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Petrologic studies of process interactions in metamorphic systems : deformation and metamorphism in the… Marchildon, Marie-France Nathalie 1999

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P E T R O L O G I C S T U D I E S O F P R O C E S S I N T E R A C T I O N S I N M E T A M O R P H I C S Y S T E M S : D E F O R M A T I O N A N D M E T A M O R P H I S M I N T H E S E L K I R K A L L O C H T H O N O R O G E N I C W E D G E ; A N D F E E D B A C K M E C H A N I S M S D U R I N G R E A C T I V E F L U I D F L O W by M A R I E - F R A N C E N A T H A L I E M A R C H I L D O N B . S c , Univers i te de Mon t r ea l , 1989 M . S c . , Rensselaer Poly technic Institute, 1992 A T H E S I S S U B M I T T E D I N P A R T I A L F U L F I L M E N T O F - T H E R E Q U I R E M E N T S F O R T H E D E G R E E O F D O C T O R O F P H I L O S O P H Y i n T H E F A C U L T Y O F G R A D U A T E S T U D I E S (Department o f Ear th and Ocean Sciences; Facu l ty o f Science) W e accept this thesis as conforming to the required standard T H E U N I V E R S I T Y O F B R I T I S H C O L U M B I A December 1999 © Nathal ie M a r c h i l d o n , 1999 In presenting this thesis in partial fulfillment 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 Earth and Ocean Sciences The University of British Columbia Vancouver, Canada Date _7>4,/g?9 ABSTRACT The first part of this thesis focuses on the tectono-metamorphic evolution of the Selkirk Allochthon, a welt of supracrustal rocks deformed and metamorphosed during Mesozoic to early Cenozoic time as a result of terrane collision at the western edge of North America. The studied area is part of the Selkirk Allochthon. Pelitic rocks in the area record a Barrovian metamorphic field gradient. Detailed mapping of metamorphic assemblage distribution and textural analysis of pelitic rocks has unraveled two distinct metamorphic events (Ml and M2). Emplacement of the Bigmouth Creek stock, which intrudes rocks of the Big Fish Creek area was dated as mid-Jurassic, with evidence for substantial reheating in early Cretaceous time. Based on regional correlation, M l is assigned a mid-Jurassic age, synchronous with emplacement of the Bigmouth Creek stock. M2 is interpreted to have resulted from the early Cretaceous thermal event. The development of early-M2 kyanite and late-M2 andalusite indicates decompression over at least 3 kbar, implying at least 10 km of denudation. Detailed analysis of structural elements suggests that the area was part of a long-lived, mid-crustal zone of intense distributed strain. Uplift and denudation of metamorphic rocks by tectonic extension and coaxial thinning were synchronous with continued contraction. A tectono-thermal model is proposed for the Allochthon as a long-lived, dynamic orogenic wedge within which deformation accommodated contraction due to terrane convergence and accretion, and polycyclic metamorphism resulted from thickening and maintenance of the wedge over at least 20 My. Significant early Cretaceous thinning of the wedge recorded by extensional strain and late-M2 decompression resulted from waning of far-field contraction, and/or from rheological modification of the wedge due to thermal relaxation. The second part of this thesis reports on a study of heterogeneous reactive transport of H2O-CO2 fluids during contact metamorphism of siliceous dolomites. Comparison of mineral assemblage distributions resulting from numerical simulations of heterogeneous flow with assemblage distribution in the Alta (Utah) contact aureole indicates that mineral reactions record flow through highly heterogeneous permeability. This is interpreted to result from positive feedback between reaction enhancement of permeability and flow-focusing. Comparison with other contact aureoles suggests that this is a common mechanism of permeability evolution during contact metamorphism of carbonate rocks. More homogeneous permeability recorded in regional metamorphic rocks may reflect fundamental differences in the relative rates of fluid flow and compaction in regional and contact metamorphic environments. i i T A B L E OF CONTENTS Abstract ii List of Figures v List of Plates vii Acknowledgments viii Chapter 1 INTRODUCTION 1 Chapter 2 P O L Y C Y C L I C M E T AMORPHIC EVOLUTION OF THE BIG FISH C R E E K AREA, NORTHERN SELKIRK MOUNTAINS, SOUTHERN OMINECA BELT, BRITISH COLUMBIA 4 2.1 Introduction 4 2.2 Regional setting V 2.3 Previous work V 2.4 Methodology 10 2.5 Local geology 10 2.6 Metamorphic events and metamorphic parageneses in pelitic rocks 12 2.6.1 M l metamorphism 13 2.6.2 M2 metamorphism 15 2.6.2.1 Chlorite zone 16 2.6.2.2 Biotitezone 17 2.6.2.3 Garnet zone , 19 2.6.2.4 Staurolite zone 21 2.6.2.5 Kyanite + staurolite zone and beyond 26 2.6.2.6 Late-M2 andalusite 29 2.7 Timing of Bigmouth Creek intrusion and metamorphism 31 2.8 Discussion 34 2.9 Conclusions.... 38 Chapter 3 OROGEN-PARALLEL AND OROGEN-TRANSVERSE SYN-COMPRESSIONAL EXTENSION, AND LONG-LIVED STRUCTURAL EVOLUTION OF THE SELKIRK A L L O C H T H O N OROGENIC WEDGE, BRITISH COLUMBIA, C A N A D A 40 3.1 Introduction 40 3.2 Geological setting 43 3.2.1 Tectonic environment 43 3.2.2 Stratigraphy 43 3.2.3 Metamorphism 46 3.3 Structural geology 46 3.3.1 Previous work 46 3.3.2 Structural elements in the Big Fish Creek area 48 3.3.2.1 Dominant foliation, S T 48 3.3.2.2 "F T", tight to isoclinal folds associated with S T 53 3.3.2.3 Mineral/intersection lineation, L T and transient foliation, ST' ... 55 iii 3.3.2.4 Stretching lineation in quartzo-feldspathic rocks 59 3.3.2.5 Fractured porphyroblasts and porphyroclasts 60 3.3.2.6 Extensional crenulation and shear bands 60 3.3.2.7 Southeast-trending crenulation and related folds 61 3.3.2.8 Northeast-trending crenulation 62 3.4 Discussion 62 3.4.1 Intensity of deformation 62 3.4.1 Nature of deformation 64 3.4.1.1 Orogen-transverse shear 64 3.4.1.2 Orogen-parallel shear 65 3.4.1.3 Sj-normal shortening (sub-vertical "flattening") 67 3.4.1.4 Transport direction '. 67 3.4.2 Timing of deformation 68 3.4.2.1 Relative timing of deformation and metamorphic recrystallization 68 3.4.2.2 Absolute timing of deformation 71 3.4.3 Tectonic implications 72 3.5 Conclusions 75 Chapter 4 IRREGULAR ISOGRADS, REACTION INSTABILITIES, AND THE EVOLUTION OF PERMEABILITY DURING METAMORPHISM 78 4.1 Introduction 78 4.2 The record of fluid flow 79 4.3 Models 83 4.4 Results 85 4.5 Discussion 85 Chapter 5 CONCLUSIONS 89 Bibliography 92 Appendix I: Photographic plates 103 Appendix II: Outcrop locations and list of samples 136 Appendix III: Mineral analyses '. 152 Appendix IV: Structural measurements 155 iv LIST OF FIGURES Chapter 2 Figure 2.1 Lithotectonic map of the Southern Omineca Belt 5 Figure 2.2 Map of metamorphic zones recorded in pelitic rocks of the southern Omineca Belt in the vicinity of Revelstoke, British Columbia 6 Figure 2.3 Distribution of mineral isograds and metamorphic assemblages recorded in pelitic rocks of the study area 8 Figure 2.4 Lithologic map of the study area 11 Figure 2.5 Core (left) to rim (right) compositional zoning profile of composite garnet porphyroblast, sample 94-NM-111B 14 Figure 2.6 Distribution of mineral isograds and metamorphic assemblages recorded in pelitic rocks in the area surrounding the Bigmouth Creek stock 23 Figure 2.7 Petrogenetic grid for pelitic rocks in the K 2 0 - FeO - MgO - A1 2 0 3 - SiO 2 -H 2 0 (KFMASH) system, modified after Spear and Cheney (1989) 25 Figure 2.8 Occurrences of late, coarse-grained andalusite and retrograde pseudomorphs after andalusite in the study area 30 Figure 2.9 U-Pb concordia plot of zircon and titanite analyses for Bigmouth a Creek K-granodiorite sample 33 Figure 2.10 Proposed dynamic wedge tectonic model for the Selkirk Allochthon 37 Chapter 3 Figure 3.1 Lithotectonic map of the Southern Omineca Belt 41 Figure 3.2 Lithologic map of the study area showing position of M2 mineral isograds 45 Figure 3.3 Attitudes of structural elements in the study area 49 Figure 3.4 Schematic representation of cyclical transposition and interpreted relationship between ST, ST' , F t , and extensional shear bands 52 Figure 3.5 Schematic representation of the attitude of lineation-defining biotite porphyroblasts 56 v Figure 3.6 Schematic representation of deformation of lineation-defining biotite flakes 58 Figure 3. 7 Deformation at the frontal and lateral tips of a spreading nappe 66 Figure 3.8 Simplified structural cross-section and tectonic model for the Selkirk Allochthon and Selkirk Fan at the latitude of the present study area 69 Figure 3.9 Proposed dynamic wedge tectonic model for the Selkirk Allochthon to account for inferred metamorphic and structural evolution 74 Chapter 4 Figure 4.1 A: Map of southern portion of Alta, Utah contact aureole. B: Whole-rock carbonate 8 1 8 0 of siliceous dolomites vs. distance from contact 80 Figure 4.2 A: Isobaric (1 kbar) phase diagram for model siliceous dolomite. B : Mineral assemblage stability as function of distance from intrusive contact and time-integrated fluid flux 81 Figure 4.3 Model mineral assemblage maps for down-T flow of initially pure H 2 0 fluid into siliceous dolomites , 84 Figure 4.4 Comparison of model results with Alta observations 86 Appendices Figure A-II.l Locations of outcrops and samples 137 vi LIST OF PLATES Plate 1 Garnet morphology and evidence for two prograde garnet growth episodes separated by retrogression 104 Plate 2 Elemental composition maps of composite garnet 106 Plate 3 Chlorite-zone, and biotite-zone textures 108 Plate 4 M2 mineral textures, staurolite zone 110 Plate 5 M2 mineral textures, staurolite zone and kyanite + staurolite zone 112 Plate 6 Retrogression and deformation of prograde M2 minerals 114 Plate 7 Sillimanite-bearing rocks from western reaches of the study area 116 Plate 8 Macroscopic structural elements 118 Plate 9 Split or fractured garnet porphyroblasts 120 Plate 10 Different stages of progressive transposition of the dominant foliation in the same outcrop, staurolite zone 122 Plate 11 Lineation-defining biotite porphyroblasts in LT-normal and LT-parallel sections.... 124 Plate 12 Rotated and deformed biotite porphyroblasts 126 Plate 13 Split or fractured porphyroblasts 128 Plate 14 Extensional structural elements 130 Plate 15 Folds with LT-parallel axes 132 Plate 16 Stretching lineation 134 vii ACKNOWLEDGEMENTS I would like to thank my parents, Monic Belzile and Jean-Paul Marchildon for opening my horizons, for challenging me to do well, and for instilling in me the curiosity which has lead to my pursuing a scientific career path. Their love and moral support have been tremendous assets throughout my life, and in particular, in seeing the present endeavor to its end. I am extremely grateful to my thesis advisor, Greg Dipple, for giving me the opportunity to tackle such diverse problems as those discussed in this thesis. He has generously provided intellectual, moral and material support during the long periods of searching for answers, and the short intervals of wonderment at the new insights we had gained. I also warmly thank the other members of my committee: Lori Kennedy, who has played an essential role in shaping my ideas about the structural evolution of the Selkirk Allochthon, and who was readily available for discussion, whenever I needed clarification about stress and strain, or just a beer-and-burger buddy; and Kelly Russell, whose extremely sharp mind has provided many an intellectual challenge, and whose hockey prowess has provided many an on-ice frustration. These people have been friends as well as advisors. I thank my fellow graduate students and metamorphic petrology group members for discussion of ideas metamorphic or otherwise: Annette Bingemer, Chris Clarkson, Ben Edwards, Stuart Knoop, Maria Niermann, Maile Smith, Mathias Westphal. A number of people in the department have provided technical support invaluable to this project, and for which I am grateful: Mark Baker, Bryon Cranston, Matti Raudsepp, and Mike St-Pierre. I also thank Jim Mortenson and Rich Friedman for their analytical work in dating the Bigmouth Creek stock. I extend my thanks to Claire Floriet, Tom Clemo, Sarah Vance and Barb Davey for cheerful assistance in the field. I would like to thank Raphael Wiist and Eliane Mandli, Paulette Tercier, Mark Caplan and Isabelle Chanoine Caplan, Claire Floriet, Boyd Benson, Mike St-Pierre, Tom Clemo, Michelle Lamberson and Arne Toma, for their friendship and good company during my stay in Vancouver, and beyond. I am grateful to the generations of Rock Dogs with whom I have shared a bench for five very successful years- Let's go Hoopsters! I would also like to thank Mr. and Mrs. Pierce, my landlords at 2832 W. 3 r d Ave for being so nice and accommodating and for providing the least expensive housing in all of Vancouver! Finally, I would like to thank Bill Minarik for his love, sustained support, and encouragement throughout this endeavor, even if it meant we were three time zones apart. viii Chapter 1 INTRODUCTION The results of two separate but related petrologic studies are presented in this dissertation: 1) a study of the tectono-metamorphic evolution of a dynamic orogenic wedge, the Selkirk Allochthon of southeastern British Columbia; and 2) a study of the effect of fluid flow on the distribution of metamorphic assemblages during metamorphism of carbonate rocks. The underlying approach in both cases is to apply various petrologic methods in order to gain a better understanding of the factors which affect the evolution of metamorphic processes, and hence metamorphic terranes, through time. This approach is grounded in detailed field and laboratory characterization of metamorphic rocks. Of particular interest throughout this thesis is the spatial distribution of mineral assemblages resulting from one, or more, metamorphic events, and the history of metamorphic reactions which lead to preservation of this macroscopic spatial distribution. Such descriptive information is used to answer questions regarding the nature and temporal evolution of processes attending metamorphism, including, specifically, deformation, fluid infiltration and evolution of hydraulic properties. The first study, which is presented in Chapters 2 and 3, is concerned with the interplay of tectonic, deformational and metamorphic processes within a dynamic orogenic wedge. The area of interest is the Selkirk Allochthon, a package of supracrustal and associated intrusive rocks deformed and metamorphosed during Mesozoic to Early Cenozoic time as a result of terrane collision at the western margin of North America. Chapter 2 presents a