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Polyphase deformation and metamorphism in the western Cariboo Mountains near Ogden Park, British Columbia Lewis, Peter D. 1987

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POLYPHASE DEFORMATION AND METAMORPHISM IN THE WESTERN CARIBOO MOUNTAINS NEAR OGDEN PEAK, BRITISH COLUMBIA °y PETER D. LEWIS B.S. (Stanford University) A THESIS SUBMITTED IN PARTIAL F U L F I L M E N T O F THE REQUIREMENTS FOR THE D E G R E E OF MASTER OF SCIENCE in THE F A C U L T Y OF G R A D U A T E STUDIES Department of Geological Sciences We accept this thesis as conforming to the required standard THE UNIVERSITY OF BRITISH COLUMBIA September, 1987 ® Peter D. Lewis, 1987 In presenting this thesis in partial fulfilment of the requirements for an advanced degree at The University of British Columbia, I agree that the Library shall make it freely available for reference and study. I further agree that permission for extensive copying of this thesis for scholarly purposes may be granted by the Head of my Department or by his or her representatives. It is understood that copying or publication of this thesis for financial gain shall not be allowed without my written permission. Department of Geological Sciences The University of British Columbia 2075 Wesbrook Place Vancouver, Canada V6T 1W5 Date: September, 1987 ABSTRACT The boundary between the Omineca Belt and the Intermontane Belt in Central British Columbia represents the suture between autochthonous North America (Barkerville Terrane) and several allochthonous terranes accreted from the west In the Quesnel Lake region allochthonous sedimentary and volcanic rocks of Quesnellia Terrane, accreted in the Jurassic, are in sharp tectonic contact with underlying siliclastic and carbonate metasedimentary rocks. The Ogden Peak study area is located 10 km east of and structurally below this suture zone and is thus well situated for observing deformational styles within the autochthonous package. Rocks exposed near Ogden Peak comprise the Hadrynian(?) to Paleozoic(?) Snowshoe Group and local diabasic intrusions. These rocks record a deformational history involving four phases of folding ( D x - D 4 ) and later brittle faulting (D 5). Earliest recognizable structures consist of recumbant isoclinal folds with a well-developed transposed foliation. This foliation is tightly folded about northwest trending, southwest verging second phase structures. Northwest trending third phase structures and northeast trending fourth phase structures occur as both crenulations and open buckles. Southeast dipping faults cut all earlier structures with tens of meters of normal offset Phase 1 and Phase 2 fold styles are compatible with a flattened buckle fold mechanism of formation, associated with elevated temperatures and reduced viscosity contrasts across layering. Later fold styles are controlled by higher viscosity contrasts and detachment along layering. Al l phases of deformation are dominated by semi-brittle mechanisms of dislocation slip and glide, mechanical twinning, and microcracking. Temperature activated diffusional creep is only locally active and does not contribute appreciably to total ii strain. The mineral assemblage garnet- staurolite-kyanite defines a metamorphic peak late in D 2 . Metamorphic temperatures of approximately 530° C at 6.0 kb have been determined using garnet/biotite geothermometry. Extensive retrograde metamorphism spans D 3 and D„, overprinting prograde assemblages and providing evidence for abundant fluids late in deformation. Late phase 1 diabase dykes locally intruded an area to the southeast of Ogden Peak. Major and trace element analyses of samples from these intrusions suggest a calc-alkaline, volcanic arc affinity. i i i i TABLE OF CONTENTS Abstract ii List of figures .-. vi List of plates viii Acknowledgements .. ix 1. INTRODUCTION 1 1.1. Location 3 1.2. Previous work 3 1.3. Purpose . 5 2. STRATIGRAPHY 6 2.1. Introduction .. 6 2.2. Snowshoe Group rocks 8 2.3. Intrusive rocks 18 2.4. Summary and discussion 22 3. MAJOR AND TRACE ELEMENT GEOCHEMISTRY 24 3.1. Introduction 24 3.2. Major element geochemistry 26 3.3. Trace element geochemistry 29 3.4. Summary and discussion 32 4. STRUCTURE 34 4.1. Introduction 34 4.2. Mesoscopic structures, metasedimentary rocks 35 4.2.1. -D, - D 4 Folding 35 4.2.2. D 5 Faulting 58 4.3. Mesoscopic structures, intrusive rocks 59 4.3.1. Western dyke set 59 4.3.2. Eastern dyke set 64 4.4. Microscopic structures ...» 68 4.4.1. Healed micTocracks 68 4.4.2. Extension fractures 69 4.4.3. Kink banding 72 4.4.4. Pressure solution features 76 4.4.5. Deformation twins 76 4.4.6. Recrystallization textures 76 4.5. Summary and discussion 81 5. METAMORPHISM 87 5.1. Introduction 87 5.2. Regional metamorphism Mj 89 5.2.1. Pelitic lithologies 89 5.2.2. Amphibole schist lithologies 101 5.2.3. Carbonate lithologies 103 iv 5.3. Retrograde metamorphism M 2 104 5.4. Geothermometry 109 5.5. Summary and discussion 113 6. REGIONAL CORRELATIONS AND TECTONIC INTERPRETATIONS 115 6.1. Regional correlations 115 6.1.1. Stratigraphy 116 6.1.2. Structure 117 6.1.3. Metamorphism 119 6.2. Tectonic implications 122 7. REFERENCES 127 8. APPENDICES 133 8.1. Appendix 1. Major and trace element analyses 133 8.2. Appendix 2: Axial dimensions of elongate pebbles 135 8.3. Appendix 3: Mineral abbreviations 137 8.4. Appendix 4: Electron microprobe analyses 138 v List of Figures 1. Major tectonic belts of the Canadian Cordillera 2 2. Regional geology of the Quesnel Lake region 4 3. Map of the Ogden Peak area 7 4. Structural succession of the Ogden Peak area 9 5. Stratigraphic contact between unit 3 and unit 4 13 6. Lithologic variation in subunit 5a 15 7. Meta-diabase dyke intruding Snowshoe Group metasedimentary rocks 19 8. Distribution and surface traces of meta-diabase dykes 20 9. Alkaline/subalkaline discriminant diagram .27 10. A F M diagram 28 11. Ti , Zr, Y ternary discriminant diagram 30 12. Ti , Zr binary discriminant diagram 31 13. Phase 1 isoclinal folds 36 14. Phase 1 isoclinally folded quartz vein 38 15. Stereographic plot of D i data 40 16. Simple shear of S0 along Sj surfaces 41 17. Stereographic plot of D 2 data 43 18. D 2 folds in quartzite and marble 44 19. t versus a plots of phase 2 folds 45 20. D 2 crenulations of biotite and muscovite 47 21. S 2 muscovite growth 48 22. Quartz vein oriented along phase 2 axial plane 49 23. Stereographic plot of D 3 data 51 24. Stereographic plot of D„ data 53 25. t versus a plot of phase 4 folds 54 26. Mesoscopic phase 4 buckle 55 27. Disharmonic phase 4 crenulation 56 28; Structural relationships of intrusive diabase dykes to country rocks 60 29. Diabase dyke of western set 61 30. Phase 2 fold of inuusive contact 62 31. Phase 1 isoclinal fold outlined by compositional layering within dyke 63 32. Contact between eastern set diabase dyke and metasedimentary country rock 65 33. Folds in metasedimentary rock along intrusive contact 66 34. Healed microcracks in quartz 70 35. Extension fractures in feldspar porphyroclasts 71 36. Kink bands in biotite porphyroblast 73 37. Subgrain developement in quartz 74 38. Deformation lamellae in quartz 75 39. Deformation twins in calcite 77 40. Quartz recrystallization texture 78 41. Recrystallized quartz mylonite 80 42. Summary of fold style and geometry, phase 1 - phase 4 82 43. Incremental strain ellipsoids showing phase 2 progressive deformation 84 44. Muscovite defining phase 1 and phase 2 foliations 91 45. Two stages of chlorite growth 92 46. Kyanite porphyroblasts growing across phase 2 crenulations 95 47. Sericitized staurolite growing across phase 2 crenulations 96 48. Helycitic inclusion trails in garnet 97 vi 49. A F M sketches showing stable mineral assemblages of pelitic zones 99 50. Chlorite pseudomorphs of garnet flattened into phase four axial planes 106 51. Quartz vein filling fracture along axial plane of mesoscopic phase 2 fold 107 52. P - T diagram applied to pelitic rocks of the Ogden Peak area 110 53. Garnet-biotite temperatures plotted against manganese concentration :. 112 54. Mineral growth/deformation correlation chart 114 55. Plate tectonic model for I M B / O M B boundary 125 vii LIST OF PLATES Air-p1alBS--aTe—included—in~me^ack—pocket- • 1. PLATE I: Bedrock geology of the Ogden Peak area. 2. PLATE II: Cross section through the Ogden Peak area. 3. PLATE III: D x and D 2 structural elements, mineral lineations. 4. PLATE IV: D 3 and D 4 structural elements, faults. viii A C K N O W L E D G E M E N T S I am indebted to my thesis supervisor, Dr. J.V. Ross, for his continued support and advice throughout the course of this study. Valuable assistance in the laboratory aspects of this project was provided by C. Hickson, S. Horsky, P. Michaels, and M . Piranian. Innumerable discussions with S. Garwin have contributed greatly to this project, and J.K. Russell, D. McMullin, D. Murphy, and J. Fillipone have helped clarify many issues. S. Sonnad provided able assistance in the field, and I am grateful to S. Taite for both moral support and assistance in the final preparation of this project Funding for field and laboratory expenses was provided by NSERC grant number 58-2134 to Dr. J.V. Ross. ix 1. INTRODUCTION The Canadian Cordillera is composed of five major tectonic belts (Figure 1), and numerous discrete tectono-stratigraphic terranes. In the area of interest to this study in central British Columbia, near Quesnel Lake, the boundary between the Omineca Belt (OMB) and the Intermontane Belt (1MB) represents the contact of autochthonous rocks of North America with allochthonous rocks of numerous displaced terranes accreted from the west The Omineca Belt here (Barkerville Terrane) is dominated by siliciclastic and carbonate rocks of the late Proterozoic(?) to Paleozoic(?) Snowshoe Group (Campbell and Campbell, 1970). Allochthonous Intermontane Belt rocks include the late Paleozoic(?) Crooked Amphibolite (Struik, 1986), which has. been tentatively correlated with the Slide Mountain Terrane, and structurally overlying Triassic and Jurassic sedimentary and volcanic rocks assigned to the Quesnellia Terrane. Extensive deformation and metamorphism of Jurassic age in both packages of rocks are believed to be related to the accretion of Quesnellia and more westerly terranes onto North America (Monger et al., 1982). The suture zone between the two belts is characterized by an extensive mylonite zone that appears to be the result of high angle convergence (Ross et al., 1985). Figure 2 is a generalized geologic map showing the spatial distribution of lithologic units in the Quesnel Lake region. 1 INTRODUCTION / 2 Figure 1. Major tectonic belts of the Canadian Cordillera and location of the Quesnel Lake area. INTRODUCTION / 3 1.1. LOCATION The Ogden Peak study area is located between the East Arm and North Arm of Quesnel Lake, on the western flank of the Cariboo Mountains in south-central British Columbia (Figure 2). An area of 50 km 2 was mapped at 1:16,000 scale during the summer of 1985 at elevations between 1,500 m and 2,500 m. The study area is completely within the Omineca Belt, approximately 10 km east of the Intermontane Belt/Omineca Belt boundary. It is easily accessible by helicopter; lower elevations may be reached by boat from Quesnel Lake. Interconnected ridges allow easy travel through the area above timberline. 1.2. PREVIOUS WORK Western Omineca Belt rocks in the Quesnel Lake region were first mapped north of the study area in the vicinity of Barkerville and Keithly Creek (Uglow, 1922; Lang, 1938; Holland, 1954; Sutherland Brown, 1957). Holland (1954) was first to assign these rocks to the Late Precambrian to Cambrian Snowshoe Formation of the Cariboo Group, with a type locality near Yanks Peak. Sutherland Brown (1957) described a 300 m thick section of Snowshoe Formation as forming the uppermost unit of the Cariboo Group, unconformably overlain by the Late Paleozoic Slide Mountain Group. Sutherland Brown outlined a structural history involving at least two phases of folding followed by brittle deformation. Based on work in the Quesnel Lake area, Campbell and Campbell (1970) suggested that the Snowshoe Formation represents a western fades of the late Proterozoic Kaza Group. This revised Sutherland Brown's interpretations which placed the Kaza Group at lower stratigraphic levels than the Snowshoe Formation. Recent mapping at 1:50,000 scale by Struik (1982, 1983) resulted in redefinition of the INTRODUCTION / 4 UJ m LU z < 1-z o 2 GC HI UJ m < o UJ z i o L. Trl»»»lc -E. Jurassic L. Palaozoic(?) Quesnellia Terrane NICOLA GROUP r£ j^ volcanlclastics QUESNEL RIVER GROUP { ,'| black phylllta, argilllto minor limsstona. volcanic* CROOKED AMPHIBOLITE E 3 metavolcanica. phyllrts L. Osvonian -E. Mlaaiaalpplan Hadrynian -Palaozoic North A m e r i c a ^ ['-. | granitic gneiaa SNOWSHOE GROUP • quartzofeldspathic schist snd gneiss, minor carbonsta Figure 2. Regional geology of the Quesnel Lake area and location of map area. Late Paleozoic to Jurassic volcanic and phyllitic rocks of Quesnellia Terrane are tectonically juxtaposed against metasedimentary rocks and granitic gneiss of North America. INTRODUCTION / 5 Snowshoe Formation as the Snowshoe Group, separated from the more easterly Kaza Group along the Pleasant Valley Fault Detailed structural analyses of Omineca Belt rocks of the Quesnel Lake region by Getsinger (1985), Montgomery (1985), Fillipone (1985), and Elsby (1985) have delineated four widespread phases of folding in Snowshoe Group rocks and intrusive granitic gneisses. Ross et al. (1985) have related the distribution and geometry of some of these structures to the Jurassic accretion of Quesnellia onto North America. 1.3. P U R P O S E The present study is one of a series of projects involving UBC graduate students over the last decade, with the overall objective of elucidating the tectonic processes related to convergence at the Intermontane Belt/Omineca Belt boundary. The Ogden Peak area is situated structurally well below the boundary, and thus spatially removed from the zone of high strain concentrated there. This location allows observation of pre-accretion structures which have been only moderately overprinted by later deformation related to convergence. Overall objectives of this study are to provide a detailed structural and stratigraphic analysis of the Ogden Peak area in order to: 1. ) correlate Snowshoe Group stratigraphy with that described across the North Arm of Quesnel Lake by Getsinger (1985) and that described across the East Arm of Quesnel Lake by Montgomery (1985), and to 2. ) relate the structural and metamorphic developement of the area to the local tectonic history, specifically those events involving the obduction of Quesnellia onto North America. 2. STRATIGRAPHY 2.1. INTRODUCTION Rocks exposed in the Ogden Peak area consist of metamorphosed siliciclastic and carbonate rocks locally intruded by mafic dykes. Metasedimentary rocks in the area were first mapped by Campbell (1963) as part of the Hadrynian Snowshoe Formation, a western fades of the Kaza Group. Regional mapping by Struik (1981, 1982) culminated in the publication of a 1:50,000 scale geologic map of the Quesnel Lake region in which he demonstrated that the Snowshoe Formation does not readily correlate with strata described elsewhere as part of the Cariboo Group (Holland, 1954; Sutherland Brown, 1957). This has led to the elevation of the Snowshoe Formation to Group status (Struik, 1984). Struik separates it from the more easterly Kaza group along the Pleasant Valley Fault (Struik, 1984; 1986a). The Snowshoe Group is included (by Struik) in the Barkerville Terrane, a subdivision of the Selkirk Terrane. Within the Selkirk Terrane, it is considered laterally equivalent to the Horsethief Creek Group as mapped by Pell (1984) east of Wells Gray Park (Struik, 1986b). Nine distinct lithologic units were mapped within the Snowshoe Group in the Ogden Peak area. Figure 3 shows the distribution of the major units across the map area. Complete transposition of bedding during polyphase deformation has obscured contacts between units, making estimates of original sedimentary thicknesses impossible. Sedimentary structures, when present, are in upright positions. However, due to the transposed nature of these rocks, one cannot assume that the complete stratigraphic succession is upright Snowshoe Group and intrusive lithologies are described below in stratigraphic 6 STRATIGRAPHY / 7 units 5a, Sb. 5c quartzite, schist 1111111 ni • 1 1 1 1 1 1 1 1 1 marble unit 3 / quartzose schist : - r-a» ~ - i amphibole schist unit 1c unit 1a marble, amphibolite quartzite, schist Ogdan Peak Figure 3. Map of Ogden Peak area showing distribution of major stratigraphic units. See figure 4 for lithologic types. STRATIGRAPHY / 8 order. Stratigraphic interpretations are then presented, along with possible correlations with surrounding areas. 