestones, and i s considered to be the lowest unit of the Cambrian. I t i s followed upward by the Yanks Peak Formation (600'-1,000') which consists of a thick-bedded quartzite. The argillaceous Midas Formation (150'-500'), the calcareous Mural Formation (400'-750'), and an unnamed argillaceous formation (300') complete the Cariboo Group. The Snowshoe Formation (Holland 1954; Sutherland Brown 1957) was o r i g i n a l l y thought to be the uppermost member of the Cariboo Group, lyi n g above the Midas Formation. However, i t has been subsequently shown (R.B. Campbell 1968, 1969; Mansy and R.B. Campbell 1970) that the Snowshoe Formation i s equivalent to the Kaza Group, and that the Mural Formation had been mapped as the Cunningham Formation. The Slide Mountain Group of Upper Devonian (?) to Lower Mississippian age rests with regional discordance on both Cambrian and Proterozoic rocks. The rocks above the discordance vary from conglomerates i n the Antler Creek area (Sutherland Brown 1957), to limestones and chert breccias i n the Black Stuart Mountain area (Mansy and R.B. Campbell 1970), to a r g i l l i t e s and basic volcanic rocks i n the Crooked Lake area (K.V. Campbell 1971). In the f i r s t two areas the Slide Mountain Group rests unconformably on Cambrian strata but i n the Crooked Lake area i t has unconformable and tectonic contacts with the underlying Proterozoic sediments. The Mesozoic sequence was deposited i n the Quesnel Trough which l i e s to the west of the Omineca Geanticline. To the south of Quesnel Lake a .largely argillaceous sequence (Upper Triassic?) i s followed by a volcanic unit (Upper Triassic and Lower Jurassic). These rocks are unconformable with the underlying Slide Mountain Group. The Mesozoic sequence was presumably eroded from the Omineca Geanticline (Tipper and R.B. Campbell 1970) during Middle or early Late Jurassic times. In the immediate v i c i n i t y of Quesnel Lake the sediments have undergone greenschist and amphibolite facies metamorphism (Fig. 2). An attempt has been made to correlate the stratigraphy of the unmeta-morphosed strata with the rock units of the metamorphic complex. This was accomplished by following known stratigraphic units and structures into the metamorphic complex and by the recognition of metamorphic rock Figure 2. Geology of the Omineca Geanticline i n the region of Quesnel Lake. LEGEND Upper T r i a s s i c and Lower Jurassic mainly v o l c a n i c - c l a s t i c rocks Upper Tr i a s s i c (?) black p h y l l i t e , s c h i s t , and gray quartzite Upper Devonian (?) - Lower Mississippian Slide Mountain Group - p h y l l i t e s , chert, quartzite, limestone, basic volcanics Lower Cambrian and Late Proterozoic Cariboo Group Unnamed Formation Mural Formation Midas Formation - a r g i l l i t e , p h y l l i t e - limestone, shale, dolomite - s i l t s t o n e , shale, p h y l l i t e Yanks Peak Formation - quartzite Yankee Belle Formation - shale, p h y l l i t e , limestone Cunningham Formation - limestone, shale Proterozoic (late) Isaac Formation - p h y l l i t e , limestone, shale, g r i t Kaza Group - feldspathic quartzite, g r i t Proterozoic Quesnel Lake Gneiss - g r a n i t i c gneiss geological contact f a u l t u anticllnorium svnclinorium i/>i/> "»1/7 Isograds - c h l o r i t e and lower b i o t i t e garnet kyanite/staurolite s i l l i m a n i t e — r GEOLOGY OF THE OMINECA GEANTICLINE IN THE REGION OF QUESNEL LAKE units which could be correlated with unmetamorphosed units. The Black Stuart Syncline (Fig. 2) can be traced into staurolite zone rocks where i t is outlined, in the Mount Watt area (Map 3), by a highly folded limestone unit. This limestone can be correlated with the Cunningham Formation and represents the only definite marker horizon i n the whole area. No large quartzite unit could be corre-lated with the Yanks Peak Formation, nor could any second limestone unit be correlated with the Mural Formation. The units of the meta-morphic complex in the Quesnel Lake area were assumed, therefore, to be older than the Yanks Peak Formation. The Yankee Belle Formation, Cunningham Formation, Isaac Formation and the Kaza Group could a l l be present within the metamorphic complex. Of these, only the Cunningham Formation can be r e a d i l y recognised. The sediments of the Kaza Group, however, can be distinguished from those of the Yankee Belle and Isaac Formations by being dominantly arenaceous and bearing minor a r g i l l i t e s and limestones. The Yankee Belle and Isaac Formations, in contrast, are dominantly argillaceous with interbedded limestones and calcareous sediments. The metasedi-ments were therefore mapped as either dominantly arenaceous, or dominantly argillaceous and calcareous. The distribution of these rock types and a cross-section of the Quesnel Lake Area are shown in Map 3, from which i t is apparent that the argillaceous and calcareous unit most probably correlates with the Isaac Formation. The Kaza Group and Isaac Formation were recognized in the meta-morphic complex by the relative proportions of the different lithologies. Even though the schists of both units appear to be identical in hand specimen and thin section, i t was considered possible that the two groups of schists might be distinct chemically. To test this over forty specimens from low grade metasediments of the Kaza Group, high grade schists of the Kaza Group and high grade schists of the Isaac Formation were analyzed (see Map 3 for locations and Appendix I for analyses). The r e s u l t s , plotted on triangular diagrams (Appendix I ) , show that the schists can not be distinguished chemically, precluding whole-rock chemical analysis as a correlative t o o l . The northern Cariboo Mountains are composed of a series of north west plunging synclinoria and a n t i c l i n o r i a , which f l a t t e n and merge to the north-west into a zone of f a u l t i n g and concentric folding (R.B. Campbell 1970). The v a r i a t i o n i n st r u c t u r a l s t y l e within these structures appears to be related to the metamorphic grade, which, i n a broad sense, may be correlated with the stratigraphic l e v e l . The structures vary from concentrically folded and faulted rocks i n the unmetamorphosed and low grade regions, through a zone of shear folding to polyphase flow folded schists and gneisses of the high grades. R.B. Campbell (1970, p. 71) wrote that, "...essentially a l l the deformation resulted from a single, perhaps long continued event." The margin of the Omineca Geanticline and the Quesnel Trough i n the Crooked Lake Area has been folded into synclinoria and a n t i c l i n o r i a (K.V. Campbell 1971). This folding i s believed to be synchronous with the metamorphism, which has l o c a l l y metamorphosed Upper T r i a s s i c sediments of the Trough to kyanite zone. This i s i n marked contrast to the r e l a t i v e l y unmetamorphosed and weakly deformed rocks i n the central part of the Trough. K.V. Campbell (1971) postulated the age of metamorph-ism and deformation to be Jurassic to early Cretaceous. 8 The structures i n the metamorphic complex are es s e n t i a l l y continuations of the major structures found i n the unmetamorphosed sediments surrounding the complex. However, i n the metamorphic complex these structures have been modified by folding coaxial with e a r l i e r f o l d s , which was associated with medium grade metamorphism. This metamorphism reached garnet zone i n the Quesnel Lake area, but to the south i n the Crooked Lake area the staurolite/kyanite zone was attained. The l a t t e r deformation was followed by a post-kinematic metamorphism which reached the s i l l i m a n i t e zone i n the Quesnel Lake area, but did not affect the rocks of the Crooked Lake area (K.V. Campbell 1971). The attitude of the isograds i s variable. In areas of high thermal gradients they are steeply dipping and close together, as i n the Penfold Creek area. In other areas, p a r t i c u l a r l y i n the northwest part of the complex, they are gently dipping and roughly concordant with the regional plunge of the folds. Post-metamorphic normal fa u l t i n g along the north arm of Quesnel Lake has displaced rock units and isograds several thousand feet. The f i n a l period of deformation produced open folds which warp and fracture the pre-existing structures. In certain areas, however, the folding i s more intense and near i s o c l i n a l folds have developed. This deformation i s associated with a period of retrograde metamorphism which i s strongest i n the areas of intense deformation ( i . e . Lynx Bay to Ogden Peak, Map 3). In these areas garnet and s t a u r o l i t e have altered to c h l o r i t e , and kyanite to s e r i c i t e . In places the retrogressive metamorphism has l e f t l i t t l e of the pre-existing mineralogy. Throughout the Quesnel Lake area c h l o r i t e i s associated with the fracture cleavage caused by this f i n a l period of deformation. 9 Relationships of structure to metamorphism are discussed in detai l for the Penfold Creek area in the next chapter. 10 F i e l d Studies The Summer of 1969 was spent as a senior assistant to Dr. R.B. Campbell of the Geological Survey of Canada i n Quesnel Lake area, studying regional geology, stratigraphy, and metamorphism and c o l l e c t i n g specimens to assist i n selection of a smaller area suitable for a c r i t i c a l study of the metamorphism and structure. Penfold Creek area was selected and s i x weeks were spent i n the summer of 1970 on detailed mapping and sample c o l l e c t i o n . Access to the area of detailed study was by d i r t road from Horsefly, B.C. to Haggens Point, by boat up the East Arm of Quesnel Lake, and then by packhorse up Niagara Creek, with the assistance of Mr. Howard Lowry of Horsefly, B.C. In the summer of 1971 an additional two weeks were spent i n the f i e l d with Mr. John Blenkinsop, mapping the gr a n i t i c Quesnel Lake Gneiss which crops out along the shores of Quesnel Lake. Specimens were collected by Mr. Blenkinsop for radiometric dating by whole-rock Rb/Sr method. Preliminary results of age measurements are included i n Part I I of this thesis, with kind permission of Mr. Blenkinsop. PART I. STRUCTURE AND METAMORPHISM OF PENFOLD CREEK AREA 1. Relationships Between Deformation and Metamorphism Penfold Creek area straddles the boundary between gently folded c h l o r i t e zone rocks of the Lanezi Arch and complexly folded s i l l i m a n i t e zone rocks of the Omineca C r y s t a l l i n e Belt (Fig. 2). The t r a n s i t i o n between these two contrasting areas occurs within a very narrow zone (Plate 15); only one and a half miles separate the b i o t i t e and s i l l i -manite isograds. Structural and metamorphic transitions can best be presented by considering f i r s t the stratigraphy of the area, then the elements of each phase of deformation, next the petrographic evidence of the r e l a t i v e timing of tectonic and thermal events, and l a s t l y relationships between fold s t y l e and metamorphic grade. Stratigraphy Rock units which have been mapped i n the Penfold Creek area are composed of the rock types presented i n Table 1. A br i e f description of these units i s given i n the map legend to 'The Geology of Penfold Creek Area' (Map 1), see also Appendix I I . The rocks of the Penfold Creek area can be divided into two li t h o s t r a t i g r a p h i c groups; one where quartzites and subarkoses pre-dominate, and the other where schists and calcareous metasediments predominate. The f i r s t of these groups may be correlated d i r e c t l y with the Kaza Group lying to the north i n the Lanezi Arch. Correlation of the l a t t e r group i s impossible within the l i m i t s of the Penfold Creek area. I t i s shown i n the Introduction to this thesis that this group ROCK TYPE VARIETIES Table 1, Description of the Rock Types from the Penfold Creek Area COLOUR MINERALOGY TEXTURE COMMENTS QUARTZITE PURE FELDSPATHIC MICACEOUS pink, l i g h t gray, l i g h t green l i g h t gray-pinkish l i g h t gray plagioclase <1%; mica <2%. plagioclase < 10%; mica < 32 plagioclase-^ 5%; mica 8-10%. equigranular, medium to f i n e grained medium grained, plagioclase coarser with i r r e g u l a r boundaries medium grained, well f o l i a t e d r e s t r i c t e d to Kaza dominantly i n Kaza i n places cross-bedded dominantly i n Isaac SUEARKOSE (QUARTZOSE SCHIST) l i g h t green gray plagioclase•» 20%; n;ica~15%; green colour due to c h l o r i t e medium to coarse grained, poor f o l i a t i o n green i n lower grades, graded beds common SIIALE (PKYLLITE-MICA SCHIST) quartz «" 35%; plagioclase ~ 152; mica ~ OXIDE RICH gray 40%; ilmenite " U ] one minor bed cont-ained 20% graphite. brown to upto 12% pyrrhotite, p y r i t e , with LSULPHIDE RICH dark gray minor ilmenite; mica-"35%; plagioclase ~10%; quartz~ 40%; grain siz e dependant on grade of metamorphism, very well f o l i a t e d medium grained, well f o l i a t e d , only rare porphryoblasts of garnet, no kyanite, s t a u r o l i t e , or s i l l i m a n i t e garnet, s t a u r o l i t e , kyanite, s i l l i m a n i t e in high grades r e s t r i c t e d to Isaac associated with c a l c -s i l i c a t e layers LIMESTONE (MARBLE) GRAPHITIC PURE IMPURE dark gray l i g h t gray l i g h t to medium brown graphite ~ 5%; quartz <1%; m i c a < l % quartz<5%; plagioclase <1%; mica< 5% graphite < .5%; pyrite*'.5%; fuchsite core to some of the mica quartz-15%; plagioclase " 3%; mica w 10%; g r a p h i t e ~ .2%; py r i t e ~1.5% f i n e grained laminated coarse to f i n e grained, c r y s t a l l i n e , mica-rich bands o u t l i n e f o l i a t i o n , but often no structure i d e n t i f i a b l e medium grained, well f o l i a t e d r e s t r i c t e d to Kaza r e s t r i c t e d to Isaac r e s t r i c t e d to Isaac c a l c - s i l i c a t e miner-a l s i n high grades CALC-SILICATE GNEISS l i g h t gray q u a r t z ~ 50%; calc-silicates« 30%; mica to green ""5%; K-spar/plagioclase 10%; opaques (pyrite, pyrrhotite, i l m e i i i t e ) ~ 6% coarse grained, poor f o l i a t i o n r e s t r i c t e d to Isaac highly v a r i a b l e i n mineral content AMPKI20LITE hornblende *• 70%; quartz 10%; mica «< dark green 3%; plagioclase " 15%; garnet where present has pla g i o c l a s e - r i c h corona medium grained, well f o l i a t e d alignment of hornblende c r y s t a l s common, beds often boudinaged r e s t r i c t e d to Isaac medium to f i n e grained, highly contains same s t r u c t -MYLONITE - l i g h t plagioclase < 1%; quartz~95%; sheared and elongate quartz c r y s t a l s , ural elements as sur-green-gray mica<5%. a few plagioclase augens rounding rocks. 1 3 can be correlated best with the Isaac Formation. The boundary between rocks of the Kaza Group and the Isaac Formation i n the Penfold Creek area, has been taken at the position of shear zone one hundred feet thick. This zone, which consists of sheared quartzites and subarkoses interleaved with mylonites, has undergone the deformational and metamorphic events recorded i n the surrounding rocks. The f a u l t i n g was therefore prior to or synchronous with the f i r s t major period of deformation, and probably was associated with movements which caused the disconformity between the Slide Mountain Group and the underlying Kaza or Cariboo Groups (see Introduction). Structure Three phases of deformation have been recognised i n Penfold Creek area. I d e n t i f i c a t i o n of each fold generation and related s t r u c t u r a l elements i s based on overprinting relationships. Separation of the various fold generations was made d i f f i c u l t by their similar orientations and the v a r i a b i l i t y of fold s t y l e s . Phase I deformation has affected a l l rocks i n Penfold Creek and surrounding areas, and i s believed to be responsible for the major synclinoria and a n t i c l i n o r i a l y i n g to the north (Lanezi Arch and Isaac Lake Synclinorium). Phase I folds plunge gently to the northwest and have steeply dipping a x i a l planes and axial-plane s c h i s t o s i t i e s , (Fig. 3B), except where refolded i n the high grade rocks. Some reversals i n d i r e c t i o n of plunge of Phase 1 folds occur, due to Phase 3 deformation, but they are not s i g n i f i c a n t enough to produce a bimodal d i s t r i b u t i o n i n the stereographic projections of Phase 1 fold axes (Fig. 3C). A strong crenulation l i n e a t i o n subparallel to F l and F2 fol d axes i s common (Fig. 3D). This l i n e a t i o n was assumed to be 1 4 Figure 3. Structural Stereograms from Penfold Creek Area In each description the contour i n t e r v a l , representing percentage per 1 percent area, i s given i n brackets, together with the number of readings. North i s marked on each stereogram. A. Poles to bedding planes from the low grade rocks, defining the orientation of F l folds - axis plunging towards 300° at 18 . (14, 7, 4, and 1%)(85) B. Poles to F l a x i a l planes (+) and axial-plane s c h i s t o s i t y (.). C. Projections of F l fold axes. (15, 10, 5, and 1%)(160). D. Projections of F l lineations. (15, 10, 5, and 1%)(86). E. Poles to sc h i s t o s i t y planes from the high grade rocks, F l a x i a l -plane s c h i s t o s i t y . Folded about F2 axis trending towards 308 plunging at 0°. (8, 5, 2, and 1%)(90). F. Poles to compositional layering from the high grade rocks. Folded about F2 axis trending towards 312 plunging at 10 . (10, 5, 2, and 1%)(92). G. Projections of F2 fold axes (.) and poles to F2 a x i a l planes (+). H. Projections of fold axes (.) and poles to fold a x i a l planes (+) of minor folds from the high grade rocks. I. Poles to F3 fracture cleavage from the low grade rocks. (30, 15, 5, and 2%) (60) . J. Poles to F3 fracture cleavage from the high grade rocks. (20, 10, 5, and 2%)(54). K. Poles to F3 kink bands from the low grade rocks. (15, 10, 5, and 2%) (52). L. Average orientations of the F3 fracture cleavage from both low and high grades, and kink bands. Also approximate positions of the maximum and minimum stress axes associated with F3. 1 5 related to Phase 1 deformation because of the absence of Phase 2 folds in the low grade rocks. In low grade rocks bedding planes have been folded by Phase 1 structures (Fig. 3A), but i n the high grades (above the garnet zone) the major schistosity (Fl axial-plane schistosity) and the compositional layering are subparallel and have been refolded by Phase 2 (Fig. 3E and F). Phase 2 folds are seen only in high grade rocks where they are represented by minor folds which have modified the Phase 1 anticlinoria and synclinoria. They are nearly coaxial with Phase 1, but the orient-ations of the axial planes are much more varied (Fig. 3G). Many minor folds in the high grade rocks could not be definitely assigned to either Phase 1 or Phase 2, and these have been plotted separately (Fig. 3H). In the lower grades no Phase 2 folds were identified. The effect of Phase 2 deformation is represented by the flattening of the Phase 1 folds, which is illustrated in Figure 4. In this figure a similarly folded p h y l l i t i c unit has been distorted in the axial region; measurements of the thickness, T, taken parallel to the axial surfaces, which should be constant for a similar fold, show a marked increase in the axial regions (Ramsay 1962). This increase has been brought about by movement along the F l axial-plane schistosity. The folds in the more quartzose units are concentric, showing maximum T on the limbs. However, there is an obvious thickening in the axial region with respect to measurements taken perpendicular to the bedding planes, which in the case of concentric folds should be constant. The deflection of the schistosity in the p h y l l i t i c units around the hinges of the quartzose units, and the sigmoidal shape of the fracture cleavage in these quartzose units could also be the result of flattening during Phase 2 deformation. Figure 4. Variation in thickness (T) of a phyl l i te unit (unit 1) and a quartzite unit (unit 2) measured para l le l to the fold axial plane. UNIT 2 The f i n a l deformation caused minor warping and b r i t t l e fracturing of the pre-existing structures. In the Penfold Creek area no Phase 3 folds can be recognized, but the cleavages associated with these folds are v i s i b l e . However, one mile to the south-east traces of F3 a x i a l planes may be drawn due to the divergence of F l and F2 fold axes, and twelve miles to the south-east F3 i s o c l i n a l folds have been recognized (see Introduction). The axes of these folds plunge very shallowly towards the north-east and have nearly v e r t i c a l a x i a l planes. This deformation was accompanied by c h l o r i t e zone metamorphism resulting i n some retrogression of the higher grade metamorphic assemblages. A closely spaced F3 fracture cleavage i s common throughout the Penfold Creek area, the orientation of which appears to change abruptly at the kyanite/staurolite isograd. However, when the orientations of the cleavages from the low and high grades, and the kink bands of the lower grades are compared to the primary stress axes of the F3 fol d s , as determined from F3 folds outside the Penfold Creek area, i t i s clear that the two cleavages are different i n o r i g i n (Fig. 3L). The b r i t t l e nature of the low grade rocks allowed two types of fracture to develop; an ac-fracture cleavage at a low angle to F l axial-plane s c h i s t o s i t y , and a series of kink bands at a high angle to F l axial-plane s c h i s t o s i t y . In the high grade rocks, where there i s a prominent s c h i s t o s i t y , only a fracture cleavage at a high angle to the sch i s t o s i t y could develop. Conjugate shear resulting from F3 deformation i s thought to be respon-s i b l e for the kink bands of the low grades and the fracture cleavage of the high grades. 