detailed characterization of the distribution of metamorphic assemblages in pelitic rocks of the Big Fish Creek area, which is underlain by rocks deformed and metamorphosed at intermediate depth in the orogenic wedge. Textural analysis of recrystallization sequence brings to light two distinct prograde metamorphic events (Ml and M2) separated in time by a period of cooling, infiltration and growth of retrogressive assemblages. Geochronological dating of the Bigmouth Creek stock, which intrudes rocks of the Big Fish Creek area allows the establishment of a mid-Jurassic age for M l , synchronous with emplacement of the stock. M2 resulted from Early Cretaceous reheating, recorded by titanite from the Bigmouth Creek stock. The spatial distribution of high- and low-pressure mineral assemblages developed during early and late stages of M2, respectively, requires that substantial exhumation affected these rocks in Early Cretaceous time. A detailed analysis of structural elements presented in Chapter 3 allows the reconstruction of a tectonic model for the wedge which accommodates the complex metamorphic history recorded in rocks of the Big 1 Fish Creek area. What permeates is a picture of a long-lived, dynamically evolving wedge controlled by the interplay of tectonic, thermo-metamorphic, and deformational processes. Far-field contraction due to collision and continued convergence led to internal deformation and thickening of the sedimentary package. This resulted in thermal relaxation, metamorphism and rheological changes in the interior of the wedge, which allowed for gravitational collapse, local extensional deformation, and decompression and cooling of the upper part of the wedge, even as overall contraction continued. Episodic fluctuations in the intensity of far-field contractional stresses and/or spatially heterogeneous internal deformation over several tens of millions of years at least, allowed for repeated heating of parts of the wedge, accounting for the polycyclic metamorphic recrystallization. Thus, the overall picture is one of "yo-yo" tectonics and metamorphism affecting a wedge active over a substantial period of time. This model may be applicable to other long-lived, highly diachronous orogenic belts such as the Appalachians and Caledonides of eastern North America and western Europe, the Hercynian belt of southern Europe, and the Himalayan chain. The second study, presented in Chapter 4, is concerned with the interplay of fluid infiltration, fluid infiltration-driven metamorphic reactions, and the resulting changes in rock permeability during metamorphism. Metamorphic permeability can not be directly measured, as rocks now exposed were probably affected by a range of permeability-altering processes during uplift and exhumation. However, the spatial variability of mineral assemblages produced by metamorphic fluid infiltration provides a picture of the spatial distribution of "time-integrated" metamorphic permeability distribution. Furthermore, by comparing natural distributions of metamorphic assemblages to quantitative numerical models of infiltration and reaction in heterogeneously permeable rocks, the heterogeneity of flow in natural systems can be estimated. Mineral assemblage distributions obtained from numerical simulations of infiltration during contact metamorphism of siliceous dolomites were compared to the spatial distribution of mineral assemblages observed in siliceous dolomites of the Alta, Utah, contact aureole. Only simulations with highly heterogeneous permeability (considerably more heterogeneous than equivalent protolith permeability) reproduced the geometry of mineral zones observed at Alta, suggesting that permeability becomes more heterogeneous during metamorphism. This is readily explained by feedback mechanisms between infiltration, mineral reaction and the resulting permeability enhancement (due to negative changes in volume during reaction), which will tend to produce progressively more focused, and hence, more heterogeneous, flow patterns. Comparison with other contact metamorphic terranes such as the Ubehebe aureole, California, suggests that this feedback mechanism is generally important during contact metamorphism of carbonate-rich rocks. On the other hand, much smaller fluid 2 flow heterogeneity inferred for some regional metamorphic terranes suggests that different factors, such as compaction, may control permeability evolution during metamorphism in these environments. Of the three main chapters which make up the bulk of this thesis, Chapter 4 has been published in a peer-reviewed journal (Marchildon and Dipple, 1998). The author contributed the larger portion of the work, including development of the two-dimensional flow and reaction algorithm and code, implementation of numerical simulations, analysis of results, drafting of figures (except Figure 4.1 A which was modified from a figure supplied by the second author), and original full draft of the text of the chapter. The second author suggested the study, provided some of the computer code, most notably the mineral reaction subroutine, as well as extensive discussion of methods, results and interpretations, and abundant useful comments on numerous generations of working drafts. Chapters 2 and 3 have not been published but are intended for eventual submission to peer-reviewed journals. But for the geochronological work, which was performed in large part by J. K. Mortensen and R.M. Friedman at the geochronology laboratory at UBC, the author is responsible for the field and laboratory results, for processing of these results into the format presented in this thesis, and for the bulk of the interpretations derived from these results and presented in this thesis. 3 Chapter 2 POLYCYCLIC METAMORPHIC EVOLUTION OF THE BIG FISH CREEK AREA, NORTHERN SELKIRK MOUNTAINS, SOUTHERN OMINECA BELT, BRITISH COLUMBIA 2.1 Introduction The southern Omineca Belt of British Columbia, Canada, is a long-lived orogenic zone formed by continued accretion and contraction at the western edge of the North American plate during Mesozoic to Early Cenozoic time (e.g., Monger et al., 1982; Parrish, 1995; Fig. 2.1). Recent geochronological studies have shown the complex nature of the Mesozoic thermal evolution of the southern Omineca Belt (Sevigny et al., 1989, 1990; Scammell, 1993; Crowley et al., 1998). However, evidence for a complex metamorphic history has remained elusive. The question of the metamorphic evolution of the Omenica Belt in the Northern Selkirk and adjacent Monashee Mountains is critical because previous work suggests that apparently continuous metamorphic zones in this area (Fig. 2.2) resulted from widely diachronous metamorphic recrystallization. On the western flank of the Selkirk structural fan (Wheeler, 1965), greenschist- to upper amphibolite-grade metamorphism has been interpreted to be of mid-Jurassic age and contemporaneous with emplacement of a suite of hornblende- and biotite-bearing granitoid bodies (Van der Leeden, 1976; Brown and Tippett, 1978; Brown and Lane, 1988; Brown et al., 1992; Colpron et al., 1996). In contrast, amphibolite- to upper amphibolite-grade metamorphism on the east side of the SFA and in the adjoining Monashee Mountains (e.g., Ghent et al., 1977, 1983; Sevigny and Ghent, 1986) is interpreted to be coeval with the emplacement of peraluminous granites at approximately 100 Ma (Sevigny et al., 1989, 1990). One interpretation of this diachroneity is that the metamorphic isograds record the displacement over time of a metamorphic "front" of regional extent, as associated deformation progressed spatially through the orogenic belt (Parrish, 1995). Alternatively, this apparent diachroneity may be due to incomplete preservation of multiple superimposed metamorphic events. This study is based on observations from the west flank of the Selkirk Fan structure (Fig. 2.1). The area is underlain by rocks which record at least two distinct prograde regional metamorphic events. Based on geochronological evidence and on textural similarity with rocks of the Big Fish Creek area, I argue that other parts of the southern Omineca Belt also underwent a polycyclic metamorphic evolution. 4 LEGEND Paleogene Kamloops Gp volcanics Middle Cretaceous granitoids Middle to Late Jurassic El granitoids Upper Trlasslc-Lower Jurassic Nicola arc volcanics Devon ian-Triassic Kaslo Gp, Harper Ranch Gp Devonian-Early Tertiary (undlf.) granitoids Upper Proterozolc-Paleozoic I Lardeau Gp, Milford Gp Eagle Bay assemblage Upper Proterozoic-Lower Cambrian Z2 Hamlll Gp, Badshot Fm Lower Proterozoic-Lower Cambrian Monashee complex, cover gneiss Upper Proterozoic Y77A Windermere Supergroup Lower Proterozoic Monashee complex, core gneiss Figure 2.1 Lithotectonic map of the Southern Omineca Belt. Thick black box shows location of present study area. Gray boxes are locations of recent studies in nearby areas: I. Scammell (1993); II. Colpron et al. (1996); III. Johnston, 1996. AB: Adamant Batholith; BS: Bigmouth Creek Stock; BRB: Battle Range Batholith; CRF: Columbia River Fault; CT: Clachnacudainn Terrane; FS: Fang Stock; GBS: Galena Bay Stock; GP: Goldstream Pluton; KB: Kuskanax Batholith; MC: Mica Creek area; MD: Monashee Decollement; MO: Monashee Complex; R: town of Revelstoke. Modified after Wheeler and McFeely (1991). 5 Intrusive rocks Jurassic w = Mid-Cretaceous Figure 2.2 Map of metamorphic zones recorded in pelitic rocks of the southern Omineca Belt in the vicinity of Revelstoke, British Columbia. Inset shows map location. Gray boxes are locations of recent studies in nearby areas: I. Sevigny et al. (1989,1990); II. Scammell (1993); III. Brown et al. (1992) and Colpron et al. (1996); IV. Johnston, 1996; V. Crowley et al. (1998). Mineral abbreviations are: Chi: chlorite; Bt: biotite; Gt: garnet; St: staurolite; Ky: kyanite; Sil: sillimanite; Ksp: K-feldspar. All other abbreviations are as in Figure 2.1. Modified from Read et al., 1991. A tectono-metamorphic model is proposed for the Omineca Belt in the area of the Northern Selkirk Mountains, as a long-lived dynamic wedge which is controlled by external forces, an evolving rheological behaviour, and a complex associated thermal evolution. This model differs from previous models that require discrete pulses of deformation and a single prograde regional metamorphic event (e.g., Van der Leeden, 1976; Leatherbarrow, 1981; Raeside and Simony, 1983; Sevigny and Ghent, 1986; Brown et al., 1992a; Colpron et al., 1996). 2.2 Regional setting The Selkirk Allochthon comprises a wedge of Late Proterozoic to Early Paleozoic continental margin sedimentary and minor volcanic rocks within the Omineca Belt. This wedge migrated eastward during crustal accretion and contraction in Mesozoic time (Monger et al., 1982; Brown et al., 1986; 1993). Eastward migration of the Allochthon is generally interpreted to have taken place primarily along a discrete crustal-scale, basement-rooted decollement surface, the Monashee Decollement (Read and Brown, 1981; Journeay, 1986; Brown et al., 1986; Brown and Journeay, 1987). Recently, however, Scammell (1993) and Johnston (1996; also see Williams and Jiang, 1999) proposed, on the basis of structural analysis of the Selkirk Allochthon in the Monashee Mountains west and south of the Big Fish Creek area (Fig. 2.2), that strain associated with eastward displacement of the Allochthon was distributed over mid-crustal high-strain zones of finite thickness. The Selkirk Fan structure (Wheeler, 1965) is part of the Allochthon. It consists of a wedge of rocks showing southwest structural vergence on the west side, and northeast structural vergence on the east side, defining a structural axis along the spine of the Selkirk and Northern Monashee Mountains. A metamorphic culmination is roughly coincident with this structural axis (Fig. 2.2). The Big Fish Creek area is located within the western flank of the Selkirk Fan structure. 2.3 Previous work Wheeler (1965) mapped several metamorphic zones which define the regional metamorphic culmination in the Northern Selkirk Mountains (Fig. 2.2). In the study area, Van der Leeden (1976) mapped metamorphic zones in pelitic rocks, including a chlorite-zone, a biotite zone, a garnet zone, a staurolite zone, and a kyanite zone, from south to north (Fig. 2.3). Leatherbarrow (1981) integrated information from different workers and using geothermobarometry, suggested that rocks on the southwest flank of the Selkirk Fan structure were metamorphosed at lower peak pressure (5 Kb) than rocks on the 7 Figure 2.3 Distribution o f mineral isograds and metamorphic assemblages recorded in pelitic rocks o f the study area. Numbers correspond to A F M assemblages as shown in A F M diagrams for the different zones. Question marks following numbers indicate that one o f the phases in the reported assemblage was completely retrogressed and that it was identified on the basis o f the pseudomorphic shape of the retrograde replacement mineral(s). Divariant assemblages observed in these rocks are shaded on the appropriate A F M diagrams. Mineral abbreviations are: C l d : chloritoid; Pr l : pyrophyllite. A l l others are as in Figure 2.2. 8 northeast flank (7 Kb). From this, he inferred post-metamorphic vertical displacement of 7 km on the poorly-constrained Argonaut Mountain normal fault, which he interpreted to trend NW-SE, between the Bigmouth Creek stock and the Adamant batholith (Fig. 2.1). The requirement for such a structural discontinuity to some extent hinges on the assumption that metamorphism is more or less contemporaneous throughout the metamorphic culmination. Perkins (1983) found that minerals developed as a result of prograde regional metamorphism across the metamorphic culmination showed similar relative timing relationships to deformation. Ghent et al. (1977, 1983) and Simony et al. (1980) constrained the sequence of metamorphic isograds in the northernmost Selkirk Mountains and Monashee Mountains north of the Columbia River, on the northeast flank on the metamorphic culmination. They mapped isograds corresponding to (1) the first appearances of kyanite, which coincides with the first appearance of staurolite, (2) the disappearance of staurolite which coincides with the first appearance of leucosome, (3) the appearance of sillimanite, and (4) the co-existence of garnet and clinopyroxene. Ghent et al. (1979, 1982, 1983) constrained peak metamorphic pressures to be in the 6-8 kbar range, for rocks of the staurolite + kyanite, to the sillimanite zones. Archibald et al. (1983) interpreted metamorphism in the southern part of the Selkirk Allochthon as synchronous with emplacement of mid-Jurassic intrusions (ca. 165-170 Ma). Van der Leeden (1976) argued that metamorphism in the present and adjacent areas must pre-date, or be synchronous with emplacement of the Bigmouth Creek stock (Figs. 2.2, 2.3), which he inferred to be of mid-Jurassic age based on similarity with known mid-Jurassic intrusive rocks in the area. Brown et al. (1992a) and Colpron et al. (1996) interpreted greenschist facies regional metamorphism and associated deformation in the vicinity of the Fang stock (Fig. 2.2) on the southwest flank of the Selkirk Fan, to have terminated by ca. 168 Ma. Sevigny et al. (1989, 1990; Fig. 2.2) interpreted regional metamorphism in the Monashee Mountains to coincide in time with emplacement of syn-kinematic granitic bodies which yielded Cretaceous U-Pb crystallization ages (ca. 100 Ma). Crowley et al. (1998; Fig. 2.2) found evidence for mid- to Late Jurassic regional thermal events in the U-Pb systematics of metamorphic monazite from the Northern Selkirk and Monashee Mountains north of the study area. Based on partial resetting of some monazite grains, and on evidence for renewed growth of monazite at ca. 70 Ma, these authors also postulated that a Late Cretaceous thermal event affected rocks in this area. In addition, they provided evidence for ductile deformation as late as Early Tertiary (ca. 61 Ma). Scammell (1993) documented 9 protracted metamorphism extending over at least 40 m.y. from ca. 140 Ma to ca. 97 Ma in the Monashee Mountains west of the Columbia River (Fig. 2.2). 