2.2. SNOWSHOE GROUP ROCKS A total present thickness of approximately 3,000 m of Snowshoe Group rocks is exposed in the Ogden Peak area. The stratigraphy is characterized by interlayered pelitic and quartzose schists, micaceous quartzites, amphibole schists, and marbles. A generalized stratigraphic column of these lithologies is illustrated in figure 4. Unit 1 minimum thickness 1,500 m This series of interlayered pelitic to quartzose schists and feldspathic quartzites are the structurally lowest rocks of the study area. The unit can be divided into three parts, la, lb, and lc on the basis of lithological variation. The base of unit 1 is not observed in the mapped area. Subunit la minimum thickness 1,000 m Subunit l a consists of equal proportions of micaceous feldspathic quartzite and pelitic schists alternating in layers ranging in thickness from 5 m to 10 m. The quartzite layers crop out as prominent ledges between areas of recessive topography underlain by schist The grey weathering quartzite has well developed stratification 2 cm to 15 cm thick and a parallel foliation defined by elongate quartz grains and planar alignment of micaceous minerals. Thin discontinuous marble lenses crop out within the quartzite layers in the western map area, but are absent to the east Up to 20% mica and feldspar minerals may be contained in the quartzite. Grain size STRATIGRAPHY / 9 Unit Thickness Lithology 5c 5b 5a 4 3 2 1c 1b 1a ^> 1000 m Quartzose schist, quartzite i ' i ' i ' i ' i 1 3 S r I I I 50 - 300 m 0 - 100 m 50 - 300 m 150 m 100 m 400 m > 1000 m Micaceous quartzite, interlayered schist Micaceous quartzite, schist, marble Impure marble, interlayered schist, amphibolite lenses Quartzose schist, thin quartzite layers Amphibolite schist Thinly bedded quartzite, interlayered schist 50 - 150 m Quartzose schist, marble, amphibolite Quartzose schist, micaceous quartzite. marble Figure 4. Structural succession of Snowshoe Group lithologies of the Ogden Peak area. STRATIGRAPHY / 10 varies from fine to medium, and large garnet porphyroblasts are typical of micaceous layers. Schist intervals have a distinctive platy appearance with yellow weathering. Thin quartzite beds (0.3 cm to 0.5 cm) locally constitute up to 25% of the succession. Mineralogy is dominated by medium to coarse grained muscovite and biotite in roughly equal proportions, with quartz, feldspar, and metamorphic garnet making up the remaining rock volume. Both lithologic types contain abundant quartz veins crosscutting compositional layering. The upper contact of this subunit is sharp and is defined by the appearance of abundant marble and amphibole schist layers indicative of subunit lb. Subunit lb variable thickness 50 m to 150 m Immediately overlying the interlayered schists and quartzites of subunit la is a sequence of mixed lithologies including impure marble, micaceous quartzite, amphibole schist, and micaceous schist This poorly exposed sequence is usually recognizable by yellow weathering of the marble component The base is sometimes marked by a friable weakly foliated feldspathic quartzite up to 10 m thick. Overlying this is a platy to massive, tan weathering impure marble up to 50 m thick. Micaceous partings and occasional quartz or amphibole rich horizons define compositional layering in this medium grained carbonate. The marble marks the top of subunit lb in the western part of the study area, while to the east, additional interbedded amphibole schists, garnet bearing pelites, and marbles continues for up to 100 m. Amphibole schist layers here range in character from fine grained hornblende-feldspar gneissic schists to very coarse grained strongly foliated schists composed exclusively of hornblende. The marble in this sequence is similar to the marble lower in subunit lb. The upper contact of STRATIGRAPHY / 11 subunit lb is gradational and marked by the gradual disappearance of carbonate and amphibole schist layers. Subunit l c thickness 400 m The uppermost division of unit 1 is a thick interval of micaceous quartzite with up to 30% pelitic schists. Alternation between these lithologies occurs over as little as 4 cm to 5 cm in the eastern part of the study area, increasing to 5 m in the west Medium grained schist is dark green and silver on fresh surfaces, weathering to a dark grey. It is more resistant than the schists lower in the section, forming steep slopes and cliffs. Highly garnetiferous intervals are common, and chlorite is locally abundant Grey weathering quartzite has well developed stratification 2 cm to 8 cm thick. Thin mica partings define a foliation and commonly contain large euhedral garnets. Unit 2 variable thickness 50 m to 200 m The top of subunit lc grades abruptly into a slope forming amphibole schist/gneiss sequence. Theoe schists weather typically to dark brown, but local areas of high carbonate concentrations weather yellow. Mineralogy consists primarily of hornblende, plagioclase, and carbonate, although the relative abundance and grain size of these components is highly variable. Plagioclase/hornblende ratios are highest in fine grained rocks, and color banded differentiated layering is slightly developed. Coarse grained rocks are homogeneous, schistose, and composed of 90% to 100% euhedral hornblende. Marble lenses up to 2.0 m thick occur at all levels within unit 2, and carbonate content increases markedly in the western study area. With these lenses, 0.1 STRATIGRAPHY / 12 cm to 1.0 cm thick amphibole/mica selvages alternate with much thicker carbonate layers. Quartz veins common to unit 1 are absent in the amphibole schist The upper contact with unit 3 is marked by the disappearence of amphibole schist and a return to pelitic lithologies. Unit 3 variable thickness 50 m to 300 m Above the amphibole schist is a resistant pelitic to quartzose schist of variable thickness. The bottom 3 m is a very fine grained fissile schist lacking visible metamorphic porphyroblasts. The remainder of the section consists of grey weathering, medium to coarse grained schist with porphyroblasts of garnet and.staurolite. Locally, up to 20% of this unit may be quartzite layers, and occasional discontinuous marble and amphibole rich horizons are intercalated within the schists in the western part of the study area. Unit 3 may contain up to 10% by volume of thin quartz veins cutting across foliation at moderate to high angles. A sharp transition to continuous marble marks the upper contact with unit 4 (Figure 5). Unit 4 variable thickness 50 m to 300 m In marked contrast to underlying lithologies unit 4 consists primarily of marble, making it an important marker horizon for regional stratigraphic correlations. Lesser amounts of pelitic schist, quartzite, and amphibole schist form discontinuous lenses within unit 4 in all parts of the study area. Schist and quartzite layers average 1 m to 2 m in thickness, and occur at all stratigraphic levels within the marble. These non-carbonates constitute up to 30% of the overall unit Relative abundances of rock types follow no consistent pattern across the area. Characteristic yellow weathering of the carbonate component facilitates easy identification of unit 4, although local areas STRATIGRAPHY / 13 Figure 5. Stratigraphic contact between units 3 and 4 north of peak 7570. Thin marble layer of unit 4 is interbedded with quartzite and schist of unit 3. STRATIGRAPHY / 14 may weather to dark brown or black. Lensoid, weakly foliated amphibolite boudins up to 5 m long and with very sharp contacts to the surrounding rock are common to the lower part of unit 4 in the central study area. Compositional layering in surrounding rocks is parallel to boudin contacts. In hand sample, the marble is coarse grained and ranges in composition from pure calcite to quartzose marble with discrete layers of mica or amphibole. Compositionally distinct layers accompany a parallel penetrative fissility. Unit 4 schists have an unusually high proportion of calcite, and weather to darker yellow than similar rocks elsewhere in the area. The upper contact, characterized by interfingering of marble with siliciclastic lithologies of unit 5, is mapped where rocks contain less than 10% marble. Unit 5 minimum thickness 1,400 m Forming the structurally highest rocks of the Ogden Peak succession is this series of alternating schists and quartzites. Unit 5 is divided into three subunits on the basis of relative abundances of constituent lithologies. Subunit 5a variable thickness 0 m to 100 m Overlying the marble of unit 4 in most of the Ogden Peak area is a section of mixed pelitic schists, quartzites, marble, and amphibole schists (Figure 6). This section is thickest just south of Ogden Peak, and thins rapidly to the east, disappearing completely over 2 km. Field evidence demonstrating either a structural or a stratigraphic origin for this thinning is inconclusive. Quartzite and pelitic schist in roughly equal abundances are the dominant lithologies in subunit 5a, with lesser amounts of amphibole schist and marble. These lithologic types alternate in layers 1 m STRATIGRAPHY / IS Figure 6. Lithologic variation in base of unit 5a at location 591a. Stratigraphy consists of pelitic schist, quartzite, carbonate, and amphibole schist alternating in layers 1 m to 3 m thick. STRATIGRAPHY / 16 to 3 m thick. Subunit 5a is overall slope forming, with quartzite forming resistant ledges. Less resistant schists and carbonates weather to tan or rust, and quartzites are typically dark grey in outcrop. Amphibole schist forms both discrete layers and discontinuous boudins within the marble. Near the top of the unit a series of repetitive marble and micaceous schist layers alternate over 5 cm to 10 cm. The top of subunit 5a is highly gradational, and marked as a gradual transition to thickly bedded quartzite and lesser pelitic schist lithologies. Subunit 5b variable thickness 50 m to 300 m Directly overlying the mixed lithologies of subunit 5a in the western study area and the marble of unit 4 in the eastern study area is a well exposed feldspathic quartzite with lesser interbedded pelitic schist The quartzite has conspicuously thick stratification of 10 cm to 1 m, weathers to a grey color, and crops out in rounded, cliff forming exposures. Visible mineralogy consists of medium to coarse grained quartz, feldspar, muscovite, and porphyroblastic garnet along micaceous partings. Pelitic schist is confined to layers 0.5 m to 2 m thick, and constitutes up to 20% of subunit 5b. Within the schist, muscovite, biotite, and quartz are medium to coarse grained, and garnet and staurolite porphyroblasts are abundant The upper contact of subunit 5b occurs as a zone of increasing pelitic schist content The contact is arbitrarily chosen as the level at which rocks are composed of 30% to 40% schist and 60% to 70% quartzite. Subunit 5c minimum thickness 1,000^  m The structurally highest rocks of the Ogden Peak area comprise a repetitive series of feldspathic quartzites and micaceous schists. Lithologically these rocks are STRATIGRAPHY / 17 similar to those of subunit 5b, with only relative abundances of each constituent changing. In addition, alternation of lithologies is on a much larger scale, with schist layers 20 m to 40 m thick common. Just west of Ogden Peak a prominent 2 m to 3 m thick marker horizon of quartz-feldspar pebble conglomerate is exposed; this layer is not found in the eastern part of the study area. The strain history of this conglomerate is analyzed in a later section. STRATIGRAPHY / 18 2.3. INTRUSIVE ROCKS In a limited portion of the study area to the southeast of Ogden Peak, foliated amphibolite bodies occur within Snowshoe Group lithologies of subunit 5b. The resistant and dark weathering nature of these mafic rocks makes them readily visible both in the field and on aerial photographs (Figure 7). They are almost completely limited to an area of 1.5 km 2 where they are clearly divisible into a western set presently oriented sub-parallel to compositional layering and an eastern set sharply truncating layering (Figure 8). In both, areas, they account for up to 20% of surface exposures. Both sets are medium grained and composed of at least 50% hornblende, with biotite, muscovite, garnet, calcite, and plagioclase making up the remainder of the rock. A poorly to well developed foliation is defined by parallel alignment of mica and hornblende and weak mineral segregations. Based on structural relations between the amphibolites and the metasedimentary rocks, together with their composition, it is inferred that they originated as diabasic intrusions. In the eastern set, dyke contacts are planar on a large scale, although local irregularites give them up to 1 meter of surface relief. The dykes are arranged in an en-echelon map pattern, and typical surface exposures of a single body measure 3 m to 20 m wide by 50 m to 150 m long. Mineral zonations occur adjacent to and parallel to the sharply defined contacts, often as a thin garnet rich band or a selvage of more coarsely crystalline hornblende. Angles between regional compositional layering and dyke contacts range from 50° to 90° . The dykes of the western set have irregular contacts, which are subparallel to compositional layering in the Snowshoe Group rocks. Thicknesses of individual dykes range from 1 m to 20 m; diffuse contacts and well developed foliation often make STRATIGRAPHY / 19 Figure 7. En-echelon diabase dykes intruding quartzite of unit 5b near location 459. Dike contacts truncate compositional layering at high angles. STRATIGRAPHY / 20 121 45' Figure 8. Map showing distribution and surface traces of diabase dykes. Dikes of western set are subparallel to compositional layering, dykes of eastern set truncate layering at high angles. STRATIGRAPHY / 21 the smaller bodies difficult to separate from the country rock. Foliation and compositional layering defined by mineral segregations are much more strongly developed here than in the eastern set No other intrusive rocks are found in significant volumes in the Ogden Peak area. 2.4. SUMMARY AND DISCUSSION STRATIGRAPHY / 22 Metasedimentary siliciclastic and carbonate rocks of the Ogden Peak area are assigned to the Downey and Ramos successions of the Snowshoe Group by Struik (1984). This assignment is based on similarities to rocks of the Downey succession rocks near Spanish Lake (Struik, 1983) and the Barkerville Gold Belt (Struik, 1982). Regional correlations of the Snowshoe Group with the Horsethief Creek Group near Kootenay Lake (Read, 1976) and the Eagle Bay Formation near Adams Lake (Schiarizza and Preto, 1984) have led to its inclusion (by Struik) in the Selkirk Terrane, a group of rocks extending to the U.S. border inferred to have similar depositional and deformational histories. . Continuous markers within the Ogden Peak succession allow stratigraphic correlations with structural successions described in adjacent areas. Across the North Arm of Quesnel Lake, Getsinger (1985) described a sequence of which the lowest units appear to be equivalent to the highest levels in the Ogden Peak succession. This correlation necessitates a fault between the two areas downdropping the north side, the existence of which is corroborated by offset in metamorphic zones (Campbell, 1970). Stratigraphy south of Quesnel Lake described by Montgomery (1985) is similar to the lowest segment of the Ogden Peak succession. Rocks in the Niagara Peak region to the east lie structurally above those in the study area (Garwin, 1987). Ages of Snowshoe Group rocks are poorly constrained. Microfossil fragments from the upper Snowshoe Group laterally equivalent to the Downey Succession have been identified as Early Paleozoic fauna (Struik, 1986). Upper age limits of Snowshoe Group rocks are approximately 340 Ma, provided by the emplacement of the Quesnel Lake Gneiss (Mortensen et al., 1987), and Early Permian, based on conodonts collected STRATIGRAPHY / 23 from the upper Snowshoe Group (Orchard and Struik, 1985). The quartzites, marbles, and schists of the Ogden Peak succession are interpreted to be metamorphic products of a series of clastic sediments, carbonates, and volcanic tuffs. Struik (1986) suggests a basalt flow protolith for some of the amphibole schists in the Downey Succession, however, the high carbonate content and preservation of relic bedding in the