18 Metamorphic Framework The zones of metamorphism displayed i n Map 2 are based on those of the c l a s s i c Barrovian sequence. A l l the rocks i n the Penfold Creek area are metamorphosed to at least c h l o r i t e zone, which gives way to b i o t i t e , garnet, kyanite/staurolite, and s i l l i m a n i t e zones to the south. Staurolite and kyanite appear together i n the sequence, and are l o c a l l y associated with a second generation rim on the f i r s t generation garnets. The sillimanite-bearing rocks have been divided into three zones. S i l l i m a n i t e i s present f i r s t i n Zone 1, where i t i s associated with large muscovite pseudomorphs after garnet and s t a u r o l i t e . Kyanite i s abundant i n Zone 1 and i s apparently stable, judging by i t s unaltered appearance and well-defined grain boundaries. A l l the f i r s t generation garnets are rimmed; the rims are related to the f i r s t appearance of small i d i o b l a s t i c second generation garnets. S i l l i m a n i t e Zone 2 marks the breakdown of kyanite to s i l l i m a n i t e , and hence the isograd between Zones 1 and 2 marks the disappearance of a phase i n contrast to the lower grade isograds which mark the appearance of a certain phase or phases. The isograd i s not perfectly sharp, but kyanite i s generally p a r t i a l l y altered to s i l l i m a n i t e or i s absent from Zone 2 rocks. The f i r s t generation garnet and s t a u r o l i t e become pro-gressively replaced by muscovite and s i l l i m a n i t e i n Zone 2. Zone 3 consists of high grade gneisses and i s characterised by the absence of the f i r s t generation garnets ( a l l altered to s i l l i m a n i t e ) ; the rimming of s t a u r o l i t e by s i l l i m a n i t e ; the replacement of pseudo-morphic muscovite by s i l l i m a n i t e ; the formation of f i b r o l i t e i n b i o t i t e ; and the abundance of second generation garnets, i n places also altered to s i l l i m a n i t e . Petrography and Timing of Mineral Growth In this section an attempt is made to correlate the periods of deformation described in the structural section with the crysta l l i zat ion of the various metamorphic minerals. In so doing the textures of each mineral species are described, together with any characterist ic relationship they may show with the other minerals. Chlorite Chlorite gives to the low grade rocks a dist inct ive green colour seen in the majority of the sediments of the Kaza Group. It i s fine grained and together with muscovite defines the schistosity in the rocks metamorphosed to garnet zone and lower. With increasing grade chlorite tends to form small knots which react with the matrix minerals to form biot i te at the b iot i te isograd. Above the garnet zone chlor ite is present only in small amounts, forming retrogressively from b io t i te , garnet and staurol i te. This retrogressive metamorphism was presumably associated with Phase 3 deformation, because F3 fracture cleavage generally contains elongate chlorite f lakes. Retrograde chlorite is not restr icted to these fractures. Biot ite Porphryoblastic b iot i te occurs in rocks of both b iot i te and garnet zones, whereas in the higher grades b iot i te is present as small laths in the matrix. The porphryoblasts contain a well defined planar internal schistosity (S^) which l ies at about 20° to the external schistosity (S e ) . They are considered to have formed after F l with the representing the axial-plane schistosity of the F l folds. The rotation of the porphryo-blasts took place during F2, which in the low grade rocks was manifest as f lattening of the F l fo lds, giving r ise to renewed move-20 merit along the axial-plane s c h i s t o s i t y of the F l folds. The matrix b i o t i t e s of the higher grades formed syn- and post- F2 and define, together with muscovite, the s c h i s t o s i t y of the higher grade rocks. Only l o c a l l y are the b i o t i t e laths seen to form polygonal arcs around microscopic folds. Muscovite Most of the muscovite, both i n the high and low grades, occurs as small laths p a r a l l e l to the s c h i s t o s i t y . However, there i s a large increase i n size across the kyanite/staurolite isograd. Polygonal arcs defining F2 folds are common i n the higher grade. In addition to the matrix muscovites, large pseudomorphic muscovites are present i n the sillimanite-bearing rocks. These muscovites progressively replace both garnet and s t a u r o l i t e . They have highly irregular grain boundaries and intertongue with the enveloping s c h i s t o s i t y (S2) i n such a manner as to leave no doubt that they formed", after this s c h i s t o s i t y and that no further movement has occurred aiong t h i s s c h i s t o s i t y . In s i l l i m a n i t e Zone 1 the pseudomorphic muscovites contain fine fibres of s i l l i m a n i t e . The proportion of s i l l i m a n i t e i n these muscovites increases as the grade increases, so that i n Zone 3 the pseudomorphs are e n t i r e l y s i l l i m a n i t e . Garnet The growth of garnet took place during two d i s t i n c t periods. The f i r s t generation garnet i s characterised by numerous inclusions which are generally recognizable as a r e l i c t s c h i s t o s i t y . These inclusions are-predominantly quartz and fine-grained ilmenite, with minor plagio-clase, b i o t i t e and muscovite. These f i r s t generation garnets a l l show some degree of rotation. The variety of in t e r n a l structures i s shown i n Figure 5, together with an interpretation of the timing of the garnet 21 Figure 5. Mineral relationships i n the Penfold Creek schists. The following drawings are sketches of the mineral relationships seen i n thin section and are representative of a l l the schists i n the area. In a l l the drawings the l a t h shaped crystals are b i o t i t e (lined) and muscovite (unlined), also the opaque mineral i s ilmenite. In the b r i e f descriptions S^ refers to the in t e r n a l s c h i s t o s i t y , and to the external s c h i s t o s i t y . The scale to the right of each drawing represents 1mm i n a l l cases. 1. B i o t i t e zone, b i o t i t e porphryoblasts with planar S. representing SI, they have been rotated during F2. The matrix, which consists of c h l o r i t e , muscovite, quartz and plagioclase, i s cut by F3 fracture cleavage. 2. Garnet zone, garnet porphryoblast with rotated S., the rotation occurred during F2. The b i o t i t e porphryoblast and the matrix are similar to (1). 3. Kyanite-staurolite zone, syn-F2 garnet with consisting of rotated F l cleavage, between which folded So may be recognized (see Fig. 7 for s i m p l i f i e d sketch of t h i s relationship). Note also the concentration of ilmenite i n the r e l i c t F l cleavage. The axis of rotation of the garnet i s perpendicular to the plane of the paper. Within the S (S2) a post-F2 kyanite (cross-hatched) c r y s t a l has c r y s t a l l i z e d . 4. Thin section from the same specimen as (3), but oriented perpen-dicular to i t . The axis of rotation l i e s within the plane of the paper (or a few degrees from i t ) , and i s indicated by the arrows. 5. F i r s t generation garnet with a planar S., which represents r e l i c t SI, surrounded by a second generation garnet rim. The other garnet shows a similar relationship but there i s no recognisable r e l i c t SI i n the f i r s t generation garnet. Note the i d i o b l a s t i c form of the post-tectonic second generation garnet rim. 6. Second generation garnet growth on a post-tectonic f i r s t generation garnet which has a S. of r e l i c t F l folds. 22 Figure 5 (cont). 7. Kyanite/staurolite zone, post-tectonic kyanite (cross-hatched) and s t a u r o l i t e (stippled) cross-cutting and mimetic i n S2. Note also the e p i t a x i a l growth of kyanite and s t a u r o l i t e . 8. S i l l i m a n i t e Zone 1, small second generation i d i o b l a s t i c garnets enclosed i n s t a u r o l i t e and kyanite which have been p a r t i a l l y altered to pseudomorphic muscovite and fibrous s i l l i m a n i t e . The pseudomorphic muscovites penetrate the S and therefore no p o s t - c r y s t a l l i n e movement has occurred along t h i s schistosity-(Plates 8, 10 and 11). 9. S i l l i m a n i t e Zone 1 f i r s t generation garnet with planar S. rimmed by a 'blocky' second generation garnet which i s cognate with the small isolated i d i o b l a s t i c second generation garnets (Plate 12). 10. S i l l i m a n i t e Zone 2, abundant fibrous s i l l i m a n i t e within pseudo-morphic muscovite, together with some c r y s t a l l i n e s i l l i m a n i t e . The a l t e r a t i o n of kyanite and s t a u r o l i t e to f i b r o l i t e i s shown by the r e l i c t cleavages i n the f i b r o l i t e . Note also the abundant s i l l i m a n i t e within the inclusions of the f i r s t generation garnet (Plate 9). 11. S i l l i m a n i t e Zone 3, nearly complete a l t e r a t i o n of a s t a u r o l i t e c r y s t a l to s i l l i m a n i t e , no pseudomorphic muscovite i s present. The small i d i o b l a s t i c second generation garnet shows no reaction relations (Plate 10). 12. S i l l i m a n i t e Zone 3, s i l l i m a n i t e formation i n b i o t i t e and a second generation garnet (Plate 13). 2 3 growth. I t i s seen that the f i r s t generation garnets started to grow after the cessation of F l and continued during the second period of deformation. Displacement of the surrounding matrix minerals and pressure shadows are common features of the f i r s t generation garnets. The second stage of garnet growth occurred after F2, with the garnets either rimming the f i r s t generation garnets or forming new c r y s t a l s . The second generation garnets are characterised by their low content of inclusions, i d i o b l a s t i c form, and thei r f a i l u r e to displace the surrounding minerals within the s c h i s t o s i t y . The rims, although l o c a l l y t h i n , convert the xenoblastic f i r s t generation garnets into i d i o b l a s t i c c r y s t a l s . The isolated second generation garnets are small (approximately 1mm) and i d i o b l a s t i c . The relationship between the rims and the isolated garnets i s w e l l i l l u s t r a t e d i n Figure 5 -9 , where the 'blo.cky' nature of the rim has i t s counterpart i n the i d i o b l a s t i c isolated crystals (Plate 12). Staurolite Staurolite c r y s t a l l i z e d after F2 and after the c r y s t a l l i z a t i o n of the second generation garnet. The timing of s t a u r o l i t e c r y s t a l l i z a t i o n i s indicated by some of the s t a u r o l i t e crystals cross-cutting the major sc h i s t o s i t y (S2) and by others including isolated second generation garnets. Staurolite also c r y s t a l l i z e d around the i d i o b l a s t i c rims of the second generation garnets i n places (Plate 11). Most of the crystals contain few inclusions, but a few are highly p o i k i l o b l a s t i c . Staurolite commonly shows e p i t a x i a l growth with kyanite. 2k Kyanite Kyanite c r y s t a l l i z e d at the same time as s t a u r o l i t e and shows many of the same textures. It has, however, fewer inclusions and i s more commonly present as isolated crystals than associated with second generation garnets. S i l l i m a n i t e Both the fibrous ( f i b r o l i t e ) and the c r y s t a l l i n e form of s i l l i -manite are present i n the high grade rocks. S i l l i m a n i t e f i r s t appears within the pseudomorphic muscovites as minute needles displayed i n a 'horse-tail' texture. These needles become more abundant with increasing grade, so that i n s i l l i m a n i t e Zone 3 large f i b r o l i t e augen are present. Within these augen i t i s not uncommon to find areas of coarsely c r y s t a l l i n e s i l l i m a n i t e . At different grades s i l l i m a n i t e i s seen also to be replacing garnet, s t a u r o l i t e , and kyanite. In the highest grades, f i b r o l i t e i s commonly found i n b i o t i t e laths. The replacement textures are displayed and described i n Figure 5. Except for the retrogressive minerals of F3, s i l l i m a n i t e c r y s t a l l i z e d l a s t i n the main sequence of metamorphic minerals. Ilmenite Ilmenite occurs as small laths and xenoblastic crystals i n a l l grades of metamorphism. I t i s also concentrated i n the f i r s t generation garnets as inclusions, where i t i s present mainly i n the r e l i c t F l s c h i s t o s i t y planes. These ilmenite inclusions tend to coalesce where the garnet has been replaced by pseudomorphic muscovite. Fl PRE- SYN-TECTONIC POST-=F F2 SYN- POST-TECTONIC F3 SYN- POST-TECTONIC CHLORITE BIOTITE MUSCOVITE GARNET KYANITE STAUROLITE SILLIMANITE I -4 I 1 4 H I 1 H IH I—I Figure 6 . Timing of mineral growth with respect to the period of deformation TECTONIC STYLE MINERAL GROWTH SYN = POST' r ^ ^ ^ x X [ BIOTITE J 0>A S „ \ \ \ > ^ ^ \ G A R N E T ( I ) r4)\ :4> \ G 4 R N E T < " - o ^ r r - \ ^ x S ° * S | F 2 P 0 S T ' GARNET(2) ^ ^ ^ T A U R O L I T C F 3 • • ^ v ~ : s 3 S 3 <5YN: ==- ^ — CHLORITE FIGURE 7. RELATIONSHIP BETWEEN PERIOD OF DEFORMATION AND MINERAL GROWTH 2 7 Relationships Between Fold Style and Metamorphic Grade The structural transitions observed in the Penfold Creek area may be correlated with increasing metamorphic grade. Prior to Phase 2 deformation the Quesnel Lake area consisted of large Phase 1 anti-clinoria and synclinoria, similar to those seen to the north at present. A strong axial-plane schistosity, which is commonly fanned, is characteristic of the p h y l l i t i c members of these folds (Plates 1 and 2). Depending on the rock competency the folding was either dominantly similar or concentric in form. This style may be considered to be that of the rocks of the chlorite zone and below. Phase 2 deformation is associated with the major regional meta-morphism of the area and is essentially restricted to the higher grade rocks. Only Phase 1 folds were recognized in the chlorite and biotite zones. These folds, however, have been flattened by the Phase 2 deformation. This flattening becomes more intense with increasing grade, so that at the garnet isograd the F l folds are iso c l i n a l with the axial-plane schistosity parallel to the fold limbs (Plates 2 and 5). A l l the rocks which have been metamorphosed to garnet zone and above have been refolded during Phase 2 deformation (Plate 4) with the F2 folds becoming more frequent with increasing grade. The F2 folds f i r s t appear as minor folds which are either concentric or similar but with no axial plane schistosity. In the rocks of sillimanite Zone 3 flow folding has become the dominant feature (Plate 6). This sequence of changing fold style is shown in Fig. 8 where representative sketches of the folds have been placed at their positions in the f i e l d . The variation, in fold style i s due to increased temperature and possibly fluid pressure towards the higher grade rocks. The increased ductility of the high grade rocks, under the stress 2*45^ B0» 33* 120*30' Figure 8. Fold Styles in the Penfold Creek Area All profiles drawn looking along the fold axes. The scale in all the profiles represents one foot. 2 9 f i e l d of F l , produced fla t t e n i n g of F l folds followed by coaxial folding. Similar variations i n fold s t y l e have been described to the south at the margin of the Shuswap metamorphic complex (Fyson 1970). This v a r i a b i l i t y i n fold s t y l e from upright folds i n the low grade p h y l l i t e s to recumbent i s o c l i n e s i n the high grade gneisses has been attributed (Fyson 1970) to the prevalence at upper and lower levels of folds of different generations. The low grade p h y l l i t e s contain small-scale discontinuous isoclines (Fl) which have been refolded by upright folds (F2). Thus the low and high grade rocks are considered to have undergone the same deformational sequence. In the Penfold Creek area no evidence has been found i n the low grade rocks for a period of folding prior to that designated F l i n this thesis. The layering contains w e l l developed sedimentary structures (cross-bedding and graded-bedding) and i s considered to be bedding. In the Black Stuart and Isaac Lake synclinoria the low grade and unmetamorphosed Lower Cambrian strata contain no evidence for a period of i s o c l i n a l folding or transposition prior to that which formed the synclinoria themselves (Mansy and R.B. Campbell 1969). 2. Mineralogy of the P e l i t i c Rocks Methods of Study Fifteen specimens from different metamorphic grades within the Penfold Creek area were selected for a n a l y t i c a l work on the constituent minerals. In choosing these specimens emphasis was placed on the sillimanite-bearing rocks. A l l the major minerals were analyzed, except 3 0 quartz, kyanite, and s i l l i m a n i t e which were assumed to have nearly stoichiometric compositions. Of the accessory minerals only ilmenite was analyzed. The analyses were carried out on the 5-channel A.R.L. microprobe at the University of Washington, Seattle, under the d i r e c t i o n of Dr. B.W. Evans. Zoning was detected only i n garnet and plagioclase, and detailed probe traverses were performed on these two minerals. The analyses of the other minerals have been averaged, where possible, from counts on at least ten different grains per polished section, with several counts on each grain. In addition, the minerals from two specimens have been analyzed i n duplicate, using different polished thin sections, so as to test for a n a l y t i c a l errors and l o c a l composi-t i o n a l variations. The instrumental conditions for the analyses were: an acceleration potential of 15 KV, specimen current between 0.05 and 0.10 microamps, integrating time of 10 sees., and an electron beam diameter of approximately 2 microns, except for plagioclase where a beam diameter of 15 microns was used. The specimen current was recorded throughout the analyses so that points with low conductivity could be discarded. Analyzed natural and synthetic minerals were used as standards. Wherever possible the standards and unknowns had similar mean atomic numbers. A l l counts were corrected for instrument d r i f t , background and dead time. Two methods were used for further corrections and computation of the percent oxide. Where the number of suitable standards was s u f f i c i e n t to graph counting rate against percent oxide, the percent oxide was read d i r e c t l y from the graph. The graph was drawn by a computer using a least squares f i t to the standard points. 3 1 Although this method was used for garnet, plagioclase, and c h l o r i t e , additional standards were used for elements of low concentration i n ch l o r i t e (see Appendix III). For a l l the other minerals this graphing technique was considered unreliable and corrections for fluorescence, atomic number and absorption were made using the Rucklidge Program EMDRO (Rucklidge 1969). The standards used for the mineral analyses are a l l from Dr. B.W. Evans' c o l l e c t i o n and are tabulated i n Appendix i l l . The d i s t i n c t i o n between ferrous and f e r r i c i r o n , and the analysis of H^ O i s impossible on the electron microprobe. In a l l the analyses iron has been calculated as ferrous. Possible v a r i a t i o n and value of the fe r r o u s / f e r r i c r a t i o i s discussed i n d i v i d u a l l y for each mineral. The weight percent H^ O has been assigned by assuming a theoretical value, except for s t a u r o l i t e where the percent has been estimated so that oxides t o t a l to 100%. Mineral Assemblages A l l the f i f t e e n specimens of p e l i t i c rocks from the Penfold Creek area contain quartz, plagioclase feldspar, muscovite and ilmenite. The different p e l i t i c mineral assemblages are i d e n t i f i e d by the name of the zone i n which they occur. The assemblages are not necessarily i n equilibrium, but merely state the minerals found within each zone. Equilibrium considerations are discussed l a t e r . Five assemblages have been recognized: 1. (c h l o r i t e zone) quartz-plagioclase-muscovite-chlorite-ilmenite-calcite 2. ( b i o t i t e zone) quartz-plagioclase-muscovite-chlorite-biotite-ilmenite-calcite 3 2 3. (garnet zone) quartz-plagioclase-muscovite-chlorite-biotite-garnet-ilmenite A. (kyanite/staurolite zone) quartz-plagioclase-muscovite-biotite-garnet-kyanite-staurolite-ilmenite 5. ( s i l l i m a n i t e zone) quartz-plagioclase-muscovite-biotite-garnet-staurolite-sillimanite-ilmenite±kyanite±pseudomorphic muscovite The sillimanite-bearing rocks have been divided into three zones dependent upon s i l l i m a n i t e replacing f i r s t kyanite and then pseudomorphic muscovite. The three assemblages of the s i l l i m a n i t e zone are: a. (zone 1) quartz-plagioclase-muscovite-biotite-kyanite-staurolite-garnet-sillimanite-ilmenite-pseudomorphic muscovite b. (zone 2) quartz-plagioclase-muscovite-biotite-staurolite-garnet-sillimanite-ilmenite±kyanite±pseudomorphic muscovite c. (zone 3) quartz-plagioclase-muscovite-biotite-staurolite-garnet-sillimanite-ilmenite. The modal analyses of the specimens under consideration (Table 2) were calculated from at least 1500 points counted on an e l e c t r i c a l point counter. I t includes, apart from the minerals mentioned above, the accessory minerals; tourmaline, apatite, zircon, epidote and r u t i l e . Table 2 Modal Analyses of the Analysed Specimens from the Penfold Creek Area. Spec, Kos. 