2.4 Methodology The work presented in this and the following chapter was based on field mapping of lithological units and metamorphic assemblages during the course of thirteen weeks over three Summers (1994 to 1996). Field work also involved description and measurement of attitudes of the various structural elements observed in the Big Fish Creek area (see Appendix IV), and collection of over 250 samples in as random a fashion as allowed by outcrop abundance and quality. (Typically, several hand samples were collected from a given outcrop, from different structural levels, in order to obtain as representative as sample as possible.) Further analytical work involved petrographic analysis with a polarizing optical microscope and a Philips XL-30 scanning electron microscope, and microstructural description of over 200 thin sections cut from over 180 samples (see Appendix II for list of samples and outcrop locations), collection of compositional maps of garnet porphyroblasts and compositional point analyses of various minerals with a Cameca SX-50 electron probe microanalyzer (operating conditions: 15 KeV accelerating voltage, 5-20 nA beam current, 1-5 um spot size, using natural standards for recalculation of raw count data; see Appendix III for results). One sample of intrusive rock (Bigmouth Creek granodiorite) was collected and dated at the Geochronology Laboratory at UBC in an attempt to constrain the timing of pluton emplacement, and its implication for the timing of metamorphism and deformation. Interpretations presented as part of this study are drawn from observations and analysis representative of an important proportion of relevant samples. 2.5 L o c a l geology A lithostratigraphic framework for rocks of the study area is presented in Section 3.2.2 of the following chapter. A very brief summary follows here. The rocks are subdivided into four broadly defined packages separated by gradational contacts (Fig. 2.4), from low to high structural level: 1) a psammitic to semi-pelitic package whose base was not mapped in this study; 2) a pelitic package, approximately 100-200 m in thickness, comprising increasingly abundant, rusty weathering calc-pelitic schist and brown-weathering impure marbles towards the top of the interval, and which grades into, 3) a calc-pelitic package characterized by sporadic intervals of impure marble, and more rarely, of massive chlorite schist or amphibolite (Fig. 2.4); and 4) a sequence 10 ^l» \\ = ^L & * v w ^ ^ ^ » \\ = ^ & ^ » \\ " ^ // wV/ow / k-n //ow/ ^' . \ i LEGEND Bigmouth Creek granitoid m\mm\ chlorite schist/amphibolite impure limestone/marble gray calcite limestone/marble 18£>1 calcic pelite pelitic schist gritty psammite Y/A semi-pelite psammite foliation (ST) strike and dip 2 km Figure 2.4 Lithologic map of the study area. Stratigraphic intervals described in text are indicated by thick brackets. 11 of brown-weathering, impure to relatively pure calcite, and/or dolomite marbles, associated with thickly layered quartzitic psammite, local amphibolite and thin intervals of talc-schist, which was observed only in the northern part of the study area. The Bigmouth Creek pluton (Wheeler, 1965) is exposed at the northern limit of the study area. It consists of K-feldspar porphyritic, hornblende + biotite granodiorite. Epidote is present in the Bigmouth Creek pluton and it occurs as subhedral, pale yellowish brown crystals in close spatial association with hornblende. It is unresorbed and interpreted to be magmatic, with habit consistent with growth late in the crystallization history of the pluton. The presence of magmatic epidote is widespread in hornblende-bearing granitoids in the southern Omineca Belt (e.g., Rice, 1941; Reesor, 1973; Ghent et al., 1991; Colpron et al., 1996). The principal structural element in the study area is a shallow to moderately-dipping foliation characterized by the alignment of minerals with shape-preferred orientation, and generally sub-parallel to compositional layering. As discussed in the following chapter, this is a transposition foliation (referred to as ST) developed over a substantial time interval and associated with intense strain. ST strikes northwest to northeast, defining a large-scale, open non-cylindrical fold. A late east-west-striking, steep crenulation foliation is commonly observed, especially in the southern half of the area. The reader is referred to Chapter 3 for an in depth treatment of the structural geology of the area. 2.6 Metamorphic events and metamorphic parageneses in pelitic rocks The Northern Selkirk Mountains and the adjoining Monashee Mountains are underlain by regionally metamorphosed rocks recording greenschist to upper amphibolite facies conditions. Figure 2.2 shows the distribution of mineral zones in pelitic rocks. The southern part of the area is dominated by greenschist facies rocks (chlorite, biotite, and locally garnet zones for pelites). These give way northward to amphibolite facies rocks (garnet zone to sillimanite + K-feldspar zone in pelitic rocks; Fig. 2.2). The metamorphic culmination is defined by a north-northwest trending sillimanite + K-feldspar zone flanked to the northeast and southwest by progressively lower-grade rocks. The axis of the metamorphic culmination coincides approximately with the axis of the Selkirk Fan (Fig. 2.2). The Big Fish Creek area is located on the southwest flank of this culmination (Fig. 2.2). It exposes a metamorphic field gradient defined by a sequence of metamorphic zones characterized by the presence of chlorite in the south, to the coexistence of kyanite and staurolite in the north (Fig. 2.3). 12 Two distinct prograde regional metamorphic events are recorded in this area. Evidence for the early metamorphic event, M1, is cryptic and consists primarily of inclusion-rich, anhedral garnet cores in rocks affected by subsequent M2 recrystallization (Plates 1A,B, 2). The latter event, M2, is better preserved, and M2 mineral assemblages and isograds define the local field gradient (Fig. 2.3). Late growth of coarse-grained andalusite is interpreted to record late-M2 decompression. These different metamorphic recrystallization events are described in detail in the following sections. 2.6.1 Ml metamorphism Early anhedral cores of garnet in composite M1+M2 garnet porphyroblasts are ubiquitous in the M2 garnet and staurolite zones in the Big Fish Creek area (Plates 1A,B, 2). These cores are usually inclusion-rich, and overgrown by inclusion-poor, euhedral rims. Garnet porphyroblasts typically preserve curved or inclined inclusion trails (Sj * Se, where S; and S e are the foliations internal and external to the porphyroblasts, respectively; Plate 1A,B,C), which locally show a structural style distinct from that of the matrix. These early garnet grains are commonly broken into two or more fragments, and an M2 overgrowth is developed more or less concentrically around the individual clasts (Plates IB, 2). A sharp compositional break is invariably associated with the textural break (Plate 2; Fig. 2.5). A few occurrences of anhedral, inclusion-rich garnets devoid of euhedral overgrowth were found, primarily in the garnet zone. Typically, these are surrounded by a chlorite-rich (± biotite) rind which commonly pseudomorphs the original garnet grain (e.g., Plate 1C). These are interpreted as early, M l garnets based on their textural similarity with cores of composite garnets. The chlorite-rich rind is interpreted to be the product of retrogression involving M l garnet resorption. M2 garnet and/or staurolite are commonly developed around M l garnet cores (see Sections 2.6.2.3 and 2.6.2.4; Plates 1A,B, 2). Often, M2 garnet and/or staurolite overgrow the chlorite-rich rind (Plates IB, 2), implying that formation of the retrograde rind around M l garnet preceded the crystallization of M2 peak metamorphic minerals (garnet and/or staurolite). Retrogression of M l garnet prior to M2 garnet growth is also supported by the anhedral, resorbed aspect of M l cores (e.g., Plates 1A,B, 2). From these observations and because no reaction sequence resulting from a single prograde evolution can explain the two texturally and chemically distinct garnet generations and the intervening production of chlorite ± biotite-rich aggregates after garnet (e.g., Spear and Cheney, 1989), growth of 13 0.5 0.4 </> Q. (0 X 0.3 (A D) X a. 0.2 a. X 0.1 0.0 M l I I I I I I I I I I I 1 I I I I I h % Fe/Fe+Mg M 2 ^Tt'llllllllllllllllll^F sps 1.0 0.9 T l CD •-^  0.8 TI CD + CQ 0.7 X cu 3 0.6 0.5 200 400 600 800 Distance (microns) 1000 v Figure 2.5 Core (left) to rim (right) compositional zoning profile of composite garnet porphyroblast, sample 94-NM-11 IB, M2 staurolite zone (this is from the left-hand-side porphyroblast of Plate 2. IB; also see Plate 2.2A). Boundary between M l and M2 garnet (dashed vertical line) is inferred based on observed sharp decrease in the grossular, and increases in the almandine, pyrope, and spessartine components of the garnet. X ^ , X ^ , X s p s , X a l m are mole fractions of the pyrope, grossular, spessartine and almandine garnet components, respectively. Note the different scales f o r Xprp. x g r s a n d x s p s > a Q d for X ^ and Fe/(Fe+Mg). 14 anhedral garnet cores and euhedral overgrowths must have been separated by an episode of retrograde growth of hydrous minerals, which required cooling and infiltration of hydrous fluid. In other words, early and late garnet growth are evidence for two distinct metamorphic or thermal pulses (Ml and M2) and an intervening episode of cooling and retrograde recrystallization. 2.6.2 M 2 metamorphism The following M2 mineral zones are encountered in the area, from south to north (Fig. 2.3): chlorite, biotite, garnet, staurolite, and kyanite + staurolite. Mineral parageneses and reaction textures in pelitic and associated rocks of these different metamorphic zones are described in detail in the following sections. It is likely that observed matrix minerals such as muscovite, chlorite, and biotite crystallized during M l metamorphism. However, as discussed in detail in Chapter 3, both M l and M2 took place during protracted, progressive deformation which resulted in sustained transposition of structural elements in these rocks. Differentiating between Ml and M2 matrix minerals based on their spatial relationship to structural elements is therefore impossible. For this reason, the biotite-in isograd in Figure 2.3 may be partly diachronous (resulting from both Ml and M2). Higher-grade metamorphic isograds (garnet-in, staurolite-in, and kyanite-in), however, are based on unequivocal identification of 1) late-M2 garnet overgrowth on M l garnet cores; or 2) other minerals (staurolite, kyanite) associated with the development of M2 garnet overgrowths. The map in Figure 2.3 shows the distribution of M2 mineral assemblages and metamorphic zones. It is based on the observed distribution of metamorphic parageneses in rocks of pelitic composition, and shows isogradic traces generally at high angle to lithologic contacts. The different mineral zones are also positioned in the appropriate AFM projections (Fig. 2.3; Thompson, 1957). The three-dimensional attitudes of M2 isograds are unconstrained. Leatherbarrow (1981) argued for steeply-dipping (> 60°) isograds on the basis of the presence of steeply-dipping isograds in the Mica Creek area (Ghent et al., 1977; 1980). No clear evidence for a steeply-dipping attitude of isograds in the Big Fish Creek area was observed in the present study. 15 2.6.2.1 Chlor i te zone The chlorite zone is located at the southern end of the study area (Fig. 2.3). The southern limit of this zone, the first occurrence of chlorite in pelitic rocks, was not encountered. The northern, and higher-grade limit of the chlorite zone is defined as the surface of first occurrence of biotite (the biotite-in isograd, Fig. 2.3). The minimum outcropping width of this zone is approximately 1 km. Pelites in the chlorite zone are characterized by variable fresh coloration, which reflects variations in the contents of chlorite, muscovite, opaques, and locally, graphite. Grain size is generally fine to medium (< 1 mm), and increases irregularly towards the biotite isograd. In addition to a dominant penetrative foliation, ST, defined by the alignment of minerals with grain-shaped preferred orientation and by compositional layering, there is a steep crenulation cleavage developed in many chlorite-grade rocks. An associated crenulation lineation (resulting from intersection of ST with this cleavage) is almost ubiquitous (see Chapter 3, Section 3.3.2.7). Chlorite-zone pelites typically comprise the assemblage muscovite + chlorite + quartz + opaques + K-feldspar ± albite ± ankerite ± calcite ± epidote (pistacite) + tourmaline + apatite ± allanite ± zircon. They generally show mm-scale compositional laminae characterized by variable amounts of quartz and phyllosilicates ± carbonates ± opaques. Muscovite and chlorite flakes lie parallel to this layering. This defines ST. Locally, relicts of an earlier foliation, defined by the preferred orientation of fine, recrystallized muscovite flakes is preserved, which shows different degrees of transposition into ST- Late folding resulted in a gentle undulation to a strong spaced crenulation cleavage overprinting S T (Plate 3A,B). Where a crenulation cleavage is well-developed, muscovite and chlorite flakes are locally aligned along this cleavage (Plate 3B). ST-parallel chlorite grain-size generally increases towards the biotite isograd (from 0.2 to 0.5 mm). Large (1-2 mm) chlorite ± muscovite porphyroblasts and/or aggregates which locally show inclusion trails (Sj) mimicking crenulated S T (Se), are present which are locally deformed by, or around which are wrapped, the crenulation cleavage planes (Plate 3B). Such chlorite + muscovite aggregates have been widely reported in low-grade metamorphic rocks and may represent prograde chlorite growth on detrital white mica clasts (e.g., van der Pluijm and Kaars-Sijpesteijn, 1984; also see Craig et al., 1982 for a review of occurrences and possible modes of origin), or may be retrograde after earlier metamorphic minerals (see Sections 2.6.2.2 and 2.6.2.3 below). 