amphibolte schist of the study area are more indicative of a volcanic sediment or dolomitic siltstone protolith. This sequence of lithologies is consistent with the continental shelf depositional setting proposed for rocks in this area. 3. MAJOR AND TRACE ELEMENT GEOCHEMISTRY 3.1. INTRODUCTION Geochemical analyses were completed for a group of intrusive rocks from the Ogden Peak area. Samples were collected from six meta-diabase dykes representing both eastern and western sets at various distances from intrusive contacts. The study was conducted using the UBC Department of Geology X-ray fluorescence spectrometer with pressed powder pellets samples. Analysis procedures and test results are summarized in Appendix 1, and sample locations are shown on Plate I. The chemical compositions of these rocks are used to characterize the igneous suites they are associated with. Specifically, the trace element concentrations of the dykes are used to infer their original tectonic environments using discriminant diagrams proposed by Pearce and Cahn (1973). A similar approach has been taken for previous workers in the Quesnel Lake region for the Quesnel Lake Gneiss and ML Perseus Gneiss (Montgomery, 1985) and the Takla Volcanics of Quesnellia Terrane (Bloodgood, 1987). Samples from Ogden Peak were analyzed in the present study in order to: 1. characterize dyke rock composition, and determine the amount of variation between intrusive bodies, 2. relate trace element abundances in dykes to tectonic setting, based on published discriminant diagrams, and 3. compare the chemistry of these dykes to that of the Quesnel Lake Gneiss, the Boss Mountain Gneiss, and the Takla Volcanics to the south. Petrogenetic relations between the intrusive bodies are established on the basis of this 24 MAJOR AND TRACE ELEMENT GEOCHEMISTRY / 25 comparison. MAJOR A N D T R A C E ELEMENT GEOCHEMISTRY / 26 3.2. M A J O R E L E M E N T GEOCHEMISTRY The lack of relic igneous textures in the diabase dykes near Ogden Peak makes classification schemes based solely on whole rock geochemistry rather than i petrographic methods necessary. The dykes are classified using the classification scheme proposed by Irvine and Baragar (1971). In this system, an initial plot of alkalies ( N a 2 0 + K 2 0 ) versus Si0 2 allows differentiation between alkaline and sub-alkaline fields. Al l dyke samples plot within the subalkaline field (Figure 9). An alternative dividing line determined from Hawaiian basalts by MacDonald (1968) places one sample near the boundary between fields. Within the sub-alkaline field, assignment to either tholeiitic or calc-alkaline suites is based on enrichment trends best illustrated on ternary A F M plots. Tholeiitic suites tend to display strong iron enrichment followed by late stage alkali enrichment, whereas calc-alkaline suites have continuous and moderate iron enrichment Al l dyke samples except one plot in Irvine and Baragar's calc-alkaline field (Figure 10). Assignment to the calc-alkaline field is tenuous, as the dykes do not represent a continuous igneous suite and actual enrichment trends are not observed. Further differentiation between calc-alkaline and tholeiitic fields is made in the following section on the basis of trace element abundances. MAJOR A N D T R A C E E L E M E N T GEOCHEMISTRY / 27 Figure 9. Alkali-silica plot of diabase dyke samples with alkaline and sub-alkaline fields delineated. Dividing line of MacDonald (1986) is derived from Hawaiian tholeiitic and alkaline rocks, division of Irvine and Baragar (1971) is from analyses of alkaline, tholeiitic, and • calc-alkaline suites. MAJOR AND TRACE ELEMENT GEOCHEMISTRY / 28 Figure 10. AFM plot of dyke samples showing tholeiitic and calc-alkaline suites. (After Irvine and Baragar, 1971). MAJOR A N D T R A C E ELEMENT GEOCHEMISTRY / 29 3.3. TRACE E L E M E N T GEOCHEMISTRY Trace element abundances in igneous rocks have been shown to vary systematically with tectonic environment (Pearce and Cann, 1973; Floyd and Winchester, 1975; Pearce, 1984; Pearce et al., 1984; see Erdman, 1985 for review). Binary and ternary discriminant diagrams showing this variation are constructed using trace element analyses of unaltered rocks from known tectonic settings. As Ti, Zr, and Y have demonstrated low mobility at moderate metamorphic grades (Pearce and Cann, 1973; Murphy and Hynes, 1986), diagrams using these elemtents as discriminators are used in classifying the altered diabase dykes. A ternary plot of Ti, Zr, and Y proposed by Pearce and Cann (1973) is used to differentiate between arc related low-potassium tholeiites and calc-alkaline basalts, within plate basalts, and ocean floor basalts (Figure 11). A l l analyses fall within the calc-alkaline basalt field, but several are in the region overlapped by the ocean-floor basalt and low-potassium tholeiite fields. An alternative plot with just Ti and Zr (after Pearce and Cann, 1973) shows five analyses in the calc-alkaline basalt field, and one analysis just inside the ocean-floor basalt field (Figure 12). MAJOR AND TRACE ELEMENT GEOCHEMISTRY / 30 Ti/100 Figure 11. Discriminant diagram using Ti/100, Zr, and Y x 3- Fields for different tectonic settings are defined by Pearce and Cann (1973). Calc-alkaline basalts plot in fields C and D, within plate basalts in field A, ocean floor basalts in field C, and low-potassium tholeiites in fields B and C. MAJOR A N D T R A C E E L E M E N T GEOCHEMISTRY / 31 15.000-10,000-Ti (ppm) 5,000-i 50 100 150 Zr (ppm) 200 Figure 12. Binary discriminant diagram using Ti and Zr. (After Pearce and Carm, 1973). Calc-alkaline basalts plot in fields A and B, low-potassium tholeiites in fields B and C, and ocean floor basalts in fields B and D. MAJOR A N D T R A C E ELEMENT GEOCHEMISTRY / 32 3.4. S U M M A R Y A N D DISCUSSION The chemical compositions of the Ogden Peak dykes suggest a calc-alkaline affinity. These results are not conclusive however and several limitations are recognized: 1. Amphibolite grade metamorphism has altered the mineralogical character of these rocks, and may have affected major and trace element abundances through fluid transfer or diffusional processes. The relatively small variation in elemental concentrations between samples from different dykes and between samples from different positions within dykes suggests this effect is small. Trace element mobility studies indicating stable concentrations of Ti, Zr, and Y even at • amphibolite grade metamorphism (Murphy and Hynes, 1986) also suggest the trace element analyses are valid. 2. Compositional fields on discriminant diagrams are empirical and are limited by available data from previous studies. 3. Composition fields are constructed from analyses of basic volcanic rocks, and applications to intrusive dykes have not been demonstrated. 4. Petrographic evidence confirming any interpretation is lacking due to overprinting of original mineralogical characteristics during amphibolite grade metamorphism. Petrogenetic relationships between the Ogden Peak dykes and other igneous rocks of the Quesnel Lake region are unclear. Studies of the Quesnel Lake Gneiss and M t Perseus Gneiss show calc-alkaline affinities and effects of crustal assimilation (Montgomery, 1985). Although bulk compositions of the Ogden Peak samples were much more mafic than either the Quesnel Lake or Mount Perseus Gneiss, trace element abundances closely resemble those of the M t Perseus Gneiss. Geochemical studies of the Takla Volcanics in the Quesnel Lake area (Bloodgood, 1987) yielded equivocal results, with characteristics of both tholeiitic and calc-alkaline suites. Structural MAJOR A N D T R A C E E L E M E N T GEOCHEMISTRY / 33 evidence suggests a time of emplacement similar to that of the Quesnel Lake and Mount Perseus Gneiss (see section 4.2). Forthcoming zircon U - P b geochronometry on samples collected from the Ogden Peak area (Mortensen et a l , 1987) may clarify these genetic relationships. 4. STRUCTURE 4.1. INTRODUCTION Four phases of folding (Di-D«) and one later phase of localized brittle deformation ( D 5 ) are recognized within the Ogden Peak area. Assignment of mesoscopic structures to specific deformational episodes was made in the field on. the basis of superposition of axial surfaces, refolding relationships, and relative orientations of structural elements. Snowshoe Group rocks record all phases of deformation; deformation within intrusive rocks is limited to developement of foliation and minor folds and warps of intrusive contacts. For simplicity, axial surfaces and axial directions are referred to as A 1 - A 4 and FJ - F A respectively, with penetrative foliations and lineations designated S1-S4 and L i - L 4 . The distribution and orientation of phase 1 and phase 2 structures in the Ogden Peak area are presented on Plate II. Phase 3, Phase 4, and late brittle features are included on Plate III. Microcracks, deformation twins, kink bands, and recrystallization textures are observed on the microscopic scale within deformed crystals. Petrographic examination of deformation fabrics allows correlation of these microstructures to the major deformational episodes in the area. Structural features in and around the metadiabase dykes suggest emplacement during D x deformation, and are discussed below separately from Dj to D 5 in metasedimentary rocks. Kinematic significance of and relationships between all structures are addressed in a concluding section. 34 STRUCTURE / 35 4.2. MESOSCOPIC STRUCTURES, METASEDIMENTARY ROCKS 4.2.1. D i - D 4 Folding Phase 1 ( D O The earliest deformational event recognized in the study area is manifested in a regional penetrative foliation and complete transposition of bedding. Associated folding occurs as very tight to isoclinal reclined or recumbent structures in quartzite and marble lithologies and as isolated floating hinges outlined by dismembered quartz seams in schists. The absence of repetition of lithologic units and mappable isoclinal fold closures indicates that phase 1 folds are limited to outcrop scale in the Ogden Peak area. However, the intensity of deformation and the abundance of isoclinal folds of bedding indicate moderate to considerable structural thickening of original stratigraphy accompanied this early folding event Evidence that compositional layering in the area is transposed includes isoclinal folding observed on outcrop scale and isolated hinges of dismembered isoclinal folds defined by quartz stringers. . Phase 1 foliation is outlined by thin monomineralic segregations and parallel alignment of micas, amphiboles, and elongate quartz and calcite grains. No mineral lineation clearly associated with D j occurs in the study area, although F i axes are often parallel to D 2 mineral lineations. D j fold limbs are typically highly attenuated and accompanied by considerable thickening of hinge regions (Figure 13). Extreme limb attenuation in schistose units has led to dismembered folds with floating hinge regions outlined by quartz segregations (Figure 14). Stereographic projections of structural elements related to phase 1 folds are STRUCTURE / 36 Figure 13. a. ) Tight phase 1 fold of compositional layering in unit l c quartzite. b. ) Mesoscopic phase 1 isocline outlined by interlayered marble and schist, near location 591a. STRUCTURE / 37 b. STRUCTURE / 38 Figure 14. Isoclinal hinge of dismembered phase 1 fold of quartz vein in pelitic schist, location 177. STRUCTURE / 39 illustrated in figure 15. Axial surfaces are defined by a penetrative foliation (Si) which is everywhere parallel or subparallel to compositional layering (So), and dips gently to the northeast or northwest This surface will hereafter be referred to as S 0 / S i . F x fold axes plunge gently to the northwest throughout the area. Vergence of D! structures is equivocal, with fold sets asymmetric both east and west Offset of lithology along Si planes in an area where S 0 and Si are at moderate angles indicates east directed shear (Figure 16). Dismembered folded quartz veins, emplaced early in D i , indicate moderate amounts of east directed shear along compositional layering. Quartz veins formed throughout D i , most abundantly in schistose lithologies. Those formed earliest are folded and transposed parallel to S]/S 0 , while those formed later in deformation maintain moderate angles to foliation. Quartz fibres, when preserved, are oriented perpendicular to vein walls. STRUCTURE / 40 Figure 15. Equal area projections of phase 1 foliation and fold axis data. Contour intervals for S0/S, plot are 1%. 3%, 5%. 7%, and 9%. 1.4% counting area determined using method outlined by Kamb (1959). STRUCTURE / 41 Figure 16. East directed offset of S 0 surfaces along Si shear planes in unit 5a micaceous quartzite, near peak 7570. S T R U C T U R E / 42 Phase 2 (D2) S 0 /Si surfaces in all lithologies are folded about upright axial planes during phase 2 deformation. D 2 is the most pronounced phase of deformation in the Ogden Peak area, with well developed structures documented in all locations. Fold scale ranges from microscopic crenulations to a map scale antiform-synform set in the eastern study area outlined by the unit 2 amphibole schist All mesoscopic phase 2 folds are outlined by S 0 /Si surfaces. Orientation data for D 2 structural elements are plotted on stereograms in figure 17. Fold axes are presently coaxial to Dr fold axes, plunging shallowly to the northwest Axial surfaces dip moderately to steeply to the northeast Angles measured between S 2 axial surfaces and compositional layering (S 0 /Si) range from 10° to 80° (See plate II for spatial distribution of S 0 /Si - S 2 angles). Interlimb angles of D 2 folds are normal to tight and hinge regions are subrounded to subangular. Marbles typically have tight angular hinges, while quartzose rocks have the most open, rounded geometries (Figure 18). Hinge zones are always thickened relative to limbs, giving most folds a "similar" shape. Graphs plotting limb dip against orthogonal layer thickness in different lithologies (after Ramsay, 1967) show clasr, 2 behavior in carbonates and schists and class lc forms in quartzite. (Figure 19). D 2 folds are strongly asymmetric in almost all locations, with a westward sense of vergence on all first order folds. Second and third order folds show expected vergence changes with position on. larger structures. Diabasic intrusions have only weak phase 2 crenulations and no S 2 foliation, and their intrusive contacts show evidence of D 2 deformation. D 2 folds have tighter forms and are more abundant in country rock surrounding the intrusive units than elsewhere in the area, possibly compensating for lesser amounts of strain in the more STRUCTURE / 43 Figure 17. Equal area projections of phase 2. axial plane and fold axis data. Figure 18. a) Normal D 2 fold in unit 5b quartzite, near location 612. b) Tight D 2 fold in unit 4 marble, location 293. Compositional layering is outlined by micaceous partings S T R U C T U R E / 45 class lb Figure 19. t' versus «< plot (Ramsay, 1967) of layer thickness as a function of limb dip. Marble and schist lithologies approach class 2 forms, quartzite displays class Ic behavior. STRUCTURE / 46 competent dykes. D 2 crenulations are ubiquitous in schistose lithologies of the study area. In carbonates and quartzites, crenulation fabrics are developed only in pelitic horizons. Two types of axial planar cleavage are associated with the crenulations. Most commonly, crenulations have tightened such that one limb is parallel to the axial plane (Figure 20). This type is primarily found in rocks of very micaceous composition. Elsewhere, axial planar growth of muscovite across layering outlines a penetrative cleavage in crenulated rocks (Figure 21). Microscopic folds and cleavage are not observed in quartzite and carbonate units, although mesoscopic folds are common. Linear features associated with D 2 include mineral orientation lineations and quartz rod lineations. Both of these lineations are penetrative and always occur parallel to F 2 . Quartz rods are round or oval in cross-section, and up to 1 m long with length-diameter ratios often exceeding 10:1. Mineral orientation lineations are defined by amphibole, mica or deformed quartz and calcite grains. A pebble conglomerate in subunit 5c has numerous quartz and feldspar clasts which are possibly deformed and now define a strong linear fabric parallel to 1 .