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 Quartz 58.2 22.4 42.1 34.0 38.2 36.5 12.8 28.2 19.3 15.3 27.6 15.8 6.5 71.7 30.8 Plagioclase 16.0 tr 9.8 8.1 16.3 5.9 10.6 15.3 9.2 18.5 5.1 19.9 49.3 12.1 10.2 Bio tite - 12.4 3.8 5.6 12.5 21.6 21.5 21.6 23.7 27.5 25,5 30.3 29.3 10.2 18.2 Muscovite11'' 10.6 58.7 29.4 34.4 22.4 11.2 7.8 19.9 20.2 18,6 13.1 7.3 4.3 2.5 28.1 Muscovite? - - - - - 4.5 2.3 8.1 3.4 9.7 8.8 - - - -Chlorite 1 9.6 5.9 12.9 14.7 - - - - - - - - - - -Chlorite 2 - - - - tr 3.6 0.8 1.5 0.4 tr 3.5 tr tr - tr Garnet - - - 1.9 8.8 5.6 26.7 2.5 7.9 7.0 5.2 5.9 4.5 3.3 3.4 Staurolite - - - - tr 8.6 1.4 1.9 0.7 1.9 1.4 2.4 0.7 - 0.4 Kyanite - - - - 0.4 2.1 15.4 0.1 10.1 - 0.5 7.5 4.0 - -Sillimanite - - - - - - - - 4.4 - 8.7 13.9 0.5 tr 7.9 Ilmenite 1.0 0.6 1.9 1.1 1.3 0.4 0.5 0.8 0.7 1.5 0.5 2.5 1.0 0.2 0.7 Tourmaline 0.1 tr 0.1 tr 0.1 tr 0.2 0.1 tr tr 0.1 - tr tr 0.3 Apatite - - tr tr tr tr - tr tr tr tr - - tr tr Zircon 0.1 - tr tr tr tr -• tr tr tr tr tr - tr tr Calcite 4.4 tr tr - - - - - - - - - - - -Epidote - - - 0.2 - - - - - - - - - - -Rutile - - - - - tr - - tr tr - tr 0.1 - -Muscovite111 - matrix muscovite. MuscoviteP - pseudomorphic muscovite. Chlorite^ - prograde chlorite. Chlorite 2 - retrograde chlorite. Note: pseudomorphic muscovite contain sillimanite, only f i b r o l i t e knots and crystalline sillimanite are registered in the sillimanite mode, no distinction has been made between chlorite' and chlorite i n specimens 1 - 4 . A l l the chlorite in these specimens has been reported as chlor i t e 1 . Mineralogy B i o t i t e Seventeen analyses were performed on b i o t i t e from fourteen l o c a l i t i e s (Fig. 9); a l l the b i o t i t e s being from the schist 'matrix,' except Specimen 7a which i s a single b i o t i t e l a t h completely enclosed i n the garnet c r y s t a l (Table 3). If a theoretical A wt.% IL^ O i s added to the oxide analyses, the totals l i e between 98.34 and 101.45 wt.%, with the majority being within 1% of 100%, Structural formulae were calculated on a basis of 22 oxygens, with the assumption that the t e t r a -hedral s i t e s were completely f i l l e d by Si"*"V and A 1 1 V with the remainder of the aluminum occupying octahedral s i t e s . The t o t a l number of ions i n octahedral site s are consistently lower than the possible s i x per formula u n i t , implying the b i o t i t e s are intermediate between t r i - and di-octahedral micas (Deer et a l . p. 61, 1962). The colour of b i o t i t e varies l i t t l e from i t s i n i t i a l medium brown colour (X = very l i g h t brown, Y = Z_ = medium brown) i n the b i o t i t e zone rocks up to the appearance of s i l l i m a n i t e . Between s i l l i m a n i t e Zones 1 and 2 b i o t i t e changes i t s pleochroism to X = light' brown, Y_ = _Z = medium to dark red-brown. A similar colour change has been shown by Hayama (1959) to be due to an increase i n the TiC^ content. He also found that high f e r r i c iron contents gave a green colour to the b i o t i t e . As no green b i o t i t e has been observed i n the Penfold Creek schists i t suggests that the b i o t i t e s have a low, and perhaps constant, Fe^O^ content. The v a r i a t i o n of b i o t i t e composition with increased metamorphic grade appears to be twofold. F i r s t the MgO wt.% i n the b i o t i t e s increases from 8.5% i n the b i o t i t e zone to a maximum of 11.4% i n Table 3 Chemical Analyses of b i o t i t e Spec. Kos. 2 3 4 5 6 7 7a 7b 8 9 10 11 12 13 14 . 13 15b S i 0 2 36.6 37.1 37.5 36.2 39.2 35.2 37.3 38,1 37.4 37.1 36.1 36.9 37.6 36.2 36.3 37.6 37.1 T i 0 2 1.75 1.77 1.38 2.00 1.68 1.84 1.57 1.61 1.77 1.82 1.82 1.99 1.93 2.04 2.35 2.02 2.38 Al 203 17.7 18.4 18.8 19.0 19.4 19.7 19.5 19.0 18.6 19.4 18.4 18.6 19.8 19.5 19.4 19.7 18.6 FeO* 20.5 20.3 21.1 18.3 17.1 17.7 17.4 17.6 17.8 16.3 19.1 17.5 17.9 18.0 19.7 20.7 20.4 MgO' 8.74 8.51 8.85 9.91 10.8 11.0 10.8 10.5 10.9 10.7 11.4 10.7 10.5 10.8 9.18 8.42 8.b0 MnO 0.12 0.12 0.09 0.01 0.02 0.01 0.C5 0.07 0.01 0.00 0.01 0.02 0.03 0.03 C.04 0.02 0.02 CaO 0.02 0.0.1 0.00 0.03 0.01 0,01 0.06 0.01 0.01 0.03 0.01 0.02 0.01 0.03 0.02 0.02 0.02 Na20 0.13 0.11 0.20 0.19 0.28 0.30 0.22 0.25 0.22 0.20 0.23 0.18 0.28 0.27 0.26 0.25 0.23 K 20 8.78 8.84 8.57 8.98 8.95 8.84 9.02 8.75 8.88 8.93 8.63 9.01 8.81 8.74 8.95 8.72 8.79 Total 94.34 95.16 96.49 94.62 97.44 94.60 95.92 95.89 95.79 94.98 95.70 94.92 96.86 95.61 96.20 97.45 96.34 +4% H 20 98.34 99.16 100.49 98.62 101.44 98.60 99.92 99.89 99.79 98.98 99.70 98.92 100.86 99.61 100.20 101.45 100.34 formulae on the basis of 22 oxygens Si 5.63 5.64 5.61 5.49 5.68 5.34 5.54 5.65 5.57 5.56 5.44 5.55 5.53 5.42 5.45 5.56 5.56 A 1 I V 2.37 2.36 2.39 2.51 2.32 2.66 2.46 2.35 2.43 2.44 2.56 2.45 2.47 2.58 2.55 2.44 2.44 A 1 V I 0.83 0.93 0.94 0.89 1.00 0.86 0.96 0.96 0.87 0.97 0.70 0.85 0.96 0.86 0.87 0.99 0.85 T i 0.20 0.20 0.16 0.23 0.18 0.21 0.18 0.18 0.20 0.20 0.21 0.23 0.21 0.23 0.26 0.22 0.27 Fe 2.63 2.57 2.64 2.32 2.08 2.25 2.16 2.18 2.21 2.10 2.40 2.21 2.20 2.26 2.48 2.56 2.56 Mg 2.00 1.92 1.98 2.24 2.34 2.48 2.39 2.32 2.42 2.38 2.56 2.40 2.30 2.40 2.05 1.86 1.97 Mn 0.02 0.01 0.01 0.00 0.00 0.00 0.01 0.01 0.00 0.00 0.00 0.00 0.00 0.00 0.01 0.00 0.00 5.68 5.63 5.73 5.68 5.60 5.80 5.70 5.65 5.70 5.65 5.87 5.69 5.67 5.75 5.67 5.63 5.65 Ca 0.00 0.00 0.00 0.00 0.00 0.00 0.01 0.00 0.00 0,00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Na 0.04 0.03 0.06 0.06 0.08 0.09 0.06 0.07 0.06 0.06 0.07 0.05 0.08 0.08 0.08 0.07 0.C7 K 1.72 1.71 1.64 1.74 1.66 1.71 1.71 1.66 1.69 1.71 1.66 1.73 1.65 1.67 1.71 1.65 1.68 1.76 1.74 1.70 1.80 1.74 1.80 1.78 1.73 1.75 1.77 1.73 1.78 1.73 1.75 1.79 1.72 1.75 FeO* - t o t a l i r o n as FeO. a - enclosed b i o t i t e . b - duplicate specimens, analyst - C.J.N.Fletcher. V_0 I ^—* 1— " I20"3S' 120*30 Figure 10. MgO content of b io t i te . (vt.'fo) 37 s i l l i m a n i t e Zone 1, and f a l l s again to 8.4% i n s i l l i m a n i t e Zone 3 (Fig. 10). This reversal i n the trend of the MgO content i n b i o t i t e with increasing metamorphic grade has been noted by Guidotti (1970), who put the maximum MgO content close to the Upper S i l l i m a n i t e isograd. In the Penfold Creek area, however, the maximum occurs near the boundary between s i l l i m a n i t e Zones 1 and 2, which i s at a s l i g h t l y lower grade than that noted by Guidotti. Work by Atherton (1968) and Lambert (1959) confirms the MgO Increase i n b i o t i t e for the b i o t i t e , garnet, and kyanite/staurolite zones. The FeO content ( t o t a l iron as FeO) of b i o t i t e shows a trend antipathetic to that of MgO (Fig. 11). The Ti02 content of b i o t i t e increases from approximately 1.4 wt.% to 2.4 wt.% with increased metamorphic grade. The change being an average increase rather than an exact relationship (Fig. 12). The changes appear to be discontinuous between the zones which are below the s i l l i m a n i t e isograd, but b i o t i t e s from the three s i l l i m a n i t e zones show a continuous v a r i a t i o n i n TiO^ content. This Ti02 v a r i a t i o n i s similar to that shown by Evans and Guidotti (Evans and Guidotti 1966; Guidotti 1970) for the b i o t i t e s from Maine. There TiO^ increases from 1.5 wt.% i n the Upper Staurolite zone to 2.5 wt.% i n the Upper S i l l i m a n i t e zone ( s l i g h t l y higher grade than s i l l i m a n i t e Zone 3 rocks i n the Penfold Creek area). MnO wt.% appears to decrease from a high value i n the b i o t i t e zone rocks to a uniformly low value i n the kyanite/staurolite zone and above. As most of the b i o t i t e analyses are from high grade rocks, this trend might be fortuitous. The included b i o t i t e l a t h i n specimen 7 i s interesting i n that a lower grade composition has been 'frozen i n ' by the garnet. This would suggest that progressive metamorphism i n i t s s t r i c t sense has taken Figure 12. TiU2 content of b iot i te (upper number) and muscovite (lower number) (wt.$) place., that i s the high grade minerals have formed from the low grade ones, and not from the minerals of the unmetamorphosed sediment. Muscovite The majority of muscovite grains that have been analyzed are from the matrix, but four pseudomorphic muscovites have also been analyzed (Table 4). If a theoretical water content of 4.5 wt.% i s added to the oxide analyses most of the totals l i e within 1% of 100%. ¥e^0^ c o n t e n t s have not been determined, but they must be considerably lower than the high f e r r i c iron muscovites (up to 6.5 wt.% Fe^O^) described by Lambert (1959) and Mather (1970) i n the low grade schists from Scotland, because the maximum t o t a l iron as FeO i n the Penfold Creek muscovites i s 2.67 wt.%. The s t r u c t u r a l formulae were calculated on a basis of 22 oxygens; i n a l l the muscovites the octahedral s i t e occupancy i s s l i g h t l y greater than 4.00, and i n general their interlayer s i t e occupancy i s less than 2.00. This conforms with several investigations on metamorphic muscovites (Lambert 1959; Evans and Guidotti 1966). Most of the compositional variations i n muscovite may be described i n terms of two s o l i d solution series; the muscovite-paragonite series involving Na replacing K, and the muscovite-phengite series where A 1 1 V A 1 V 1 = ( F e M g ) V 1 S i 1 V (phengite i s also considered to be an i n t e r -mediate member of the muscovite-celadonite s e r i e s ) . The formulae of the end members are: To calculate the molecular percentage of each of these end-members muscovite paragonite phengite " K 2 A 1 4 ( S 1 6 A 1 2 ) O 2 0 ( O H ) 4 - N a 2 A l 4 ( S i 6 A l 2 ) 0 2 Q ( 0 H ) 4 - K 2(FeMg)Al 3(Si 7Al)0 2 ( )(0H) 4 the following procedure was adopted: Table 4 Chemical Analyses of Muscovite Spec, Nos. 6p 7b 8 8p 9p 10 lOp 11 12 13 14 15 T102 FeO* MgO MnO CaO Ka20 K2O 0.44 0.36 0.40 0.37 0.71 0.41 0.50 0.54 0.49 0.64 0.41 0.77 0.45 0.73 0.21 0.67 0.68 0.56 0.78 0.86 2.21 2.21 1.61 1.32 0.00 0.00 0.03 0.02 0.57 0.57 2.28 1.58 0.01 0.02 0.54 9.35 10.3 10.1 2.67 1.12 0.00 0.04 0.66 9.97 1.33 0.87 0.00 0.01 1.04 1.19 0.73 0.00 0.01 1.80 0.84 0.59 0.00 0.01 1.97 1.11 0.74 0.00 0.04 1.29 1.11 0.78 0.00 0.02 1.66 1.24 0.77 0.00 0.01 1.85 0.97 0.58 0.00 0.00 2.15 9.44 8.67 8.42 9.33 9.54 9.08 8.71 1.11 0.74 0.00 0.01 1.48 9.00 1.00 0.62 0.C0 0.01 1.68 1.26 0.88 0.00 0.01 1.62 0.98 0.53 0.00 0.01 1.89 1.19 0.75 0.00 0.02 1.40 1.11 0.69 0.00 0.02 1.65 1.17 0.68 0.00 0.01 1.66 1.29 0.72 0.00 0.03 1.20 1.28 0.68 0.00 0.02 1.50 15b Si02 47.5 46.7 46.6 46.2 46.0 45.8 45.9 46.2 46.5 45.7 46.3 44.9 44.5 46.3 46.2 45.9 45.0 45.7 46.2 46.2 46.0 0.76 AI2O3 33.2 34.2 33.6 33.3 35.6 36.5 36.9 36.0 35.6 35.7 36.3 35.9 36.5 35.2 36.4 34.8 36.4 36.6 36.4 35.7 35.7 1.38 0.66 0.00 0.02 1.65 8.82 9.54 9.26 9.14 9.02 8.20 9.43 9.07 8.96 Tot a l 94.91 95.68 95.13 94.33 95.00 95.13 95.13 95.25 95.70 94.99 95.43 93.91 93.58 95.54 95.48 93.87 94.57 94.58 96.05 95.31 95.15 +4.5% H 20 99.41 100.18 99.63 98.83 99.50 99.63 99.63 99.75 100.20 99.49 99.93 98.41 98.08 100.04 99.98 98.37 99.07 99.08 100.55 99.81 99.65 formulae on the basis of 22 oxygens Si AlIV AlVI T i Fe K* Mn 6.32 1.68 3.52 0.04 0.24 0.32 0.00 6.22 1.78 3.58 0.04 0.24 0.26 0.00 24 76 3.53 0.04 C.25 0.32 0.00 6.24 1.76 3.54 0.04 0.30 0.22 0.00 6.25 1.75 3.63 0.07 0.15 0.18 0.00 21 79 3.71 0.04 0.13 0.15 0.00 6.20 1.80 3.74 0.05 0.09 0.12 0.00 12 88 3.73 0.05 0.12 0.15 0.00 6.15 1.85 3.69 0.05 0.12 0.15 0.00 6.08 1.92 3.68 0.06 0.14 0.15 0.00 6.11 1.39 3.76 0.04 0.11 0.11 0.00 6.04 1.96 3.73 0.08 0.12 0.15 0.00 5.99 2.01 3.74 0.05 0.11 0.12 0.00 6.14 1.86 3.64 0.07 0.14 0.17 0.00 6.11 1.89 3.78 0.02 0.11 0.10 0.00 6.17 1.83 3.69 0.07 0.13 0.15 0.00 01 99 3.74 0.07 0.12 0.14 0.00 6.07 1.93 3.79 0.05 0.13 0.14 0.00 6.08 1.92 3.72 0.08 0.14 0.14 0.00 6.11 l.*>9 3.69 0.09 0.14 0.13 0.00 10 90 3.69 0.08 0.15 0.13 0.00 4.12 4.12 4.14 4.10 4.03 4.03 4.00 4.05 4.01 4.03 4.02 4.OS 4.02 4.02 4.01 4.04 4.07 4.11 4.08 4.05 4.05 Ca Na K Muscovite. Paragonite Phengite 0.00 0.15 1.60 1.75 35.4 8.6 56.0 0.00 0.15 1.74 1.89 42. 7. 50. 0.00 0.14 1.73 1.87 35.5 7.5 57.0 0.00 0.17 1.72 1.89 39.5 8.5 52.0 0.00 0.28 1.63 1,91 52. 14. 33. 0.00 0.47 1.50 1.97 48.2 23.8 28.0 0.00 0.52 1.46 1.98 52.8 26.2 21.0 0.00 0.33 1.64 1.97 0.00 0.42 1.61 2.03 0.00 0.48 1.54 2.02 0.00 0.55 1.47 2.02 0.00 0.49 1.54 2.03 molecular percent end-members 56.2 16.8 27.0 52.3 20.7 27.0 47.2 23.8 29.0 50.8 27.2 22.0 48.9 24.1 27.0 0.00 0.44 1.52 1.96 54.6 22.4 23.0 0.00 0.41 1.61 2.02 48. 20. 31. 0.00 0.48 1.56 2.04 55. 23. 21. 0.00 0.36 1.57 1.93 53.2 18.8 28.0 0.00 0.43 1.54 1.97 50. 21. 28. 0.00 0.43 1.39 1.82 47.4 23.6 27.0 0.00 0.31 1.58 1.89 55. 16. 28. 0.00 0.39 1.53 1.92 52. 20. 27. 0.00 0.43 1.52 1.95 49.9 22.1 28.0 FeO* - t o t a l iron as FeO. p - pseudomorphic muscovite. b - duplicate specimens, analyst - C.J,N.Fletcher. 1. molecular percent paragonite equals Na/(Na + K) X 100. 2. molecular percent phengite equals Fe + Mg/1.0 X 100. 3. molecular percent muscovite equals K^ Na + KJ X 100 - phengite %. If a theoretical value of 2.00 for (Na + K) had been used instead of the observed values, the totals for the three end-members would not have been 100%. This would suggest that the proposed end-members do not describe fu l l y the muscovite compositions. However, the procedure which has been adopted conforms with methods used in other published papers. The paragonite content of muscovite is f a i r l y constant up to the kyanite/staurolite isograd, i t then increases rapidly up to sillimanite Zone 1 (Fig. 15). The paragonite content in the sillimanite zones appears to increase slightly through Zone 1 and part of Zone 2, and then decrease through the rest of Zones 2 and 3. Even though this compositional reversal cannot be proved from the data presented here; i t is consistent with the results from higher grade rocks in other areas (Evans and Guidotti 1966; Guidotti 1970). These results show a decrease in the paragonite content from 18 molecular percent in the lower sillimanite zone to 8.5 molecular percent in the sillimanite-potash feldspar zone. The phengite component of muscovite decreases with increasing metamorhpic grade (Fig. 16) up to the appearance of sillimanite; further increase in grade appears to give no systematic decrease in phengite component. In fact the sillimanite zone matrix muscovites are sur-prisingly constant in composition with respect to the phengite component. As with the paragonite end-member there is a sharp break in the phengite percentage trend at the kyanite/staurolite isograd. A l l the muscovites FIGURE 13. PLOT OF MUSCOVITE ANALYSES IN PART OF THE SAF TRIANGLE. have been plotted on two v a r i a t i o n diagrams (Figs. 13 and 14) to show how close the analyzed muscovites l i e to the theoretical j o i n between the phengite and muscovite end-members. The TiO^ content of muscovite appears to follow that of b i o t i t e i n that i t increases from 0.4 wt.% to 0.8 wt.% as the grade increases from c h l o r i t e zone to s i l l i m a n i t e Zone 3. The rapid compositional change of the muscovite end-member at the kyanite-staurolite isograd i s also seen by a rapid increase i n the TiC^ content. The four pseudomorphic muscovites which have been analyzed are consistently different i n composition from the matrix muscovites. They appear to have higher paragonite and lower phengite contents than the matrix muscovites (Fig. 14). The TiO^ content of the pseudo-morphic muscovites i s low i n comparison to the matrix muscovites. The TiC^ content suggests formation at a lower grade than the matrix muscovites while the paragonite-phengite contents suggest the reverse. The TiC^ content may be explicable i f pseudomorphic muscovites formed d i r e c t l y from garnet and s t a u r o l i t e , both of which have r e l a t i v e l y low TiC>2 contents. The prograde o r i g i n of these pseudomorphic muscovites i s i n accord with the work of Guidotti (1968) except that he found no compositional difference between them and the matrix muscovites. The difference i n composition between the two types of muscovites i n the same rock strongly suggests a lack of chemical equilibrium i n these schists. Chlorite Analyses of four c h l o r i t e s , ranging from c h l o r i t e to garnet zone, are presented i n Table 4. If a theoretical water content of 11.4 wt.% i s added to the oxide analyses the totals a l l l i e within 1% of 100%. Table 5 Chemical Analyses of Chlorite Spec. Nos. 1 2 3 4 s i o 2 26.7 26.0 26.8 26.2 T i 0 2 0.09 0.09 0.09 0.09 A1 2 U 3 21.8 21.8 21.4 22.1 FeO* 26.6 27.2 27.1 26.3 MR0 13.0 12.2 12.4 12.9 MnO 0.18 0.22 0.26 0.22 CaO 0,21 0.19 0.37 0.17 Na20 0.01 0.01 0.01 0.02 K20 0,08 0.09 0,08 0.06 Total 88.67 87.80 88,51 88,06 +11.5% H20 100.17 99.30 100.01 99,56 formulae on the basis of 28 oxypen; Si 5.57 5.71 5.63 5.52 AlIV 2.42 2.29 2.37 2,48 A 1 V 1 2,93 3.35 2.92 3,00 Ti 0,01 0.01 0.01 0.01 Fe 4.62 3.99 4.73 4.71 MS 4.07 4.02 3.90 4.06 Mn 0.04 0.04 0.05 0.04 Ca 0.05 0.04 0,09 0.04 Na 0.00 0.00 0,00 0,00 K 0.02 0.02 0.02 0.02 11.74 11.47 11.72 11.88 FeO* - total iron as FeO. analyst - C,J .N.Fletcher. k6 Structural formulae have been calculated on a basis of 28 oxygens, with the tetrahedral s i t e s being completely f i l l e d by S i 1 V and A 1 1 V . The remainder of A l was assumed to occupy the octahedral s i t e s . If these chl o r i t e s are assumed to be 'unoxidized c h l o r i t e s ' they l i e within the f i e l d s of r i p i d o l i t e and brunsvigite composition according to the nomenclature of Hey (1954). The ch l o r i t e s have a highly r e s t r i c t e d composition, and show no systematic variations with grade. The compos-it i o n s are very similar to those reported from rocks of c h l o r i t e , b i o t i t e , and garnet zone from Scotland (Atherton 1968; Mather 1970). The analyzed chlorites were from the matrix and not from fracture planes. The chlorites so analyzed were, therefore, assumed to have been formed during prograde metamorphism. Garnet A l l the garnets studied were traversed.for S i , A l , Fe, Mn, Mg and Ca and representative plots of these traverses, together with the drawings of the traversed garnets, are given i n Figure 17. Both Si and A l were found to be r e l a t i v e l y constant within each specimen and the traverses of these elements are not given. The analysis of T i was performed, but the standards used were garnets of low T i content and thus give a low degree of accuracy. The results suggest, however, that the garnets from garnet and kyanite/staurolite zones have a T i 0 2 content of approximately 0.07%, while those from higher grades have approximately 0.03% Ti02- From the traverses the maximum compositional range for each specimen has been taken and tabulated, together with the calculated formulae and molecular percent end-members, i n Table 6. The s t r u c t u r a l formulae were calculated on the basis of 24 oxygens. The tetrahedral s i t e occupancy was assumed to be s i x , which has resulted i n some of the ^ 7 Figure 17. Microprohe traverses of selected garnets In a l l the traverses FeO i s taken to be t o t a l i r o n , also note that, except for FeO, the scales for the oxide percentages are the same. The scale to the right of each drawing of the traversed garnets repre-sents one millimeter i n a l l cases. The traverses of the garnets are i d e n t i f i e d by c a p i t a l l e t t e r s (A, B, C, D, E, F), whereas minerals are i d e n t i f i e d by small l e t t e r s ( s - s t a u r o l i t e , k-kyanite, s i - s i l l i m a n i t e , m-pseudomorhpic muscovite, q-quartz). Each specimen i s b r i e f l y des-cribed below: Spec. 4. Garnet zone, syn-F2, no second generation growth. Spec. 5. Kyanite/staurolite zone, syn-F2, l i t t l e or no second generation growth. Spec. 6. Zone 1 s i l l i m a n i t e zone, f i r s t generation garnet surrounded by large pseudomorphic muscovites. ^ 8 Figure 17 (cont). Spec. 7. Zone 1 s i l l i m a n i t e grade, thick second generation rim around f i r s t generation garnet core (traverse A-B), also t o t a l l y second generation garnet (traverse C-D). Spec. 8. S i l l i m a n i t e Zone 1, small second generation euhedral garnets surrounded by s t a u r o l i t e and pseudomorphic muscovite. Spec. 9. S i l l i m a n i t e Zone 2, large unzoned f i r s t generation garnet with thi n second generation rim, a l t e r a t i o n to s i l l i m a n i t e present. Spec. 10. Si l l i m a n i t e Zone 2, thin second generation rim on f i r s t generation garnet. k9 Figure 17 (cont) Spec. 11. S i l l i m a n i t e Zone 2, f i r s t generation garnet with a planar in t e r n a l s c h i s t o s i t y rimmed by a thick second generation growth. Spec. 12. S i l l i m a n i t e Zone 2, large f i r s t generation garnet with a thin second generation rim (traverse E-F). Also remnants of an altered garnet i n a s i l l i m a n i t e knot (traverses A-B, C-D). Spec. 13. S i l l i m a n i t e Zone 2, large second generation garnet with possible f i r s t generation core. Spec. 