16 2.6.2.2 Biotite zone Immediately north of the chlorite zone is the biotite zone (Fig. 2.3). Its low-grade boundary corresponds to the surface of first appearance of biotite in pelitic rocks. The biotite zone is approximately two kilometers in width, and is bounded at its northern, high-grade limit by the first occurrence of M2 garnet in pelitic rocks. Pelitic schists in the biotite zone are generally coarser-grained and paler in color than their equivalent in the chlorite zone. There is an irregular increase in grain size towards the garnet isograd. Coarser-grained rocks, especially in the upper part of the biotite zone, have a sheen due to the abundance of coarse muscovite flakes. As in the chlorite zone, a strong penetrative foliation, ST, is ubiquitous, and a steep crenulation cleavage is also locally developed, although not as extensively as in the chlorite zone. An associated intersection lineation is still relatively widespread. Typically, pelites of the biotite zone are characterized by the assemblage biotite + muscovite + chlorite + quartz + opaques ± plagioclase ± epidote + tourmaline + apatite ± allanite ± zircon. The first appearance of biotite is as basal cleavage-controlled intergrowth or overgrowth on medium-grained chlorite flakes in mm-scale chlorite ± muscovite aggregates, or as discrete, small, randomly oriented, equant flakes in these aggregates. Up-grade from the first appearance of biotite, biotite flakes in chlorite + biotite clots are still the dominant mode of occurrence of biotite (Plate 3C); this biotite becomes coarser and the clots become progressively more biotite-rich, at the expense of chlorite. Fine (100-200 um) isolated biotite flakes parallel to ST are rare at the low-grade end of the biotite zone, but become pervasive as the garnet isograd is approached. Biotite flakes showing random to Sj-parallel orientation become coarser and more abundant up-grade from the first appearance of biotite. Of particular interest are coarse biotite porphyroblasts which, from the upper part of the biotite zone onward, are widely developed (Plate 3D). The relationship of these porphyroblasts to matrix structures is not simple, but as discussed in detail in Chapter 3, provides important constraints on the nature of deformation in these rocks. Muscovite and chlorite in biotite-grade rocks generally show textures similar to those observed in chlorite-grade rocks. In general, no carbonate is present in biotite-bearing pelites of this zone. Medium-grained epidote porphyroblasts which overgrow ST as well as the later crenulation cleavage are locally abundant. 17 The typical pelitic paragenesis in biotite-zone rocks is biotite + chlorite + muscovite + quartz. Chloritoid was not identified. The close spatial relationship between biotite and chlorite flakes or aggregates in the lower half of the biotite zone suggests that a probable biotite-producing reaction was (e.g., Wang and Spear, 1991; Spear, 1993): chlorite + K-feldspar -> biotite + muscovite + quartz + H 2 0 (2.1) A limited amount of carbonate-bearing rocks occurs in the chlorite zone suggesting that subordinate reactions involving the consumption of a carbonate phase, such as: quartz + muscovite + siderite + H 2 0 -> biotite + chlorite +C0 2 (2.2A) or, quartz + muscovite + chlorite + ankerite -> biotite + plagioclase + C 0 2 + H 2 0 (2.2B) (e.g., Ferry, 1984; Symmes and Ferry, 1991) may also have lead to biotite production in the biotite zone. The biotite "isograd" is defined as the first appearance of biotite in pelitic rocks. Chlorite-grade (i.e., biotite-free) pelitic assemblages occur sporadically up-grade of the biotite isograd (Fig. 2.3). This may due to: 1) biotite stability attained at different conditions depending on the nature of the biotite-in reaction (reactions 2.1 or 2.2; e.g., Ferry, 1984); 2) the continuous nature of reactions 2.1 and 2.2, which will proceed at lower temperature in Fe-rich rocks than in Mg-rich rocks; 3) the effect of spatially variable H 2 0 and C 0 2 activities on reactions 2.1 and 2.2 (e.g., Ferry, 1984; Symmes and Ferry, 1991); or 4) kinetic factors. As described above, incipient development of biotite in the low grade portion of the biotite zone is usually as basal cleavage-controlled intergrowth or overgrowth on chlorite flakes, suggesting that early biotite growth is controlled in part by the ability to nucleate on pre-existing chlorite flakes, and that material transfer to biotite growth sites may be a limiting factor. In summary, the isogradic surface for first appearance of biotite is not a true reaction isograd but rather, is probably defined by a series of reactions governed primarily by bulk composition. The close spatial association of biotite and chlorite in the lower part of the biotite zone could also result from retrogression of biotite (e.g., Veblen and Ferry, 1983), and some chlorite in the biotite zone 18 can be shown to be retrograde after biotite and/or garnet (see section 2.6.2.3 below). The irregular nature of the biotite-in isograd could therefore in part represent the effects of heterogeneous retrogression. 2.6.2.3 Garnet zone The M2 garnet zone is located immediately north of the biotite zone (Fig. 2.3). The spatial extent of this zone is poorly constrained but appears to be relatively restricted (~1 km in width; Fig. 2.2). The low-grade boundary of this zone is defined by the appearance of M2 garnet overgrowth on M l garnet cores. The upper limit of the M2 garnet zone is defined by the surface of first appearance of staurolite in pelitic rocks (Fig. 2.3). There is a gradual increase in the grain size of matrix phyllosilicates from the chlorite zone onward. From the upper part of the biotite zone through to the kyanite + staurolite zone, matrix micas are generally relatively coarse-grained, giving pelitic schists a relatively pale color and a bright sheen (e.g., Plate 4A,B). As in the chlorite and biotite zones, a strong penetrative foliation, ST, is everywhere observed. A steep crenulation cleavage is rarely observed, and a crenulation lineation is less extensively developed than in lower grade rocks. Pelitic rocks in the garnet zone are characterized by the assemblage garnet + biotite + chlorite + muscovite + quartz + plagioclase ± epidote + opaques ± ilmenite + zircon + allanite + tourmaline. Much of the biotite in the garnet zone is texturally similar to biotite found in the upper part of the biotite zone. In addition, biotite is also commonly found as a replacement product after garnet, with or without chlorite (Plate 1). Chlorite is ubiquitous in garnet zone rocks and occurs as ST-parallel to randomly-oriented flakes. Late, randomly- or concentrically-oriented chlorite is found in replacement rims around garnet (with or without biotite; Plate 1C). Anhedral, inclusion-rich garnet porphyroblasts interpreted to be of M l origin are common in rocks between the garnet isograd and the first appearance of staurolite (Fig. 2.3). These are generally partially to completely pseudomorphed by retrograde chlorite ± biotite (Plate 1C,D). The development of M2 garnet (euhedral, inclusion-poor) is generally limited to thin overgrowths on these early porphyroblasts. Occurrences of these overgrowths define the M2 garnet zone (Fig. 2.2). 19 No chloritoid was found in the study area. On the other hand, biotite and chlorite are abundant in pelitic rocks of the biotite and garnet zones. Therefore, the M2 garnet producing reaction in the garnet zone was probably: chlorite + biotite + quartz + muscovite ± plagioclase -> muscovite + garnet + H 2 0 (2.3) (e.g., Spear and Cheney, 1989). Several points are worth noting concerning garnet occurrences in the M2 garnet zone: 1) M2 garnet growth is restricted to overgrowths on M l porphyroblasts; 2) the spatial range of M l garnet porphyroblasts more or less coincides with that of M2 garnet (although some chlorite + biotite pseudomorphs in the M2 biotite and chlorite zones may be after early garnet); and 3) the width of the M2 garnet zone is quite restricted (~ 1 km). The limited extent of the M2 garnet zone, and the restricted occurrences of M2 garnet in the garnet zone may in part be due to the paucity of appropriate lithological types in the area corresponding to the garnet zone. Outcrops in this area predominantly expose quartzo-feldspathic, psammitic to semi-pelitic rocks, and where pelites do occur, it is usually as thin, dip-slope intervals. M2 garnet overgrowths were found in relatively thin pelitic layers in otherwise dominantly quartzo-feldspathic intervals in four outcrops (Fig. 2.3). Constraining the position of the garnet-in isograd along the main north-south trending ridge in Figure 2.3 is hampered by the fact that psammitic rocks crop out over several hundred meters across strike of the isograd (Fig. 2.4). Hence the coincidence of M l and M2 garnet occurrences may be due to bulk compositional controls restricting the development of garnet rather than actual spatial coincidence of P-T conditions for development of garnet during M l and M2 metamorphic recrystallization. Another possible cause for some of the foregoing observations is the restricted stability of garnet + chlorite in bulk compositions typical of pelites in the study area. M l and M2 garnets tend to be relatively Mn-poor (Fig. 2.5). Spear and Cheney (1989) noted that the temperature range within which garnet + chlorite are stable is related to the Mn content of garnet (or the bulk rock). The more Mn-poor the garnet, the more limited is the stability of the garnet + chlorite paragenesis. 20 Yet another explanation is that the surface corresponding to the first appearance of M2 garnet in these rocks, the M2 "garnet isograd", may in fact be a surface defining the limit of total retrograde resorption of garnet subsequent to peak M2 metamorphism. A number of instances of completely retrogressed garnet (i.e., chlorite ± biotite pseudomorphs after garnet) were encountered in the garnet zone of Figure 2.3 (e.g., Plate ID). In addition, chlorite ± biotite clots are commonly encountered in rocks of the biotite and chlorite zones (Plate 3C). Phyllosilicates in these clots are ST-parallel to randomly-oriented, suggesting that the timing of formation of these clots relative to matrix structures is variable. Some of these clots may be of diagenetic origin (e.g., Craig et al., 1982). However, those which appear to be late with respect to deformation, and occur in the high-grade part of the biotite zone, may have originated by retrograde replacement of M2 and/or M l garnet porphyroblasts. The locally extensive retrogression of M2 mineral assemblages in higher-grade rocks (see below) is consistent with this interpretation. Alternatively, the restricted spatial extent of the M 2 garnet zone and the coincidence of M l and M2 garnet occurrences may be due to a kinetic control on the development of M2 garnet. Occurrences of M2 garnet in the M2 garnet zone are exclusively as overgrowths on M l cores. Even in the staurolite zone (and possibly also in the kyanite + staurolite zone), M2 garnet occurs predominantly as overgrowths on early M l cores, suggesting that nucleation may be a limiting factor in the development, and hence spatial distribution of M2 garnet. This, in turn, may reflect the short-lived nature of M2 metamorphism. The question of the spatial coincidence of M l and M2 garnet remains without an unequivocal answer at this time. It may consist of a combination of the possible causes discussed above, but insufficient information is presently available to evaluate their relative importance. 2.6.2.4 Staurolite zone The M2 staurolite zone is a little over two kilometers wide in map view (Fig. 2.3). Its lower-grade limit is the first appearance of staurolite (which everywhere coexists with biotite). Its higher-grade limit is constrained by the appearance of kyanite in staurolite-bearing rocks. In the western part of the area shown in Figure 2.3, rocks of staurolite grade are the highest-grade pelitic rocks encountered. This is probably due to the lack of exposure in general and in particular, to the paucity of outcrops of pelitic rocks in that area. However, rocks characterized by staurolite + biotite parageneses were found, in reconnaissance mapping, to crop out west of Big Fish Creek, to the northwest of the area shown in Figure 21 2.2 (see Fig. 2.6). Northward, these give way to staurolite + sillimanite parageneses. The significance of this will be addressed in Section 2.6.2.5. Staurolite-grade pelites are generally coarser-grained than lower-grade rocks, and show more discrete compositional layering at the millimeter scale, i.e., thin (1-2 mm, typically) mica-rich laminae separating thicker (-3-10 mm), more quartzo-feldspathic layers. This layering coincides broadly with the foliation defined by the preferred orientation of anisotropic minerals, ST, although, as demonstrated in Chapter 3, the nature of this foliation is very complex in detail. A crenulation lineation on S T is observed which is more or less parallel to intersection lineations in lower-grade rocks, suggesting the development of a steep crenulation cleavage. Pelitic rocks in the staurolite zone typically include the assemblage staurolite + garnet + biotite + muscovite + chlorite + quartz + opaques + plagioclase ± zoisite ± allanite ± tourmaline ± zircon ± apatite ± monazite ± xenotime. Through most of the staurolite zone, garnet typically occurs as composite porphyroblasts consisting of inclusion-rich, anhedral M l cores, and inclusion-poor, euhedral M2 rims (Plates IB, 2, 5A). The porphyroblasts commonly consist of fragments of split M l cores overgrown on all sides by M2 garnet. Euhedral, inclusion-poor grains devoid of obvious M l cores also occur throughout the staurolite zone. In the low-grade half of the staurolite zone, these are usually small (< 0.5 mm) and almost exclusively spatially associated with much larger composite garnet porphyroblasts (Plate 5A). Locally, they are included in staurolite (Plate 5 A). In the higher-grade portion of this zone, subhedral, inclusion-poor garnet porphyroblasts are also observed which are not spatially associated with composite garnets. These are also interpreted to be of M2 generation. Where staurolite has developed adjacent to M2 garnet, the garnet is embayed, suggesting that garnet consumption is involved in staurolite production (Plates IB, 5 A). In the higher-grade part of the staurolite zone, garnet porphyroblasts are locally very deeply embayed, commonly forming atolls. These porphyroblasts are spatially associated with staurolite, and embayments are typically filled in large part by biotite (e.g., Plate 5C). Biotite occurs in the staurolite zone as medium- to coarse-grained, ST-parallel to randomly-oriented flakes. The causes and implications of this distribution will be addressed in detail in Chapter 3. In addition, biotite flakes are developed around garnet in close association with staurolite (Plates IB, 5A). 