^ Commonly associated with D 2 folds are planar quartz veins formed along axial planes, especially in quartzose and schistose lithologies (Figure 22). These are composed of microcrystalline milky quartz and rare microcrystalline muscovite, and measure up to 20 cm thick and several meters in length. STRUCTURE / 47 Figure 20. D 2 crenulations defined by muscovite and biotite in unit 3 schist, in sample 104. STRUCTURE / 48 Figure 21. Axial planar muscovite growth defining S2 cleavage. STRUCTURE / 50 Phase 3 (D3) D 3 is the least well developed phase of deformation in outcrop of the Ogden Peak area. Phase 3 upright open warps with wavelengths of several kilometers are mapped in the field primarily by changes in bedding orientation. These folds are responsible for the overall map pattern and the great-circle stereonet distribution of S o / S j surfaces. Mesoscopic folds are warps with "parallel" forms and open profiles. Upright crenulations in D 3 orientations are found only in hinge zones of the map scale structures. Orientations of structural elements related to D 3 folds are displayed in figure 2 3 . Fold axes are coaxial to those of D j and D 2 , plunging shallowly to the northwest with steeply dipping to vertical axial planes. Symmetric fold profiles preclude sense of vergence determinations. A spaced joint set is oriented perpendicular to F 3 axes in outcrop. These fractures are folded during D « but are unaffected by phase 2 folds, suggesting association with D 3 . No offset is observed along these fractures, and their time of formation with respect to other D 3 features is inconclusive. STRUCTURE / 51 Figure 23. Equal area projections of D3 axial plane and fold axis data. STRUCTURE / 52 Phase 4 (D4) D , crenulations and open warps refold earlier structures at nearly right angles. Both crenulations and buckles occur in all parts of the study area; these structures are most strongly developed in schists surrounding the intrusive rocks. Kink bands coaxial with F 4 crenulations are believed to be related to this same deformational event F 4 fold axes have shallow to moderate northeast plunges and steeply southeast or northwest dipping axial surfaces (Figure 24), allowing easy differentiation from earlier folds. Axial surfaces are nearly at right angles to S 0 / S i . Type III "dome and basin" interference patterns (Ramsay, 1967) result from the superposition of D 4 on D 2 and D 3 structures. Hinge angles are open to normal on folds with wavelengths of less than 0.5 m and open on folds with greater wavelengths. In all lithologies, D 4 folds have "parallel" class lb to lc profiles with little or no thickening in hinges, as illustrated by plots of layer thickness against limb dip (Figure 25). Slickenside lineations may occur along bedding surfaces at high angles to F 4 fold axes. These are only found on limbs of mesoscopic F 4 buckle folds and probably indicate layer-parallel movement during folding. Slip along layer boundaries may create mesoscopic disharmonic folds in areas (Figure 26). D 4 folds are symmetric with the exception of parasitic structures which show appropriate vergence changes with position on larger structures. Where D 4 crenulations are most strongly developed, a spaced crenulation cleavage is defined by the alignment of mica along fold limbs. Mica growth did not accompany the formation of this cleavage. In micaceous rocks, disharmonic crenulations often form as a result of slip along foliation planes during deformation (Figure 27). A joint set developed during or after D 4 dips steeply to the northeast and STRUCTURE / 53 Figure 24. Equal area projections of D 4 axial plane and fold axis data. STRUCTURE / 54 f t' versus^, plot (Ramsay, 1967) of layer thickness as a function of limb dips for D 4 folds. A l l lithologies display class lb to lc behavior. STRUCTURE / 55 Figure 26. Layer parallel detachment along mesoscopic phase 4 buckle in unit 1 micaceous quartzites, near location 391. STRUCTURE / 56 Figure 27. D 4 disharmonic crenulation showing effects of layer-parallel slip during deformation, sample 523. STRUCTURE / 57 southwest In the field, these joints are perpendicular to F 4 fold axes. STRUCTURE / 58 4.2.2. D5 Faulting High angle faults offset a l l lithologies in the Ogden Peak area and are designated Ds. They strike north-northeast and dip steeply to the southeast, and are easily mapped in the field based on offset of stratigraphy and the substantial topographic expression associated with them. The most continuous faults occur with a regular spacing of 0.5 km to 1 km. Plate IV shows the locations and orientations of the major faults in the field area. Net slips on several faults were calculated using slip direction indicators and stratigraphic offsets measured in the field. Mullions on fault surfaces indicate a dominant down-dip component of slip, and later slickensides suggest minor strike-slip offset Calculated net slip values of 20 m to 50 m are obtained assuming normal offset S T R U C T U R E / 59 4.3. MESOSCOPIC STRUCTURES, INTRUSIVE ROCKS Meta-diabase dykes in the study area are divided into a western set sub-parallel to compositional layering in the country rock and an eastern set which truncates layering at high angles. Each set. has unique deformation features which will be discussed separately below. Figure 2 8 is an interpretive cross section illustrating overall structural relationships of each set with the surrounding rocks. 4.3.1. Western dyke set Dikes of this set truncate S 0 / S i surfaces at low angles and dip shallowly to the north (Figure 2 9 ) . Intrusive contacts may be sharp or diffuse and gradational over 0.5 m, and are deformed by D 2 folds (Figure 3 0 ) . Structures internal to the dykes record a strain history similar to that of the surrounding Snowshoe Group rocks suggesting pre- to syn- Dj emplacement A well developed penetrative foliation, parallel to intrusive contacts, is outlined by mineral grain orientations and weakly developed compositional layering. Associated mineral orientation lineations plunge northwest parallel to Lj in the country rocks. Rare isoclines are outlined by compositional layering, but the intrusive contacts are not likewise folded (Figure 31 ) . Phase 2 folds of compositional layering within the dykes have wavelengths of up to 0.5 m and tight geometries. STRUCTURE / 60 Figure 28. Structural relationships of intrusive diabase dykes to surrounding country rocks. Dikes of western set are sub-parallel to compositional layering and folded during phase 2 deformation. Eastern set dykes are at high angles to compositional layering and intrusive contacts are not visibly deformed. STRUCTURE / 61 Figure 29. Diabase dyke of western set with sharp contact subparallel to compositional layering. Foliation within dykes is parallel to that in metasedimentary country rocks. STRUCTURE / 62 Figure 30. Phase 2 fold of intrusive contact in dyke of western set near location 591. STRUCTURE / 63 Figure 31. Phase 1 isoclinal fold outlined by compositional layering within dyke of western set, location 591. STRUCTURE / 64 4.3.2. Eastern dyke set Dikes of the eastern set are large tabular bodies striking north-northwest with sub-vertical dips. Intrusive contacts are sharp, planar, and unfolded. Structures in this set of dykes indicate a strain history lacking the earliest features found in the western set of dykes and the Snowshoe Group rocks. Oriented mineral grains and weakly developed compositional segregations define a penetrative foliation parallel to the S o / S i enveloping surfaces in adjacent country rocks and at a high angle to intrusive contacts. Amphibole grains lying in the foliation outline a northwest plunging mineral lineation, parallel to the L 2 mineral lineation in surrounding country rocks. Folds internal to the eastern set of dykes are limited to crenulations of the dominant foliation within the outermost 20 cm to 30 cm. These folds have styles and orientations similar to D 2 stuctures in the adjacent country rock. Metasedimentary rocks adjacent to the dykes contain structural features unique to this location. Figure 32 illustrates the style, orientation, and spatial distribution of these structures along an intrusive contact S 0 /S x surfaces in the country rock are nearly perpendicular to dyke contacts, but may be folded into sub-parallelism with the dykes by folds with axial planes parallel to the intrusive contacts (Figure 33). A weak foliation defined by oriented biotite/chlorite aggregates may form across the S 0 / S i surface in a direction parallel to the dominant foliation in the dykes. A third planar surface defined by leucocratic seams crosscuts both of these foliations in an orientation parallel to the dyke contacts. Although the structure within these seams is obliterated during subsequent metamorphism, their orientations suggest a hydraulic fracture mechanism of formation synchronous with dyke intrusion. The present orientation of dyke contacts and the penetrative foliation within the S T R U C T U R E / 65 Figure 32. Contact between eastern set diabase dyke and metasedimentary country rock at location 651c. Planar surfaces described in text are defined by: 1. S o / S j compositional layering in country rock, 2. Mica foliation crossing S 0 / S i at low angles and parallel to foliation in intrusive rocks, and 3. leucocratic seams parallel to intrusive contacts. STRUCTURE / 66 Figure 33. Folds in metasedimentary rock along intrusive contact, eastern dyke set near location 582. Compositional layering is locally parallel to intrusive contact. STRUCTURE / 67 dykes suggest emplacement and subsequent deformation in a strain field similar in orientation to that associated with the So/Sj foliation. Lack of transposition of intrusive contacts limits the possible time of formation to very late in Di, but prior to the onset of D2. STRUCTURE / 68 4.4. MICROSCOPIC STRUCTURES Textures observed in thin section allow the determination of mechanisms by which deformation was accomplished. Microstructures may be correlated with specific phases of deformation by their orientations, relations to mesoscopic fabrics, and the relative ages of metamorphic grains in which they occur. Earliest formed structures may be overprinted or obliterated by later deformation and annealing. Knowledge of deformation mechanisms operative on a crystalline scale helps verify the temperature/pressure conditions which are inferred to accompany deformation. Microstructures observed in rocks of the Ogden Peak area include healed microcracks, filled microscopic extension fractures, kink bands, deformation twins and lamellae, pressure solution features, and recrystallization textures. 4.4.1. Healed microcracks Microcracks form during deformation as microscopic planes of brittle extension. If these cracks are limited in extent, they may undergo later healing through a diffusional process as outlined by Wanamaker and Evans (1985). Partially healed microcracks may be detected in crystals as planar arrangements of cylindrical or spheroidal pores. These pores represent gas filled inclusions which formed in response to lattice mismatch in the healing crystal. Cylindrical inclusions formed early in the crack healing process may divide into isolated spheroids under temperature induced increased diffusion (Wanamaker and Evans, 1985). Cylindrical pore shapes, together with the tendency for inclusion planes to cross grain boundaries, allow differentiation between healed microcracks and "Boehm" lamellae inclusion planes formed through dislocation slip processes (Fairbairn, 1949). STRUCTURE / 69 In rocks of the Ogden Peak area, healed microcracks commonly occur in quartz grains and occasionally in feldspar porphyroclasts of feldspathic quartzites. Unhealed microcracks filled with foreign crystalline material are common in garnet, hornblende, and feldspar, and are discussed in the following section. Rocks with a greater than 50% mica content contain much higher microcrack densities than do more silicious rocks. Trails of spheroidal to cylindrical inclusions may be continuous through several adjoining quartz grains (Figure 34). Orientation measurements from several samples show a majority of microcrack planes dipping steeply to the northwest, with the remainder randomly oriented. This preferred orientation is always at a high angle to the S 0 / S i foliation, and 50° to 70° from F 2 and F 3 fold axes. 4.4.2. Extension fractures Mineral filled extension fractures may be used to determine local strain axis orientations during deformation. Infilled mineral phases, if not recrystallized, display preferred growth orientations and reflect the stable mineral assemblage at the time of formation. Extension fractures are commonly observed in hornblende, in garnet, and in feldspar porphyroclasls. Garnet porphyroblasts are fractured along planes perpendicular to the So/Sx foliation. Mica growth is common within fractures, and later retrogression appears to nucleate along them. Hornblende within the unit 2 amphibole schist is broken and extended along fractures perpendicular to S 0 / S i . Fractures are infilled with equant quartz grains which are presently randomly oriented but may have been elongate fibres at their time of formation. Feldspar porphyroclasts are dissected by extension fractures infilled with equant STRUCTURE / 70 Figure 34. Healed microcracks in quartz defined by planar arrangements of elongate bubble inclusions, sample 447. Healed cracks are often continuous across several grains of different orientations. STRUCTURE / 71 Figure 35. Extension fractures in feldspar porphyroclasts of sample 188. STRUCTURE / 72 quartz grains. These occur at moderate or high angles to S0/Si and do not continue into adjacent feldspar clasts (Figure 35). Fracture orientations are often sub-parallel to healed microcracks in the feldspar and in surrounding quartz grains. 4.4.3. Kink banding Kink bands in mica and undulose extinction in quartz provide evidence for deformation through dislocation slip processes. These features are most common in sections cut normal to F 4 fold axes. Biotite porphyroblasts have kink band boundaries at high angles to S 0 / S i where they are deformed about phase 4 crenulations (Figure 36). Cleavage planes within kink bands are rotated up to 40° with respect to the surrounding crystal in a direction which brings them closer to the A 4 orientation. Under higher strains grains may fail brittlely along kink band boundaries to form new separate grains. Undulose extinction in quartz occurs either as distinct subgrain regions of slightly differing optical orientation or as a warping of the quartz structure with no distinct subgrain boundaries developed. Rocks with evidence of high D 2 strain are associated with subgrain developement (Figure 37). Subgrains are equant or elongate parallel to kink band boundaries, and optical orientations are rotated less than 5° from adjacent subgrains. Undulose extinction without subgrain developement is characteristic of rocks deformed dominantly by phase 4 folds. Narrow "Boehm" deformation lamellae defined by changes in quartz relief and birefringence are re-oriented across subgrain boundaries (Figure 38). STRUCTURE / 73 Figure 36. Kink bands in biotite porphyroblast deformed about a phase 4 microfold in sample 369. Kink band boundaries are subparallel to or symmetric about axial planes of folds. STRUCTURE / 74 Figure 37. Polygonization and subgrain developement in quartz, sample 463, evidence of diffusion climb deformation mechanism. STRUCTURE / 75 Figure 38. "Boehm" deformation lamellae in quartz, sample 130. Lamellae are re-oriented across kink band boundaries. STRUCTURE / 76 4.4.4. Pressure solution features Sutured grain boundaries, hydraulic fractures, and quartz pressure shadows provide evidence of deformation by pressure solution processes (Elliot 1973). Sutured grain boundaries, common between quartz grains, occur in all orientations and cannot be positively correlated with any specific deformational episode. However, mesoscopic quartz-filled hydraulic fractures, derived probably in part from pressure solution, are correlated with all phases of deformation. Quartz pressure shadows about garnet and staurolite porphyroblasts indicate pressure solution was active during or following metamorphic mineral growth. 4.4.5. Deformation twins Several orientations of twin lamellae are observed in calcite grains of carbonate rocks. Twin lamellae are narrow relative to grain size, and are sharply bounded by planar surfaces extending across the full width of the host grain (Figure 39). They occur at all angles to the S 0 / S i foliation, and form irregularities in grain boundaries. Lamellae comprise between 10% and 40% of individual calcite grains. 