14. S i l l i m a n i t e Zone 3, small i d i o b l a s t i c second generation garnets. Spec. 15. S i l l i m a n i t e Zone 3, small i d i o b l a s t i c second generation garnets. Table 6 Chemical Analyses of Garnet Kos. A 4 4 5 5 6 6 7 7 8 8 9 9 SiO 38.5 38.8 38.8 37.7 37.7 37.2 37.2 36.7 36.7 36.2 36.2 36.7 36.7 A I 2 6 3 21.4 21.0 21.0 21.2 21.2 20.6 20.6 21.0 21.0 20.7 20.7 20.8 20.8 FeO* 23.5 32.5 24.5 34.6 29.2 34.8 33.5 35.9 25.5 35.4 34.5 34.6 34.0 MnO 8.5 1.2 7.0 0.3 2.9 1.9 3.0 0.3 7.5 1.0 3.0 1.9 2.6 MgO 1.2 2.5 1.5 3.8 1.4 3.2 3.2 4.7 1.3 3.7 3.7 3.2 3.2 CaO 8.5 5.1 6.0 3.1 8.0 2.7 1.7 1.3 8.4 2.6 1.5 2.1 1.7 Total 101.6 101.1 98.8 100.7 100.4 100.4 99.2 99.9 100.4 99.6 99.6 99.0 99.5 formulae on the basis of 24 oxygens SI 6.06 6.12 6.22 5.99 6.02 5.98 6.03 5.90 5.91 5.88 5.69 5.96 5.98 A 1 I V 0.00 0.00 0.00 0.01 0.00 0.02 0.00 0.10 0.09 0.12 0.11 0.04 0.02 6.06 6.12 6.22 6.00 6.02 6.00 6.03 6.00 6.00 6.00 6.00 6.00 6.00 A 1 V I 3.97 3.90 3.96 3.96 3.99 3.89 3.94 3.88 3.88 3.84 3.86 3.94 3.97 F e 3 + 0.03 0.10 0.04 0.04 0.01 0.11 0.06 0.12 0.12 0.16 0.14 0.06 0.03 F e 2 + 3.09 4.29 3.28 4.60 3.90 4.68 4.54 4.83 3.43 4.81 4.69 4.70 4.63 Mn 1.13 0,16 0.95 0.04 0.39 0.26 0.41 0.04 1.02 0.14 0.41 0.26 0.36 Mg 0.28 0.59 0.36 0.90 0.33 0.77 0.77 1.13 0.31 0.90 0.90 0.77 0.78 Ca 1.43 0.86 1.03 0.53 1.37 0.47 0.30 0.22 1.45 0.45 0.26 0.37 0.30 5.93 5.90 5.62 6.07 5.99 6.18 6.02 6.22 6.21 6.30 6.26 6.10 6.07 molecular percent end-members Almandine 52.1 72.7 58.4 75.8 65.1 75.8 75.4 77.6 55.2 76.4 74.9 77.0 76.4 Andradite 0.8 2.4 0.9 1.0 0.3 2.7 1.6 2.8 2.7 3.7 3.4 1.4 0.8 Grossular 23.3 12.2 17.4 7.7 22.5 4.9 3.4 0.8 20.6 3.5 0.8 4.6 4.1 Pyrope 4.7 10.0 6.4 14.8 5.6 12.4 12.8 18.1 5.0 14.2 14.3 12.7 12.8 Spessartine 19.1 2.7 16.9 0.7 6.5 4.2 6.8 0.7 16.5 2.2 6.6 4.3 5.9 Feo* - t o t a l iron as FeO. A - b i o t i t e grade garnet (see assemblage map for location) Fe3+ - calculated by making the t o t a l octahedral site occupancy equal to 4.00. The two analyses for each specimen represent the maximum compositional range of garnet within that specimen, the analyses being taken from Figure 17. analyst - C.J.N.Fletcher. O Table 6 (cont) Chemical Analyses of Garnet , Nos. 10 10 11 11 12 12 13 13 14 14 15 15 S1O2 36 .1 36.1 36 .2 36.2 37.8 37.8 37.1 37.1 37.2 37.2 36 .5 36 .5 AI2O3 20.8 20.8 20.6 20.6 20.7 20.7 21.0 21.0 20.8 20.8 20.9 20.9 Feo* 35.5 33.6 35.1 32.5 35.5 34.8 33.6 31.7 35.1 33.9 36.9 34.8 MnO 1.0 2.8 0.6 3.6 2.9 2.2 2.5 1.7 2.1 2.9 1.2 2.7 MgO 2.8 3.4 3.3 3.8 3.0 3.8 4.0 4.6 3.3 3.7 3.1 3.5 CaO 3.3 2.0 3.2 1.8 1.2 1.4 1.1 1.9 •2.3 1.5 1.0 0.8 To t a l 98.7 99.0 98 .5 101.1 101.1 100.7 99.3 98 .0 100.8 100.0 99.6 99.2 formulae on the basis of 24 oxygens Si 5.89 5.91 5.91 5.92 6.04 6.03 5.99 6 .01 5.96 5.99 5.94 5.94 A l " 0.11 0.09 0.09 0.08 0.00 0.00 0.01 0.00 0.04 0.01 0.06 0.06 6 .00 6 .00 6 .00 6 .00 6.04 6.03 6 .00 6 .01 6 .00 6 .00 6 .00 6.00 A 1 V I 3.88 3.92 3.87 3.90 • 3.90 3.89 3.98 4.01 3.89 3.93 3.95 3.95 F e 3 + 0.12 0.08 0.13 0.10 0.10 0.11 0.02 0.00 0.11 0.07 0.05 0.05 F E 2 + 4.84 4.60 4.79 4.45 4.75 4.65 4.54 4.30 4.71 4.56 5.02 4.74 Mn 0.14 0.39 0.08 0.50 0.39 0.30 0.34 0.23 0.29 0.40 0.17 0.37 Mg 0.68 0.83 0.80 0.93 0.71 0.90 0.96 1.11 0.79 0.89 0.75 0.85 Ca 0.58 0.35 0.56 0.32 0.21 0.24 0.19 0.33 0.40 0.26 0.17 0.14 6.24 6.17 6.23 6 .20 6.06 6.09 6.03 5.97 6.19 6.11 6 .11 6 .10 molecular percent end-members Almandine 77.6 74.6 76.8 71.9 78.3 76.3 75.2 72.0 76.2 74.8 82.2 77.7 Andradite 2.9 2.0 3.1 2.5 2.5 2.6 0.4 0.0 2.6 1.6 1.3 1.2 Grossular 6.4 3.7 5.9 2.6 0.9 1.3 2.7 5.5 3.8 2.6 1.5 1.1 Pyrope 10.9 13.4 12.9 15.0 11.8 14.9 16.0 18.6 12.8 14.5 12.3 13.9 Spessartlne 2.2 6.3 1.3 8.0 6.5 4.9 5.7 3.9 4.6 6.5 2.7 6.1 52 garnets having these sites completely f i l l e d or overfilled with Si, and others with the tetrahedral site occupancy being made up by the 3+ addition of Al . Fe has been calculated by assuming a theoretical octahedral site occupancy of 4.00. The compositional variations of the garnets, both within a particular crystal and also among specimens are complex, but the nature of the substitution between the divalent cations appears to be twofold: Mn substitutes for Fe, and Ca substitutes for both Mg and Fe. This may be deduced from the specimens showing no Ca or Mg zoning, but a reciprocal variation in FeO and MnO (specimen 15), and also from those specimens showing marked CaO zoning where the antipathetic variation in MgO combined with an increase in the FeO zoning balances the substitution (specimen 7). Zoning in garnet has been attributed to Rayleigh fractionation (Hollister 1966) and segregation during freezing (Atherton 1968). Two assumptions are made in both these models: f i r s t that the element considered has relatively low concentration in both the garnet and the surrounding minerals, and second that there is no zoning of the element in the co-existing minerals. Mn appears to satisfy both conditions and the observed zonal patterns compare well with the calculated curves. However, Ca, Fe, and Mg zonal patterns are not explained by either model, due to the abundance of each element in the co-existing minerals and effect of Ca zoning in plagioclase. Some of the f i r s t generation garnets show no zoning (e.g. specimen 9), this could be due to the lack of MnO in the surrounding matrix, as suggested by Hollister for unzoned garnets. However, there appears to be no appreciable difference in the MnO contents of biotite, muscovite or staurolite of these specimens and those containing zoned 53 garnets. The homogeneous composition could be explained by the zoning being obliterated by d i f f u s i o n within the garnet. That d i f f u s i o n occurs i n the garnets i s shown by specimen 7 where the boundary between the f i r s t and second generation garnets has been smeared out by d i f f u s i o n (Fig. 15; also see Appendix IV on d i f f u s i o n i n garnet). As has been shown i n the section on petrography there were two de f i n i t e periods of garnet growth; these periods may be i d e n t i f i e d chemically i n some specimens, but i n others they are i n d i s t i n c t due to the sim i l a r composition of the f i r s t and second generation garnets. Staurolite Eight staurolites have been analyzed from specimens within the s i l l i m a n i t e zone (Table 7). In these analyses the t o t a l iron content has been reported as FeO following the data from Mossbauer studies which suggest that the iron i n s t a u r o l i t e i s a l l i n the ferrous state (Bancroft et a l . 1967; Smith 1968). The water content has been estimated by difference so that the oxide t o t a l was 100%. The str u c t u r a l formulae have been calculated on a basis of 48(0,OH), which conforms with the 2+ s t a u r o l i t e formula ((0H)„Fe. (0,0H) AT o S i o 0 . , ) suggested by Deer et a l . A H 2. l o o 4 4 (1962 p.153). The staurolites were a l l traversed i n several directions to test for r a d i a l or sectoral zoning. A l l were found to be homogeneous. The already well-documented increase i n ZnO with grade i s also observed within the Penfold Creek staurolites (Fig. 19). This increase i n ZnO i s balanced by a decrease i n MgO (Fig. 20) on nearly a 1 : 1 basis, with 2+ 2+ l i t t l e v a r i a t i o n i n either Fe or Mn . i s highly variable but does not r e f l e c t the ZnO or MgO trends, suggesting that (Zn, Mg) occupy a different s i t e to T i i n the s t a u r o l i t e structure. 54 Table 7 Chemical Analyses of Staurolite Nos. 6 7 8 9 10 11 12 15 S i0 2 27.6 28.0 27 A 27.4 27.4 28.7 28.0 27.5 Ti02 0.61 0.83 0.37 0.49 0.41 0.68 0.84 0.26 A I 2 O 3 53.5 54.3 54.1 53.8 53.6 53.3 53.6 54.2 FeO* 13.8 13.6 14.2 13.8 14.0 13.2 14.1 13.5 MgO 1.83 1.83 1.94 1.98 1.64 1.59 1.58 1.30 MnO 0,15 0.10 0.07 0.06 0.09 0.05 0.17 0.09 ZnO 0.31 0.61 0.65 0,66 0,88 0.71 0.70 1.57 H20 2.20 0.73 1.27 1.81 1.98 1,77 1.01 1.58 formulae on the basis of 48(0,OH) Si 7.67 7.91 7.71 7.66 7.66 7.98 7.90 7.72 Al 17.53 18.08 17.95 17,73 17.67 17.48 17.82 17.66 Ti 0.13 0.18 0.08 0.10 0.09 0.14 0.18 0.06 17.66 18.26 18.03 17.83 17.76 17.62 18.00 17.72 Fe 3.21 3.09 3.34 3.23 3.27 3.07 3.32 3.17 Mg 0.76 0.77 0.81 0.82 0.68 0.66 0.66 0.54 Mn 0.04 0.02 0.02 0,01 0.02 0.01 0.04 0.02 Zn 0.06 0.13 0.14 0.14 0.18 0,15 0.16 0.33 4.07 4.01 4.31 4.20 4.15 3.89 4.18 4.06 OH 4.01 1.37 2.38 1.69 3.69 3.28 1.90 2.96 FeO* - total iron as FeO. H20 calculated to make the analyst - C,J .N.Fletcher, oxide total equal 100%. Figure 18, Average anorthite. content of plagioclase. Figure 19. ZnO content of staurolite.(wt.^) 56 Fifcure 2 0 . Atomic proportion of Mg against atomic, proportion of Zn i n s t a u r o l i t e . 0-9-1 0-8 CO c o O)07-2 theoretical, line for 7.n»M« substitution (gradient = -1) X 0-6H 0 5 - 1 — r 010 0 20 Zn ions 0-30 5 7 Plagioclase The presence of f i v e spectrometers on the EMX model of the ARL microprobe makes the plagioclase analyses highly r e l i a b l e . This i s substantiated by the closeness of the totals to 100%, and X and s i t e occupancy to 4.00 and 16.00 respectively on a basis of 32 oxygen atoms. At least s i x plagioclase grains from each specimen were traversed because of the high v a r i a b i l i t y of the plagioclase composition within in d i v i d u a l grains and among grains. In Table 8 the maximum compositional range among grains within each specimen i s given, together with the type and extent of the zoning. The zoning, where present, i s reverse i n character and may be i d e n t i f i e d o p t i c a l l y where the composition varies by four or more percent anorthite component. In the majority of grains, however, the compositional v a r i a t i o n i s inconsistent with a regular zoning pattern. The plagioclase compositions vary among grains from the same specimen, even when comparing the outermost rim. This could be due to the actual outermost rim not being analyzed, owing to the r e l a t i v e l y large diameter of the electron beam required for the analysis of sodium. However, the va r i a t i o n can be as great as 10% anorthite component among grains, suggesting equilibrium has not been attained. Compositional differences occur i n the range An greater than 15% and thus can not be due to the influence of the p e r i s t e r i t e solvus. The anorthite content of plagioclase increases from almost pure a l b i t e i n the ch l o r i t e zone to oligoclase/andesine i n the kyanite/ s t a u r o l i t e zone. In s i l l i m a n i t e Zones 1 and 2 the anorthite content i s f a i r l y constant, but decreases i n Zone 3. The increases between b i o t i t e Table 8 Chemical Analyses of Plagioclase Spec. Nos. 1 1 3 3 4 4 5 5 6 6 7 7 7b 7b 8 ti S i 0 2 68.9 68.6 68.5 67.1 63.1 62.3 63.2 60.0 64.5 60.2 62.2 60.3 61.6 59.8 63.1 61.0 A 1 2 0 3 19.8 20.1 20.1 20.3 22.8 22.9 23.0 25.3 22.7 25.6 24.2 25.5 24.1 25.5 22.7 24.0 CaO 0.12 0.20 0.16 0.48 3.50 3.97 4.76 7.65 2.98 6.16 4.64 6.37 5.22 7.10 3.89 5.53 Ka 20 11.7 11.6 12.0 11.8 9.64 9.48 9.12 7.49 10.0 8.23 9.23 8.07 8.83 8.01 9.40 8.68 K 20 0.00 0.05 0.06 0.05 0.04 0.03 0.12 0.08 0.07 0.05 0.10 0.06 0.06 0.04 0.07 0.03 Total 100.52 100.55 100.82 99.73 99.08 98.68 100.20 100.52 100.31 100.24 100.37 100.30 99.81 100.45 99.16 99.24 Z 16.02 16.03 16.09 16.00 16.03 16.00 15.96 15.95 16.04 16.05 16.04 16.03 16.00 15.97 16.00 15.97 X 3.95 3.94 4.09 4.11 4.01 4.07 4.06 4.06 4.00 4.01 4.06 4.00 4.04 4.12 3.99 4.08 molecular percent end-members A l b i t e 99.4 98.8 99.0 97.5 83.1 81,1 77.1 63.6 85.8 70.5 77.8 69.4 75.1 67.0 81.1 73.8 Anorthite 0.6 0.9 0.7 2.2 16.7 18.8 22.2 35.9 13.8 29.2 21.6 30.3 24.5 32.8 18.5 26.0 Orthoclase 0.0 0.3 0.3 0.3 0.2 0.1 0.7 0.5 0.4 0.3 0.6 0,3 0.4 0.2 0.4 0.2 Zoning none ? ? rev (11%) rev (4%) ? rev (3%) ? Nos. 9 9 10 10 11 11 12 12 13 13 14 14 15 15 15b 15b S102 62.7 60.4 61.4 59.6 61.4 60.9 64.9 62.7 63.3 61.8 62.3 61.0 66.0 66.0 64.8 64.8 Al 2 0 3 23.4 25.1 24.5 25.1 23.9 24.3 22.4 23.9 22.8 23.4 23.9 25.2 21.7 22.5 22.2 22.2 CaO 4.69 6.09 5.44 6.62 5.36 6.16 2.83 4.38 3.46 4.30 4.78 5.88 2.11 2.78 2.68 3.01 Na 20 9.16 8.23 8.50 8.07 8.77 8.31 10.4 9.49 9.56 9.12 8.94 8.20 10.7 10.3 10.3 10.1 K 20 0.02 0.04 0.02 0.06 0.01 0.05 0.09 0.05 0.06 0.09 0.07 0.08 0.09 0.06 0.06 0.03 T o t a l 99.97 99.86 99.86 99.45 99.44 99.72 100.62 100.52 99.18 98.71 99.99 100.36 100.60 101.64 100.04 100.14 Z 15.99 16.02 16.03 15.99 15.98 15.97 16.00 16.01 16.04 16.02 16.02 16.05 16.00 16.02 16.01 16.01 X 4.04 4.01 3.97 4.08 4.06 4.06 4.08 4.07 3.97 4.02 4.00 3.94 4.04 3.99 4.03 4.02 A l b i t e Anorthite Orthoclase Zoning molecular percent end-members 77.9 22.0 0.1 70.8 29.0 0.2 73.8 68.6 26.1 31.1 0.1 0.3 74,7 25.2 0.1 70.7 29.0 0.3 85.5 13.0 0,5 79.4 20.3 0.3 83.0 16.6 0.4 78 20 0, 76.9 71.3 22.7 28.2 0.4 0.5 89.7 9.8 0.5 86.8 12.9 0.3 rev (7%) rev (4%) rev (2%) 87.1 85.7 12.6 14.1 0.3 0.2 none rev - reverse zoning. 7b, 15b - duplicate specimens, analyst - C.J.N.Fletcher. Note: the two analyses given for each specimen show the maximum compositional v a r i a t i o n found between d i f f e r e n t grains, the maximum extent of zoning found within a pa r t i c u l a r grain for each specimen i s given i n anorthite percent, the molecular percent end-members were calculated from the r a t i o s of the cations Na, Ca, and K r e s p e c t i v e l y . 00 59 and garnet zones, and garnet and kyanite/staurolite zones appear to be sharp. The composition of plagioclase from s i l l i m a n i t e Zone 3 l i e s within the p e r i s t e r i t e range, which implies that the top of the p e r i s -t e r i t e solvus i s below the conditions of Zone 3. This i s i n agreement with the results of Crawford (1966). Ilmenite The composition of ilmenite varies among grains i n a l l of the samples (Table 9). This v a r i a t i o n i s most marked i n MnO, where a threefold increase can occur among different grains within the same specimen. Each grain i s , however, homogeneous. The ranges i n composition are tabulated i n Table 9 for selected specimens. These analyses are con-sidered to have a low degree of r e l i a b i l i t y due to the few grains of ilmenite within each polished section. Numerous sections from each specimen would have to be prepared to ascertain the f u l l extent and nature of the variations. No systematic v a r i a t i o n i n composition with increase of metamorphic grade could be detected. This i s i n accord with the results of K.V. Campbell (1971), who analyzed ilmenite from varying grades within rocks of the Kaza Group to the south of Penfold Creek. No ch a r a c t e r i s t i c compositional difference was found between the ilmenites associated with the pseudomorphic muscovites and those i n the schistose matrix. 3. Attainment of Equilibrium In polymetamorphic ter r a i n s , such as the Penfold Creek area, i t i s important to ascertain whether a par t i c u l a r mineral assemblage attained equilibrium during the f i n a l period of metamorphism, or whether there are two or more mineral generations present which are i n d i s -equilibrium. Petrographic evidence for two or more generations includes; 60 Table 9 Chemical Analyses of Ilmenite Spec. Nos. T i 0 2 FeO* MgO MnO 1 50.6 -52.1 45.5 -45.7 0.02- 0.03 0.38- 0.79 3 51.7 -52.2 45.9 -46.3 0.09- 0.10 1.29- 1.51 5 50.6 -50.8 47.5 -47.9 0.04- 0.05 0.28- 0.42 7 49.4 -49.8 46.4 -46.7 0.06- 0.08 0.47- 1.40 Spec. Nos. 12 15 T i 0 2 FeO* MgO MnO 51.6 -52.3 44.0 -44.3 0.08- 0.08 0.41- 1,29 49.4 -51.1 45.5 -46,8 0.12- 0.31 0.38- 0.62 51.4 -51.7 46.4 -47.3 0.08- 0.08 0.19- 0.44 FeO*- total iron as FeO. analyst - C.J .N.Fletcher. Note: analyses show the maximum range of composition for each oxide from several grains within each specimen, no totals or structural formulae are given due to the extreme variation of composition among grains within the same specimen. 6i early generation minerals being replaced ( i .e . pseudomorphic muscovites after garnet and staurol i te) , rimmed (I.e. garnet), or cross-cut ( i . e . muscovite), and the presence of two size fractions within the same rock ( i .e . garnet). Most of these textures have been observed in the high grade schists from Penfold Creek, and have been described in the section of petrography. To determine whether the two generations are in equilibrium or disequilibrium requires a comparison of the mineral chemistries. This is the only way disequilibrium can be proved (Zen 1963). The scarcity of analyzed specimens from the low grade schists makes a chemical appraisal of the equilibrium conditions uncertain. The chemical variations of the minerals between zones do however conform with the proposed isograd reactions, (see section on reactions). Also similar chemical variations have been reported from other metamorphic terrains. The low grade schists , up to the kyanite/staurolite isograd, show no petrographic evidence of two periods of prograde metamorphism. It thus appears that the low grade schists can be considered to have reached chemical equilibrium. The compositions of muscovite and plagioclase from rocks of s i l l i -manite zone indicate that these rocks did not attain complete equ i l -ibrium. The muscovites pseudomorphic after garnet and staurol ite have a chemical composition consistently different from the muscovites of the matrix. This is well i l lustrated in Figure 14 where the four pseudo-morphic muscovites have a lower phengite component, they are also lower in TiC^ and generally have a higher paragonite content. Plagioclase is weakly zoned in many of the s i l l imanite schists , however there also appears to be a wide variation in composition among grains within the same specimen. These two factors strongly suggest disequilibrium but might be explained by having equilibrium domains smaller than the size of 62 a polished section. Although garnet i s strongly zoned, rimmed and i s i n places present as two size fractions, the outermost parts of a l l the garnets within the same specimen have ess e n t i a l l y the same composition. Thus they appear to have been i n equilibrium with the surrounding minerals when they ceased to c r y s t a l l i z e . The matrix micas within a pa r t i c u l a r section are r e l a t i v e l y constant i n composition. In both b i o t i t e and muscovite the standard deviations for (S i , A l ) , (K, Fe), Mg, and Na average at 1.2%, 1.7%, 2.6% and 4.0% respectively. The increase i n the standard deviations for Mg and espec-i a l l y Na i s probably due to the a n a l y t i c a l technique. A l l the elements are less homogeneous than those described by Evans and Guidotti (1966). The schists from the Penfold Creek area may therefore be considered to have been i n equilibrium In the lower grades of metamorphism, and i n disequilibrium i n the higher grades. Reactions occurring i n the p e l i t i c rocks Reactions which formed the index minerals of the p e l i t i c rocks of Penfold Creek have been deduced from petrographic and chemical evidence. Each isograd i s discussed separately but emphasis i s placed on s i l l i m a n -ite-forming reactions. Low grade assemblages receive less attention due to limited data.,. I t w i l l be shown that minor bulk compositional variations can determine whether a part i c u l a r index mineral could have formed from a part i c u l a r specimen. This sensitive compositional control makes modes of matrix mineralogy unreliable i n postulating or v e r i f y i n g possible reactions. 63 Linear Regression Technique The linear regression technique* (Greenwood 1968) has also been used to determine possible reactions between the specimens from d i f -ferent metamorphic zones. This technique uses a least-squares method to find any l inear dependence that can be written among the minerals found in one or more assemblages. Consider two equilibrium assemblages A and B, consisting of the minerals A^, , A^ and B^, B^, B^  respectively. Modelling one of these minerals, say A^, against a l l the others results in an equation of the form: (1) A± = c 2 A 2 + c 3 A 3 + c ^ + c 5 B 2 + c 6 B 3 where c is a coeff icient which may be either negative or posit ive. If the signs of the coeff icients are such that eq.( l ) can be rewritten: (2) A1 + c 2 A 2 + c 3 A 3 = c ^ + c 5 B 2 + c B 3 then for both assemblages A and B to be equilibrium assemblages they must have formed under different physical conditions, as only assemblage A or assemblage B can be stable under any specified set of conditions. If on the other hand the coefficient signs are not consistent with eq.