22 Figure 2.6 Distribution of mineral isograds and metamorphic assemblages recorded in pelitic rocks in the area surrounding the Bigmouth Creek stock. Mineral isograds are shown as heavy lines, and locations are either known (solid lines) or inferred (dashed lines). Location o f main study area is indicated by box. Compiled from data by Franzen (1974), V a n der Leeden (1976), Leatherbarrow (1981), Gibson and Hoy (1985), and the present study. Mineral abbreviations are as in Figure 2.2. 23 Chlorite is generally present in rocks of the staurolite zone. Locally, it occurs as Sx-parallel, medium-grained flakes which are interpreted to be prograde with respect to M2 metamorphism. More commonly, it is found surrounding garnet porphyroblasts, but overgrown by staurolite. Abundant late chlorite of probable post-peak M2, retrograde timing, occurs as flakes which either cross-cut matrix structures or replace M2 prograde minerals (e.g., Plate 6A,B). The development and/or preservation of staurolite in rocks of the staurolite zone is extremely heterogeneous. Throughout this zone, staurolite is commonly found adjacent to garnet porphyroblasts, associated with biotite ± chlorite (Plates IB, 4A, 5A, 6A). Staurolite appears to post-date garnet in these instances, and this is confirmed by the local occurrence of small, euhedral, M2 garnets included in staurolite porphyroblasts (Plate 5A). Some staurolite occurs in the matrix away from garnet porphyroblasts (Plate 5D). Staurolite generally forms subhedral to euhedral, poikiloblastic grains. Grain size increases irregularly from about 1 mm in the lower staurolite zone, to over 1 cm in the upper part of this zone. Inclusion trails within staurolite porphyroblasts are generally straight, but the SJSe relationships are irregular: some staurolite grains show relative rotation, others do not, and these variations may be observed within a single outcrop (Plate 5C,D). Staurolite is commonly partially to completely retrogressed to an intergrowth of fine muscovite ± chlorite flakes. Often, retrograde phyllosilicates after staurolite are aligned parallel to ST (Plate 6B). In many pelitic outcrops in the staurolite zone no fresh staurolite was found. However, widespread occurrences on the outcrop or in thin section of staurolite-shaped mats rich in muscovite ± chlorite, typically in close association with garnet porphyroblasts which are more or less chloritized (Plate 6B) are interpreted to represent retrogressed, staurolite porphyroblasts. These occurrences are indicated in Figures 2.3 and 2.6 with a question mark. The general association of staurolite and biotite overgrowing garnet with which chlorite is commonly associated, suggests that the staurolite producing reaction in these rocks is related to (Fig. 2.7) garnet + chlorite + muscovite -> staurolite + biotite + quartz + H 2 0 (2.4A). This is the typical staurolite isograd univariant reaction in pelitic rocks metamorphosed at intermediate pressure for an ideal bulk composition in the K2O - FeO - MgO - A I 2 O 3 - Si0 2 - H 2 0 (KFMASH) system. In this system, equilibrium 2.4A is univariant so co-existing staurolite + garnet + biotite + chlorite should be restricted to the staurolite + biotite isograd surface. The widespread co-existence of these four phases in the study area is likely due to the increased thermodynamic variance 24 Figure 2.7 Petrogenetic grid for pelitic rocks in the K 2 0 - FeO - M g O - A 1 2 0 3 - S i 0 2 - H 2 O ( K F M A S H ) system, modified after Spear and Cheney (1989). Short-dashed lines are equilibria in the K F A S H sub-system, and long-dashed lines are equilibria in the K A S H sub-system. Black squares are K F M A S H invariant points with absent phases shown in brackets. Inferred pressure-temperature paths for rocks of the M 2 staurolite zone (gray) and of the kyanite + staurolite zone (black) are shown. See text for discussion. Abbreviations are: A i m : almandine; A n d : andalusite; A n n : annite; A s : alumino-silicate; Crd: cordierite; FeCld : Fe-chloritoid; FeSt: Fe-staurolite; M s : muscovite; Qtz: quartz; Tic : talc; others are as in Figures 2.3. 25 resulting from additional components, such as Mn and Ca, consistent with the observation of significant spessartine and grossular components in garnet (Fig. 2.5), and widespread occurrence of plagioclase. A more realistic staurolite producing reaction in the study area is therefore: garnet + chlorite + muscovite -> staurolite + biotite + quartz + plagioclase + H 2 0 (2.4B). Whitney and Ghent (1993) interpreted the co-existence of staurolite and garnet, as well as garnet compositional discontinuities, in pelitic rocks of the Solitude Range to the northeast of the Big Fish Creek area, to result from prograde reactions involving the consumption of chloritoid, which is preserved in these rocks as inclusions in garnet. Because chloritoid was not observed in the Big Fish Creek area, and because textural evidence strongly suggests a link between staurolite growth and garnet consumption, this scenario probably did not play a significant role here. 2.6.2.5 Kyanite + staurolite zone and beyond Rocks in which kyanite and staurolite coexist were observed in a single outcrop in the northeast corner of the mapped area (Fig. 2.3). These rocks are similar to staurolite zone rocks, except for the sporadic occurrence of kyanite porphyroblasts spatially associated with garnet and staurolite (Plate 5B). As with other M2 minerals, kyanite commonly bears fractures at high angle to the dominant matrix foliation, S T (Plate 5B). Kyanite is found in some pelitic layers and not in others, and where kyanite is absent, kyanite-grade rocks are similar to staurolite-grade rocks. The reader is referred to the previous section for a detailed description of phases other than kyanite. Pelites in the kyanite + staurolite zone (Fig. 2.3) are characterized by a paragenesis including kyanite + staurolite + garnet + biotite + muscovite + quartz ± chlorite ± plagioclase. The persistence of staurolite in the kyanite stability field implies the isograd reaction (e.g., Spear and Cheney, 1989; Fig. 2.7): staurolite + chlorite + muscovite + quartz -> kyanite + biotite + H 2 0 (2.5A). However, the reaction (Fig. 2.7), 26 staurolite + muscovite + quartz -> kyanite + garnet + biotite + H 2 0 (2.6A) can not be ruled out because kyanite is encountered only in one outcrop (the isograd-defining outcrop), in an assemblage which may represent the terminal A F M staurolite-breakdown reaction defined by reaction 2.6A. On the other hand, the lack of evidence for late garnet growth (e.g., Plate 5B), and the presence of a kyanite + staurolite zone of finite width immediately to the southeast of the study area (Figs. 2.2, 2.6; Leatherbarrow, 1981) is consistent with reaction 2.5A being the kyanite-in reaction in these rocks. Kyanite was not found in the western part of the study area (Fig.2.3). West of Big Fish Creek, a staurolite + biotite zone gives way northward to a zone where staurolite and fibrolitic sillimanite co-exist in apparent equilibrium (Fig. 2.6). The spatial association of staurolite, chlorite (after garnet), biotite and fibrolite, and in particular, intergrowths of sillimanite and biotite (Plate 7A), suggest the reaction: staurolite + chlorite + muscovite + quartz -> Alternatively, the occurrence of fibrolite and biotite intergrowths mimicking pre- to syn-kinematic garnet shapes (Plate 7A,B,C) could indicate that fibrolite grew on biotite pseudomorphs after garnet, produced by reaction 2.4. Kerrick (1987) postulated that fibrolite production after biotite is possible in the presence of Cl-rich fluids via a reaction such as: and that reaction 2.7 can occur at temperatures below coarse ("crystalline") sillimanite stability. Further north along Big Fish Creek, and north of the Bigmouth Creek stock are pelitic rocks characterized by the assemblage sillimanite + garnet + biotite. Staurolite is absent, and euhedral garnet overgrows fibrolitic to coarse-grained sillimanite + biotite aggregates (Plate 7B,C). At least some of these aggregates are pseudomorphs after early garnet as indicated by the preservation of inclusion-rich cores to these euhedral garnet porphyroblasts, which are locally rotated with respect to the external foliation, ST, and are interpreted at least in part to be of M l origin based on textural similarities with M l garnet elsewhere, such as splitting (Plate 7B) and the inclusion-rich, anhedral nature of these cores (Plate 7C). Other aggregates of sillimanite and biotite may be pseudomorphs after staurolite. Sillimanite is sillimanite + biotite + H 2 0 (2.5B). biotite + HC1 -> fibrolite + Si0 2 + KC1 + (Mg xFe,. x)Cl 2 + H 2 0 (2.7). 27 coarser in these rocks. It coexists with biotite as pseudomorphs of garnet ± staurolite and is a common matrix constituent (Plate 7D). These observations suggest the reaction (Fig. 2.7): staurolite + muscovite + quartz -> sillimanite + garnet + biotite + H 2 0 (2.6B). The observed mineral assemblages are inferred to be synchronous with M2 mineral assemblages found throughout the map area (Fig. 2.3) based on two lines of evidence. First, occurrences of S-r parallel, deformed sillimanite and biotite were encountered (Plate 7D), indicating that these assemblages developed before the end of ductile deformation. The second line of evidence hinges on the continuity of the mineral isograds they define with mineral isograds in adjacent areas. Based on the distribution of assemblages west of Big Fish Creek, isograds have been extrapolated to the northwest from the main study area, as illustrated in Figure 2.6. The mineral zones thus defined in the present study area tie in well with those of Gibson and Hoy (1985) to the northwest (Fig. 2.6). The spatial continuity of the minerals zones shown in Figure 2.4 suggests that they are related to the same metamorphic recrystallization episode. In the Mica Creek area (Fig. 2.2), the sillimanite isograd is sharp, defined by a zone a few tens of meters-wide at most, where kyanite and sillimanite co-exist (Ghent et al., 1977, 1980). In contrast, to the west of that area, kyanite and fibrolitic sillimanite co-exist over a 15 km-wide zone (Doucet et al., 1985; Sevigny and Ghent, 1986; and Digel et al., 1989; Digel et al., 1998; Fig. 2.2). One type of fibrolite occurrence described is "small knots intergrown with biotite" (Digel et al., 1989; p. 96), a texture very similar to that observed in the west and north of the Bigmouth Creek stock (Plates 7A,B,C). These authors postulated that fibrolite development may be related to temperature increase due to sporadic emplacement of small granitic bodies. Although kyanite was not observed to co-exist with fibrolite in the present study area, it occurs in the kyanite + staurolite zone more or less along isogradic strike of the first fibrolitic sillimanite occurrence found to the west of the Bigmouth Creek stock (Fig. 2.6). It is possible that recrystallization, resulting either from a temperature increase (e.g., Digel et al., 1989) or decompression, of kyanite-bearing rocks to produce fibrolite resulted in the complete resorption of kyanite, during M2 metamorphism. Alternatively, the presence of fibrolitic sillimanite rather than kyanite in rocks west of the Bigmouth Creek stock may reflect a regional pressure gradient along strike of the isograds during M2 metamorphism. 