4.4.6. Recrystallization textures Recrystallization textures are observed in quartzose lithologies where fine grained neoblasts form along grain boundaries at the expense of larger grains. Scalloped recrystallized grain boundaries have irregularities with constant wavelengths and amplitudes. In one example recrystallization nucleates along boundaries between highly elongate quartz ribbon grains (Figure 40). This type of recrystallization is localized in a quartzose segregation forming the core of a tight microscopic D 4 fold. A quartzose mylonite displays similar textures (Figure 41) where original grains have partially STRUCTURE / 77 Figure 39. Deformation twins in calcite from location 261. Twin lamellae generate irregularities in grain boundaries STRUCTURE / 78 Figure 40. Quartz recrystallization texture in core of tight phase 4 microfold, sample 523. Fine grain neoblasts have nucleated along ribbon grain boundaries. STRUCTURE / 79 recrystallized to small equant shapes with equilibrium grain boundaries. Ribbon grains are still visible in hand sample. STRUCTURE / 80 Figure 41. Recrystallization of ribbon grains in quartz mylonite, leading to equant, equigranular texture, sample 2401. S T R U C T U R E / 81 4.5. S U M M A R Y A N D DISCUSSION A structural sequence in which intensity of folding and ductility of rocks decrease over time is documented by structures in the Ogden Peak area. Structural styles accompanying deformation are strongly influenced by rheologic constraints dependant mainly on lithologic type. Styles and geometries of folds representative of each deformational episode are illustrated in figure 42. Earliest deformation is in the form of a regional transposed foliation and mesoscopic isoclinal folds. Quartz veins initiated in D i are inferred to form by a hydraulic fracture mechanism fed by fluids derived from metamorphic dewatering and pressure solution processes (Norris and Henley, 1976, Etheridge et al., 1984). Increasing metamorphic grades are implied by mica foliation growth and continued dewatering. Sedimentary compositional differences were enhanced by dissolution and solution transfer processes to form the well developed compositional layering (Durney, 1972). Dismembered folds of quartz veins within the transposed layering indicate substantial shear along S 0 / S i surfaces. Diabase dykes which intruded during D i may have been transposed into orientations sub-parallel tc local compositional layering. Later dykes have a foliation parallel to S 0 / S i in surrounding rocks, but contacts are not transposed. This dyke foliation, together with S i mica growth across isoclinal fold hinges, suggests that high shear strains during early folding were followed by a flattening strain later in D i . This may be a consequence of increased loading as higher structural levels were tectonically thickened. Folds formed early in D : were flattened into present isoclinal forms under conditions of high strain and low viscosity contrast Tight class l c to class 2 folds with axial planar mica growth indicate a continuation of high temperature, low viscosity contrast conditions during D 2 folding. STRUCTURE / 82 M E S O S C O P I C F O L D D E V E L O P M E N T Phase 1: SW Phase 2: SW Phase 3: SW Phase 4: NW 10 m Figure 42. Summary of mesoscopic fold style and geometry, phase 1 - phase 4. STRUCTURE / 83 Phase 1 folds were further tightened during this event, modifying them to isoclinal forms. Phase 2 folds initiated as buckle or flexural slip folds controlled by slight viscosity contrasts across layering and were subsequently flattened into their present class lc to class 2 forms (Ramsay, 1962). The extent of flattening is strongly controlled by rock type, leading to tight class 2 folds in marble and schist and open class lc forms in quartzites. Consistant asymmetry of folds indicates non-coaxial strain with a component of west-directed simple shear during deformation. Deformed pebbles in one location and quartz rods throughout the area imply a constrictional component of strain with elongation resolved in the F 2 direction. This strain history can be approximated by a series of incremental strain ellipsoids in which the flattening ( X - Y ) plane rotates away from the maximum compressive stress direction. If quartz rods and elongate clast axes truly represent the regional elongation direction, this rotation would be about the X strain axis, as illustrated in figure 43. Lesser amounts of D 2 strain in meta-diabase dykes are juxtaposed against anomalously high strain in immediately adjacent country rocks. This suggests that the dykes behaved more competently and controlled deformation in surrounding lithologies to allow for strain homogeneity on a large scale. Phase 3 broad buckles indicate a transition following D 2 to lower temperature, higher viscosity contrast conditions of folding. Deformation was probably a continuum through phases 2 and 3, as D 2 folds were refolded coaxially about more upright D 3 axial planes. This refolding is slight and not easily discernible on stereonet projections of phase 2 data. Phase 4 open (class lb) folds imply continued folding at high viscosity contrasts. Detatchment along S o / S j surfaces on all scales indicate a deformation mechanism dominated by flexural slip between buckled layers. As with phase 2 Figure 43. Incremental strain ellipsoids showing possible phase 2 progressive deformational history. Z and Y axes rotate about X axis in non-coaxial deformation. STRUCTURE / 85 structures, lesser amounts of strain in dykes are reflected by high strains in adjacent metasedimentary rocks. Latest deformation occurs as steeply dipping faults with no observed associated folding. Brecciation zones along faults indicate a majority of strain at this time was accommodated by brittle translation. The formation of abundant quartz-filled hydraulic fractures with varying relations to mesoscopic structures provides evidence of ample pore fluid availability during all phases of deformation. Strain was accommodated on a microscopic scale through a combination of ductile and semi-brittle mechanisms. The dominant deformation mechanisms operating appear to vary in different types of minerals. Within a given mineral species, the effects of various mechanisms are dependant on: 1. The constitutive minerals of the overall aggregate, and relative proportions thereof, 2. the relationship of the specimen observed to mesoscopic structures, and 3. the timing of deformation within the specimen. Quartz, mica, and calcite are the major mineralogical constituents of metasedimentary rocks of the study area, and thus accommodated a majority of the strain in most rocki. Quartz deformed through a combination of pressure solution, microcracking, and dislocation glide. Earliest deformation was accommodated by pressure solution and dislocation glide in quartzose rocks, with the addition of microcracking in micaceous rocks. Strain softening through dynamic recovery (polygonization) and dynamic recrystallization along grain boundaries has obscured many of the earliest features. Preservation of bubbles along healed microcracks and "Boehm" lamellae indicate that later intracrystalline diffusion is limited, and suggests recrystallization is accomplished STRUCTURE / 86 through grain boundary diffusion. Later deformation (D 4) in quartz is mainly by dislocation glide and kinking, and is limited to pelitic rocks. Recrystallization overprints phase 4 features only in regions of high D 4 strain, where the small size of neoblasts implies falling temperatures and high differential stresses (Carter, 1976). Thus, it is inferred that temperature induced recovery mechanisms in quartz are unique to D i and D 2 , and only slight amounts of stress induced dynamic recrystallization are associated with D 4 . Muscovite and biotite deform at all times through dislocation glide and kinking. Under higher strains during D 4 , kink band boundaries may fail brittlely to form rotated subgrains. The lack of kink banding or even undulose extinction in mica outlining phase 2 crenulations suggests recrystallized grains have replaced older strained grains. Carbonate minerals have been deformed dominantly by mechanical twinning. The narrow width and sharp boundaries of twin lamellae are typical of low strain, low temperature conditions of deformation (Groshing, 1972; Nicolas and Poirier, 1976). Kink banding associated with dislocation glide would not be expected at the temperature and pressure conditions inferred to have accompanied deformation in the Ogden Peak area (see section 5.4). Deformational episodes are more specifically related to metamorphic mineral growth and regional tectonic events in succeeding sections. 5. METAMORPHISM 5.1. INTRODUCTION The Ogden Peak area is located on the western flank of the northwestern extension of the Shuswap Metamorphic Complex (SMC), which is defined arbitrarily as a metamorphic and plutonic complex bounded by the sillimanite isograd (Reesor, 1970; Brown and Read, 1983). The northwestern SMC is said to comprise late Proterozoic to early Paleozoic paragneiss structurally overlying the Monashee Decollement (Okulitch, 1984). Metamorphic peak conditions in the Quesnel Lake region range from sillimanite grade within the SMC to chlorite grade along the North Arm of Quesnel Lake and immediately adjacent to the I M B / O M B suture. The Ogden Peak area contains garnet to kyanite grade metamorphic rocks situated between low grade rocks to the west and sillimanite grade rocks to the east Metamorphic zone boundaries within and adjacent to the area do not follow lithologic or structural controls. Pelitic rocks of the Ogden Peak area contain indicator minerals typical of a Barrovian-type metamorphic sequence of the middle amphibolite fades. Prograde mineral assemblages have been subject to an extensive retrograde event resulting in the replacement of primary porphyroblasts with chlorite and muscovite. Campbell (1971) mapped two metamorphic zones in the area based on first appearances of garnet and staurolite + kyanite in pelitic rocks. The present study locally redefines Campbell's zone boundaries and furthur subdivides his staurolite + kyanite zone into separate staurolite and staurolite + kyanite zones. Microscopic textures are used to establish the relative timing of mineral growth and deformational events. Metamorphic peak conditions were attained late in D 2 followed by extensive retrograde metamorphism through D 4 related to waning 87 METAMORPHISM / 88 temperatures and abundant fluids. Metamorphic conditions during D 3 are uncertain due to poorly developed microstructures related to this deformational event Microprobe analyses of garnet-biotite pairs allow determination of metamorphic temperatures and pressures. Textures and timing of mineral growth are discussed separately below for each pelitic metamorphic zone and for each non-pelitic rock type. Garnet-biotite geothermometry and retrograde metamorphism are addressed in later sections. Constitutive mineralogy typical of each metamorphic zone and lithology is summarized at the beginning of each section. Abbreviations of mineral names used in this chapter are listed in appendix 3. METAMORPHISM / 89 5.2. REGIONAL METAMORPfflSM M , Amphibolite grade mineral assemblages in the Quesnel Lake area formed during a regional metamorphic event that was synchronous with accretion of Intermontane Belt terranes in mid-Jurassic time (Monger, 1982). Pelitic, carbonate, and amphibolitic lithologies of the Ogden Peak area each have unique mineral assemblages related to this event 5.2.1. Pelitic lithologies Three metamorphic zones are delineated in pelitic lithologies of the Ogden Peak area. These zones are defined by first appearences of garnet, staurolite, and kyanite. Individual staurolite and kyanite occurences appear to be largely controlled by local bulk composition, and these metamorphic zones transect major lithologic units. Metamorphic zone boundaries and sample locations referred to in the text are identified on plate I. METAMORPHISM / 90 Garnet-biotite zone Q z - B t - M u - G t - P l - C h ± To, Sph, Op, Ap, Ru, Zr The garnet-biotite zone is located in the most northwesterly part of the study area, and is defined in pelitic rocks by the presence of garnet and biotite and the absence of staurolite and kyanite. Garnet textures range from idiomorphic poikiloblasts to highly retrograded chlorite knots. Porphyroblasts truncate the S o / S i foliation, which may be slightly deflected around the largest garnets. Planar quartz inclusion trails (Sp are subparallel to the external S0/St foliation. Quartz and chlorite pressure shadows about garnet are elongate parallel to the mineral lineation. Microcracking of garnet along fractures normal to the S„/Si foliation is common. Textural evidence delineates several generations of biotite growth. Fine grained reddish-brown biotite intergrown with earliest muscovite defines the S 0 / S i foliation. Later porphyroblastic biotite crosscuts the foliation and S2 crenulation planes and is kinked by D 4 crenulations. Muscovite growth lasts through all early phases of deformation. Earliest muscovite is present as a medium to fine grained phase outlining the So /S j foliation. Syn-D 2 muscovite defines S2 surfaces and may be crosscut by later randomly oriented muscovite (Figure 44). Sericite is a common retrograde product of garnet Two distinct chlorite textures are present in garnet grade rocks. Coarse unoriented blades of chlorite crosscut the So /S j foliation defined by oriented biotite and muscovite. Later growth includes retrogression of garnet porphyroblasts and mimetic replacement of biotite by chlorite (Figure 45). METAMORPHISM / 91 Figure 44. Muscovite defining S „ / S i and S 2 foliations, sample 193 M E T A M O R P H I S M / 92 Sketch of photomicrogaph showing two stages of chlorite growth in sample 527. Early chlorite is randomly oriented across the So/Si foliation. Later retrograde cholrite is a mimetic replacement of biotite. METAMORPHISM / 93 Staurolite Zone Qz-E-t-Mu-Gt-St-Pl ± Ch, Sph, Op, Ap, Ru, Zr, To The staurolite zone occupies a narrow east-west trending area just south of the garnet + biotite zone. Indicator minerals are staurolite, garnet, and biotite, in the absence of kyanite. Because occurences of staurolite are compositionally restricted to aluminous rocks which are spatially limited, the exact location and orientation of the boundary between staurolite and garnet + biotite zones is unknown. Staurolite forms 2 cm to 4 cm elongate porphyroblasts with long axes randomly oriented in the S0/S] plane. Extensive retrogression of porphyroblasts makes textural relationships to deformational fabrics and other minerals uncertain, but cross sections of partially preserved grains truncate the S 0 / S i foliation. Sericite pseudomorphs of staurolite mantle garnet porphyroblasts. A lack of pressure shadows or preferred alignment of elongate porphyroblasts indicates growth late or post- phase 2 deformation. Subidiomorphic garnet porphyroblasts have poikilitic cores surrounded by inclusion free rims, indicating changing chemical conditions during two stage growth. Inclusions of quartz and ilmenite in cores define a fabric which is weakly crenulated, and is discordant to the external foliation. Porphyroblasts truncate S„/Su and pressure shadows of quartz and muscovite parallel to 12 indicate some garnet growth occured pre- or syn- D 2 . Biotite and muscovite exhibit similar textural relationships to those observed in the garnet + biotite zone. Sericite is an additional phase generated by retrogression of staurolite. Chlorite occurence is limited to replacement of other phases, with no evidence for stability during peak metamorphic conditions. METAMORPHISM / 94 Staurolite + Kyanite Zone Q z - B t - M u - G t - S t - K y - P l ± Kspar, Ch, Sph, Op, Ap, Ru, Zr, To The staurolite + kyanite zone covers most of the study area. Kyanite occurence is sparse within the zone, and zone boundaries are consequently only approximately located. Kyanite occurs in two distinct settings. Most commonly it is found in quartz-kyanite knots as 4 cm to 8 cm long randomly oriented blades. These knots are limited to pelitic schists, and minerals within them may be kinked by D 4 crenulations. Kyanite also forms small porphyroblasts in muscovite schists. Extensive retrogression to sericite makes textural relationships uncertain, but in one example kyanite clearly crosscuts both D 2 crenulations and S2 muscovite (Figure 46). This latter mode of occurence is used in defining the staurolite + kyanite zone boundaries. Staurolite is texturally similar to that previously described in the staurolite zone, with long axes