(2), then one could infer that no amount of change in the physical conditions could change assemblage A into assemblage B, and therefore different bulk chemistries must be responsible in part. Error Propagation The significance of the equation calculated using the linear regression technique may be estimated by comparing the residuals to the error accumulated in the calculation. The combined standard error (e^) of each component of the modelled mineral calculated from the regression * the regression equations were calculated using the University of Br i t ish Columbia l ibrary program - stepwise regression BMD:02R. 6k equation: i s A l = C 2 A 2 + °3 A 3 + / , 2 2 , 2 2 , . e l " ^ ^ 2 + C 3 e 3 + } where e^, e^ are the standard errors of the components in minerals A^, A^, . . . . respectively (Braddick 1954). If the combined standard error in the calculated modelled mineral is greater than the residual then the equation may be considered s igni f icant . However, large mineral coefficients (c) and large numbers of independent variables could increase the errors to such an extent to make the equation appear s ign-f icant. Obviously, the best equation is the one which gives the least residual to the calculated modelled mineral. The calculated reactions discussed in the following paragraphs are given in Table 10 together with residuals and standard errors of the calculated mineral. In a l l reactions the number of independent variables was kept to a minimum. Thus the effect of plagioclase sol id solution has not been considered. However, an indication of the variation in plagioclase composition to be expected from the proposed reactions can be gained from the Ca and Na residual of the modelled mineral. Reactions were calculated both with and without ilmenite. The ones which gave the lowest residuals are those given in Table 10. The mineral compositions entered in the regression equations were the cation proportions as given in Tables 3, 4, 5, 6 and 9. Water was not included. Biotite Isograd Biot i te f i r s t appears within chlorite knots which formed just below the b iot i te isograd. As no potash feldspar has been identif ied in the metasediments of the Kaza Group in the Penfold Creek area the two possible biotite-forming reactions are: Table 10. Regression Equations for the Mineral Assemblages from Penfold Creek Area Assemblages g Regression Equations Si T i A l Fe Mg Mn Ca Na 1 and 2 1 1.00BIOT(2) + 16.25MUSCU) - 0.65CHL0RU) + 15.73MUSC(2) + 6.83QTZ RES ERR 0.00 1.70 0.21 0.02 -0.12 1.43 -0.26 0.10 0.45 0.19 -0.01 0.00 -0.03 0.00 0.12 0.14 0.33 0.64 1 and 3 2 1.00BIOTO) + 14.46MUSC(1) = 0.48CHLORU) + 14.35MUSC(3) + 4.80QTZ + 0.10ILM RES 0.00 ^0.01 -0.01 0.00 -0.01 -0.02 -0.02 0.19 0.01 ERR 1.54 0.02 1.64 0.09 0.18 0.01 0.00 0.12 0.58 1 and 4 3 1.00GNT(4) + 23.31HUSCU) + 1.48BI0T(4) = 1.45CHLOR(l) + 23.07MUSC(4) + RES 0.00 -0.01 0.00 0.00 -0.02 0.11 0.79 -0.34 0.02 9.71QTZ + 0.12ILM ERR 2.48 0.02 2.07 0.20 0.29 0.00 0.00 0.21 0.93 2 and 4 4 1.00GNT(4) + 34.81MuSC(2) + 0.22BIOT(4) + 0.65QTZ = 0.61CHLOR(2) + 35.41M)SC(4) + 0.65ILM RES 0.00 -0.13 0.03 0.12 -0.17 0.13 0.83 -0.79 -0.01 ERR 3.71 0.06 3.13 0.24 0.32 0^00 0.00 0.32 1.45 3 and 4 5 1.00CNT(4) + 28.73MUSC(3) + 29.15BI0T(4) + 0.52ILM - 1.70CHLOR(3) + 28.30MUSC(4) + 28.45BI0T(3) + 1.87QTZ RES 0.00 0.03 -0.01 -0.03 0.03 0.11 0.71 0.09 0.02 ERR 4.09 0.13 3.03 1.82 2.09 0.01 0.01 0.26 1.56 4 and 5 6 1.00STAUR(6) + 4.79BIOT(5) + 3.75MUSC(5) + 6.02QTZ - 2.08CHX(4) + 1.07BIOT(4) + 7.18MUSC(4) + 0.49ILM RES 0.00 0.00 0.00 0.00 0.00 -0.08 0.08 0.00 0.00 ERR 0.69 0.04 0.57 0.26 0.36 0.00 0.00 0.06 0.27 1.00CARN(6) + 14.81MUSC(6) + 6.3SBI0T(6) + 0.48ILM = 3.25SILL + 13.16PMUSC(6) + 7.95BI0T(14) + 6.06QTZ RES 0.00 0.00 ERR 1.62 0.06 0.00 -0.01 1.38 0.41 0.00 0.58 0.27 0.00 0.39 -0.01 0.00 0.00 0.39 0.59 Note: 1. 2. 3. 4. 5. 1.00STAUR(6) + 10.22MUSC(6) + 3.69BI0T(6) + 0.25ILM 9.07PMUSC(6) + 4.80BI0T(14) + 0.03QTZ 9.65SILL + RES ERR 0.00 1.09 0.00 0.03 0.00 0.94 0.00 0.24 0.00 0.34 0.00 0.00 0.00 0.00 Assemblage number refers to specimen number i n which, the assemblage ocours. Equations have been computed using the molecular proportion of the cations f o r the minerals, water i s not included. matrix muscoviti; PMUSC —.pseudomorphic muscovite; STAUR - s t a u r o l i t e ; BI0T - b i o t i t e ; CHL0R - c h l o r i t e ; GNT SILL - s i l l i m a n i t e ; QTZ - quartz; ILM garnet; MUSC ilmenite. The number beside each mineral refers to the assemblage i n which i t occurs. RES - r e s i d u a l « actual mineral composition of modelled mineral- computed composition, according to regression equation. ERR - Standard error of computed model. The f i r s t mineral i n each equation i s the modelled mineral. 0.00 0.27 0.00 0.39 • 66 K-r ich muscovite + chlor i te + quartz = b iot i te + less K-rich muscovite + V (Ti l ley 1926) and phengite + chlor ite = muscovite + b iot i te + quartz + H^ O (Mather 1970) The regression reactions (Table 10) between assemblages 1 and 2, and 1 and 3*, using b iot i te as the modelled mineral, suggest that the f i r s t reaction is the most reasonable. Very low residuals have resulted where the phengite components of the muscovites are nearly ident ical (assemblages 1 and 3), whereas the low phengite muscovite in assemblage 2 has resulted in large residuals. In both cases, however, the muscovites are on the wrong side of the equation due to the muscovites of the b iot i te zone being more potassic. This indicates that neither assemblage 2 nor assemblage 3 could have formed from the minerals of assemblage 1, and that there must be a signif icant difference in their bulk chemistries. A large compositional difference between specimens 1 and 2 can be inferred from the regression reactions, because the standard error of the modelled b iot i te i s smaller than the residual for several of the components. Garnet Isograd The c lass ica l garnet-forming reaction: (Fe, Mg) chlorite + Quartz = Almandine + Mg chlorite could not have occurred in the Penfold Creek schists , where the chlorite composition remains relat ively constant throughout the ch lor i te , b iot i te and garnet zones. If one assumes that garnet formed from chlor ite (and there is no petrographic evidence to support this) then the resulting . * the assemblages refer to the minerals in the regression equations of Table 10, and the numbers to the specimens of Table 2. 67 reaction must be of the form: 2Chlorite + 2Phengite = 2Garnet + 2Biotite + Muscovite + 2Quartz + Three reactions have been computed between the assemblages 1 and 4, 2 and 4, and 3 and 4. The standard error in a l l these cases is high for most of the components, so the reactions must be treated with some scepticism. Low residuals are observed where there is a difference in the phengite component of the muscovites ( i .e . assemblages 1 and 4, 3 and 4), and high residuals are observed where the phengite components are pract ical ly identical ( i .e . assemblage 2 and 4). This would indicate that phengite is involved in the garnet-forming reaction. The computed reactions also show that the amount of b iot i te increases with the formation of garnet. Both these factors would support the reaction proposed above. The muscovites appear on the wrong side of the reactions due to bulk compositional differences between the specimens. This means that assemblage 4 could not have formed from assemblage 1, 2, or 3. A l l the computed reactions have large positive residuals in Ca and Mn, and in two of the reactions large negative residuals in Na. This would suggest that plagioclase should become strongly a lb i t i c in the garnet zone. That the opposite occurs can only be attributed to the presence of ca lc i te in the chlorite and b iot i te zones. The high Mn r e s i -dual indicates that the Mn in garnet cannot be explained by the reactions. It is possible that ilmenite was the source of Mn, which might account for the extreme var iab i l i t y in the composition of ilmenite between grains in the same specimen. The ilmenite used in the regression equations is only an average of a l l the ilmenites in the Penfold Creek area (Ti - 2.08, Fe - 2.04, Mg - 0.01, Mn - 0.05 on a basis of six oxygens). 68 Staurolite/Kyanite Isograd No reaction relationships have been observed for the formation of either kyanite or staurol i te , except that kyanite is commonly associated with the muscovite-rich portions of the rock. This suggests that the kyanite forming reaction i s : This reaction, however, would oppose the observed compositional var ia -tions of the matrix minerals across the isograd. The absence of prograde chlorite from the staurolite-bearing schists strongly suggests that staurol ite formed from chlor i te . The formation of garnet and staurol ite do not appear to have been contemporaneous. Thus i t is probable that garnet was not involved in the staurolite-forming reaction. Therefore, the regression equation calculated between assembla-ges 4 and 5 contains chlor ite but not garnet. The computed reaction indicates that assemblage 5 could have formed from assemblage 4, and that the two assemblages must have developed under different physical conditions. The very low residuals and low errors make the regression equation extremely re l iab le . The computer reaction explains the compos-i t iona l variations in the matrix minerals. Expressed in terms of muscovite end-members, the reaction may be .written*: 12Chlorite + 17Phengite = 2Staurolite + 23Biotite + HMuscovite + Because the computed reaction does not involve plagioclase and there are very low residuals for Ca and Na, the marked increase in the anorthite component of plagioclase in the staurol ite zone cannot be * The stoichiometric formulae for the minerals in this equation are: Chlorite - (Fm) AISi A10 (OH) ; Phengite - K (Fm)Al Si A10 (OH) ; Staurolite - ( F m ) . A l , Q S i „ 0 . , ( O H ) _ ; Biot ite - K (Fm)„AIS i^0 ^ f u H ) „ ; 4 ±0 Q 40 Z J 3 i U Z Muscovite - K A l J U S i - O . , n ( 0 H ) „ . paragonite + quartz = albite + kyanite + HO 2' 3 10 explained by the staurolite-forming reaction. It is thought that the compositional change in plagioclase occurred during the f i r s t meta-morphic period when the maximum grade of metamorphism reached was the upper garnet zone. The anorthite content of plagioclase stable under these conditions was also stable under the conditions of the staurol ite zone of the second metamorphic period. The staurolite-forming reaction apparently had no effect on the plagioclase composition. Si l l imanite Isograds Three reactions are considered to have occurred in the s i l l imanite zone rocks: Kmuscovite garnet s i l l imanite b iot i te (1) KAl 2 Si 3 A10 1 ( ) (OH) 2 + (Fm) 3 Al 2 (S i0 4 ) 3 = 2Al2SiC>5 + K(Fm) 3Si 3A10 1 ( ) (0H>2 quartz + S i0 2 paragonite quartz albite s i l l imanite (2) 'NaAl oSi,A10, n(0H)_ + S i0 o = NaAlSi O 0_ + Al o Si0_ + H„0 Z J 1U z z o o z _> z (3) kyanite = s i l l imanite Reaction (3) indicates that a temperature gradient was present across the three zones of the s i l l imanite grade. Reactions (1) and (2) depend not only on this temperature gradient but also on the P n , P T and the composition of the muscovite (ss)"*". The relat ive total positions of reactions (1) and (2) in P -T space have been drawn in Figure 21 based on calculations presented in Appendix V. Both reactions generate a family of reaction curves dependent on the composition of the throughout the following discussion Kmuscovite refers to the pure end-member potassium muscovite, while muscovite (ss) refers to the sol id solution ser ies. 70 Figure 21. Effect of muscovite-paragonite- sol id solution on the si l l imanite-forming reactions: 1. Muscovite + almandine = 2si l l imanite + b iot i te + quartz 2. Paragonite + quartz = albi te + s i l l imanite + H^ O The positions of these two reactions in P u n - T space and the effect of sol id solution have been calculated from thermodynamic data (see Appendix V), however, only reaction (2) gave reasonable results , these have been plotted on the diagram (.) for P .. = 4000 bars. total Where the muscovite (ss) is the same for both reactions a combined reaction may be written: 3. Muscovite (ss) + almandine = s i l l imanite + b iot i te + albite + quartz. The position of this curve is determined by the points (o) of in ter -section of the curves for reactions (1) and (2). The paragonite content of the muscovite (ss) increases along this curve from X to B to C. An explanation of the changes of P n and muscovite composition with increasing temperature is given in the text. 7 1 muscovite (ss) . The sequence of the curves may be deduced from the equilibrium constants for each reaction: a. Biot ite 1 (for constant composition of qtz, s i l l , b iot , gnt) reaction (1) x. Kmusc' gnt Kmusc reaction (2) x (for constant composition of ab, s i l l , qtz) parag parag (K - equilibrium constant; a - act iv i ty ; f - fugacity; x - mole fraction) An increase in K for both reactions, resulting from an increase in the paragonite and Kmuscovite component respectively, w i l l favour the formation of muscovite (ss) (Appendix V). Curves have been drawn in Figure 21 to represent x = 1 . 0 , 0.9, 0.7, 0 .5 , 0.3 and 0.1 for muse Reactions (1) and (2) can be combined into one reaction where the composition of the muscovite (ss) is the same in both reactions (20% paragonite molecule in the example below): (4) 5Muscovite (ss) + 4garnet = 9Sill imanite + 4Biotite + albite + The thermodynamic calculations for reactions (1) and (2) are given in Appendix V. Lack of data for almandine resulted in a highly improbable position for reaction (1). Chemical and petrographic evidence suggests that reaction (4) takes place in Penfold Creek area. For this reason the position of reaction (1) ha,s been arb i t rar i l y moved to intersect reaction (2) as shown in F ig . 21. The reason for choosing a total pressure of 4Kb instead of a pressure above the alumino-si l icate t r ip le point (Richardson et a l . 1969) was because the maximum thickness of sediments above the Penfold Creek schists i s 40,000 feet which approximates to 4Kb. It is considered that the effects of order-disorder (Zen 1969) reaction (1), and x parag = 1.0, 0 .9 , 0.7, 0 .5 , 0.3 and' 0.1 for reaction (2). 3Quartz + H20 72 and grain size (Cameron and Ashworth 1 9 7 2 ) may be suff ic ient to lower the t r ip le point below 4Kb. Although no disorder has been found in f i b r o l i t e using infrared traces (Cameron and Ashworth 1 9 7 2 ) , they do not state the l imits of detection of the method employed, and so a small amount of d i s -order could have been present. As has been pointed out by Greenwood (pers. comm.) i t is the f i r s t 5% of disorder which would affect the equi l ibr ia the most. Curve XC (Fig. 2 1 ) at the intersection of the degenerate curves ( 1 ) and ( 2 ) represents reaction ( 4 ) . Not only do P Q and T vary along this curve but so does the muscovite composition. The curve for reaction ( 4 ) represents, therefore, a P n buffer but is not a buffer for the muscovite (ss) composition. In s i l l imanite Zone 1 i t is considered that P T T „ was less than P , . H2O total Because of the arbitrary placement of the curve for reaction ( 1 ) the pres-sure difference cannot be estimated, thus position A in Figure 2 1 has been chosen only to i l lus t ra te the changes with increasing temperature. As reaction ( 4 ) proceeded the muscovite (ss) became progressively enriched in the paragonite component, with an increase in s i l l imani te ,a lb i te and Fe-b iot i te components at the expense of garnet and Kmuscovite. When P ^ equals P , the f „ n remains constant and continued reaction requires loss M total 1 ^ 0 of R^ O from the system followed by the loss of garnet (this occurred at 7 C ' B in F ig . 2 1 ) . When garnet has been consumed a further increase in temp-erature w i l l decrease the paragonite content of muscovite (ss), and increase the albite component of plagioclase because reaction ( 2 ) is the only one operating (along CD in F ig . 2 1 ) . This condition is realized in s i l l imanite Zone 3 . The proposed model accounts for compositional changes which have been described for muscovite (ss), for the iron enrichment of b io t i te , 73 and for the increase i n the a l b i t e component of plagioclase. The two b i o t i t e s used i n the regression equation for the breakdown of garnet to pseudomorphic muscovite and s i l l i m a n i t e i n assemblage 6 were those from specimens 6 and 14 (Table 3). I t was considered that the b i o t i t e compositions tended towards that of b i o t i t e (14) as the grade increased through the s i l l i m a n i t e zones. Thus, b i o t i t e (14) represents the end-product of the reaction. This approach had to be used because the lack of equilibrium i n the sillimanite-bearing assemblages. The regression reaction for the breakdown of garnet has both low residuals and errors. It has the same form as reaction (4) discussed previously, except for the absence of a l b i t e i n the regression equation. However, the residuals of the regression equation would suggest that the anorthite component of plagioclase should increase, rather than decrease, as the reaction continued. The observed plagioclase compositions are r e l a -t i v e l y constant throughout Zone 1 and 2 indicating that the small amount of the grossular component i n the garnet does not r a d i c a l l y affect the plagioclase composition. The regression reaction for the breakdown of s t a u r o l i t e (Table 10, reaction 8) was derived i n the same way as the reaction for garnet. There are no residuals to the equation and the errors are low. The equation i s very similar to that of the garnet breakdown. Because of this i t i s suggested that the s t a u r o l i t e breakdown could be discussed from the standpoint of combining the reactions: (6) 3Staurolite + 4Kmuscovite + 7Quartz = 31Sillimanite + 4Biotite + 3H20 (2) Paragonite + Quartz = A l b i t e + S i l l i m a n i t e + H20 Lack of thermodynamic data prevent the calculation of reaction (6). However, because i t i s a dehydration reaction i t w i l l have a positive 7k slope on the P n - T p lot , and must have a different slope to the curve for reaction (2) to intersect i t . Staurolite is not completely consumed in s i l l imanite Zone 3. This may be due to the mantling effect of s i l l imani te . 4. The Petrology of the Ca lc -S i l i cate Bearing Assemblages The occurrence of the c a l c - s i l i c a t e bearing assemblages in the Penfold Creek area is total ly restr icted to the Isaac Formation, where they are associated with thin impure calcareous bands in the dominantly p e l i t i c suite of rocks. Metamorphosed limestones, containing a small percentage of quartz, also occur in the Isaac Formation. However, they are stable to the highest grades of metamorphism found in the Penfold Creek area. The majority of the c a l c - s i l i c a t e assemblages are present in s i l l imanite zone rocks and thus the i n i t i a l unmetamorphosed mineral assemblage can only be inferred. Two isograds have been mapped for the c a l c - s i l i c a t e assemblages based on the reactions: tremolite + 3calcite + 2quartz = 5diopside + 3C0^ + E^O - (1) 5phlogopite + 6calcite + 24quartz = 3tremolite + 5Kspar + 6CC>2 + 2H20 - (2) To define these isograds i t is necessary to plot the positions of the assemblages with co-existing reactants, products, and reactants and products. For instance, in reaction (1) the zone below the isograd is defined to be where the reactant assemblage tremolite + ca lc i te + quartz co-exist . Above the isograd the product assemblage diopside is present without a l l three of the reactant minerals. Those assemblages which contain a l l four minerals are defined as lying on the isograd. Theoretically, the latter assemblage should be highly restr icted in 75 d i s t r i b u t i o n , however, natural assemblages may depart s i g n i f i c a n t l y from the i d e a l (Thompson 1957). In the Penfold Creek area zones containing reactants and products are present for both reactions, the zone for reaction (1) being at least three-quarters of a mile wide, (Fig. 22). This could be due to the assemblage buffering the state of the system throughout the zone. The two isograds are approximately p a r a l l e l to the p e l i t i c isograds and are not considered to have any s i g n i f i c a n t cross-cutting relationships, because of their wide reaction zones and the scarcity of data points. The minerals of reaction (2) are contained within the system KA102-CaO-MgO-Si02-C02-H20. However, with the addition of other components phlogopite and tremolite w i l l become b i o t i t e and Ca-amphibole respectively. No d i s t i n c t i o n has been made between phlogopite and b i o t i t e , and tremolite and Ca-amphibole i n the positioning of the isograds i n Figure 22. Plagioclase (>An30), c l i n o z o i s i t e , scapolite, grossular garnet, and sphene are a l l stable phases under the conditions i n which the c a l c - s i l i c a t e s of reactions (1) and (2) are found, that i s from above the kyanite/staurolite isograd to above the Zone 3 s i l l i m a n i t e isograd. The opaque phases present are ilmenite, pyrite and pyrrhotite. Experimental work on the s t a b i l i t y of the c a l c - s i l i c a t e minerals has been extensive, and the reactions relevant to this study have been plotted i n Figure 23. The position of the scapolite s t a b i l i t y f i e l d i s only approximate. However, i t s r e l a t i v e position with respect to the other reactions i s correct (O r v i l l e 1970 lecture). The position of reaction (2) has to l i e i n the s t a b i l i t y f i e l d of tremolite + c a l c i t e and thus l i e s between the reactions: 76 Figure 22. Ca lc - s i l i ca te assemblages in the Penfold Creek area. The isograds are based on the reactions: Q tremolite + 3calcite + 2quartz = 5diopside + 3C02 + H20 | | 5phlogopite + 6calcite + 24quartz = 3tremolite + 5K-spar 6C02+-2H20 legend Q equation products (above reaction curve) equation reactants (below reaction curve) ^ equation reactants + products (on reaction curve) _L. scapolite present O grossular garnet present 7 7 Figure 23. T - X diagram for the c a l c - s i l i c a t e reactions relevant to the Penfold Creek assembalges. The curves represent the calculated and experimentally deter-mined phase equi l ibr ia for a constant total pressure of 5000 bars. (Gordon and Greenwood 1971; Metz 1970; Orv i l le 1970 - unpublished) The univariant l ine A^B represents where the assemblage: calc i te quartz - plagioclase - c l inozois i te - scapolite - grossular, i s stable. The shaded area represents the s tab i l i t y f ie ld for the assemblage: quartz - plagioclase - c l inozois i te - dlopside - Ca-amphibole -scapolite. Abbreviations: Ca - ca lc i te ; Q - quartz; Ta - ta lc ; Tr - tremolite; Phg - phlogopite; Ksp - potash feldspar; Di - diopside An - anorthite; Gr - grossular; Wo - wollastonite; Zo - zo is i te ; PI - plagioclase feldspar; Mz - mizzonite Sc - scapolite. Important reactions: Tr + 3Ca + 2Q = 5Di + 3C02 + H20 5Pg + 6Ca + 24Q = 3Tr + 5Ksp + 6C02 + 2H20 5Ta + 6Ca + 4Q = 3Tr + 6C02 + 2H?0 A l l complete curves shown are stable. A l l complete reaction curves experimentally determined or calculated. A l l complete reaction curves positioned by Schreinmakers analyses or estimated from other equi l ib r ia . ^— (short l ine at invariant points) metastable extensions. 