28 2.6.2.6 L a t e - M 2 andalusite Splays of large muscovite ± chlorite pseudomorphs after andalusite (Plates 4B) are widespread in the M2 staurolite and, locally, garnet zones (Fig. 2.8). Andalusite retrogression is generally complete, but fresh andalusite was observed in a handful of outcrops. On the outcrop, the splays of porphyroblasts appear relatively undeformed (Plate 4B); they are well exposed in mica-rich layers parallel to the dominant foliation, ST, but also cross-cut this foliation. In thin section, andalusite porphyroblasts include grains of earlier-formed M2 mineral assemblages (Plate 4C,D). Locally, these porphyroblasts are boudinaged, with resulting fractures at high angle to Sj (Plate 6C). Matrix phyllosilicates locally wrap around the porphyroblasts. Muscovite ± chlorite pseudomorphs after andalusite consist of aggregates of fine flakes which are aligned with S T to variable extents. In many cases, retrograde phyllosilicates showed strong alignment parallel to S T (Plate 4D). Where fresh andalusite and other M2 minerals are preserved, no evidence of intervening retrogression is observed (e.g., Plate 4C,D). Therefore, andalusite is interpreted to have developed in the late stages of M2 metamorphism. This interpretation is also consistent with the spatial correlation of occurrences of andalusite or pseudomorphs after andalusite, and the M2 staurolite zone (in particular, the fact that the appearance of andalusite coincides approximately with the M2 staurolite isograd; Fig. 2.8). Given the close spatial (Fig. 2.8) and textural (Plate 4C,D) association of andalusite and staurolite, the probable andalusite-forming reaction in these rocks is: staurolite + chlorite + muscovite + quartz -> andalusite + biotite + H 2 0 (2.5C). This reaction is shown in Figure 2.7 to proceed at pressure below 2.5 kbar. In contrast, M2 kyanite production through reaction 2.5A took place at pressure in excess of 6 kbar (Fig. 2.7). This implies a pressure differential from the low-grade half of the staurolite + biotite zone to the kyanite-in isograd (i.e., approximately 1500 m) of almost 4 kbar, or a minimum pressure gradient of 2.6 kbar/km, which is unreasonable for crustal rock densities. Therefore, production of andalusite in the staurolite zone via reaction 2.5C implies substantial decompression (in excess of 3 kbar) during the late stages of M2 metamorphism (Fig. 2.7), following attainment of conditions leading to kyanite production in higher-29 Figure 2.8 Occurrences o f late, coarse-grained andalusite and retrograde pseudomorphs after andalusite in the study area. 30 grade rocks. Figure 2.7 shows schematic M2 pressure-temperature paths for rocks of the staurolite and of the kyanite + staurolite zones which are consistent with this interpretation. As shown in this figure, some staurolite-zone rocks, especially in the western part of the study area, may have reached M2 peak temperature in the sillimanite stability field, consistent with the observation of co-existing staurolite and sillimanite west of Big Fish Creek (Fig. 2.6; Plate 7A). An alternative explanation for the presence of andalusite developed late in the metamorphic recrystallization history is that it is related to emplacement of an igneous body, such as the Bigmouth Creek pluton, within rocks of the Big Fish Creek area. Given the limited outcropping extent of the pluton (with a diameter of approximately 5 km; Fig. 2.6), it is unlikely that heat advected by this body would have effected substantial prograde recrystallization 3 km away, where andalusite is encountered. This conclusion is also supported by the lack of evidence for recrystallization to low-pressure assemblages of kyanite-bearing rocks in the M2 kyanite + staurolite zone, between the outcropping area of andalusite and the Bigmouth Creek stock. The interpretation of the presence of an intrusive body below (or above, and now eroded) the area characterized by andalusite, although somewhat ad hoc, can not be ruled out. However, occurrence of andalusite with kyanite inclusions at a locality approximately 10 km southeast of the Big Fish Creek area (M. Colpron, personal communication, 1994), suggests that andalusite development was regional in scale, and not restricted to the immediate vicinity of a hidden intrusive body. 2.7. Timing of Bigmouth Creek intrusion and metamorphism The timing of regional metamorphism on the west flank of the Selkirk Fan structure has been interpreted as mid-Jurassic based on observed timing relationship between deformation, metamorphism and emplacement of intrusive bodies which have yielded mid-Jurassic crystallization ages (e.g., Wheeler, 1965; Archibald et al., 1983; Brown et al., 1992a; Colpron et al., 1996). Wheeler (1965) and Van der Leeden (1976) have interpreted the age of emplacement of the Bigmouth Creek stock, exposed in the northern part of the study area, to be ca. 170 Ma based on compositional and textural similarities with other mid-Jurassic intrusions in the area. Van der Leeden (1976) interpreted the peak of regional metamorphism in the area to pre-date or to be synchronous with emplacement of the Bigmouth Creek stock, and based on this, inferred that peak regional metamorphism in the area was mid-Jurassic or older. 31 A sample of K-feldspar-megacrystic, hornblende- and biotite-bearing quartz monzonite from the Bigmouth Creek stock was dated by U-Pb geochronology of magmatic zircon and titanite. Instrumental analysis was performed by Richard Friedman and Jim Mortensen in the geochronology laboratory at the University of British Columbia and results are presented in Figure 2.9. Six zircon fractions define a linear trend with an upper intercept at 1.75 ± 0.02 Ga, and a lower intercept at 157 ± 3.3 Ma (MSWD = 182). The lower intercept age corresponds to the magmatic zircon crystallization age during emplacement of the Bigmouth Creek pluton; the upper intercept represents a component of inherited zircon that was present in all fractions. The two concordant titanite fractions gave z u o P b / Z J ° U ages of 140.5 ± 0.8 Ma and 137.4 ± 1.4 Ma. These results indicate that emplacement and crystallization of the Bigmouth Creek stock occurred in latest Middle to earliest Late Jurassic (i.e., 157 Ma). The younger titanite ages could be interpreted in two ways: 1) the pluton cooled very slowly to reach titanite U-Pb closure temperature (ca. 600 °C) approximately 17-20 m.y. after crystallization of zircons; 2) the pluton cooled to below 600 °C prior to 140 Ma, and titanite ages record resetting of the U-Pb system by a heating event at ca. 140 Ma. Although the data are permissive of both interpretations, the second scenario is tentatively favored for a number of reasons. Firstly, the presence of magmatic epidote in Bigmouth Creek granodiorite suggests that initial emplacement of the intrusive body was at depths well in excess of 20 km (i.e., pressure in excess of 6 kbar; Zen and Hammarstrom, 1984; Zen, 1985). Magmatic epidote is commonly observed in mid-Jurassic hornblende-bearing intrusions in the southern Omineca Belt, and has been interpreted by previous authors to be consistent with crystallization at substantial depth (Ghent et al., 1991; Colpron et al., 1996). Magmatic epidote in the Bigmouth Creek pluton implies that country rocks were at mid-crustal temperature and pressure conditions in late mid-Jurassic time. Because mid-Jurassic metamorphism is common in other parts of the SFA (e.g., Archibald et al., 1983; Brown et al., 1992; Colpron et al., 1996), M l metamorphism recorded in rocks of the study area is tentatively interpreted to reflect these mid-crustal conditions attained at about the time of emplacement of the Bigmouth Creek pluton (i.e., M l metamorphism is mid-Jurassic). Retrograde assemblages developed on M l garnet porphyroblasts, and on which M2 garnet overgrowths were overprinted (see Section 2.6.1 above) indicate that a significant cooling interval affected rocks in the area between M l and M2 prograde recrystallization events. M2 metamorphism therefore represents a distinct thermal pulse which is interpreted to be related to resetting of titanite U-Pb systematics in rocks of the Bigmouth Creek stock. In other words, the M2 metamorphic overprint in the area is interpreted as Early Cretaceous, the timing of this resetting. 32 Figure 2.9 U-Pb concordia plot o f zircon and titanite analyses for Bigmouth Creek K-feldspar megacrystic, hornblende and biotite-bearing granodiorite sample. Straight line is best fit linear regression to results obtained from five zircon fractions (A-F , a sixth fraction, C , was excluded); lower intercept is at 157 ± 3.3 M a , upper intercept is at 1.75 ± 0.02 Ga . Titanite results for two fractions are also shown ( T l and T2). They define ^ P b / 2 3 ^ ages o f 140.5 ± 0.8 M a and 137.4 ± 1.4 M a , respectively. 33 It should be noted that the above discussion and ensuing conclusions regarding the timing of M l and M2 metamorphic events are based on limited absolute timing constraints on an intrusive body, and that direct ages for metamorphic recrystallization in the country rocks are as yet unavailable. Therefore, interpretations based on these conclusions remain tentative until more extensive absolute timing constraints are available. 2.8 Discussion Rocks of the study area record a polycyclic Mesozoic metamorphic history. As discussed at length in the following chapter, ductile deformation associated with long-lived development and maintenance of a regional transposition foliation, ST, was attended and outlived by both M l and M2 metamorphic pulses. This is evidenced by the Sr-parallel inclusion trails in M l porphyroblasts (e.g., Plate 1A), and S T parallel alignment of retrograde minerals after M2 porphyroblasts (Plate 6B,D). Geochronological constraints on crystallization and subsequent thermal resetting of minerals of the Bigmouth Creek granitoid intrusion, as well as correlation with timing of metamorphism in nearby areas places a Middle to Late Jurassic age (ca. 157 Ma, the U-Pb age of zircon crystallization) on M l metamorphism, and suggests that M2 prograde recrystallization occurred in latest Jurassic to Early Cretaceous time (ca. 140 Ma, age of final U-Pb closure of titanite). This implies that rocks in and around the study area were part of a long-lived orogenic belt affected by episodic deformation and metamorphism over several tens of millions of years at least, in Jurassic to Cretaceous time. Evidence for multiple garnet growth episodes have been reported from the Northern Selkirk Mountains and nearby areas by previous works (immediately to the northeast, Perkins, 1983; in the North Adams river area to the north, Sevigny and Ghent, 1986; in the Cariboo Mountains to the northwest, Currie, 1988; in the Solitude Range to the east, Gal and Ghent, 1990; Whitney and Ghent, 1993). None of these authors related these textures to distinct, overprinting thermal pulses. On the other hand, geochronological arguments have been presented which suggest a complex and prolonged thermal history for rocks of the Selkirk Allochthon. Based on geochronological evidence, Scammell (1993) argued that rocks in the hangingwall of the Monashee Decollement, in the Monashee Mountains to the west of the present study area (Fig. 2.2) record diachronous metamorphism and deformation over a time interval in excess of 40 m.y., from ca. 140 Ma to ca. 100 Ma. More recently, Crowley et al. (1998) presented geochronological evidence for two distinct thermal pulses, of mid-Jurassic and mid-Cretaceous age, respectively, having affected rocks in the Mica Creek area, to the north of the study area. In addition, 34 Colpron (personal communication, 1996) obtained Late Cretaceous (ca. 75 Ma) Ar-Ar ages on micas from the Groundhog Basin, at the southern end of the study area. Based on barometry and Ar-Ar geochronometry, Colpron et al. (1996) argued that rocks in the vicinity of the Fang stock, to the south of the Big Fish Creek area (Fig. 2.2), underwent substantial and rapid exhumation immediately after regional metamorphism around 170-168 Ma. In contrast, development of kyanite-bearing assemblages during M2 in Early Cretaceous time in the Big Fish Creek area implies that if these rocks were substantially exhumed in mid-Jurassic time, they subsequently underwent significant burial between the Middle Jurassic and the Early Cretaceous. Alternatively, the discrepancy may represent spatially variable tectonic evolutions for rocks in different parts of the Selkirk Allochthon. Parrish (1995) integrated available geochronological information on the timing of regional metamorphism in the Selkirk Allochthon and argued that there exists a consistent relationship between timing of metamorphism and structural level in the Selkirk Allochthon: progressively structurally deeper rocks appear to record progressively younger regional metamorphic recrystallization. The results of this study are consistent with, and reinforce this interpretation of the spatial diachroneity of regional metamorphism by showing that the timing of M2 metamorphism in the northern portion of the southwest flank of the Selkirk Fan structure is intermediate between that of metamorphism in areas to the south, at higher structural levels (Figs. 2.1, 2.2; Archibald et al., 1993; Brown et al., 1992a; Colpron et al., 1996), and to the north and west, at lower structural levels (Figs. 2.1, 2.2; Sevigny et al., 1989, 1990; Scammell, 1993). However, this study shows that in addition to being spatially diachronous, metamorphism in the Selkirk Allochthon was episodic and polycyclic. Crowley et al. (1999) present geochronologic evidence for spatially limited and diachronous thermal events in the Monashee Mountains north of the Big Fish Creek area. These authors argue for local heat sources (e.g., plutonism) to account for this restricted spatial extent of thermal overprints. The Bigmouth Creek pluton is the only exposed magmatic body of volumetric significance in the area, and its emplacement may have contributed to the thermal event which led to prograde metamorphic recrystallization (e.g., M l , in mid-Jurassic time). On the other hand, despite the apparent geometric relationship between M2 metamorphic zones and isograds, and the outcropping area of the Bigmouth Creek pluton, it is unlikely that a pluton 5-6 km in diameter (see Fig. 2.6) would have a significant thermal effect on country rock located > 4 km away, which is the case for rocks which define M2 metamorphic zones and isograds in the Big Fish Creek area. It is possible that the Bigmouth Creek pluton 3 5 extends at depth below these rocks, or that there is a distinct intrusive body in that location, but the lack of surface exposure of these potential heat sources makes it impossible to assess their importance. On the other hand, spatial continuity of M2 metamorphic isograds in the area, with those mapped in adjacent areas to the southeast and northwest (Fig. 2.6; Van der Leeden, 1976; Leatherbarrow, 1981) suggests a regional cause for the thermal overprint. A tectonic model, based on numerical models of a critically-tapered, doubly-vergent orogenic wedge (Willett et al., 1993), is presented in Fig. 2.10 which integrates the foregoing observations and interpretations of metamorphic evolution of rocks of the study area. Docking of the Quesnellia Superterrane at the western edge of North America (e.g., Monger et al., 1982), and continued contraction, in Early to mid-Jurassic time, resulted in thickening of the North American miogeoclinal sequence into a critically-tapered, active tectonic wedge (Fig. 2.10A). Middle to Late Jurassic M l metamorphism resulted from thermal relaxation of this thickened wedge (e.g., England and Thompson, 1984), coeval with emplacement of the Bigmouth Creek pluton at depths well in excess of 20 km. Thermal relaxation at depth lead to establishment of a ductile crustal root, which may have allowed the maintenance of a stable plateau area, during continued contraction and eastward translation of Quesnellia and of the continental margin sedimentary sequence (Fig. 2.10B). Evidence of retrogression between M l and M2 prograde metamorphic events suggests that growth of the orogenic wedge was episodic, with intervening periods during which erosion (tectonic or otherwise) might have prevailed over contraction. Significant relaxation of far-field stresses, and/or expansion of the area of ductile lower crust due to M2-related heating in latest Jurassic-Early Cretaceous resulted in gravitational collapse of the wedge, accommodated at upper to middle crustal levels by orogen-parallel and orogen-transverse spreading, as demonstrated in the following chapter (Fig. 2.10C). In addition to abundant structural evidence presented in Chapter 3 for extensional ductile deformation, late-M2 development of andalusite supports the interpretation of tectonic erosion by extensional deformation. This model of a long-lived, dynamical orogenic wedge affected by compressional deformation due to overall crustal contraction, and by extensional deformation associated with gravitational collapse of the overly thickened wedge is in agreement with similar interpretations proposed recently for rocks of the Selkirk Allochthon in nearby areas of the Monashee Mountains (Scammell, 1993; Johnston, 1996; also see Brown et al., 1993). 