of porphyroblasts randomly aligned in the S o / S i foliation plane. Sphene and quartz inclusions are abundant, but no internal fabric is defined. Several relationships suggest staurolite growth in part accompanied D 2 . In one example where staurolite mantles an earlier garnet, the mineral pair has a quartz pressure shadow elongate parallel to 12. Elsewhere, staurolite grew across phase 2 crenulations (Figure 47). Garnet forms idioblastic to xenoblastic poikiloblasts with varying degrees of retrogression to chlorite. Two stage growth is indicated by inclusion free rims mantling poikilitic cores. The following textural relationships suggest garnet growth throughout phase 2 deformation: 1. helycitic S^ . quartz inclusion trails where crenulation fabrics are subparallel to those in the surrounding rock (Figure 48), METAMORPHISM / 95 Figure 46. Kyanite porphyroblasts growing across phase 2 crenulations, with later retrogression to sericite, sample 193. METAMORPHISM / 96 METAMORPHISM / 97 Figure 48. Helycitic inclusion trails in garnet, sample 293. Sz- crenulations are parallel to phase 2 crenulations in surrounding rock. METAMORPHISM / 98 2. deflection of S2 surfaces about garnet prophyroblasts, and 3. pressure shadows about garnet aligned parallel to 12 lineations. Biotite and muscovite textures are similar to those in other pelitic zones. No evidence is seen for stability of chlorite in the prograde assemblage of the staurolite-kyanite zone. Discussion Pelitic rocks of the Ogden Peak area contains Barrovian prograde mineral assemblages of up to kyanite/staurolite grade. Reactions marking first appearances of indicator minerals are uncertain due to strong compositional controls on mineral appearance and extensive retrogression of prograde reaction textures. Possible reactions may be modelled using observed stable assemblages and assuming a six component system of A1 2 0 3 , FeO, K 2 0 , H 2 0 , MgO, and Si0 2 . These generalized reactions are approximate in that they are derived from observed changes in stable mineral assemblages rather than reaction textures, and microprobe analyses indicate significant amounts of CaO and MnO are present (see section 5.4). The first appearance of staurolite marks the boundary between garnet-biotite± chlorite and garnet-biotite-staurolite stable assemblages, suggesting chlorite may break down to form staurolite. A possible balanced reaction involving one change in A F M diagram topology (Figure 49) is propsed by Albee (1965): 95Gt(Al) + 499Mu + 324Ch = 204St + 449Bt + 441Qz + 1194H20 An alternative reaction not involving garnet is presented by Hoscheck (1969) and is proposed for the "staurolite in" isograd to the south in the Boss Mountain area (Fillipone, 1985): METAMORPHISM / 99 A Staurolite - Kyanite zone AFM sketches showing stable mineral assemblages of pelitic metamorphic zones. METAMORPHISM / 100 3Mu + Ch = St + 3Bt + 7Qz + 14H 2 0 As garnet breakdown is not observed in the staurolite zone, the latter reaction is more viable. The transition to the staurolite + kyanite zone is somewhat more problematic. Formation of kyanite may have been at the expense of staurolite, according to the reaction: St + Mu + Qz = Gt + Bt + Ky + H 2 0 (Carmichael, 1978). Extensive retrogression of both staurolite and kyanite make the feasibility of this reaction uncertain. This reaction involves two changes in A F M diagram topology in which first the staurolite-chlorite tie line must be broken to form a biotite-ASK tie line (Figure 49). Alternatively, biotite, garnet, staurolite, and kyanite may all be stable phases of a seven or more component system as suggested by appreciable amounts of calcium and manganese in garnet (see appendix 4). The above generalized reactions should be considered speculative and lack firm supportive textural evidence. METAMORPHISM / 101 5.2.2. Amphibole schist lithologies H b - B t - Q z - C h - P l - M u + Cc, II, Sid Amphibole schists form all of unit 2, and occur occasionally as layers within other stratigraphic units. They are comprised of of 80% to 100% hornblende, except westernmost exposures of unit 2 which have 35% to 40% carbonate minerals. Hornblende forms large subidiomorphic grains which are mostly elongate in the S 0/S! plane. It shows faint to moderate green pleochroism, and larger crystals are extremely poikilitic. Ilmenite and quartz inclusions in hornblende define an internal foliation (S.) which in elongate grains is subparallel to the external foliation. In equant grains S^ . may be slightly discordant to the dominant foliation. Pre- to syn- phase 2 growth is indicated by porphyroblasts kinked about D 2 crenulations. Biotite comprises up to 10% of the volume of amphibole schists. It occurs both as an alteration product of hornblende and as earlier prograde biotite growth. Earliest biotite may be completely enclosed by later hornblende porphyroblasts. Radiating fibrous chlorite is common both as alterations of biotite and hornblende and isolated knots within the quartz/feldspar matrix. Accessory minerals include clinozoisite, sphene, and rutile. Clinozoisite occurs as idiomorphic fine grains within chlorite aggregates or as inclusions within hornblende porphyroblasts. Rutile and sphene are abundant idiomorphic grains with long axes oriented in the dominant S 0/Sj foliation. They are commonly associated with chlorite, but isolated grains may be completely enclosed by hornblende. Carbonate rich amphibole schists in the western study area are composed of 30% to 40% siderite, with lesser amounts of calcite. Siderite forms anhedral grains which surround more euhedral phases, and often have calcite rims. Hornblende has a higher inclusion density here than in non-carbonate schists, and is also rimmed by METAMORPHISM / 102 calcite. Clinozoisite is more abundant than in non-carbonate lithologies, and forms euhedral grains partially or completely enclosed in hornblende. Elongate clinozoisite is aligned within the So/Sj foliation and is bent about phase 2 crenulations. METAMORPHISM / 103 5.2.3. Carbonate lithologies Cc-Bt-Par(?)-Ep-Qz ± PI. Hb, Sph, II Carbonate lithologies are limited to the east-west trending unit 4 marble and occasional lenses in siliciclastic units. Mineralogy consists of calcite, mica, epidote, quartz and hornblende. The marble occurs in both staurolite and staurolite + kyanite pelitic zones, but the distribution of mineral phases reflects compositional variations rather than changes in metamorphic grade. Coarse grained calcite comprises 80% to 95% of all unit 4 carbonates. It forms a coarse granoblastic matrix with a weak foliation defined by elongate calcite grains. Dark green xenomorphic hornblende phenocrysts are aligned subparallel to the calcite foliation. They are poikilitic with randomly oriented inclusions of quartz, epidote, sphene, and ilmenite, and are typically embayed by calcite. Hornblende is most abundant in the westernmost exposures of unit 4, where it makes up 10% to 20% of lithology. The micaceous component of the carbonates consists of green to brown biotite in hornblende bearing marbles and white mica in more pure carbonates. Both occur as small euhedral flakes aligned parallel to foliation, either completely enclosed by calcite or growing across grain boundaries. Fine grained epidote is commonly associated with small quartz rich pods in the marble. It also occurs rarely as isolated grains within the calcite matrix, or as inclusions in hornblende porphyroblasts. METAMORPHISM / 104 5.3. RETROGRADE METAMORPHISM M 2 Rocks of the Ogden Peak aiea have been subject to a regional retrograde metamorphism evidenced by the replacement of prograde mineral assemblages by lower grade hydrous phases. Biotite is replaced by mimetic growth of chlorite and small amounts of quartz. Original grain boundaries and intragranular kink bands are preserved throughout chlorite growth. Garnet porphyroblasts are primarily replaced by unoriented fibrous chlorite along rims with lesser sericite, quartz, and ilmenite concentrated near cores. Sericite pseudomorphs of staurolite have small amounts of chlorite and ilmenite restricted to core zones of pre-existing porphyroblasts. Kyanite is replaced by a randomly oriented matrix of sericite with rare larger muscovite grains. The completeness of retrogression is determined by local composition on both microscopic and macroscopic scales. Porphyroblasts in pelitic schist show greatest retrogression, commonly with complete replacement of kyanite, staurolite, and garnet Retrogression is less complete in structurally lower pelitic rocks (unit la), and increases higher in the succession. Amphibole schist lithologies lack primary kyanite and staurolite, but partial chloritization of biotite is common. Carbonate rocks and quartzites both lack retrograde textures. On a microscopic scale, the degree of retrogression is strongly determined by the identity of the phases present Kyanite and staurolite never have original growth faces preserved, and are 60% to 100% replaced even when in contact with garnet only slightly mantled by chlorite. Garnet is anywhere from unaffected to 100% replaced, but is never replaced to a greater extent than adjacent staurolite and kyanite porphyroblasts. Likewise, biotite may be up to 100% replaced by chlorite, but never more than garnet staurolite, and kyanite in the same rock. METAMORPHISM / 105 Retrograde textures are primarily developed in layers along the S 0 / S i foliation plane, occasionally resulting in a rock with alternating retrograde and non-retrograde compositional layers. Timing of retrograde metamorphism relative to deformation is roughly constrained by microscopic textures. Chlorite pseudomorphs of garnet are flattened parallel to D 4 axial planes in a region of high D 4 strain where unaltered garnets retain original crystal shapes (Figure 50). In another example, biotite replacement preferentially occurs along phase 4 crenulation axial planes. Discussion Extensive retrograde metamorphism in the Ogden Peak area indicates abundant fluids accompanied reduced temperatures and pressures (Turner, 1981). Prograde reactions in underlying rocks may have provided a source for fluids, which were then transported through regional advective flow coupled with local diffusion (Mohr et al., 1986). Local variations in completeness of retrogression appear to be controlled strongly by composition and may result from variable rates of advection and diffusion. Limited availability of fluids would have resulted in preferential retrogression of higher grade porphyroblasts which would be out of equilibrium earliest, and might explain why replacement of staurolite and kyanite is generally more complete than that of garnet and biotite. M 2 retrograde metamorphism closely followed Mi prograde metamorphism and continued through D 4 folding. Independent evidence for high fluid activity at this time exists in the formation of quartz filled extension fractures along axial planes of D 2 mesoscopic folds (Figure 51). These fractures would not have been in an extensional field until D 4 . The following speculative reactions are consistent with the proposed retrogression sequence and observed textures: METAMORPHISM / 106 Figure 50. Chlorite pseudomorphs of garnet flattened into phase four axial planes, from sample 593. METAMORPHISM / 107 Figure 51. Quartz vein filling fracture along axial plane of mesoscopic phase 2 fold, near location 280. METAMORPHISM / 108 Kyanite breakdown: 3Ky + 3Qz + 2K + + 3 H 2 0 = 2Mu + 2H + Staurolite breakdown: St + 3Bt + 7Qz + 14H 2 0 = 3Mu +Ch (Hoscheck, 1969) Garnet breakdown: 4Alm + Bt + 12H 2 0 = Ch + Mu + 3Qz (Fillipone, 1985) Biotite breakdown: Bt + H 2 0 + 2H + = Ch + Qz + II 4- 2K + These are generallized reactions derived from observed mineral assemblages. In high grade pelites, excess K + ions generated from biotite breakdown may be consumed by kyanite breakdown in the high grade pelites. Electron microprobe analyses of specimens might allow determination of exact balanced metamorphic reactions by which retrogression occurs. METAMORPHISM / 109 5.4. G E O T H E R M O M E T R Y Metamorphic temperatures have been calculated in rocks of the Ogden Peak area from Fe-Mg partitioning between co-existing garnet/biotite pairs using the experimental calibration of Ferry and Spear (1978). Newton and Haselton (1981) published a modification of the Ferry and Spear geothermometer to correct for mixing of Ca in garnet An empirical model proposed by Ganguly and Saxena (1984) further corrects for both Ca and Mn components in garnet In the present study, temperatures are calculated using the original Ferry and Spear calibration and both of the above correction models. Pressure estimates for use with the geothermometer are constrained by aluminosilicate stability fields (Holdaway, 1971) and the G R A I L geobarometer of Bohlen et al. (1983). The G R A I L geobarometer is based on the univariant reaction garnet + rutile = kyanite + quartz + ilmenite, which shifts to higher pressures with increasing almandine component in garnet The assemblage kyanite + quartz + ilmentite + garnet(Alm 77) is stable in kyanite + staurolite zone rocks, indicating maximum pressures of 6.5 kb in the kyanite stability field (Figure 52). Maximum pressures for a sample from the garnet + biotite zone are less well constrained. Pressure estimates for the Ogden Peak area of 7 kb (McMullin and Greenwood, 1986) determined using the Ghent (1976) geobarometer fall outside the pressure range indicated by GRAIL geobarometry, and may represent conditions earlier in the metamorphic history. The Ferry and Spear geothermometer has only a small pressure dependence (3°C to 4°C per kbar) and will not be notably affected by the above range of pressures. Elemental compositions of coexisting garnet-biotite pairs were obtained using the U B C Department of Geology electron microprobe. Testing procedures and results are METAMORPHISM / 110 500 600 700 800 Temperature f^c) Figure 52. P - T diagram applied to pelitic rocks of Ogden Peak. (1) G R A I L geobarometer of Bohlen et al., modified for almandine(77); (2) from Hoscheck (1969). ASK stability fields of Holdaway (1971). METAMORPHISM / 111 detailed in appendix 4, and sample locations are identified on plate I. Temperature estimates were determined using the Ferry and Spear geobarometer and the Newton and Haselton and Ganguly and Saxena calibrations. The calibration of Newton and Haselton consistently yielded temperatures higher than either of the other geothermometers. Although the Newton and Haselton calibration does not consider the effects of manganese in garnet, the temperatures obtained appear to be independent of manganese content and are most consistent with mineralogical assemblages observed. Thus, they are preferred over temperatures determined from other geothermometers. Temperatures determined from three pelitic schist samples are plotted as a function of manganese concentration in figure 53. Sample 598 from the garnet + biotite zone yielded an average temperature of 530 ± 24° C. This is almost identical to the 540 + 20° C temperature from the kyanite + staurolite zone (sample 233). An anomalously low temperature of 463° C from sample 447 may result from local chloritization of garnet rims, and probably represents a retrograde temperature. Samples collected from the diabase dykes (651-2, 651c) yield temperatures consistently higher than those from pelitic rocks (562 ± 23° C, 585 ± 13° C). In summary, kyanite grade pelitic rocks of the Ogden Peak area record metamorphic temperatures of approximately 540° C at 6 kbar. This temperature is consistent with mineralogical assemblages and with temperatures determined by McMullin and Greenwood (1986) for the same area (545 ± 4°C). Retrograde metamorphism involving re-equilibration of garnet and biotite has reset temperatures in many rocks of the area. METAMORPHISM / 112 Ferry-Spear O Newton- Haaetton • Ganguly-Saxena A 9 o a. E 0 600-550 500-I co CM 00 CB IO T - a I O C 0 CM A O « 2 co IO IO I CO CO CM O A 450-o A p»6 kb l I I I i i 1.0 2.0 3.0 G a r n « t - M n wt.