7 8 5talc + 6ca l c i t e + Aquartz = 3tremolite + 6C02 + 2H 0 tremolite + 3ca l c i t e + 2quartz = 5diopside + 3C0 2 + H20 The composition of the evolved vapor phase for both these reactions and for reaction (2) i s the same, and, therefore, the shapes of the reaction curves i n P-T-X space must be s i m i l a r . Figure 23 refers only to the pure Mg phases ( i . e . diopside, tremolite, phlogopite). The minerals i n the Pnefold Creek cal c -s i l i c a t e rocks ce r t a i n l y contain appreciable amounts of Fe, judging by their o p t i c a l properties. The d i s t r i b u t i o n of Fe between the various phases w i l l effect the position of the equilibrium curves i f there i s prefe r e n t i a l take up of Fe into one of the phases. For instance, i f Fe i s p r e f e r e n t i a l l y taken up i n pyroxene rather than the amphibole i n reaction (1) then the s t a b i l i t y f i e l d of the pyroxene assemblage w i l l become larger at the expense of the amphibole f i e l d . It might be possible, therefore, to reverse the positions of certain equilibrium curves i n T-X space i f one considered the Fe component. This p o s s i b i l i t y i s explored i n Figure 24, which i s a X -T pl o t , which assumes that Fe i s prefer-r e e n t i a l l y taken up i n pyroxene i n reaction (1) and i n mica for reaction (2). The two reaction curves would then meet at an invariant point and the position of the potash feldspar-forming reaction would be reversed with respect to reaction (1). Without actual analyses i t i s d i f f i c u l t to say whether this s i t u a t i o n could occur. However, referring to the l i t e r a t u r e i t i s found that i n most metamorphic rocks Fe i s p r e f e r e n t i a l l y incorporated into the amphibole for both the mineral pairs pyroxene: amphibole and mica:amphibole (Leelanandam 1970; Carmichael 1970). This would indicate that reactions (1) and (2) would diverge i n X - ^  s P a c e » rather than the s i t u a t i o n shown i n Figure 24. Thus the equilibrium curves 79 Figure 24» The effect of Fe sol id solution on the c a l c - s i l i c a t e reactions The symbols are the same as in Figure 23. However Di , Tr, and Phg are now Fe-bearing and not pure end-members. It has been assumed that Fe is preferential ly incorporated into pyroxene (diopside) for reaction ( l ) , and into mica (phlogopite) for reaction (2). R e a c t i o n s » 3T;r + 5Ksp + 6C0 ? + 2H20 5Di + 3 C 0 2 + H20 Phg + 12Di + 6C02 + 2H20 Phg + 2Di • 4Q Ksp + 3Di + H20 + 3C02 5Phg + 6Ca + 24Q Tr + 3Ca + 2Q 3Tr + Ksp + 6Ca Ksp + Tr Phg + 3Ca + 6Q absent phases Di] Ksp,Phg] Ca] Tr] 8 0 i n Figure 23 can be used for the discussion of the phases found i n the Penfold Creek assemblages. The position of reaction (2) r e l a t i v e to reaction (1) has been confirmed i n the f i e l d by Carmichael (1970). In the area described by Carmichael (1970) as one approaches the highest grade rocks the isograd defined by reaction (2) i s encountered f i r s t , and followed by the isograd defined by reaction (1), but i n the Penfold Creek area the isograds are reversed (Fig. 22). This apparent anomaly can be explained i n two ways: f i r s t i f P.... . , increased r e l a t i v e to P with increasing grade fluxd t o t a l (see Fig. 25), and second i f the composition of the f l u i d phase changed systematically r e l a t i v e to the grade of metamorphism (Fig. 26). Changes i n either of these two variables, P f - j ^ ^ a n c^ f l u i d composition, could account for the d i s t r i b u t i o n of the c a l c - s i l i c a t e isograds i n the Penfold Creek area. A variable P r i . however, requires P.... to be less than fluxd f l u i d P with the pressure difference being supported by the strength of tot 3 . j -the mineral grains. This condition i s unlikely to occur i n a metamorphic environment, where certain phases appear to be highly soluble ( i . e . quartz). The composition of the f l u i d phase i n the Penfold Creek calc-s i l i c a t e s must have had a X of less than 0.2 because of the s t a b i l i t y L.U2 of either grossular garnet, z o i s i t e , or scapolite i n most of the mineral assemblages (Fig. 22 and 23). Thus an increase of X toward the C 0 2 higher grade would account for the apparent reversal of reactions (1) and (2). The cause of this increase i s discussed i n a l a t e r section. The importance of minor changes i n the composition of the f l u i d phase i s w e l l - i l l u s t r a t e d i n a layered c a l c - s i l i c a t e where the following two assemblages were found: 8 1 Figure 25 The effect of f l u i d pressure v a r i a t i o n on the c a l c - s i l i c a t e assemblages The reaction shown i n the ^ f i u ^ ~ T pl°t ( x C Q ~ 0.2) are: 1. 5talc + 3calcite + 4quartz = 3tremolite + 6C0 2 + 2^0 2. 5phlogopite + 6calcite + 24quartz = 3tremolite + 5Kspar + 6C0 2 + 2H20 3. tremolite + 3calcite + 2quartz = 5diopside + 3C0 2 + H20 The figure i l l u s t r a t e s how a different ^f-^u^ between the high and low grades could have resulted i n the d i s t r i b u t i o n of the c a l c -c i l i c a t e isograds now seen i n the Penfold Creek area. Figure 26 The effect of f l u i d composition v a r i a t i o n on the c a l c - s i l i c a t e assemblages The reactions shown i n the X„„ - T plot (P.... . , = 5Kb) are the C0 2 c f l u i d same as those i n Figure 25. The figure i l l u s t r a t e s how a different f l u i d composition between the high and low grades could have resulted i n the d i s t r i b u t i o n of the c a l c - s i l i c a t e isograds now seen i n the Penfold Creek area. FIGURE 25 FIGURE 26 calcite - quartz - plagioclase - clinozoisite - scapolite - grossular (1) quartz - plagioclase - clinozoisite - diopside - Ca-amphibole - scapolite (2) In Figure 23 assemblage (1) li e s on the univariant line A-B, and assem-blage (2) is located in the shaded divariant f i e l d . Temperature can not be responsible for the different assemblages because of their close proximity. However, a small change in X would explain the differences c u2 It is important to realize that the univariant line on which assemblage (1) lies becomes multivariant when other components are considered. The composition of the original sediment can only be inferred, The present assemblages suggest a range of compositions from an a r g i l l -aceous limestone to a calcareous a r g i l l i t e . Although no dolomite is found in any of the present assemblages i t is probable that i t was present originally and was responsible for part of the calc-silicate paragenesis. In support of this dolomite has been found in low kyanite/ staurolite zone metasediments of the Isaac Formation near Lynx Bay. Controls of Metamorphism Two basic models can be proposed to account for the rise in temperature necessary to cause regional metamorphism. The f i r s t consists of a sedimentary pile being downwarped into a zone of higher temperatures and pressures. The second model requires the heat to have been transferred by either convection or conduction into the sedimentary pi l e . The heat source in this latter case was probably in the form of an igneous body. In Penfold Creek area the close spacing of isograds negates the possibility of a purely depth controlled metamorphism. Thus, the energy driving the metamorphism must have been derived from a more or less localized heat source, presumably igneous. The shapes of isotherms in regional metamorphic areas have been discussed by Greenwood (Abstr. 1971). 8 3 He suggested that i f the heat transfer resulted from convective mass transport of heat by pore f lu ids , then the metamorphic area would tend to have closely spaced and nearly ver t ica l isograds along i ts margins. This geometry holds in the Penfold Creek and surrounding areas where only one and a half miles separate the b iot i te and s i l l imanite isograds, and the isograds are essential ly ver t ica l over several tens of miles. Before the arr iva l of the heat source which in i t iated the meta-morphism the rocks were probably under the stress system imposed on the rocks during the f i r s t phase of deformation. Increase in temperature, and possibly f lu id pressure, would have made the rocks more ducti le and the intensity of the ensuing deformation would be direct ly related to the grade of metamorphism. The change in fold style observed in the Penfold Creek area conforms with this model. The second period of metamorphism occurred after the second period of deformation in an essentially hydrostatic stress f i e l d . This meta-morphism took place, therefore, after the stress f ie ld of the f i r s t two periods of deformation had ceased to operate, possibly due to annealing by the f i r s t thermal event. Whether there was a long period of time between the two periods of metamorphism is open to question. However, the two generations of garnet growth are d is t inc t , and no apparent retro-grade metamorphism occurred between the two metamorphic episodes. It i s suggested, therefore, that the two episodes are pulses of essential ly the same metamorphism, and that the stress f ie ld diminished between the two pulses. The reaction isograds which have been proposed for the s i l l imanite zone metasediments of the Penfold Creek area require that with increasing grade P. H2O increases in the pe l i t i c assemblages and P c o 2 increases 8 4 in the carbonate assemblages. This means that the two assemblages acted as closed systems, and that there was a substantial difference between the f lu id compositions of the p e l i t i c and calcareous assemblages. The major reactions occurring in the p e l i t i c rocks in the s i l l imanite zone buffer the P^O' e - ' - i m ^ n a t ^ n S t n e n e e d for an external source of water. The carbonate assemblages on the other hand requires that CC^ enrichment of the f lu id phase came from some source external to the exposed rocks. This source of CO^-rich f lu id is postulated to be the decarbonation of underlying calcareous sediments, at higher temperatures than the s i l l i -manite zone rocks now observed. When the CO^- and H^O-rich f luids entered the calcareous and p e l i t i c rocks which are now observed they changed the composition of the f lu id phase which was or ig inal ly present. Consider the decarbonation reaction in Figure 27. The composition of the f lu id phase on the decarbonation reaction curves is fixed (A ) at a particular temperature (T). If a water-rich f lu id is added to the system, the composition of the f lu id phase w i l l change to, say, B (A--, greater than B_. in F ig . 27). This ^^ 2 results in the part ia l pressure of CO^ decreasing and thus this f lu id is out of equilibrium with respect to the decarbonation reactions. The decarbonation w i l l proceed unt i l enough CC^ has been produced to bring the f lu id composition back to A . This system is a ' se l f - f lush ing ' arrangement whereby any water entering the system is immediately swamped by CO^. This w i l l continue unt i l the particular carbonate assemblage has been exhausted at which point the f lu id composition w i l l move to another decarbonation curve. The scheme proposed suggests that in the pelites the f lu id phase must have been primarily rL^ O and in the carbonates a mixture of tL^ O and CG^ 8 6 in proportions buffered by the mineral reactions. Penfold Creek area marks the northern extremity of the Shuswap metamorphic complex. At present, none of the workers in this area have suggested a cause for the metamorphism. However, i t has been pointed out (Reesor 1970, p.86) that a narrow zone of heat r ise was necessary to have produced the present distr ibut ion of isograds. Previous authors have been understandably reticent in proposing theories of the source of energy in Shuswap metamorphism. The following hypothesis thus is at best highly speculative. The southern Shuswap metamorphic complex is intruded by several Cretaceous granit ic plutons ( i .e . Kuskanax batholith) which are generally post-tectonic. They are associated, however, with a late non-metamorphic period of folding in some places • The age of the Shuswap meta-morphism has been dated at post -Tr iassic by Hyndman (1968) and Jurassic to early Cretaceous by K.V. Campbell (1971). Thus, the plutons were intruded only a short time after the metamorphism. It is suggested here that the metamorphism was caused by the heat derived from the plutons prior to their intrusion. The association of the plutons with the higher grade metamorphic rocks in the southern Shuswap metamorphic complex would support this hypothesis. The number of Cretaceous plutons decreases towards the north, so that at the northern l imits of the complex in the Quesnel Lake area there are no post-tectonic plutons. It i s possible that the Coast Crystal l ine Complex and the Shuswap Complex have similar or igins. The Shuswap may represent the roof of a batholithic complex similar to that of the Coast Crystal l ine Complex. In support of this the roof pendants in the Coast Range Batholith have been regionally metamorphosed to assemblages typical of the c lass ica l Barrovian metamorphism as has the Quesnel Lake Area. 8 ? PART I I . THE QUESNEL LAKE GNEISS Introduction The Quesnel Lake Gneiss crops out mainly along the north shore of the East Arm of Quesnel Lake, where i t forms an elongate ridge r i s i n g to 25 00' above the lake l e v e l . I t consists of quartzo-feldspathic gneisses and g r a n i t i c gneiss, and was f i r s t recognized by R.B. Campbell (1961, 1963), who considered i t an exotic body of unknown age. More recently, he has suggested (1969) that " i t represents sheets of c r y s t a l l i n e basement upon which the sedimentary sequence was deposited, though i t may possibly be the product of transformation i n s i t u of some otherwise unrecognized u n i t . " Structure The gneiss i s s t r u c t u r a l l y concordant with the surrounding meta-sediments of the Kaza Group. Both the gneiss and the surrounding sediments have undergone three phases of deformation. The orientations of the s t r u c t u r a l elements associated with these deformations compare d i r e c t l y with those described from the Penfold Creek area. The gneiss has been folded into a tight antiform that i s overturned towards the south-west. This structure has been warped by an F2 p a r a s i t i c fold and arched during F3 deformaton. These deformations have resulted i n the gneiss having i t s present ' f i n g e r - l i k e ' outcrop pattern with the length of the body being p a r a l l e l to the major structural trend of the area, and the disappearance of the gneiss to the north-west and south-east as i t plunges beneath surrounding meta-sediments. Major folds are not observed within the gneiss because of the 8 8 lack of compositional layering. Quartz-feldspar augen and some quartz grains are elongate within the major fo l ia t ion . These have imparted a strong l ineation to the gneiss, which is sub-paral lel to the axes of folds of the surrounding metasediments. The major fo l ia t ion is cut by two sets of F3 fracture cleavage and a series of sub-paral lel joints and faults (Plate 18). A l l the structures within the gneiss can be di rect ly correlated with structures in the surrounding metasediments; no structures indicate that the gneiss has undergone a deformation earl ier than the f i r s t phase of deformation found in the surrounding metasediments. Fine-grained apl i te dykes are found both within the Quesnel Lake Gneiss and surrounding metasediments. They are folded and are generally concordant with the enveloping fo l ia t ion . Where they show cross-cutting relationships the gneissic fo l ia t ion can be traced through the ap l i te . They are assumed to be oogenetic with the gneiss. Near Peninsula Bay several elongate mafic inclusions have been recognized. At the water l ine they have been d i f ferent ia l ly weathered into oval p i t s . They are elongate within the major fo l ia t ion and aligned para l le l to the dominant l ineation of the gneiss. Metamorphism The metasediments surrounding the gneiss contain an abundance of Barrovian sequence index minerals. The metamorphic grade increases along the gneiss from garnet zone in the north-west to kyanite/staurolite zone in the south-east. The isograds l i e approximately perpendicular to the str ike of the gneiss. Associated with the highest grades of metamorphism are numerous pegmatite dykes which cut both the gneiss and surrounding metasediments. The texture of the gneiss varies with grade of metamorphism. In the garnet zone i t is a poorly fo l ia ted , coarse-grained, augen gneiss (Plate 16) while in the higher grades i t is a we l l - fo l ia ted , medium-grained, equigranular gneiss (Plate 18). No change in mineralogy could be related to the grade of metamorphism. Along one of the contacts of the gneiss there is a patchy develop-ment of a r e l i c t skarn, which has subsequently undergone regional meta-morphism. The c lass i f i cat ion of this rock as a re l i c t skarn was made primarily on i t s mineral assemblage: andraditic garnet - 40% (R.I . , 1.86), epidote - 29%, aegirine augite - 19%, ferrohastingsite - 10%, and calc i te - 10%. A similar assemblage was found in abundance near Dutchman Mountain by K.V. Campbell (1971) which is at the south-eastern extremity of the Quesnel Lake Gneiss. He reports a chemical analysis of this rock which has high total iron (FeO - 11.0%), high calcium (CaO - 17.0%), low s i l i con (Si0 2 - 45.1%) and low potassium (K20 - 0.7%). This analysis is consistent with the skarn c lass i f i ca t ion . Petrology The Quesnel Lake gneiss is pink to l ight grey, coarse to medium grained, and, except for the more mafic parts near the contacts, the proportions of the constituent minerals vary l i t t l e within the body. The major minerals, in order of abundance, are: quartz, microcline, plagioclase, epidote, b io t i te , and muscovite. The minor constituents include chlor i te , hornblende, sphene, apatite, orthite and opaque minerals. The microcline is perthi t ic in places and myrmekite is not uncommon. The b iot i te character ist ical ly shows green to l ight brown pleochroism in contrast to purely brown pleochroic biot i tes of the surrounding schists . Oriented medium-grained b iot i te flakes give 90 the gneiss an easily v i s ib le fo l ia t ion in the higher grades of meta-morphism. Epidote crystals are subhedral, zoned, and rarely cored by orthite. Augen from lower grades are composed of coarse-grained aggregates of quartz, microcline and plagioclase. The apl i te dykes are composed of f ine-grained, equigranular quartz, microcline and plagioclase. The deformed mafic inclusions have a mineral assemblage similar to the enclosing gneiss except that horn-blende and epidote are more abundant (Plate 17). Petrochemistry Chemical analyses of the Quesnel Lake Gneiss (Table 11) and some representative metasediments of the Kaza Group have been plotted on triangular variation diagrams (Figs. 28 and 29), The gneiss analyses include analyses from Dutchman Mountain Gneiss (K.V. Campbell 1971) which is part of the same gneissic body. The most str iking feature of these plots is the overall chemical difference between the gneisses and the metasediments. This would suggest that the gneiss is not part of the Kaza" Group. The gneisses also appear to l i e on a chemical trend which is ident ical to that quoted by Larsen (1948) as evidence for magmatic or igin of a granitic bathol ith. It has been argued that any gneissic terrain would show a similar chemical variation due to metamorphic d i f ferent iat ion. However, this would not explain why the surrounding schists , which appear to have undergone the same metamorphic episodes, do not show any such variations. Age and Origin of the Gneiss It may be suggested from the foregoing sections that the Quesnel Lake Gneiss could represent: 9 1 Table 11 Chemical and Modal Analyses of the Quesnel Lake Gneiss , Nos. 1 lb 2 3 4 5 6 7 8 I. Nos. 23 23 156-1 156-3 161 162 307-3 309-1 309-2 S i0 2 57.5 57.6 65.9 64.4 70.5 69.8 66.4 67.6 73.3 T i 0 2 0.80 0.83 0.47 0.35 0.24 0.30 0.40 0.40 0,17 Al 2 03 18.0 17.9 14.7 18.9 14.3 14.8 14.4 17.0 13.7 Fe203 3.50 3.40 2.50 1.70 0.90 2.00 1,50 1,40 0.60 FeO A.70 4.70 1.50 1.00 0.80 1.10 2.30 1.90 0.80 MgO 2.70 2.40 1.50 0.30 1.30 1.80 2.50 1.70 1.10 CaO 6.50 6.40 3,20 2.40 1.10 2.80 3.60 3.00 1.70 MnO 0.13 0.14 0.07 0.08 0.06 0.04 0.12 0.07 0.02 K20 2.90 2,90 3.60 6,10 5.40 3.30 2.70 4.40 4.00 Ka20 2.00 2.20 2.80 4.30 3.50 2.90 2.70 3.40 2.90 H20 1.70 1.70 1.20 0.60 0.50 0.90 0.90 O.bO 0.40 P 2°5 0.18 0.30 0.13 0.02 0,07 0.10 0,11 0.23 0,06 C0 2 0.20 0.10 0.10 0.20 0.10 0.10 0.10 0.10 0.10 Total 100.8 100.6 97.6 100.1 98.7 98.0 97.7 98.9 102.0 modal analyses Quartz M M A M A M A M A Microcline tr tr M A A M m A M Plagioclase A A M M m M A A M Epidote M M M m m m m m m Biotite M M m m m M M m m Muscovite m m m tr m tr m tr tr lb - duplicate analysis. A - abundant >30%. M - major 10-30%, m - minor 2-10%. tr - trace<2%. analyst - S.Courvi l le , Geological Survey of Canada. Note: for specimen locations see Map 4. 