36 A. Early to Middle Jurassic B. Middle to Late Jurassic C. Early Cretaceous j Continental crust j Ductile crustal root Continental mantle — - Instantaneous flow line Figure 2.10 Proposed dynamic wedge tectonic model for the Selkirk Allochthon to account for inferred metamorphic and structural evolution. Modif ied after Willett et al. , 1993. A : Early accretion of westerly-derived terrane to the western edge of North Amer ica and associated contraction leads to crustal thickening, thermal relaxation of thickened wedge, and results i n M l metamorphic event. B : A stable wedge geometry is established which is maintained by the balance o f far-field contractional stresses and expansion o f ductile crustal root due to continued thermal relaxation. Parts o f the orogenic wedge undergo burial while others are uplifted, accounting for l imited retrogression between M l and M 2 metamorphic event. Continued growth o f the wedge and associated thermal relaxation locally (at least) produce heating and result i n M 2 metamorphism C Relaxation o f far-field contractional stresses and/or expansion of the ductile crustal root leads to substantial gravitational collapse o f the dynamic wedge, during the late stages o f M 2 metamorphism. Resulting extension produces denudation and decompression as recorded by development o f late-M2 andalusite. 2.9 Conclusions This chapter presents the results of a field and laboratory study of the thermal and metamorphic evolution of rocks of the Selkirk Allochthon on the western flank of the Selkirk Fan structure of the southern Omineca Belt of British Columbia. Implications of these results are discussed in the previous section. The main conclusions derived from these results are listed below. 1. Two distinct metamorphic events affected rocks in the Big Fish Creek area: - M l , the earlier event, is recorded by relict porphyroblastic minerals which were subjected to cooling and retrogression prior to development of M2 prograde mineral assemblages. M l was probably regional in extent; - M2 is characterized by a Barrovian sequence of metamorphic zones in the Big Fish Creek area. From south to north, this field gradient includes the following metamorphic zones in rocks of pelitic composition: chlorite, biotite, garnet, staurolite, kyanite + staurolite. Late-M2 andalusite is widespread in rocks of the garnet and staurolite zones, suggesting substantial late-M2 decompression. 2. The Bigmouth Creek stock was emplaced in late Middle Jurassic as indicated by magmatic zircon U-Pb ages (157 ± 3.3 Ma). Magmatic epidote in Bigmouth Creek granodiorite suggests that emplacement depth was substantially greater than 20 km. Earliest Cretaceous U-Pb ages (140.7 ± 0.8 Ma and 137.4 ± 1.4 Ma) from titanite from the granodiorite are interpreted to be due to thermal resetting at that time. 3. Based on evidence for widespread mid-Jurassic metamorphism in other areas of the Selkirk Allochthon, and for deep emplacement of the Bigmouth Creek pluton at that time, M l metamorphism is interpreted to be of mid-Jurassic age. 4. M2 is interpreted to be the result of the same thermal pulse which caused resetting of U-Pb systematics of titanite in the Bigmouth Creek pluton in Early Cretaceous time Thus, timing of M2 metamorphism is Early Cretaceous. 5. Ductile deformation associated with a long-lived transposition foliation was ongoing, in that part of the Selkirk Allochthon, from prior to the peak of M l metamorphism in mid-Jurassic time, to after the peak of M2 in Early Cretaceous time. 38 6. Consistent with the foregoing conclusions, the Selkirk Allochthon in and around the study area is interpreted to represent a dynamic orogenic wedge which actively reacted to changes in external and internal parameters (far-field stresses and rheological modifications) over at least 20 m.y. in Jurassic to Cretaceous time, and probably longer, as suggested by constraints from adjacent areas. This model is supported by the structural analysis of these rocks presented in the following chapter. 39 Chapter 3 O R O G E N - P A R A L L E L AND O R O G E N - T R A N S V E R S E S Y N - C O M P R E S S I O N A L E X T E N S I O N , AND L O N G - L I V E D S T R U C T U R A L E V O L U T I O N O F T H E S E L K I R K A L L O C H T H O N O R O G E N I C W E D G E , BRITISH C O L U M B I A , C A N A D A 3.1 Introduction In the last 15 years, the concept of a fold and thrust belt as a dynamic, critically-tapered wedge (Davis et al., 1983; Dahlen et al., 1984; Dahlen, 1984) has been applied to whole orogenic belts (e.g., Malavieille, 1984; Stockmal et al., 1986; Jamieson and Beaumont, 1988; Willett et al., 1993). In parallel, evidence for syn-compressional extension in a number of orogenic belts has been assigned to maintenance of a stable ("critical") geometry in response to crustal thickening by compression, and associated rheological changes due to thermal relaxation following thickening (e.g., Burchfield and Royden, 1985; Jamieson and Beaumont, 1988; Holdsworth and Grant, 1990; Hodges and Walker, 1992; Hodges et al., 1993; Diez Balda et al., 1995). Recently, Scammell (1993) brought to light evidence for long-lived deformation resulting in sustained transposition of structural elements in rocks at the base of the Selkirk Allochthon (Read and Brown, 1981), an orogenic wedge exposed in the southeastern Canadian Cordillera (Fig. 3.1). Scammell (1993) interpreted extensional structures in these rocks as syn-compressional, and proposed a tectonic model involving extrusion (e.g., Burchfield and Royden, 1985; Thompson et al., 1997) at the base of the Allochthon by contemporaneous east-directed thrusting linked to development of the Rocky Mountain Foreland belt to the east, and extension along a west-vergent ductile shear zone (also see Johnston, 1996). Scammell (1993) argued that the Selkirk Allochthon represents a long-lived dynamic wedge, and that syn-compressional extension recorded near the base of the wedge is the result of gravitational collapse following overthickening. This extrusion model for the base of the Selkirk Allochthon, however, contrasts with interpretations of the deformation at higher levels in the Selkirk Allochthon wedge in terms of discrete events linked to specific changes in the geometry of the wedge boundaries (e.g., Brown et al., 1986; Price, 1986; Brown and Lane, 1988; Colpron et al., 1998). Long-lived, "dynamic" maintenance of an active orogenic wedge is also difficult to reconcile with a recent interpretation by Parrish (1995) of eastward, 40 20 km 'Allochthon., •c r M3BS KB i 40 LEGEND Paleogene Kamloops Gp volcanics Middle Cretaceous I.'. 1 granitoids Middle to Late Jurassic 1 x "~ I granitoids Upper Trlasslc-Lower Jurassic I v N Nicola arc volcanics Devonian-Triassic Kaslo Gp, Harper Ranch Gp Devonian-Early Tertiary (undif.) I+ + l granitoids Upper Proterozoic-Paleozoic Lardeau Gp, Mitford Gp Eagle Bay assemblage Upper Proterozoic-Lower Cambrian I- • -I HamillGp, BadshotFm Lower Proterozoic-Lower Cambrian Monashee complex, cover gneiss Upper Proterozoic Windermere Supergroup Lower Proterozoic Monashee complex, core gneiss Figure 3.1 Lithe-tectonic map of the Southern Omineca Belt. Thick black box shows location o f present study area. Gray boxes are locations o f recent studies in nearby areas: I. Scammell (1993); II. Colpron et al. (1996); HI. Johnston, 1996. A B : Adamant Batholith; B S : Bigmouth Creek Stock; B R B : Battle Range Batholith; C R F : Columbia River Fault; C T : Clachnacudainn Terrane; F S : Fang Stock; G B S : Galena Bay Stock; G P : Goldstream Pluton; K B : Kuskanax Batholith; M C : M i c a Creek area; M D : Monashee Decollement; M O : Monashee Complex; R: town o f Revelstoke. Modif ied after Wheeler and McFee ly (1991). 41 and downward progression, over time, of metamorphic and associated deformational activity in the Selkirk Allochthon. Of particular importance in resolving these apparent conflicts is the nature and timing of deformation in the internal parts of the orogenic wedge. Recognition of long-lived deformation, and in particular, cycles of compression and thickening, and resulting heating and gravitationally-driven extension, would favor a long-lived "dynamic wedge" model for the Selkirk Allochthon. On the other hand, evidence for discrete, and stylistically distinct deformation events, and for a spatial progression in the temporal distribution of metamorphic recrystallization events (i.e., structurally downward younging trend of metamorphism; Parrish, 1995) would tend to favor a model dominated by downward, and foreland migration of deformation and metamorphic "fronts" over time. This chapter presents a study of the structural elements observed in the Big Fish Creek area of the Selkirk Allochthon, in the Northern Selkirk Mountains of British Columbia (Fig. 3.1). Rocks in this area are of particular interest because they preserve the record of a protracted and complex history of metamorphic recrystallization (see previous chapter) which can be used to unravel the associated deformation events. This is in contrast to other areas in the Selkirk Allochthon where the record of deformation and metamorphism is less complete, either because of early cessation of deformation and metamorphic activity, or because of continued reworking until the late stages of orogenic evolution which obliterated the record of early evolution of the orogenic wedge. Structural elements and style observed at intermediate levels of the Selkirk Allochthon suggest a history of protracted ductile deformation in a regionally significant shear zone from as early as the Middle Jurassic, to at least as late, and probably later than the Early Cretaceous. As argued below, ductile deformation involved non-coaxial strain associated with eastward thrusting of the wedge, and with west-verging syn-'compression extensional strain internal to the wedge, as well as substantial sub-vertical shortening (i.e., orogen-parallel and orogen-transverse coaxial extension). The style of deformation in the Big Fish Creek area suggests that the Selkirk Allochthon orogenic wedge was a long-lived, dynamic entity affected by external and internal forces (terrane convergence, gravity), and by deformation (compressional and extensional) of the wedge in response to these driving forces. 42 3.2 Geological setting 3.2.1 Tectonic environment The study area is located within the southern Omineca Belt, an orogenic core zone developed as a result of collision and accretion to the western margin of North America of westerly-derived terranes during Mesozoic to Early Cenozoic time (Fig. 3.1; Monger et al., 1982). Rocks underlying the study area are part of the Selkirk Allochthon, a package of Late Proterozoic to Paleozoic North American continental margin rocks which was thrust eastward as much as 250-300 km during accretion and continued convergence (e.g., Price and Mountjoy, 1970; Read and Brown, 1981; Brown et al., 1993). Displacement of the Selkirk Allochthon has been interpreted to have taken place in part along the Monashee Decollement, a zone of intense ductile shear exposed to the west of the study area (Read and Brown, 1981; Journeay, 1986; Brown and Journeay, 1987). The Monashee Decollement is interpreted to separate the basement and cover sequence of the Monashee Complex structural culmination from the overlying Selkirk Allochthon (e.g., Read and Brown, 1981; Brown and Read, 1983; Journeay, 1986; Scammell, 1993), and to extend as the root zone of the basal decollement to the Rocky Mountain thrust and fold belt to the east (Brown et al., 1992b). Brown et al. (1992b) have also interpreted the Monashee Decollement itself as representing the roof thrust to a series of basement-rooted, east-directed thrusts, the combined displacement of which matches the inferred shortening in the Rocky Mountain foreland belt (also see Cook et al., 1992). Recently, however, some workers have questioned the existence of the Monashee Decollement as a discrete shear zone (e.g., Williams and Jiang, 1999). Regionally, rocks of the Selkirk Allochthon record a complex and protracted deformation and metamorphic history associated with Mesozoic convergence, and regional extensional deformation at least as early as the Late Paleocene (e.g., Parrish et al., 1988; Carr, 1991, 1992; Parrish, 1995). The stratigraphic and metamorphic framework of rocks in and around the study area are briefly reviewed below. The structural evolution of these rocks is the main focus of this chapter. 