% 4.0 5.0 Figure 53. Garnet-biotite temperatures from samples 447, 598, and 233 plotted as a function of manganese concentration in garnet METAMORPHISM / 113 5.5. S U M M A R Y AND DISCUSSION Metamorphism in the Ogden Peak area ranges from chlorite grade to kyanite grade. A metamorphic peak of approximately 540° C at 6.0 kb was attained in kyanite grade pelitic schists late in or just after phase 2 deformation. Relative timing of mineral growth and deformational episodes is summarized in figure 54. Mica foliation growth accompanied phase 1 and phase 2 deformation, but later mineral' growth is unoriented except for mimetic replacement textures. Relations between phase 3 deformation and mineral growth are inconclusive due to the lack of microscopic structures. Complete retrogression of metamorphic porphyroblasts along with syn- and post metamorphic quartz vein intrusions provide strong evidence for abundant fluids at all stages of metamorphism. Fluids present during retrogression may have been generated at presently exposed and lower structural levels and trapped by overlying semi-impermeable rocks. Lithologic boundaries which project above the Ogden Peak area and might have provided possible fluid traps include an Upper Snowshoe Group marble and the Omineca Belt/Intermontane Belt boundary. Ethoridge et al. (1983) propose that a fluid cap at shallow structural levels is generated as a natural consequence of regional prograde metamorphism. The cap would develop at the boundary between hot rising metamorphic fluids and cooler meteoric water, as mineral phases precipitated from the cooling metamorphic fluid into existing pore spaces, thereby reducing permeability. METAMORPHISM / 114 Phase 1 syn post Phase 2 syn post Chlorite Muscovite Biotite Garnet Staurolite Kyanite Phase 4 pre syn post P r o g r a d e Retrograde c o n s t r a i n e d i n f e r r e d Figure 54. Mineral growth / deformation correlation chart 6. REGIONAL CORRELATIONS AND TECTONIC INTERPRETATIONS 6.1. REGIONAL CORRELATIONS . The Intermontane Belt/Omineca Belt boundary is a suture zone of crustal proportions which strongly influences deformational and metamorphic styles in rocks of both belts. Previous studies in the Quesnel Lake area have concentrated on areas either adjacent to and straddling the suture zone (Montgomery, 1978; Fillipone, 1985; Elsby, 1986; Bloodgood, 1987; Carye, 1985) or closer to the central Omineca Belt in high grade rocks of the Shuswap Metamorphic Complex (Fletcher, 1972; Klepacki, 1980; Engi, 1984; Getsinger, 1985). Regional workers have described a general stratigraphy comprising rocks of both belts (Campbell, 1972; Struik, 1984), but have not studied in detail the structural and stratigraphic progression eastwards from the margin. The current geologic studies of the Ogden Peak area allow correlation of structural styles and stratigraphy described along the suture zone with that to the northeast, as well as add detail to existing regional work. Additionally, metamorphic studies of the Ogden Peak area help constrain regional metamorphic styles. In this section, the stratigraphic succession and the progression of deformational and metamorphic styles eastward from the margin are described with specific references to the Ogden Peak area. A discussion of possible relations to the tectonostratigraphic framework of the Intermontane and Omineca tectonic belts follows. 115 R E G I O N A L CORRELATIONS A N D TECTONIC INTERPRETATIONS / 116 6.1.1. Stratigraphy A 10 kilometer thick structurally continuous section of cratonic metasedimentary rocks has been described from the accretion boundary/suture eastward into the Shuswap Metamorphic Complex (Ross et al., 1987). These rocks comprise members of the Snowshoe Group, which ranges in age from Hadrynian to Early Paleozoic (Campbell, 1978; Struik, 1983). Upper and lower successions can be delineated within the Snowshoe Group on the basis of lithology in the Ogden Peak area and in the Three Ladies Mountain area (Getsinger, 1985). The lower succession is dominated by siliciclastic metasedimentary rocks. An upper succession of siliciclastic and carbonate metasedimentary rocks is separated from the lower by major marker units of marble and amphibole schist that occur close to its base and are traceable over a distance of 70 kilometers. No evidence is seen for layer-parallel thusts within the sequence, as inferred by Struik (1986a) from stratigraphic observations. R E G I O N A L CORRELATIONS A N D TECTONIC INTERPRETATIONS / 117 6.1.2. Structure Earliest recognizable structures in all Snowshoe Group rocks of the Quesnel Lake region are mesoscopic isoclinal folds outlined by compositional layering. They are commonly rootless, and together with a strong axial planar foliation define a regional transposed foliation which outlines all later structures. Map scale features related to this deformation are not as yet documented in the Ogden Peak and surrounding areas. Second phase folds are west verging in the Ogden Peak area and east of Ogden Peak, but westwards from Ogden Peak, closer to the margin, second phase folds have east vergent asymmetries. Temporal relations to metamorphic mineral growth and other deformational episodes suggest these structures formed synchronously throughout the region. East verging D 2 folds are the oldest structures common to both autochthonous and allochthonous packages. Phase three upright to west verging folds are syn- to post- metamorphic throughout the Quesnel Lake region, and associated strain increases toward the west Amplitudes of major folds decrease eastward from the suture zone while wavelengths remain constant and related mesoscopic folds become less abundant Third phase folds are the earliest structures which deform the margin, and are responsible for the present regional map pattern. Fold forms are controlled by contrasting rheology across the IMB/OMB suture. Synforms are typically cuspate forms in which slivers of less competent Triassic phyllite of Quesnellia are drawn down into shear zones within more competent Snowshoe Group rocks (Fillipone, 1985). Lobate antiforms are cored by competent Snowshoe Group rocks. Late northeast trending folds (phase 4 at Ogden Peak) are ubiquitous throughout the Quesnel Lake region, but at the margin are antedated by a fourth northwest trending fold set absent in the Ogden Peak area. R E G I O N A L CORRELATIONS A N D TECTONIC INTERPRETATIONS / 118 Post-folding brittle faulting is cornrnon in the Ogden Peak area and to the northeast, but becomes increasingly less conspicuous nearer the suture zone. Absolute ages of deformational events are roughly constrained by several plutonic bodies in the Quesnel Lake region. The Boss Mountain Gneiss and Quesnel Lake Gneiss, deformed by all phases described above, yield zircon U - P b dates of 338 Ma to 375 Ma (Mortensen et al., 1987). This is consistent with a zircon U - P b age of 335 Ma to 450 Ma on a p r e -F i dioritic gneiss in the Three Ladies Mountain area (Getsinger, 1985). Phase 2 folds are believed to have formed during the accretion of allochthonous terranes of the Intermontane Belt onto North America (Ross et al., 1985). Stratigraphic evidence elsewhere indicates a Jurassic age for this event (Monger, 1982). Timing of convergence is constrained in the Kootenay Arc region, some 150 km south of Quesnel Lake, by syn- to post- kinematic granitic plutons of 177 Ma to 155 Ma which intrude both continental rocks and rocks of Quesnellia (Archibald et al., 1983). Third phase deformation follows obduction, but precedes emplacement of the Hobson Pluton which was dated using Rb-Sr whole rock analyses at 138 + 12 Ma and 163 ± 7 Ma (Pigage, 1977). In addition, the undeformed Raft Batholith, dated using Rb-Sr whole rock analysis at 104.3 ± 33 Ma, cuts the suture between Omineca Belt and Intermontane Belt rocks (Jung, 1986), although recent K - A r dates on this intrusion indicate a Jurassic age (R. L. Armstrong, oral comm., 1987). These constraints place the timing of phase 1 deformation at Late Paleozoic to Early Mesozoic, and the timing of phase 2 and phase 3 as Jurassic. However, the plutonic bodies used to constrain the age of phase 3 are located in areas where this deformational episode is not strongly developed and may not have affected a competent plutonic body. Thus, the above minimum age assigned D 3 is suspect R E G I O N A L CORRELATIONS A N D TECTONIC INTERPRETATIONS / 119 6.1.3. Metamorphism Metamorphic rocks in the Quesnel Lake region range from chlorite to sillimanite grade. Metamorphic zones delineated by early workers (Campbell, 1962; Campbell & Campbell, 1972) have been shown to represent in some cases isograds marked by discontinuous reactions, and elsewhere changes in mineral assemblages resulting from variations in bulk composition (Getsinger, 1985; Fillipone, 1985; Fletcher, 1972). Sillimanite grade rocks are localized in three distinct nodes. Metamorphic zone boundaries about these nodes are extremely asymmetric, resulting in wide metamorphic zones through most of the area but very narrow zones adjacent to the Omineca Belt/Intermontane Belt suture and along the eastern boundary of the Shuswap Metamorphic Complex. Zone boundaries are broadly folded by major phase 3 structures. The Ogden Peak area is located on an antiformal metamorphic culmination in an area of relatively wide metamorphic zones. Most parautochthonous rocks are chlorite grade, but staurolite and garnet zone boundaries locally crosscut the suture zone in the Eureka Peak area. Metamorphic zones are sharply offset across the North Arm of Quesnel Lake northwest of the study area where a probable late brittle fault places sillimanite grade rocks against chlorite grade rocks. Metamorphic textures, including porphyroblastic inclusion trails, pressure shadows, and kinked grains, as well as the above map scale relations have been used to determine the relative timing of metamorphism and deformational events. Metamorphic grade increased through phase 1 and phase 2, peaking late in or following phase 2 but prior to the onset of phase 3. Absolute timing of metamorphism has been inferred from cooling ages in several locations. Sphene from the Quesnel Lake Gneiss near Isoceles Peak yields a U-Pb date of 174±4 Ma for cooling through the 500° C isotherm (Mortensen et al., 1987). A less reliable date of 114 Ma to 117 Ma is R E G I O N A L CORRELATIONS A N D TECTONIC INTERPRETATIONS / 120 determined from monazite from the Boss Mountain Gneiss, and may represent cooling through a lower temperature (Mortensen et al., 1987). An anomalously young 83 Ma K - A r age determined from a muscovite/plagioclase pegmatite in the Three Ladies Mountain area (Getsinger 1985) probably represents cooling through 400° C. It is not clear whether peak metamorphic conditions were reached everywhere in the Quesnel Lake region simultaneously, and the above range of cooling ages suggests metamorphism may have been diachronous through the region. Widespread hydrothermal retrogression observed in rocks around Ogden Peak is also documented to the northwest in the Three Ladies Mountain Area (Getsinger, 1985) and along the East Arm of Quesnel Lake near Isoceles Peak (Montgomery, 1985; McMullin, oral comm., 1987). These three areas outline a northwest trending zone of retrogression which roughly follows stratigraphy and the suture zone but cuts sharply across metamorphic zone boundaries. Synmetamorphic veins, especially common in high grade rocks, have composition and structural relations which indicate a hydraulic fracture mechanism of formation during dewatering of the sedimentary pile (Garwin, 1987; Ross et al., 1987). These veins, together with the asymmetry of metamorphic zones about thermal nodes, suggest that advective heat transfer played a strong role in influencing metamorphic styles in the Quesnel Lake region. The superposition of retrograde metamorphism on wide metamorphic zones is evidence of advection across, zone boundaries (England and Thompson, 1984). Advective heat transfer would be prominent in zones of high directional fluid permeability, which would be localized according to pre-existing structural or lithologic anisotropics. Distribution of retrograde metamorphism in a single continuous zone suggests fluids ponded in that area below a semi-impermeable barrier and were present while temperatures and pressures waned following metamorphic peak REGIONAL CORRELATIONS AND TECTONIC INTERPRETATIONS / 121 conditions. Continuous carbonate units in the Upper Snowshoe Group and the OMB/IMB suture zone both project above the region of retrogression and dictate a geometry that may have provided such a barrier. REGIONAL CORRELATIONS AND TECTONIC INTERPRETATIONS / 122 6.2. TECTONIC IMPLICATIONS Observations of structural and metamorphic styles in the Ogden Peak area add detail to existing models which outline the geologic evolution of the Intermontane Belt/Omineca Belt boundary in central British Columbia. Prior to the accretion of allochthonous terranes from the west, siliciclastic and carbonate sediments had accumulated along the western shore of North America in a marginal basin or shelf setting (Struik, 1982). Sediments comprising the Snowshoe Group included siliclastic and volcanic sediments, as well as carbonates. Granitic bodies now forming the Quesnel Lake and Boss Mountain gneisses intruded into this package during the Devonian and Mississippian. Volcanic and sedimentary rocks of Quesnellia Terrane formed in the Late Paleozoic to Jurassic in a volcanic arc setting off the coast of North America (Monger, 1977). It is unclear whether mafic and ultramafic rocks of the Crooked Amphibolite represent oceanic basement to this arc, or whether they are part of a separate tectono-stratigraphic terrane. Earliest recognizable structures (Di) occur only in the autochthonous rocks, and are thus thought to be associated with deformation prior to or very early in convergence. During phase 1 deformation, Snowshoe Group and intrusive granitic rocks underwent dewatering and mesoscopic folding at the greenschist grade of metamorphism. The earliest structures common to both allochthon and autochthon consist of easterly verging folds concentrated along the suture zone. These folds do not deform the margin and decrease in intensity away from it, and are thus believed to be related to Jurassic convergence and subsequent overthrusting of Quesnellia onto the craton. Rocks at the margin and along several detachment surfaces higher in the R E G I O N A L CORRELATIONS A N D TECTONIC INTERPRETATIONS / 123 allochthon are strongly mylonitized during this thrusting. Fold styles are controlled by lithology and position within the structural succession. In the allochthon, folds are tight to isoclinal at the margin and open upwards, and axial planes are marked by pressure solution and slaty cleavage developement In rocks of the Snowshoe Group tight to isoclinal folds are accompanied by discrete zones of mylonitization and brittle faulting immediately below the suture (Fillipone, 1985), but at deeper levels brittle detachment surfaces are lacking. Phase 1 folds in Snowshoe Group rocks are tightened during initial convergence. Obduction of Quesnellia resulted in elevated temperatures and pressures in underlying Snowshoe Group rocks, contributing to the increased metamorphic grade inferred to occur at this time. Present exposures of the suture zone near Crooked Lake suggest a minimum of 70 km to 80 km of shortening by overthrusting. The absence of easterly verging folds in the Ogden Peak area suggests that either: 1) the thrust surface has ramped high above present exposures or 2) overthrusting is much less than that observed at Crooked Lake, and the suture did not extend above the Ogden Peak area. Easterly verging folds and the suture zone itself are deformed by westerly verging folds following initial obduction. These structures indicate continued convergence which is no longer taken up by movement along the tectonic boundary. West verging map scale folds of the boundary have forms which are strongly controlled by the viscosity contrast between allochthonous and autochthonous rocks. Synclines are cuspate zones riddled with quartz filled hydraulic fractures and may have served as conduits for advective heat flow into higher level rocks during metamorphism (Ross et al., 1985). The absence of these cuspate zones around Ogden Peak is again evidence that the area is structurally far removed from the zone of convergence. R E G I O N A L CORRELATIONS A N D TECTONIC INTERPRETATIONS / 124 Several models have been proposed to explain the change from early east verging structures to later west verging structures (Ross et a l , 1985; Brown et al., 1986). A simple model proposed by Ross et al. relates the change in vergence to a switch in subduction direction (Figure 55). Following initial obduction, the oceanic slab carrying Quesnellia is subducted beneath the leading edge of North America, producing the observed change to westerly vergences. This model requires that west verging phase 2 folds situated structurally well below the margin (near Ogden Peak) formed synchronously with phase 3 folds at the margin, with the attendant implication that metamorphic peak conditions are also attained later to the east An east-stepping metamorphic front would be expected with thrusting induced tectonic loading occuring later to the east A recent model proposed by Brown et al. (1986) has initial convergence followed by west directed "back thrusting" from the craton, with associated west verging folds. This model may not apply to the Quesnel Lake region, where detailed mapping has not shown the existence of back thrusts. However, the back thrust model is attractive in that it allows contemporaneous metamorphic peak conditions and phase 2 folding throughout the area. A third possibility is that east directed subduction occurs throughout convergence, and that obduction of Quesnellia represents a short lived accretion event in which high level volcanic and sedimentary rocks are decoupled from the subducting slab. East verging folds would be concentrated along the presently exposed suture zone in this model, and west verging folds would form throughout the autochthonous package at all stages of convergence. Additional study of the nature of mafic and ultramafic sequences exposed along the margin will help in evaluating this proposition. Regional models proposed by Monger et al. (1982) involve large scale lateral R E G I O N A L CORRELATIONS A N D TECTONIC INTERPRETATIONS / 125 1) P r e - J u r a s s i c Snowshoe Group sediments 3) L a t e J u r a s s i c (?) r e v e r s a l Figure 55. Plate tectonic model for evolution of I M B / O M B boundary region. Modified from Dewey (1976), Ross et al. (1985). 