92 Figure 28. T R I A N G U L A R P L O T O F M E T R S E D I M E N T A N D G N E I S S A N A L Y S E S ( M O L E P E R . ) AL203 93 Figure 29. T R I A N G U L A R P L O T O F M E T A S E O I M E N T A N D G N E I S S A N A L Y S E S ( M O L E P E R . ) FEO^MGO 9 ^ (1) a previously unrecognized unit within the Proterozoic sediments (2) basement exposed in the core of a tight ant ic l ine (3) a thin s l i ce of basement thrust into the overlying sediments (4) a granit ic Igneous intrusion. The remarkable difference both chemically and petrographically between the gneiss and the surrounding sediments, the proven chemical and petrographic difference between the potash feldspar-bearing schists of the Kaza Group and the granitic gneiss (K.V. Campbell 1971), and that no similar unit has been recognized in the Kaza Group of the Cariboo Mountains appear to exclude the f i r s t suggestion. The structure which has been proposed for the Quesnel Lake Gneiss (Fig. 30) would exclude the poss ib i l i ty of the gneiss representing a basement core to a recumbent ant ic l ine. The alternative structure shown in Figure 28b would, however, conform with the suggestion of a basement core. This type of structure is consistent with the outcrop pattern and would compare di rect ly with the structure and or ig in , as suggested by Ross (1968), for a similar gneissic body near Revelstoke. This structure would require the postulation of an extra period of deformation for which evidence has not been observed. No earl ier fo l ia t ion or l ineation has been seen which might correlate with this early deformation. Although the stratigraphy is d i f f i c u l t to interpret near the gneiss, the distr ibution of the 'skarn' rocks is consistent with the structure shown in Figure 28a. No 'skarn' rocks have been seen along the metasediment/gneiss contacts which bound the narrow infold of metasediment along the north shore of the East Arm of Quesnel Lake. A s l i ce of basement thrust into the overlying sediments and sub-sequently folded is a possible or igin for the gneiss. In support of 9 5 Figure 30. Two possible cross-sections of the Quesnel Lake Gneiss, 9 6 this severe shearing has occurred i n some of the schists surrounding the gneiss, and that travertine breccia and serpentine boulders are present near one of the gneiss contacts. These might be taken to imply a zone of weakness within the crust during recent and possibly older times. The contact between the gneiss and schist shows no sign of shearing but instead a zone of contact metamorphism and metasomatism. I f , as suggested, the contact rocks are genetically related to the gneiss, then the thrust s l i c e would have to include the contact rocks. This i s a p o s s i b i l i t y but the thrust plane must be near the gneiss because the surrounding rocks, except i n the contact zone, appear to be i d e n t i c a l to known metamorphosed sediments of the Kaza Group. The question whether the gneiss could represent basement i s nearly impossible to answer as no structures e a r l i e r than Phase 1 have been recognized. The major f o l i a t i o n of the gneiss i s concordant to i t s contacts and the s c h i s t o s i t y of the surrounding rocks. The intensity of the gneissic f o l i a t i o n i s related to the grade of metamorphism (R.B. Campbell 1970) suggesting that a strong f o l i a t i o n was not present prior to the major metamorphic episode seen i n the surrounding schists. If the gneiss represents base-ment, i t would have had to possess only a weak f o l i a t i o n at the time of emplacement. F i n a l l y , there are several lines of evidence which suggest that the Quesnel Lake Gneiss represents a metamorphosed igneous intrusion: the recognition of regionally metamorphosed skarns, a p l i t e dykes i n the gneiss and surrounding sediments, mafic inclusions within the gneiss which could represent xenoliths i n the granite intrusion, and the igneous trends to the analyzed specimens. Further indirect evidence i s that the gneiss appears to have been r e l a t i v e l y unfoliated before 9 7 the deformation and metamorphism associated with Phases 1 and 2. A preliminary age for the Quesnel Lake Gneiss which has been obtained from whole rock Sr/Rb ratios is 750 ± 50 mi l l ion years, (Blenkinsop 1971 — Pers. Comm. ). The rocks from the Purcell (Belt) System to the south range in age from 1500 - 850 m.y. (Obradovich and Peterman 1968, Ryan and Blenkinsop 1971) and are separated from the overlying Windermere System by a widespread unconformity and a period of igneous intrusion and folding. This orogeny has an age between 675 - 745 m.y. and has been named the East Kootenay Orogeny. As yet no ages have been obtained from the Windermere System. The Quesnel Lake Gneiss would therefore l i e approximately within the time l imits of the East Kootenay Orogeny, which would favour an igneous origin for the gneiss. Unfortunately, i t has been inferred that the surrounding rocks belong to the Kaza Group which is part of the Windermere System. No Purcell rocks have been identif ied in the Cariboo Mountains. However, in areas of high grade metamorphism positive ident i f icat ion would be d i f f i c u l t . It i s also possible that the age of the East Kootenay Orogeny was older in the Cariboo Mountains, so that a thick sedimentary p i le of rock of the Windermere System could have been deposited prior to the intrusion of a later granite. The va l id i ty of this suggestion would depend on the age of the Racklan Orogeny in the Yukon, which is correlated with the East Kootenay Orogeny. However, no radiometric age has been published for this orogeny. The depth of bur ial need not have been great because the structure of the gneiss suggests that i t was intruded as a s i l l . It could be argued that the age of 750 m.y. for the Quesnel Lake Gneiss confirms the suggestion that the gneiss is basement and that the age was reset during the East Kootenay Orogeny due to the loss of radiogenic strontium. If this were the case, i t is peculiar that the major period of folding and metamorphism now seen in the surrounding rocks does not appear to have caused a signif icant loss of radiogenic strontium. Conclusive evidence that the Quesnel Lake Gneiss represents either a basement s l i ce thrust into the overlying sediments or a metamorphosed granit ic s i l l has not been found. However, the author considers that the weight of evidence supports the latter suggestion that the gneiss represents a granite s i l l . APPENDIX I Chemical analyses of metasediments from the Kaza Group and Isaac Formation Chemical analyses of metasediments from the Kaza Group and Isaac Formation are presented i n Table 12. The location of these specimens i s given in Map 3* A l l the specimens are considered to have been either a r g i l l i t e s or arenites. Triangular plots of the bulk rock analyses are given in Figure 31. Table 12 Chemical Analyses of the Schists from the Quesnel Lake Area F i e l d Nos. A 5-1 105-1 108-1 109-1 117 140 301-2 143-3 14(i 148-2 149-2 181 230-2 232 311-8 312-3 252-1 226-5 229-1 Spec. Kos. 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 S i 0 2 73.8 60.8 52.4 63.3 65.1 55.9 56.7 62.0 60.:! 55.6 58.8 57.8 58.0 59.5 46.3 56.3 59.8 76.9 59.4 T i 0 2 0.58 0.86 0.67 0.92 0.88 1.14 0.96 0.92 0.U9 1.52 1.59 1.61 0.96 1.04 0.92 1.14 0.94 0.54 0.65 A l 2 0 3 12.8 18.6 24.1 19.1 17.2 21.1 21.7 21.3 18.5 20.9 22.5 18.5 18.1 17.8 17.7 21.0 16.8 9.5 20.1 FC2O3 1.2 1.6 1.5 1.1 1.3 3.6 1.1 1.6 1.:! 3.4 1.4 1.2 3.5 1.4 1.7 2.4 1.1 1.1 3.2 FeO 3.0 5.8 4.6 5.0 4.7 5.0 7.3 5.9 7.3 7.3 4.7 4.3 5.9 6.0 7.9 7.9 6.8 3.8 2.9 MgO 1.2 2.9 3.5 2.3 3.0 4.6 3.7 2.8 3.6 1.9 2.5 3.0 6.4 2.7 2.8 3.3 3.4 2.3 3.2 CaO 0.3 0.6 1.0 0.7 1.0 0.6 0.6 1.1 1.1 2.8 0.5 1.3 0.5 3.8 8.8 0.6 1.8 0.6 1.5 MnO 0.03 0.05 0.04 0.07 0.07 0.09 0.09 0.21 0.08 0.13 0.04 0.11 0.06 0.15 0.12 0.13 0.C8 0.08 0.22 K 20 2.6 3.8 6.0 4.3 4.5 3.B 4.9 3.9 4.3 2.8 4.3 3.5 3.1 2.2 2.0 4.8 2.8 0.8 4.6 Na 20 1.3 1.1 2,5 1.0 1.4 1.5 0.8 1.7 2.0 3.4 1.2 1.7 0.5 2.9 4.4 0.4 3.4 1.8 1.2 l l 2 0 1.7 2.1 3.1 2.6 2.6 3.5 3.1 2.8 2.o 1.6 3.0 2.6 4.7 1.8 2.4 4.2 1.4 2.0 3.0 ? 2 0 5 0.03 0.07 0.08 0.06 0.06 0.07 0.80 0.14 0.20 0.17 0.08 0.05 0.12 0.17 0.29 0.11 0.09 0.06 0.12 C0 2 <0.1 <ro.i <0.1 <0.1 <0.1 <0.1 0.1 <0.1 <o.:i <0.1 <0.1 <0.i <0.1 <0.1 4.4 <0.1 <0.1 0.4 0.7 Total 99.0 98.3 99.5 100.5 101.8 100.9 101.8 100.9 101. 3 101.5 101.8 99.2 101.8 99.5 99.7 102.3 98.4 99.9 100.6 modal analyses Quartz A A hi A A M M A A M A A A A A A A A A Plagioclase M M M M M m m A ro A M M in M m tr M tt -B i o t i t e m M M M M m M M D H M M ro M M tr M - -Muscovite n M A M n A A n tn M A m M m - A m ro A C h l o r i t e n tr t r t r t r t r t r - t r t r t r t r ro t r ro M tr M M Garnet TO t r ro m m ro n in ro M tn H ro t r ro - -S t a u r o l i t e - a - - - m - t r tr t r a - - - - tr - -S i l l i m a n i t e - - - - - - - - - - - - • - - - t r - -Kyanite - M - - - t r m m - - - - - - - - - - -Opaques t r t r - t r t r t r t r tr t r a tr t r t r ro tr t r t r t r t r Rock Unit K K K K K K K K 1 I I I H I 1 I I K. K A - abundant > 30Z; M - aajor 10-30Z; m - minor 2-10Z ; t r - trace <2Z. K - Kaza Group rocks; I - Isaac Formation rocks analyst - S.Courville, Geological Survey of Canada. Note: for specimen lo c a t i o n see Map 3. Table 12 (cont) F i e l d Nos. 238-2 Spec. Nos. 20 S10 2 39.3 T i 0 2 0.76 AI2O3 16.8 Fe?0 3. 1.7 FeO 4.3 KgO 5.6 CaO 2.1 Mr.O 0.21 K^O 2.7 Na 20 2.7 HoO 1.7 P2O5 0.12 CO2 <0.1 To t a l 98.0 Cuartz A Plagioclase to B i o t i t e m Muscovite tn C h l o r i t e K Garnet m St a u r o l i t e S i l l i m a n i t e Kyanite Opaques t r Rock Unit K 256-1 259 276-1 21 22 23 79.9 52.0 51.7 0.40 0.86 1.31 8.50 24.2 27.0 0.9 1.2 •1.4 1.9 5.3 4.8 2.0 6.0 1.4 0.3 0.7 1.0 0.04 0.21 0.13 1.6 6.8 5.8 1.4 0.9 1.6 1.2 3.7 3.9 0.03 0.08 0.05 0.2 0.1 <0.1 98.4 102.1 100.1 A M M n m m tr M tn m A A tr m m - n n — in - t r t r K K I Chemical Analyses of 297-2 282-1 300-1 24 25 26 54.3 56.5 58.5 1.02 1.19 1.56 20.7 19.0 18.8 2.0 4.0 1.3 6.6 6.2 6.2 2.8 3.0 4.6 1.1 0.7 1.1 0.06 0.19 0.19 5.3 3.6 4.0 1.0 2.5 1.6 3.5 3.1 2.9 0.04 0.14 0.11 <0.1 <0.1 <0.1 98.4 100.1 100.9 A M A Q H a m a ut A K M a m m M m M t r m tr 1 I I the Schists from the 195-2 211-1 211-2 27 28 29 58.1 56.7 61.8 1.12 1.24 1.01 21.0 16.5 15.0 1.4 1.3 1.4 7.4 9.5 7.2 4.3 5.S 5.7 1.5 0.3 1.4 0.09 0.06 0.12 3.1 4.8 2.7 2.0 1.0 3.0 1.8 2.8 2.1 0.10 0.06 0.04 <0.1 <0.1 <0.1 101.9 100.6 101.5 modal analyses M M A M m M M A M m m m m M m a m m - m tr m tr t r t r tr t r I I I Quesnel Lake Area 334-1 333-2 331-5 30 31 32 79.0 58.9 56.5 0.51 1.10 0.73 8.60 18.4 19.4 0.5 1.8 3.7 2.8 7.9 3.6 2.2 4.1 3.8 0.5 1.8 1.9 0.02 0.14 0.12 1.4 3.1 4.4 1.3 2.9 1.5 1.2 2.1 2.8 0.32 0.09 0.12 <0.1 <0.1 0.3 98.4 102.3 ' 98.9 A A A m A m m M a m m M m - in m a -- t r -m t r -m - -t r t r m I 1 K 329-1 329-2 329-4 33 34 35 72.6 53.5 54.4 0.61 0.87 1.04 11.6 23.1 25.3 0.9 2.2 2.2 3.5 5.0 4.6 2.4 3.6 2.7 0.9 0.4 0.3 0.07 0.03 0.05 2.0 5.4 6.4 1.8 0.5 1.0 1.8 5.0 4.3 0.06 0.11 0.11 0.1 <0.1 <0.1 98.3 99.7 102.4 A A A m m tr m - -A A m m t r t r : -t r t r t r K K K 331-3 239-1 239-2 36 37 38 80.6 77.0 66.5 0.37 0.54 0.92 8.60 10.1 15.0 1.0 1.5 2.5 1.9 2.8 5.3 1.9 2.4 3.1 0.6 0.3 0.4 0.04 0.30 0.08 0.6 1.0 1.9 2.1 • 2.0 1.9 1.0 2.0 3.8 0.02 0.06 0.14 <0.1 <0.1 <0.1 98.7 100.0 101.5 A A A M H M m - -m m a m m a t r - -t r t r t r K K V-102 Figure 31 Triangular plots of metasediment analyses from the Isaac Formation and Kaza Group. RL203 K20+NR20+CR0 • LOW GRADE KRZR GROUP A HIGH GRRDE KRZR GROUP + HIGH GRRDE ISRRC FORMATIC FEO+MGO K20 • LOW GRADE KRZR GROUP A HIGH GRRDE KRZR GROUP + HIGH GRADE ISRRC FORMAT IC 103 APPENDIX II Rock unitB of Penfold Creek area This short section is meant as a supplement to Table 1, which describes the rock types, and to the legend of Map 1, which br ie f ly describes the rock uni ts . 1) Rock units of the Kaza Group A. Relatively homogeneous l ight green subarkose containing several good examples of graded bedding. The unit i s well bedded. B. Interlayered gray phyl l i te (C) and feldspathic quartzite. The unit is composed of approximately 80% feldspathic quartzite which i s interbedded with thin (0.5' - 20') beds of phy l l i te . The feldspathic quartzite i s well bedded and in a few places has well preserved sedimentary structures such as graded beds and cross-bedding. D. Thickly bedded sequence of dominantly pink to pale gray quartzite (90%). The quartzite beds range in thickness from 0.5' to 50» and are interbedded with fine-grained b io t i te -bearing phyl l i tes . The phyl l i te beds average 21 in thickness, with a maximum of 10'. E. Fine-grained dark gray laminated limestone. Graphite i s abundant. Thin layers of brown coarsely crystal l ine l ime-stone (approximately 0 .3 inches thick) are interbedded with the graphitic variety. 1 0 4 P. Dominantly a thickly bedded l ight gray quartzite, with two f i f teen foot inter-beds of black graphite-r ich phy l l i te . In places the phyl l i te has a ' c l i n k e r - l i k e ' appearance. G. Light green mylonite containing mainly highly sheared f e l d -spathic quartzites and some more micaceous members. Very elongate strained quartz crystals with the occaisional rotated feldspar grains may be seen in thin section. This rock unit exerts prominant control on physiography, forming a c l i f f one hundred feet high along the north side of the headwaters of Penfold Creek. 2) Rock units of the Isaac Formation H. Well laminated dark green amphibolite. One of the amphibolite bands has been severely boudinaged leaving a series of amphib-o l i te pods. There i s only a s l ight tendency for elongation of the amphibole crystals to paral le l the fo ld axes. A decussate texture i s more common. Garnets, where present, are red-brown and invariably have a white plagioclase-r ich corona surrounding them. I. Interlayered dark brown mica schist and micaceous quartzite. This unit i s the most abundant in the Isaac Formation of Penfold Creek. The rocks have a banded appearance, with the l ight gray quartzites rarely being more than 3' thick. The proportion of quartzite to schist varies up to a maximum of f i f t y percent quartzite. 105 J . The mica schist i s exactly the same as the schist in unit ( i ) . In this unit , however, there i s l i t t l e or no quartzite. K. A re lat ive ly d is t inct and homogeneous unit containing large rotated garnets (up to 1" in diameter). The mica schist appears to be more feldspathic than the schists of units ( i ) and ( j ) . L. Gray marble ranging from coarsely crystal l ine homogeneous to fine-grained layered var iet ies . This unit forms a 200' c l i f f i n one place. Any structures in the marble are outlined by very thin layers of mica. However, large sections of the marble are devoid of v is ib le structures. Quartz i s present as rounded blebs scattered at random throughout the rock. Muscovite, commonly with a core of fuchsite, pyrite cubes, and f ine ly disseminated graphite make up the balance of the mineralogy. Where this unit has been thinned to less than 50' i t i s commonly indistinguishable from unit (M). M. Brown weathering medium-grained schistose marble. The percent of mica varies so that in places there are nearly pure marble bands. These la t te r bands are discontinuous and rarely more than a few feet thick. N. Well layered sequence of quartzites with less than 5% schistose layers. The quartzites range from feldspathic to micaceous and from very l ight gray to l ight pink in colour. 0. A unit defined by an association of sulphide-rich schists , ca lc -s i l i c a t e , carbonate, and minor amphibolite. The unit i s easi ly recognisable in the f i e ld by being weathered to an iron-stained s i l t . The highly variable mineralogy depends on grade of meta-morphism and composition of the original sediment. 106 APPENDIX III Microprobe Standards The following standards were used for analyses of the minerals from Penfold Creek area. A l l the standards, with the exception of kyanite, are from the col lect ion of Dr. B.W.Evans at the University of Washington, Seattle. To calculate oxide percentages either a graph-. : ing technique (computer program UPRoT^University of Washington) or a complete correction program (Rucklidge EMDRO, University of B r i t i s h Columbia) was used. GARNET BIOTITE U/W id nos #9-12 #9-1 #9-4 #7.5 Source Cal . Tech. U/C Berk Cal . Tech. U/C Berk Location Nuevo Gore Mtn Sturbridge Crown Pt S i0 2 36.2 39.0 37.5 34.4 T i 0 2 0.05 0.08 0.00 3.06 A1 2 0 3 20.0 22.1 21.4 13.3 FeO* 21.5 22.0 32.7 30.4 MnO 21.1 0.48 0.59 O.38 MgO 0.31 11.5 6.50 4.56 CaO 0.48 4.00 1.24 0.02 Na20 - - - 0.16 K 20 - - - 9.20 Total 99.64 99.16 99.93 95.40 Technique Graph Rucklidge FeO* - total iron as FeO. 1 0 7 FELDSPAR Tj/W id nos #1.12 #1.3 #1.1 #2.1 Source Lindsley Crawford Smith Location syn. An90 Norway Tiburon S i 0 2 45-4 62.4 68.2 63.9 A1 2 0 3 35-0 23.0 19.9 18.0 FeO* 0.04 0.06 - 0.01 CaO 18.3 4.70 - -Na20 1.05 8.70 11.75 0.13 K 20 - 0.65 0.03 16.3 Total 99.79 99.51 99.88 93.34 Technique Graph MUSCOVITE CHLORITE ILMENITE U/W id nos #8.1 #11.6 #11.2 #13.3 Source U/C Berk Troramsdorf Albee Anderson Location New York Switzerland S i 0 2 44-7 33.2 26.9 -T i 0 2 0.37 0.00 0.11 49.7 A1 2 0 5 31.5 15.7 22.2 -FeO* 5.12 2.52 23.9 46.6 MnO 0.03 0.01 0.48 0.68 MgO 0.99 36.2 15.3 1.28 CaO 0.02 0.00 0.00 -Na20 0.67 - 0.10 -K 20 10.45 - 0.23 -Total 94.35 85.63 89.22 98.26 Technique Rucklidge Graph Rucklidge Additional biot i te #7. 5 (Ti.K) Standards garnet #9.12 (Mn) An90 #1.12 (Ca) albite #1.1 (Na) 108 STAUROLITE U/W id nos #12.9 #5-11 #16.3 Penfold Creek Source Roeder Frisch Hodgson kyanite, assumed Location basalt glass clinopyroxene sphalerite stoichiometric Si02 49-4 44.9 Fe 9.2 Al 33-4 Ti0 2 2.75 1.97 Mn 0.16 Si 17.4 A1203 12.0 10.23 Zn 56.3 0 49.2 FeO* 11.6 12.2 S 34.1 MnO 0.16 0.37 MgO 10,26 8.11 CaO 10.02 18.8 Na20 2.21 2.27 K20 0.58 0.05 Total 98.93 98.90 Technique Rucklidge No staurolite standards were available so the following standards were used in analysing for the different elements: av. atomic wt. Si • - kyanite 10.71 Ti -- clinopyroxene 13.29 Al -- kyanite 10.71 Fe -- basalt glass 12.31 Mn -- garnet (#9.12) 15.35 Mg -• basalt glass 12.81 Zn - sphalerite 24.78 109 APPENDIX IV Diffusion i n Garnets The inset i n Figure 32 shows a CaO microprobe traverse across a garnet i n Specimen 7. This i s a r e p e t i t i o n of the CaO traverse i n Figure 17 which also shows the MnO, MgO and FeO traverses. This part i c u l a r garnet i s considered to have c r y s t a l l i z e d i n two stages: an inner syn-tectonic garnet and an outer post-tectonic garnet. The v boundary between the two generations of garnet i s sharp and i t may be assumed that an i n i t i a l chemical discontinuity would also be sharp. The CaO curve shows a marked jump from approximately 2.0 to 8.0 weight percent across the boundary. However, a detailed traverse of this jump has the characteristics of a d i f f u s i o n curve (Fig. 32). If i t i s assumed that the present arrangement was due to d i f f u s i o n i t should be possible to estimate the duration of the second metamorphism. To do this i t i s necessary to assume that d i f f u s i o n only occurred at or above 500°C, and that l i t t l e or no d i f f u s i o n occurred below these temperatures. The calculated duration of the metamorphism would therefore be the time the rocks were subjected to temperatures i n excess of 500°C. The temperature of 500°C was chosen because i t was considered the lowest possible temperature at which garnet could have formed. A higher temperature w i l l shorten the calculated duration. Unfortunately, there i s no published data on d i f f u s i o n i n garnets, so i t i s necessary to use data on o l i v i n e . Simpkin (1966) estimated a -11 -12 2' d i f f u s i o n c o e f f i c i e n t (D) of 10 to 10 cm /sec for Mg i n o l i v i n e i n a p i c r i t e s i l l at 1100°C. This i s i n good agreement with the experimental result of D Mf. . = 3.7 x 1 0 _ 1 1 cm2/sec at 1100°C (Misener 1972). o l i v i n e Figure 32. A d i f f u s i o n curve from a microprobe traverse of a garnet lOOO-i 800H <0 o V O a o>600-e O o 400-200 9 8 o 7 CO o 6 •*-> «• £ 5 position of detailed traverse N 1-0 01 — i — 2-0 \ \ * A \ V V V \ \ \ * X 0 2 3 0 4-0 5-0 Radial Distance in mm 1 03 0-4 0-5 I l l Misener has also calculated the activation energy (AH) - 47 ± 4 Kcals and D - 0.34 ± 0.1 x 10 - 2 cm /sec for this d i f f u s i o n process. These o results allow the calculation of D for any temperature using the re-lationship : —AH/RT D = D e where R = gas constant T = temperature Using the maximum values allowed by the stated errors a maximum D M?. . at 500°C of 1.0 x 10~ 1 7 cm2/sec has been calculated, olxvme 2 Using the relationship x = 2Dt (Simpkin 1966) where x i s the width of the d i f f u s i o n zone (0.04cm i n this case), D i s the di f f u s i o n coefficient and t the time i n seconds, i t was found that the maximum duration of the second metamorphism was 2.5 m i l l i o n years. If the section through the garnet was non-radical then the apparent duration of the metamorphism would be longer than the actual duration. Thus a maximum value would s t i l l be obtained. 