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Tectonic setting of basic volcanic rocks determined using trace element analyses. Earth and Planetary Science Letters, 19, pp. 290-300. Pell, J. 1984. Stratigraphy, structure and metamorphism of Hadrynian strata in the southeast Cariboo Mountains, British Columbia. Unpublished PhD. thesis, University of Calgary, Calgary, Alberta. Ramsay, J. G . 1962. The geometry and mechanics of "similar" type folds. Journal of Geology 68, pp. 75-93. Ramsay, J. G . 1967. Folding and fracturing of rocks. McGraw-Hil l Book Company, 568 P-Read, P. B. 1976. Geology, Lardeau west half, British Columbia. Geological Survey of Canada, Open File 432. Reesor, J. E. 1970. Some aspects of structural evolution and regional setting in part of the Shuswap Metamorphic Complex. In: Structure of the Canadian Cordillera, editor J. O. Wheeler. Geological Association of Canada Special Paper 6, pp. 73-86. Ross, J. V., Fillipone, J. A., Montgomery, J. R., Elsby, D. C. and Bloodgood, M . A. 1985. Geometry of a convergent zone, Central British Columbia, Canada. Tectonophysics 119, pp. 285-297. Ross, J. V., Lewis, P., and Garwin, S. 1987. Geology of the Quesnel Lake Region, central British Columbia: Geometry and implications. In: Geological Society of America, Abstracts with programs, Cordilleran Section Meeting, Hilo, Hawaii, 19. Schiarizza, P. and Preto, V.A. 1984. Geology of the Adams Plateau-Clearwater area. British Columbia Department of Energy, Mines, and Petroleum Resources, preliminary map 56. Struik, L. C. 1982. Snowshoe Formation, central British Columbia. In Current Research, part B. Geological Survey of Canada, Paper 82-IB, pp. 117-124. REFERENCES / 132 Struik, L. C. 1983. Bedrock geology of Spanish Lake and parts of adjoining map areas, central British Columbia. Geological Survey of Canada, Open File map 920. Struik, L. C. 1984. Geology of the Quesnel Lake and part of Mitchell Lake map area, central British Columbia. Geological Survey of Canada, Open File map 962. Struik, L. C. 1986a. Imbricated terranes of the Cariboo Gold Belt with correlations and implications for tectonics in Southeastern British Columbia. Canadian Journal of Earth Sciences, 23, pp. 1047-1061. Struik, L. C. 1986b. A regional east dipping thrust places Hadrynian onto probable Paleozoic rocks in Cariboo Mountains, British Columbia. In Current Research, part A. Geological Survey of Canada, Paper 86-IA, pp. 589-594. Sutherland Brown, A . 1957. Geology of the Antler Creek area, Cariboo District, British Columbia. British Columbia Department of Mines, Bulletin 38. Thompson, A. B. 1976a. Mineral reactions in pelitic rocks: I. Predictions of P - T - X (Mg-Fe) phase relations. American Journal of Science, 276, pp. 401-424. Thompson, A. B. 1976b. Mineral reactions in pelitic rocks: II. Calculation of some P - T - X (Mg-Fe) phase relations. American Journal of Science, 276, pp. 425-454. Thompson, A. B., Lyttle, P. T., and Thompson, J. B. Jr. 1977. Mineral reactions and A - N a - K and A - F - M fades types in the Gassets schist, Vermont American Journal of Science 277, pp. 1124-1151. Turner, F. J. 1981. Metamorphic Petrology, 2nd edition. McGraw Hil l , New York, 524 P-Uglow, W. L. 1922. Bedrocks and quartz veins of Baikerville map area, Cariboo district, British Columbia. Geological Survey of Canada, Summary report 1922, part A, pp. 82-87. Wanamaker, B.J. and Evans, B. 1985. Experimental diffusional crack healing in olivine. In Point defects in Minerals, Robert N . Schock, Editor. American Geophysical Union geophysical monograph 31. pp. 194-210. Wheeler, J. O. and Gabrielse, H . 1972. The Cordilleran structural province. Geological Assodation of Canada, Spedal Paper 11, pp. 1-81. 8. APPENDICES 8.1. APPENDIX 1. MAJOR AND TRACE ELEMENT ANALYSES Major and trace element concentrations were analyzed for the samples listed below using the U B C Department of Geology X-ray fluorescence spectrometer. Pressed powder pellets were prepared using a tungsten-carbide ring mill and agate mortar to minimize contamination introduced during sample preparation (Hickson and Juras, 1986). Samples analyzed Meta-diabase dykes: 442j 444 297 293b 459 591 Metasedimentary Rocks: 246n (quartzite) 372 (pelitic schist) 188 (feldspathic quartzite) 340 (amphibole schist) 293a (marble) 133 APPENDICES / 134 T A B L E 1: MAJOR E L E M E N T CONCENTRATIONS Si0 2 A1203 FeO MgO CaO Na 20 K 20 T i 0 2 MnO P 20 : 442-J 57.46 17.38 8.13 4.01 7.58 3.15 1.34 0.47 0.12 0.36 444 53.05 18.26 10.44 3.78 8.44 3.99 0.92 0.18 0.15 0.16 297 59.76 16.63 7.63 2.82 6.86 3.69 1.17 0.65 0.11 0.66 293-b 59.05 16.55 7.66 . 2.53 7.62 4.66 0.39 0.66 0.11 0.79 459 58.83 17.01 9.08 3.26 5.50 4.21 1.09 0.61 0.17 0.25 591 55.63 16.91 8.39 4.50 8.26 3.15 2.05 0.53 0.12 0.45 246-n 93.95 3.11 0.56 0.47 0.41 0.10 1.19 0.06 0.00 0.16 372 66.55 17.03 4.74 1.91 0.54 0.59 7.95 0.48 0.09 0.13 188 86.05 8.41 1.13 0.58 0.23 1.42 2.00 0.10 0.02 0.05 340 55.78 12.90 9.35 9.00 10.22 0.75 0.65 0.86 0.14 0.35 293a 5.65 2.90 1.08 2.92 85.07 1.25 0.65 0.06 0.07 0.38 T A B L E 2: T R A C E E L E M E N T CONCENTRATIONS Cr Nb Rb Sr V Y Zr 442-j 26 6 BDL 811 233 24 72 444 28 6 B D L 537 319 21 155 297 - 6 15 477 - 26 102 293-b - 7 BDL . 820 - 24 105 459 37 3 10 412 307 22 90 591 36 6 47 586 275 23 78 372 36 17 180 111 82 17 198 246-n 15 6 9 9 60 17 41 188 12 BDL BDL BDL 23 BDL 31 340 879 38 87 126 239 35 276 293a 67 42 47 3130 25 40 104 B D L = Below detection limits APPENDICES / 135 8.2. APPENDIX 2: AXIAL DIMENSIONS OF ELONGATE PEBBLES Dimensions of elongate quartz and feldspar clasts were measured in the field at location 602. Measurements were taken on surfaces parallel to foliation and perpendicular to 12 lineation. Long and intermediate axes were measured on foliation surfaces, short axes on lineation-normal surfaces. In each case, the largest 10% of clasts in a defined area were recorded: The arithmetic mean of each group is plotted on a Flinn diagram in figure 23. APPENDICES / 136 Long axes (nun) Quartz 56 26 38 31 20 17 46 18 15 21 18 22 22 23 39 18 32 28 32 18 17 24 16 29 Feldspar 11 21 9 5 12 12 9 18 12 8 9 9 13 10 11 11 6 8 8 11 . 17 22 13 15 Intermediate Axes (mm) Quartz 7 6 4 7 8 8 6 6 7 7 7 5 11 3 3 3 4 2 12 3 7 8 3 5 Feldspar 9 9 5 3 4 5 5 5 6 3 5 4 2 7 2 3 7 6 3 5 4 3 4 6 Short Axes (mm) Quartz 2 3 2 2 2 3 1 3 2 3 2 4 2 3 2 2 3 4 3 1 2 3 3 2 Feldspar 3 2 3 4 6 4 4 3 3 4 3 5 4 3 2 3 3 4 3 3 4 3 2 3 8.3. APPENDIX 3: MINERAL ABBREVIATIONS APPENDICES / An Andalusite Ap Apatite Aim Almandine Bt Biotite Cc Calcite Ch Chlorite Cz Clinozoisite Ep Epidote Gt Garnet Hb Hornblende 1 1 Ilmenite Kspar Potassium Ky Kyanite Mt Magnetite Mu Muscovite Op Opaques Par Paragonite PI Plagioclase Qz Quartz Ru Rutile Sid Siderite Ser Sericite Sill Sillimanite Sph Sphene St Staurolite To Tourmaline Zo Zoisite Zr Zircon APPENDICES / 138 8.4. APPENDED 4: ELECTRON MICROPROBE ANALYSES Microprobe analyses of garnet and biotite were completed using the C A M E C A SX-50 microprobe at the UBC Department of Geology. Five samples were collected from pelitic schists and meta-diabase intusions in garnet + biotite and staurolite + kyanite metamorphic zones in the Ogden Peak area. Operating conditions were similar for all analyses: Accelerating potential = 15kV Specimen current = 50 nanoamps Counting time = 10 seconds Beam diameter = 10-15 microns Garnet and biotite were analyzed separately and the microprobe was recalibrated for each mineral species. Standards used consisted of natural and synthetic minerals from the U.B.C. Department of Geology collection (table 1). This work represented the initial use of the newly installed C A M E C A microprobe. APPENDICES / 139 T A B L E 1: STANDARDS FOR G T - B T G E 0 7 H E R M O M E T R Y Element Standard Locality UBC id no. GARNET Si Al Fe Mg Mn Ca Ti Pyrope Pyrope Fayallite Pyrope Pyroxmangite Pyrope Hornblende Victor Mine, S. Africa G278 Victor Mine, S. Africa G278 Synthetic O250 Victor Mine, S. Africa G278 Shidara-Machi, Japan P245 Victor Mine, S. Africa G278 Kakanui, New Zealand A229 BIOTITE Si Hornblende Al Orthoclase Ti Hornblende Mg Forsterite Fe Fayallite Mn Pyroxmangite Ca Hornblende Na Albite Kakanui, New Zealand A229 St. Lawrence. N. Y. F96 Kakanui, New Zealand A229 San Carlos, Arizona 0275 Synthetic O250 Shidara-Machi. Japan P245 Kakanui, New Zealand A229 Langlois, Orezon F20 APPENDICES / 140 TABLE 2: SAMPLE LOCATIONS Number 233 Latitude (N). 52° 35' 38" Longitude (W) 120° 43' 23" Rock type St + Ky pelitic schist zone 447 52° 36' 36" 120° 46' 19" St + Ky zone pelitic schist 598 52 a- 37' 23" 120° 48' 28" Gt + Bi zone pelitic schist 651-2 52° 36' 31" 120° 46' 39" St + Ky zone meta- diabase 651-c 52 : 36' 29" 1203 46* ^2" . St + Ky zone meta-diabase APPENDICES / 141 Sample 851-c 6 5 1 C - 3 6 5 1 C - 4 6 5 1 C-6 6 5 1 C - 7 Garnet S10. 3 7 . 10 37 . 6 0 36 . 9 3 37 . 47 Al .0, 21 . 10 2 1 . 2 3 2 1 . 38 2 1 . 0 5 MgO 1 . 98 2 . 0 5 2 . 1 7 2.20 MnO 0. 6 7 0 . 79 0 . 6 9 0 . 6 5 CaO 6 . 0 8 6 . 3 4 6 . 4 3 6 . 3 9 T 1 O 1 0 . 0 8 0 . 0 3 0 . 0 4 0.01 T o t a l 9 9 . 46 1 0 0 . 0 8 9 9 . 4 0 9 9 . 4 2 B i o t i t e S10, 34 . 8 0 36 . 57 35 . 78 46 . 53 Al ,0. 18 2 3 18 . 6 5 17 . 76 1 5 . 8 0 T10, 1 . 5 3 2 . 0 8 1 4 3 1 . 6 2 MgO 9 . 3 0 1 0 . 0 8 1 0 . 6 1 8 . 77 FeO 2 1 . 20 19.86 19 . 56 17 20 MnO 0. 13 0 . 0 6 0 . 0 3 0.02 CaO - 0 . 0 4 0 . 0 5 0.02 Na.O 0 . 0 3 0 . 0 3 0.02 -K,0 8 . 73 9 . 1 6 8.81 7 . 1 4 H't 0 3 . 8 6 4 OO 3 . 9 0 4 . 25 T o t a l 9 7 . 8 0 1 0 0 . 5 4 97 . 96 1 0 1 . 6 0 APPENDICES / 142 Sample 447 447-1 Garn e t S l O r 37.62 &l>Oi 21.34 FeO 34.52 MgO 2.40 MnO 0.36 CaO 4.70 TIOi 0.04 T o t a l IOO.96 B l o t 1te S10i 36 . 73 Al .0, 17 . 97 TIOi 1 . 76 MgO 12.80 FeO 16.83 MnO 0.05 CaO 0.01 Na.O 0.17 K ,o 9.17 HiO 4.01. T o t a l 99 . 49 APPENDICES / 143 TABLE 3: SAMPLE COMPOSITIONS Sample 233 - 2 3 3 - 2 Garnet S i 0 2 3 7 . 2 4 A 1 2 0 3 21 . 4 8 FeO 3 2 . 3 6 MgO 3 . 4 2 MnO 1 . 3 8 CaO 2 . 5 3 T i 0 2 -T o t a l 9 8 . 4 1 B i o t i t e S i 0 2 3 3 . 2 0 A 1 2 0 3 1 8 . 2 6 T i 0 2 1 . 3 0 MgO 1 4 . 4 1 FeO 1 7 . 9 7 MnO 0 . 0 4 CaO -Na 20 0 . 0 6 K 2 0 6 . 7 0 H 2 0 3 . 8 6 T o t a l 9 5 . 8 0 233 4a 233 4b 3 6 . 6 8 3 7 . 1 6 2 0 . 9 9 2 1 . 1 1 3 2 . 0 6 3 1 . 7 0 3 .21 2 . 9 9 2 . 3 1 3 . 0 0 3 .61 3 . 9 1 9 8 . 8 8 9 9 . 8 7 3 5 . 6 4 3 7 . 3 8 1 8 . 6 8 1 9 . 0 7 I. 32 1.42 II. 78 1 3 . 7 3 1 7 . 4 9 1 4 . 5 9 0 . 0 9 0 . 10 0 . 0 1 0 . 0 8 0 . 2 0 8 .21 9 . 4 4 3 . 9 5 4 . 0 8 9 7 . 9 2 1 0 0 . 0 1 2 3 3 - 5 2 3 3 - 6 3 7 . 0 5 2 0 . 5 4 2 1 . 0 6 2 0 . 7 9 3 1 . 2 9 3 2 . 3 1 3 . 2 0 3 . 4 6 4 . 3 2 1.31 2 . 6 4 3 . 5 5 0 . 0 7 9 9 . 6 3 9 8 . 4 4 3 4 . 7 7 3 4 . 0 4 1 8 . 3 9 1 9 . 2 4 1.34 0 . 9 4 1 2 . 7 8 1 3 . 8 0 1 8 . 5 6 1 8 . 4 5 0 . 1 5 0 . 1 0 0 . 0 2 0 . 0 5 0 . 0 4 8 . 0 4 7 . 4 3 3 . 9 2 3 . 9 4 9 8 . 0 2 9 7 . 9 7 APPENDICES 7 144 Sample 59S 598- 1 5 9 8 - 2 598-3 G a r n e t S i O » 36 . 3 6 3 7 23 ; 37.23 A 1 , 0 i 2 1.07 J 1 17 . 2 1 . 38 F e O 3 0 . 8 1 3 1 2 9 30 . 88 M g O 2 .09 1 3 3. 1 .82 M n O 2 . 30 1 7 9 3 . 00 C a O 6 18 6 6 0 6 . 27 T i O , ' 0. 12 C 0 3 0 . 0 7 T o t a 1 98 . 9 3 • 0 0 . 6 0 1 0 0 . 10 B i o t i t e S I O , 36 . 70 3 5 8 5 36 . 46 A l , 0 , 18.12 17 60 ' 18 . 55 T i O , 1 . 74 1 . 60 1 . 49 M g O 9 . 89 9 . 2 1 10. 25 F e O 19.01 17 . 3 6 18 . 46 M n O 0. 10 0 . 0 6 0. 10 C a O 0 03 0 .03 N a , 0 0 0 2 0 . 0 3 0 .04 K , 0 9 . 55 9 3 6 9 . 98 H , 0 3 . 95 3 . 85 3 96 T o t a 1 99 07 9 6 . 54 99 . 32 Saapla 681-2 6512-1 G a r n e t S IO i 3 5 . 0 5 A l i O . 21 .17 FeO 30 .94 MgO 2 .64 MnO 0 . 77 CaO 5 . 7 9 TIOi 0 . 0 5 T o t a l 97 .02 B l o t 1te S IO i 3 5 . 5 0 A 1,01 17 40 T IOi 1.91 MgO 9 . 7 0 FeO 19.75 MnO 0 . 0 6 CaO O 02 N a i 0 O .03 K10 9 30 H.O 3 .66 T o t a l 9 7 . 5 3 6512-2a 6512-2b 3 6 . 1 0 36 55 21.21 21 .79 31 .64 31 .77 2 .05 2 .35 O 66 0 .67 6 .37 6 .26 0.01 96 .04 99 .04 34 .46 32 .85 17.87 18.37 1.84 1.51 9.91 10.61 21 .02 20 .90 0 . 09 0 . 08 O 06 O.14 0 .03 0 .02 8 .57 6 .84 3 .85 3 .77 9 7 . 7 0 95 .10 6512-3 6512-5 36 .38 36 .67 21.31 2.1.50 30.42 31 .66 2 .02 2 .26 0 .69 0 73 7.01 6 .23 0 .05 0.01 97 92 99 .05 34 .65 34 57 17.18 17.31 1.97 1.17 9 .28 10.40 20 89 21 .08 0 .05 O.06 0 .06 0 . 05 0 .03 0 .03 8 .66 7 .67 3 . 80 3 .80 96 .55 9 6 . 1 0 6 5 1 2 - 6 a 36 .48 21 .23 31 .45 2 .35 0 .66 6 .46 O 04 98 .56 6512 -6b 35.81 21 .60 31 .34 2 :33 0 . 6 5 6 .33 0 .05 9 9 . 0 5 6512-7 35 .57 21 . 16 31 .84 2 .27 0 .67 6 .07 0.01 97 .58 34.61 18.11 1 .50 9 .54 20 .63 0 . 0 5 0.01 0 .02 8 .97 3.84 97 . 27 35 .62 17.53 1.51 9 . 5 9 19.49 0 .05 0 .02 9 .68 3 .85 9 7 . 3 5 38 .52 16 .89 1 .22 9.61 19.31 0 . 0 3 0 .03 0 .04 7.64 3 .93 9 7 . 2 0 Compositions Temperatures G t X - M g G t X - F e G t X - C a G t X - M n B t X - M g B t X - F e F & 5 N & H G ft 2 3 3 - 2 0 8 2 4 0 4 . 3 7 4 0 0 4 3 S O O 1 8 9 0 3 . 3 3 8 0 2 . 3 3 6 0 4 9 t 5 2 0 4 6 6 2 3 3 - 3 0 . 7 4 9 0 4 . 34 3 0 O . 7 4 6 0 0 2 4 3 0 2 . 7 BOO 2 . 9 8 0 0 5 9 6 6 4 6 5 9 2 2 3 3 - 4 0 . 7 7 7 0 4 . 3 4 9 0 0 6 2 8 0 0 31 BO 2 8 3 5 0 2 . 2 2 0 0 4 0 8 5 4 8 4 9 9 2 3 3 - 4 O . 7 170 4 . 2 6 0 0 0 . 6 7 2 0 O . 4 0 9 0 3 . 0 too 1 . 7 9 5 0 4 2 4 464 4 2 6 2 3 3 - 5 0 . 77 10 4 2 180 O 4 5 6 0 O 5 9 0 0 2 9 120 2 . 3 7 2 0 5 2 6 5 5 6 5 1 9 2 3 3 - 6 0 . 7.590 4 . 3 3 0 0 0 6 8 4 0 0 2 9 4 0 3 1350 2 3 5 1 0 4 9 0 5 3 4 4 8 5 2 3 3 - 7 O . 8 3 8 0 4 . 3 8 5 0 0 6 1 7 0 0 18 10 3 O 1 10 1 7 2 5 0 4 4 4 4 8 2 4 2 5 65 1 C - 3 O. 4 7 7 0 4 3 7 5 0 1 0 5 0 0 0 0 9 2 0 2 1550 2 . 7 5 8 0 5 0 7 5 7 5 544 6 5 1 C - 4 0 . 4 9 0 0 4 . 2 8 3 0 1 0 8 6 0 0 1070 2 . 2 5 1 0 2 : 4 8 8 0 4 8 0 5 5 0 5 1 5 6 5 1 C - 6 0 . 5 2 2 0 4 . 2 7 B O 1 1 IOO 0 0 9 4 0 2 . 4 3 3 0 2 5 1 7 0 4 8 0 5 5 0 5 10 6 5 t C - 7 0 . 5 2 8 0 4 . 2 5 5 0 1 10OO o. 0 8 9 0 1 . 8 4 7 0 2 . 0 3 2 0 501 572 53 1 6 5 1 2 - 1 0 . 2 8 0 0 2 . 1500 0 . 5 4 0 0 0 . 0 5 0 0 2 . 2 4 5 0 2 . 5 6 6 0 5 2 6 5 9 6 5 5 3 6 5 1 2 - 2 O . 2 7 9 9 2 1400 0 . 5 2 0 0 0 . 10O0 2 . 2 9 9 0 2 7 3 B O 5 3 9 6 0 7 5 6 9 6 5 1 2 - 2 0 . 4 9 9 0 4 . 33 10 1. 1 ISO 0 0 9 10 2 5 1 5 0 2 . 7 B O O 4 8 2 5 5 3 5 16 65 1 2 - 3 0 . 5 3 7 0 4 0 8 6 0 1. 0 8 7 0 o 1050 2 . 1800 2 . 7 5 3 0 5 5 9 6 3 4 5 8 9 6 5 1 2 - 3 0 4 9 0 0 4 . 1510 1 2 3 3 0 0 0 9 5 0 3 1 170 3 6 130 501 5 8 2 5 4 0 6 5 1 2 - 5 0 . 5 4 4 0 . 4 . 2 7 9 0 •V. 0 7 9 0 0 0 9 9 0 2 . 4 3 7 0 2 7 7 7 0 5 1 8 5 8 8 5 4 6 6 5 1 2 G 0 54 10 4 . 27 4 0 1 1240 0 0 9 10 2 2 2 3 0 2 6 9 7 0 5 3 5 6 0 9 5 6 6 6 5 1 2 - 6 0 5 6 7 0 4 2 7 5 0 1 1060 0 . 0 9 0 0 2 2 2 2 4 2 5 3 7 0 531 6 0 3 5 5 7 6 5 1 2 - 7 o. 5 4 8 0 4 . 3 2 6 0 1. 0 9 6 0 o. 0 9 6 0 2 2 2 8 0 2 . 3 8 0 0 4 9 8 5 6 8 527 6 5 1 2 - 7 O 5 5 8 0 4 3 8 8 0 1 : 07 10 o 0 9 3 0 2 . 1860 2 . 4 6 5 0 5 15 5 8 3 542 4 4 7 - 1 0 . 2 8 0 0 2 . 3 0 0 0 0 . 40OO 0 . 0 2 0 0 2 . 7 8 6 0 2 . 2 0 2 0 4 1 4 4 6 3 432 5 9 8 - 1 0 . 2 5 0 0 2 . lOOO 0 5 4 0 0 o. 16O0 2 . 2 3 8 0 2 4 140 4 8 4 5 5 2 5 2 2 5 9 8 - 2 0 2 3 0 0 2 . 1000 0 5 7 0 0 0 . 1200 1 3 2 0 0 1 4 IOO 4 6 0 531 5 0 2 5 9 8 - 3 0 . 2 2 0 0 2 . 0 7 0 0 0 . 5 4 0 0 0 . 2 0 0 0 2 . 3 1 4 0 2 . 3 3 9 0 4 3 9 5 0 6 4 8 4 


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