112 APPENDIX V Thermodynamic calculations of the s i l l imani te -forming reaction curves 1) Definitions of the thermodynamic parameters used in the calculations. T - temperature ° K . P s - pressure acting on sol id phases. R - gas constant. f n 2 - fugacity of oxygen. V - molar volume. A V B - molar volume change in sol id phases for a reaction, x^  - mole fraction of component i . ^ ^react " change in enthalpy for a reaction in standard state. ^ ^ ? , 2 9 8 ~ heat of formation from the elements. A.G° . - change in Gibbs free energy for a reaction in reacx standard state. A G ° " Gibb s free energy of formation from the elements. S ° 9 8 - entropy at 298.15°K A S ° - change in entropy for a reaction in standard state. 1 1 3 2) Calculation of the reaction curve for : Garnet + Muscovite = Biot i te + 2Sill imanite + Quartz The thermodynamic data for these minerals (Robie and Walbaum 1968) are: V S ° g Q A H f , 2 9 8 (cc/mol) (cals deg/gfw) (Kcals/gfw) Quartz 22.69 9*88 -217.6 Muscovite 140-71 69-00 -1421.2 Si l l imanite 49-90 22.97 -618.6 The data for biot i te (annite) and garnet (almandine) have been obtained in the following manners a) V almandine - 115.28cc/mol based on a = 11.526$ (Skinner 1956). annite - 153-l6cc/mol based on the ce l l parameters given by Eugster and Wones 1962. b) ^298 calculated from the entropy of the oxides using the method described by Fyfe et a l . (1958, p. 28-34). almandine (Fe^A^Si^O-^) ~ 76.76 cal/mol deg. annite (KFe^AlSijO-^tOH^ - 101-39 cal/mol deg. c) AH°^298 annite - -1235«0 Kcals/gfw (Eugster and Wones 1962, p. 115) There are no data available for almandine so that i t was neccessary to calculate H° ^ 99 ^ o r almandine from experimentally determined curves for the reaction* Ilk 3Almandine + 0 2 = 3 Hercynite + 2Magnetite + 9Qnartz at an oxygen pressure defined by the fayalite-magnetite-quartz buffer (Hsu 1968 p. 61). To determine AH of reaction from the pressure-temperature curve for this reaction the method of Orv i l le and Greenwood (1965) was used: d(log f 0 2 ) -AV S dP s A H r e a c t d(l/T) 2.303BT d(l/T) 2.303R where log (fn 2 ) = (0.09374(P - 1.0) - 25738.0)/T + 9.00 (Wones and Gilbert 1969). From which: A H 2 g 8 of reaction = 94«6 Reals (at 1 bar) A H ° 2gg for magnetite (-267400 cals) and quartz (-204646 cals) are known (Robie and Walbaum 1968), but there are no data for hercynite. To obtain this data i t was neccessary to use the reactions: i . 2Fe + 2A1 20 5 + 0 2 = 2FeAl204 for which A G ° e a c t = -142900 + 36.OT cals (Turnock and Eugster 1962) i i . 2A1 + 3/202 = A1 2 0 5 for w h i c h A G ° e a c t = -378082 cals (Robie and Walbaum 1968) These g i v e A G £ j 2 C . 8 for hercynite = -444165.3 cals (Fe + 2A1 + 2O2 = FeAl 2 0^). It i s known that Sggs hercynite = 25.4 cals (Robie et a l . 1968) thus: A S r e a c t = 2 5 , 4 " S Fe " 2 S A 1 " 2 S 0 2 = "92»46 cals (Fe + 2A1 + 20 2 . FeAl 2 0 4 ) A H r e a c t = A G ? e a c t * K e a c t Thus: A H ° „ (hercynite) « -471732 ca ls . Returning to the reaction: jJAlmandine + Cv, = 3Hercynite + 2Magnetite + 9Si0 2 we have: AH° n r , D (almandine) = -1263.9 Kcals. 1 In summary the following data have now been calculated: Y S296 K > 2 9 e (cc/mol) (cals deg/gfw) (Kcals/gfw) almandine 115.28 76.76 -1263.9 annite 153-16 101.79 - 1 2 J 5 . 0 These figures permit calculation of equilibrium curves for the reaction: Almandine + Muscovite = Annite + 2Sillimanite + Quartz for which: A V r e a c t = + 1 9 ' 6 6 cc/n>ole. AS r e a c t = +11.45 cals. A H r e a c t " " 4 ' 7 K c a l s « Gradient of reaction curve = AS/AV = +2437 bars/1000 Equilibrium temperature at 1 bar = -410°K The steep gradient for the reaction curve i s consistent with other expsimentally determined equilibria (Thompson 1955, p. 74). The position of the curve in P - T space seems to be unreasonable. Lack of reliable enthalpy data for almandine may be the cause. 116 3) Calculation of the reaction curve fori Paragonite + Quartz => Albite + Sillimanite + HgO The thermodynamic data for these minerals are: Sillimanite* Albite (low)* Quartz* Paragoni te e V v (cc/mol) 49.90 100.07 22.69 132.53 S298 A Gf,298 (cal deg/gfw) (cals/gfw) -583600 H, 22.97 50.20 9.88 63.7 24.498 31.208 -883988 -204646 -1328500 AH f,298 (cals/gfw) -618600 -937100 -217600 -1420200 * - Robie and Walbaum 1968 1 - Chatterjee 1970 The equilibrium curves in Pf^ u^^ - T space were calculated for X J J ^ Q = 1«0» 0.8, 0.5, 0.3, and 0.1 using the method described by-Fisher and Zen (1971), and the tables for the fugacity of water (Burnham et a l . 1969). From these a plot for P J J ^ O " ^ W A S T A * C E N ^ O R a total pressure of 4000 bars (Fig. 21). The effect of muscovite solid solution on this reaction was calculated using the relationships ^ f i n a l = ^ i n i t i a l + ' l 0 g paragonite derived by Gordon (1970"» P« 97-99). Curves have been calculated for paragoni t - 0.9, 0.7, 0.5, 0,3, and 0.1, 117 Acknowledgements This study was made p o s s i b l e by the means' of grants from the N a t i o n a l Research C o u n c i l (grant nos. A-4222) and the G e o l o g i c a l Survey of Canada (grant nos. EMR 2550), held by Dr. H.J.Greenwood, and U n i v e r s i t y of B r i t i s h Columbia Grad-uate F e l l o w s h i p s , 1970-72. The w r i t e r i s indebted to Dr. H.J.Greenwood who supervised t h i s t h e s i s and whose guidance, i n t e r e s t and suggestions are hereby g r a t e f u l l y acknowledged. S p e c i a l thanks are due to Dr. R.B.Campbell of the G e o l o g i c a l Survey of Canada, who suggested the area of study and f r e q u -ented the author with the geology of the r e g i o n . The Geolog-i c a l Survey o-f Canada i s a l s o to be thanked f o r supplying 50 bulk reck analyses and numerous t h i n s e c t i o n s . E l e c t r o n microprobe analyses were conducted at the .Univ e r s i t y of Washington under the h e l p f u l d i r e c t i o n of Dr. E.W.Evans. Mrs. L . L e i t z k i n d l y prepared the microprobe data f o r the computer. F i n a l l y I would l i k e to give s p e c i a l thanks to my w i f e , Helen, who not only was a most e x c e p t i o n a l f i e l d a s s i s t a n t , but a l s o weathered the t r i a l s and t r i b u l a t i o n s of being a graduate student's w i f e . 118 REFERENCES Atherton, M.P., 1968. The v a r i a t i o n i n garnet, b i o t i t e , and c h l o r i t e composition i n medium grade p e l i t i c rocks from the Dalradian, Scotland, with particular reference to the zonation i n garnet. Contr. Miner. P e t r o l . 18, 347-71. Bancroft, G.M., Maddock, A.G., & Burns, R.G., 1967. Application of the Mossbauer effect to s i l i c a t e mineralogy-1. Iron s i l i c a t e s of known c r y s t a l structure. Geochim. Cosmochim. Acta, 31, 2219-46. Braddick, H.J.J., 1954. Physics of Experimental method, Wiley and Sons, New York. Burnham, C.W., Hollaway, N.F., & Davis, N.F. Thermodynamic properties of water to 1000°C and 10,000 bars, G.S.A. spec, paper Number 132. Cameron, W.E. & Ashworth, J.R., 1972. F i b r o l i t e and i t s relationship to s i l l i m a n i t e . Nature Phys. S c i . Vol. 235 Feb. p.134-6. Campbell, K.V. & Campbell, R.B., 1969. Quesnel Lake map-area, B r i t i s h Columbia. Report of A c t i v i t i e s , Pt. A, A p r i l to October, 1969; Geol. Surv. Can., Paper 70-1, p.32-5. Campbell, K.V., 1971. Metamorphic petrology and structural geology of the Crooked Lake area, Cariboo Mountains, B r i t i s h Columbia. Unpublished Ph.D. thesis, University of Washington, Seattle, U.S.A. Campbell, R.B., 1961. Quesnel Lake, west h a l f , B r i t i s h Columbia. Geol. Serv. Canada, Map 3-1961. 1963. Quesnel Lake, east h a l f , B r i t i s h Columbia. Geol Surv. Canada, Map 1-1963. 1970. Structural and metamorphic transitions from infrastructure to suprastructure, Cariboo Mountains, B r i t i s h Columbia. Geol. Assoc. Canada, Special paper nos. 6, 67-72. Carmichael, D.M., 1970. Intersecting isograds i n the Whetstone Lake Area, Ontario. Jour. P e t r o l . 11, 147-81. Chatterjee, N.D., 1970. Synthesis and upper s t a b i l i t y of paragonite. Contr. Min. and Pe t r o l , v. 27 pp. 244-57. Crawford, M.L., 1966. Composition of plagioclase and associated minerals i n some schists from Vermont, U.S.A., and South Westland, New Zealand, with inference about the p e r i s t e r i t e solvus. Contr. Miner. P e t r o l . 13, 269-94. Deer, W.A., Howie, R.A., & Zussman, J . , 1962. Rock forming minerals: Vol. 1. Ortho- and ring s i l i c a t e s , p. 153. New York: John Wiley & Sons. i , 1962. Rock forming minerals: Vol. 3. Sheet s i l i c a t e s , p. 61. New York: John Wiley & Sons. Douglas, R.J.W., Gabrielse, H., Wheeler, J.O., Scott, D.F., & Belyea, H.R., 1969. Geology of Western Canada, i n Geology and Economic Minerals of Canada, R.J.W. Douglas Editor, 366-488. Eugster, H.P. & Wones, D.R., 1962. S t a b i l i t y of the ferruginous b i o t i t e annite. Jour. Pet. v. 3 part 1, pp. 82-129. Evans, B.W., & Guidotti, C.V., 1966. The sillimanite-potash feldspar isograd i n Western Maine, U.S.A. Be i t r . Mineral. Petrogr. 12, 25-62. Fisher, J.R. & E-An Zen, 1971. Thermodynamic calculations from hydro-thermal phase equilibrium data and free energy of H^ O. A.J.Sc. v. 270, pp. 297-314. Fyfe, W.S., Turner, F.J. & Verhoogen, J . , 1958. Metamorphic reactions and metamorphic facies. Mem. 73, G.S.A. pp. 28-34. 120 Fyson, W.K., 1970. Structural relations i n metamorphic rocks, Shuswap Lake area, B r i t i s h Columbia. G.S.C. Spec, paper Number 6. pp. 107-22. Gordon, T.M., 1968. Some silicate-carbonate phase relations i n tL^ O - CO^ mixtures. Unpubl. Ph.D. thesis Princeton Univ. pp. 97-99. Gordon, T.M., & Greenwood, H.J., 1971. The s t a b i l i t y of grossularite i n H20 - C0 2 mixtures. Am. Min. 56, p. 1674-88. Greenwood, H.J., 1961. The system NaAlSi20g-H20-argon: t o t a l pressure and water pressure i n metamorphism. Jour. Geophys. Res. 66, 3923-46. , 1968. Matrix methods and the phase rule i n petrology, XXIII Int. Geol. Congr. (Prague), 6. pp. 267-79. , 1971. Mass transport of heat i n metamorphism. Metamorphism i n the Canadian C o r d i l l e r a , Geol. Soc. Canada, Cordilleran Section, Abstracts, p. 12. Guidotti, C.V. 1968. Prograde muscovite pseudomorphs after s t a u r o l i t e i n the Rangeley - Oquossoc areas, Maine. Am. Min. 53. p. 1368-76. , 1970. The mineralogy and petrology of the t r a n s i t i o n from the lower to upper s i l l i m a n i t e zone i n the Oquossoc area, Maine. Jour. Pe t r o l . 11, 277-336. Hayama, Y., 1959. Some considerations on the colour of b i o t i t e and i t s r e l a t i o n to metamorphism. Jour. Geol. Soc. Japan, 65, 21-30. Hey, M.H., 1954. A new review of the chl o r i t e s . Mineralog. Mag. 30, 277-92. Holland, S.S., 1954. Yanks Peak - Roundtop Mountain area, Cariboo D i s t r i c t , B r i t i s h Columbia. B.C. Dept. Mines, B u l l . 34. H o l l i s t e r , L.S., 1966. Garnet zoning: an interpretation based on the Rayleigh fractionation model. Science, 154, no. 3757, 1647-51. Hsu, L.C., 1968. Selected phase relationships i n the system A l -Mn - Fe - Si — 0 — H: a model for garnet e q u i l i b r i a . Jour. Pet. v. 9, part 1, pp. 40-83. Hyndman, D.W., 1968. Mid-Mesozoic multiphase folding along the border of the Shuswap metamorphic complex. G.S.A. B u l l . v. 79, pp. 575-88. Lambert, R.St.J., 1959. The mineralogy and metamorphism of the Moine schists of the Morar and Knoydart d i s t r i c t s of Inverness-shire. Trans. R. Soc. Edinb. 63, 553-88. Larsen, E.S., 1948. Batholith of Southern C a l i f o r n i a , G.S.A. Mem. 29. Leelanandum, C«» Chemical mineralogy of hornblendes and b i o t i t e s from charnockitic rocks of Kondapalli, India. Jour. Pet. v. 11, part 3, pp. 475-505. Mansy, J.L., & Campbell, R.B., 1969. Stratigraphy and structure of the Black Stuart synclinorium, Quesnel Lake map-area, B r i t i s h Columbia. i n Report of A c t i v i t i e s , Pt. A, A p r i l to October, 1969; Geol. Surv. Can., Paper 70-1, p. 38-41. Mather, J.D., 1970. The b i o t i t e isograd and the lower greenschist facies i n the Dalradian rocks of Scotland. Jour. P e t r o l . 11, p. 253-75. Misener, D.J., 1972. Cation d i f f u s i o n i n Fe-Mg Olivine at elevated pressures and temperatures. Am. Geop. Union. Abstr. 53 Annual Meeting. Obradovich, J.D., & Peterman, Z.E., 1968. Geochronology of the Belt Series, Montana. Can. Jour. Earth S c i . 5, p. 737-47. O r v i l l e , P.M. & Greenwood, H.J., 1965. Determination of AH of reaction from experimental pressure-temperature curves. Am. Jour. Sci. v. 263, pp. 678-83. Ramsay, J.G., 1962. The geometry and mechanics of formation of " s i m i l a r " type folds. Jour. Geol. 70, p. 309-27. Reesor, J.E., 1970. Some aspects of s t r u c t u r a l evolution and regional setting i n part of the Shuswap metamorphic complex, G.S.C. Spec. Paper Number 6, pp. 73-86. Richardson, S.W., G i l b e r t , M.C., & B e l l , P.M., 1969. Experimental determination of kyanite-andalusite and andalusite-sillimanite e q u i l i b r i a ; the aluminum s i l i c a t e t r i p l e point. Am. Jour. Sci. v. 267, pp. 259-72. Robie, R.A. & Waldbaum, D.R., 1968. Thermodynamic properties of minerals and related substances at 298.15°K and 1 atmosphere pressure and at higher temperatures. Geol. Surv. B u l l . 1259. Roddick, J.A., Wheeler, J.O., Gabrielse, H., & Souther, J.C., 1967. Age and nature of the Canadian part of the Circum P a c i f i c Orogenic Belt. Tectomorph. 4, p. 319-37. Ross, J.V., 1968. Structural relations at the eastern margin of the Shuswap Complex, near Revelstoke, southeastern B r i t i s h Columbia. Can. Jour. Earth S c i . 5, 831-49. Rucklidge, J . , & Gasparrini, E.L., 1969. Specification of computer program for processing electron microprobe a n a l y t i c a l data. EMPADR VII. Dept. Geol., Univ. of Toronto, Toronto, Ontario. Ryan, B.D., & Blenkinsop, J . , 1971. Geology and geochronology of the Hellroaring Creek stock, B r i t i s h Columbia. Can. Jour. Earth S c i . 8, p. 85-95. Simpkin, T., 1966. Zoned ol i v i n e s and the cooling history of a p j _ c r i t i c s i l l i n Scotland. Abstr. G.S.A. meeting San Francisco, Spec. Paper Number 101. Skinner, B.J., 1956. Physical properties of the end-members of the garnet group. Am. Min. v. 41, pp. 428-36. Smith, J .V . , 1968. The c r y s t a l structure of s t a u r o l i t e . Am. Min. 53, 1139-55. Sutherland Brown, A., 1957. Geology of the Antler Greek area, Cariboo d i s t r i c t , B r i t i s h Columbia. B.C. Dept. Mines, B u l l . 38. , 1963. Geology of the Cariboo River area, B r i t i s h Columbia. B.C. Dept. Mines, B e l l . 47. Thompson, J.D., 1955. The thermodynamic basis of the mineral concept. Am. Jour. S c i . v. 253, pp. 65-103. Thompson, J.B., 1957. The graphical analysis of mineral assemblages i n p e l i t i c s c h i s t s . Am. Min. 42, 842-58. Tipper, H.W., & Campbell, R.B., 1970. Geology and mineral exploration potential of the Quesnel Trough, B r i t i s h Columbia. CIM Trans., LXXIII, p. 174-9. Turnock, A.C. & Eugster, H.P., 1962. Fe-Al oxides: phase relationships below 1,000°C. Jour. Pet. v. 3, pp. 533-65. wheeler, J .O . , 1970. Summary and discussion - Structure of the Southern Canadian C o r d i l l e r a , G.S.A. Spec. Paper, Number 6, p. 155-66. Wones, D.E. & Gi l b e r t , M.C., 1969. The fayalite-magnetite-quartz assemblage between 600° and 800°C. Am. Jour. S c i . v. 261-k, Schairer Vol. pp. 480-9. Zen, E-An, 1963. Components, phases, and c r i t e r i a of chemical e q u i l i b r i a i n rocks. Am. Jour. S c i . 261, p. 929-42. , 1969. The s t a b i l i t y relations of the polymorphs of aluminum s i l i c a t e : a survey and some comments. Am Jour. Sci. v. 267, pp. 297=310. 124 PLATES 1. Flattened F l sim i l a r folds i n a sequence of interbedded p h y l l i t e s and quartzites. B i o t i t e zone. Scale - 1'. 2. Fanned axial-plane cleavage i n F l syncline. B i o t i t e zone. Scale - 10'. 3. F l isoclines i n quartzose s c h i s t . S i l l i m a n i t e Zone 1. 4. Refolded F l isoclines i n a sequence of micaceous quartzites and mica schists. S i l l i m a n i t e Zone 2. Scale - 6". 5. F l isoclines i n a sequence of micaceous quartzites and mica schists. S i l l i m a n i t e Zone 3. 6. 'Flow folds'. S i l l i m a n i t e Zone 3. 125 PLATES (cont), 7. R e l i c t F l folds and s t r a i n - s l i p cleavage i n a f i r s t generation garnet which has been surrounded by a thin i d i o b l a s t i c rim of a second generation garnet. Staurollte/kyanite zone. Plate edge = 3mm. 8. Replacement of s t a u r o l i t e by pseudomorphic muscovite and fibrous s i l l i m a n i t e . Note interfingering of the pseudomorphic muscovite with external s c h i s t o s i t y . S i l l i m a n i t e Zone 1. Plate edge = 3mm. 9. Alteration of kyanite to c r y s t a l l i n e s i l l i m a n i t e . S i l l i m a n i t e Zone 2. Plate edge = 1mm. 10. Replacement of garnet by pseudomorphic muscovite and fibrous s i l l i m a n i t e . S i l l i m a n i t e Zone 1. Plate edge = 3mm. 11. Second generation garnets enclosed by s t a u r o l i t e (s) which has been p a r t i a l l y replaced by pseudomorphic muscovite and s i l l i m a n i t e . Staurolite has grown e p i t a x i a l l y with kyanite (k). S i l l i m a n i t e Zone 1. Plate edge = 3mm. 12. F i r s t generation garnet with r e l i c t F l axial-plane s c h i s t o s i t y rimmed by 'blocky' second generation garnet (G). An isolated second gener-ation garnet has been enclosed i n a s t a u r o l i t e c r y s t a l (S). Both the garnet and s t a u r o l i t e have been p a r t i a l l y replaced by pseudomorphic muscovite and s i l l i m a n i t e . S i l l i m a n i t e Zone 1. Plate edge = 2mm. 126 PLATES (cont). 13. Small i d i o b l a s t i c second generation garnets. B i o t i t e has been p a r t i -a l l y replaced by f i b r o l i t e i n the lower part of the plate. S i l l i m a n i t e Zone 3. Plate edge = 4mm. 14. S i l l i m a n i t e augen i n mica schist. Note small remnants of s t a u r o l i t e i n the centre of some of the augen. .Sillimanite Zone 3. 15. General view of the Penfold Creek area looking north-west. The pyramid-shaped mountain i n the background consists of unmetamor-phosed f l a t - l y i n g sediments of the Cariboo Group within the Isaac Lake Synclinorium. The s i l l i m a n i t e isograd l i e s just below the white limestone c l i f f s i n the foreground. 16. Poorly f o l i a t e d granite gneiss from the low staurolite/kyanite grade. Augen composed of microcline-quartz intergrowths. Scale - 6". 17. Mafic inclusions i n an equigranular granite gneiss. The f o l i a t i o n of the gneiss i s p a r a l l e l to the length of the inclusions. 18. Well f o l i a t e d granite gneiss from high staurolite/kyanite zone. F3 j o i n t s (pencil l i e s along one of them) are subparallel to the major faults which cross-cut the gneiss. Note one of the F3 fracture cleavages at 20° to the major s c h i s t o s i t y . MAP 4 GEOLOGY OF THE QUESNEL LAKE GNEISS 121° OO* LEGEND 0 L Scale 2 I PLEISTOCENE AND RECENT g l a c i a l deposits and recent a l l u v l u a TERTIARY AND QUATERNARY ba s a l t i c breccia and tu f f UPPER TRIASSIC AND LOWER JURASSIC mainly v o l c a n i c - c l a s t i c rocks UPPER TRIASSIC (?) black p h y l l i t e , schist, gray quartzite MISSISSIPPIAN ( i n part) greenstone, amphibolite geological contact f a u l t ""-.^ F l f o l d axes F l lineations F3 f o l d axes F3 fracture cleavage F l s c h i s t o s i t y or *>, compositional layering bulk rock analyses (sample locations) (defined) -— "* ~" >» _ (approximate) Isograds c h l o r i t e and lower miles b i o t i t e garnet kyanite/staurolite C R O S S - S E C T I O N X © \ PROTEROZOIC Kaza Group - feldspathic quartzites, s c h i s t s , g r i t , minor limestone Kaza Croup - microcline-bearlng schists Kaza Group - c a l c - s i l i c a t e skarn Quesnel Lake Gneiss - gra n i t i c gneiss A + 10,000-1 5000H sea level-\ -5000 J tu < \ I -5 N h-CC < B i T 53° OO'r LEGEND UPPER TRIASSIC AND LOWER JURASSIC mainly volcanic-clastic rocks UPPER TRIASSIC (?) black p h y l l i t e , schist, gray quartzite MISSISSIPPIAN (in part) greenstone, amphibolite, a r g i l l i t e , quartzite, limestone, conglomerate LOWER CAMBRIAN AND PROTEROZOIC PROTEROZOIC Cariboo Group (from oldest to youngest) -Cunningham Limestone, Yankee Belle Fm, Yanks Peak Quartzite, Midas Pm, Mural Fm, a r g i l l i t e unit (unnamed) Isaac Formation - p h y l l i t e , shale, quartzite, limestone Kaza Croup Quesnel Lake Gneiss - granitic gneiss geological contact ~ - — (defined) (approximate) fault isograds a n t i c l i n o r i a l axis synclinorial axis F l and F2 fold axes F3 fold axes bedding or compositional 52°30' GEOLOGY BY R.B.CAMPBELL AND C.J .N.FLETCHER MAP 3 121 °30' DIAGRAMMATIC CROSS-SECTION sea level -10,000-" 120*35* 1 2 0 ° 30' KAZA UNITS L E G E N D ISAAC UNITS SYMBOLS A Light green subarkose Inter layered gray p h y l l i t e ^ ^ and feldspathic quartzite Fink quartzite, minor p h y l l i t e Graphitic limestone Interlayered gray quartzite and graphitic p h y l l i t e Amphibolite Interlayered mica schist and micaceous quartzite 52° 45' -I MAP 1 120° 3 5 ' 1 2 0 ° 30 

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