3.2.2 Stratigraphy Rocks in the study area have been alternatively interpreted as belonging to the Upper Proterozoic Windermere Supergroup (e.g., Wheeler, 1965; Brown et al , 1977, 1978; Brown and Tippett, 1978) or the Lower Paleozoic Lardeau Group (e.g., Brown and Lane, 1988; Logan and Colpron, 1995). The controversy stems in large part from the fact that these rocks are strongly deformed, recording intense 43 isoclinal folding, non-coaxial shearing and transposition, imparting significant uncertainty to attempts at regional correlation. In addition, few stratigraphic top indicators have been observed, and because of widespread isoclinal folding, it is extremely difficult to ascertain the large-scale stratigraphic facing in the area. In any case, detailed stratigraphic correlation is beyond the scope of this study and the "stratigraphic" framework described below does not rest on any interpretation of relative age: rock units are distinguished in terms of their respective structural, not stratigraphic, levels. The rocks mapped can be subdivided into four broadly defined packages separated by gradational contacts (Fig. 3.2). The structurally lowest package comprises centimeter- to meter-thick psammitic and semi-pelitic layers typically separated by thin (mm) mica-rich laminae. The lower boundary of this package was not identified within the bounds of the study area. Of particular note is the presence, near the top of this package, of several intervals of coarse gritty psammite, up to several meters in thickness, characterized by quartz and feldspar porphyroclasts which are commonly highly strained. These gritty intervals are found across the study area and hence are useful markers (Fig. 3.2). This psammite/semi-pelite package is structurally overlain by 100-200 meters of dominantly pelitic schist compositionally layered on a cm- to dm-scale. Calc-pelitic schist and brown-weathering impure marble become abundant towards the top of this pelitic interval. This transition defines a gradational contact with the overlying package, which consists of approximately equal amounts of pelitic, and calc-pelitic schist, with sporadic occurrences of cm- to dm-thick impure marble layers, and more rarely, of massive chlorite schist or amphibolite, depending on metamorphic grade (Fig. 3.2). In the northern part of the area, this calc-pelitic package gives way up-section to a sequence of brown-weathering, impure to relatively pure calcite, and dolomite marbles, associated with more amphibolite and locally, thin intervals of talc-schist (Fig. 3.2). Particularly noticeable are thick (several tens of meters) intervals of rusty brown-weathering dolomitic marble. The Bigmouth Creek pluton (Wheeler, 1965) is exposed at the northern limit of the study area. It consists of K-feldspar porphyritic, magmatic epidote-bearing, hornblende granodiorite. Late Jurassic to Early Cretaceous U-Pb ages on zircon and titanite from the Bigmouth Creek granodiorite were obtained as part of this study (see previous chapter). As discussed in the latter part of this chapter, the presence of magmatic epidote, indicating emplacement of the granodiorite body at substantial depth, along with the U-Pb ages provide a lower limit on the timing of a high-pressure metamorphic event which affected rocks of the study area (see next section). This has implications for the timing of extensional deformation and associated decompression in the area. These points are discussed in Section 3.4. 44 I u "5. \ \y LEGEND Bigmouth Creek granitoid chlorite schist/amphibolite impure limestone/marble gray calcite limestone/marble calcic pelite pelitic schist l~"J gritty psammite semi-pelite psammite ^ foliation (ST) strike and dip Figure 3.2 Lithologic map of the study area showing position of M2 mineral isograds. Stratigraphic intervals described in text are indicated by thick brackets. 45 3.2.3 Metamorphism The previous chapter presents a detailed description of the distribution of metamorphic assemblages, the metamorphic reaction history, and an interpretation of the metamorphic evolution of rocks of the Big Fish Creek area. The salient points are summarized here. An early metamorphic event, M l , is preserved as cores to composite, polycyclic garnet porphyroblasts in pelitic rocks. A later, medium-pressure prograde recrystallization event (M2) resulted in a sequence of metamorphic zones developed in rocks of pelitic composition, ranging from a chlorite zone in the south, to a kyanite + staurolite zone in the northern part of the Big Fish Creek area (Fig. 3.2). Extensive retrograde replacement of M l assemblages preceded M2, indicating that the two metamorphic events were separated by an episode of cooling. Because of the extensive retrograde destruction of M l assemblages prior to M2, the intensity and spatial distribution of prograde M l recrystallization can not be accurately established. The M2 metamorphic overprint is well preserved. Andalusite is observed to overgrow M2 mineral assemblages in the north-central part of the area, and is inferred to represent decompression near or following the thermal peak of M2 metamorphism. Integration of these and other observations on the metamorphic history of the study area suggests that it was complex, polycyclic and protracted. The temporal and spatial relationships between deformation and metamorphic recrystallization are discussed in Section 3.4.2.1 below, where it is shown, based on textural criteria, that ductile deformation attended and outlasted prograde metamorphism in the area. In addition, there appears to be a close spatial and temporal correlation between intensity of metamorphism and style and intensity of deformation. 3.3 Structural geology 3.3.1 Previous work The Big Fish Creek area lies on the southwest flank of the Selkirk Fan axis (SFA), a zone of structural divergence defined by a NW-trending axis separating rocks of southwest structural vergence to the southwest, from rocks of northeast structural vergence to the northeast (Fig. 3.1; Wheeler, 1963, 1965; Brown and Tippett, 1978). Since Wheeler's (1963, 1965) early work, a number of authors have described the structural geology and interpreted the deformational history of the Selkirk Fan in terms of discrete, spatially correlatable, deformation phases (e.g., Franzen, 1974; Van der Leeden, 1976; Tippett, 1976; Brown et al., 1977, 1986; Brown and Tippett, 1978; Read and Brown, 1979; Simony et al., 1980; Raeside 46 and Simony, 1982; Perkins, 1983). In this scheme, the overprinting of two generations of folds of opposite vergence (northeast and southwest) and of variable spatial intensity produces the structural fan (Brown and Tippett, 1978; Price, 1986; Brown et al., 1986). A significant problem with this approach is that structures assigned to a given deformation phase based on similarity of style are known to be widely diachronous. For instance, on both flanks of the Selkirk Fan, D2, the second phase of deformation, is interpreted to pre-date, or be coeval with the main phase of metamorphic recrystallization (Van der Leeden, 1976; Brown and Tippett, 1978; Leatherbarrow, 1981; Simony et al., 1980; Perkins, 1983; Sevigny and Simony, 1989). However, this metamorphic event is thought to be of Jurassic age on the west flank of the SFA (Archibald et al., 1983; Brown et al., 1992a; Colpron et al., 1996), whereas on the east flank, it is interpreted to be of mid- to Late Cretaceous age (e.g., Sevigny et al., 1989, 1990). Parrish (1995) integrated geochronological data for the Selkirk Allochthon and the underlying Monashee Complex and argued that metamorphism and associated deformation are progressively younger with structural depth, consistent with foreland-propagating deformation in the Selkirk Allochthon wedge. However, as described in the previous chapter and briefly summarized above, the history of prograde metamorphism and associated deformation in the study area is complex and probably spanned a substantial temporal range from mid-Jurassic to Cretaceous time. The relatively extensive record of metamorphism preserved in rocks of the Big Fish Creek area provides an excellent opportunity to address the question of the relative timing of different structural elements. The fact that they record at least two distinct prograde regional metamorphic pulses is atypical of rocks in nearby areas. This is either because later metamorphism in other areas was sufficiently intense to mask evidence of earlier metamorphism (e.g., Scammell, 1993; Crowley et al., 1998); or because these areas did not see as complex a metamorphic history and were only affected by one metamorphic event (e.g., Colpron et al., 1996). Whereas the mineralogical signature for polycyclic metamorphism may be cryptic, recent geochronological evidence suggests that high grade metamorphic rocks of the Selkirk Allochthon north and west of the present study area were, in fact, affected by at least two distinct thermal events during Mesozoic time (Sevigny et al., 1989; 1990; Scammell, 1993; Crowley et al., 1998). As discussed at length below, many structural elements encountered in the study area were produced by progressive deformation and transposition over a significant period of time, as constrained by the syn-deformational metamorphic recrystallization history. This poses an additional problem when attempting the spatial correlation of structures based on style because similar structural elements may 47 have developed at widely different times (e.g., Williams and Zwart, 1977; Williams, 1985; Tobisch and Paterson, 1988, 1990). Furthermore, the diachronous nature of both deformation and metamorphism tempers the significance of temporal correlations of deformation and metamorphic events based on textural arguments. These problems are further addressed in the following sections. 3.3.2 Structural elements in the Big Fish Creek area The structural grain of the Big Fish Creek area is controlled by the distribution of a moderate- to shallow-dipping penetrative foliation (Fig. 3.3 A; see Appendix IV for listing of all field-based structural measurements). Closely associated with this foliation are tight to isoclinal folds, mineral and stretching lineations, extensional crenulation and shear bands. These early structures are associated with the main phase of deformation in this area. They are overprinted by at least two distinct sets of late crenulations. The following sections present a description of these different structural elements. Although no obvious structural discontinuity was encountered in the field, it will be argued that rocks of the study area represent part of a zone of intense distributed strain of probable regional significance. As discussed below, the dominant foliation is a transposition foliation developed during protracted non-coaxial deformation. In the following sections, distinction is made between structures formed prior to or during, and those formed after, the development of the transposition foliation. 3.3.2.1 Dominant foliation, ST The dominant outcrop-scale structural element in the Big Fish Creek area is a penetrative foliation defined by compositional layering and the alignment of minerals with grain-shape preferred orientations (e.g., Plate 8). Previous workers (Ghent et al., 1977; Simony et al., 1980; Raeside and Simony, 1983; Dechenes and Simony, 1984; Doucet et al., 1985; Sevigny and Simony, 1989) have ascribed the development of this foliation in areas in and around the Northern Selkirk Mountains to transposition resulting from superposition of two temporally distinct, but coaxial folding events and used an Si+2 notation to highlight its composite nature. In the Big Fish Creek area, the dominant foliation is locally discontinuous and is commonly observed to be axial planar to tight to isoclinal, commonly intrafolial, and locally refolded folds with attenuated to rootless limbs (Plates 8A, 9A). As discussed below, ST is associated with intense and protracted deformation. For these reasons, it is interpreted to represent a transposition foliation because of its complex and composite nature (e.g., Plates 9A,D, 10; Tobisch and Paterson, 1988, 1990). Following Williams and Campagnoni (1983), it is referred to herein as S T- Observations presented in this and the following sections further support this interpretation. 48 Figure 3.3A Attitudes of structural elements in the study area. Lithological contacts and distribution of ST attitudes are shown on map. Stereographic projection of transposition related structures: poles to ST and ST', FT, LT and stretching lineations. 49 Figure 3.3B Attitudes of structural elements in the study area. Late structures: intersection lineations associated with late crenulations, and late fold axes. 50 Scammell (1993) reached similar conclusions regarding the nature of the dominant foliation in higher-metamorphic-grade rocks of the Monashee Mountains to the west (Fig. 3.1). Measured attitudes of S T are shown in Figure 3.3A. In general, the strike of S T varies from northeasterly in the south to northwesterly in the north, and the dip is moderate to shallow (10-40°). This spatial variation defines a broad, markedly non-cylindrical, approximately upright antiform (Fig. 3.3A) which had been assigned to a regional third phase of folding by previous authors (e.g., Van der Leeden, 1976; Brown et al., 1977; Brown and Tippett, 1978). The antiformal trend is not regular: a relatively abrupt change in attitude of S T is observed around the position of the M2 staurolite-in isograd (Figs. 3.2, 3.3A). North of this line, the attitude of S T is much less variable than south of it. As discussed below, S T resulted from continued transposition of a variety of fabric elements as a result of protracted and intense general shear. These fabric elements, which comprise tight to isoclinal asymmetric shear folds and associated axial planar foliation (ST', preferred alignment of platy minerals, crenulation, and crenulation cleavage), as well as extensional shear bands (or S-C'-type fabric), developed at different stages of superimposed cycles of transposition of S T as illustrated schematically in Figure 3.4. In thin section, the dominant foliation is typically highlighted by the preferred orientation of platy minerals, and compositional layering (Plates 10, 11C,D). It is commonly irregular to anastomosing, and forms an S-C'-type fabric where anastomosing segments show continuity across the main trend of the foliation, or where discrete shear bands are developed (Plates 9D, 11C,D, 12C). Often, a folded pre-existing foliation is preserved in microlithons bounded by laminae defining a later foliation (Plate 10). The degree of overprinting of the earlier foliation by the later foliation is variable, the later foliation varying in style from gentle crenulation (Plate 10A,B) to well-developed crenulation cleavage, to essentially completely transposed foliation (e.g., Plate 10C,D; c.f., Bell and Rubenach, 1983; Tobisch and Paterson, 1988, 1990). The co-planarity, at the scale of the outcrop and larger, of the dominant foliation in rocks which show different stages of the transposition cycles at the fine scale indicates that strain involved continued rotation of newly formed foliations towards a plane parallel to the dominant foliation (Tobisch and Paterson, 1988), which is interpreted to parallel the local shear plane (e.g., Williams 1983). These observations highlight the spatially heterogeneous nature of finite strain resulting from progressive deformation. Tobisch and Paterson (1988) found similar features to those described above in rocks from a ductile shear zone in the Lachlan Fold Belt and interpreted them to be the result of spatially heterogeneous development of overprinting transposition cycles during progressive deformation. These authors described the complex nature of such foliations in terms of consecutive cycles of evolving 51 Figure 3.4 Schematic representation of cyclical transposition and interpreted relationship between ST, S x \ FT T , and extensional shear bands. A: General overview of geometry. The structural elements are developing in a macroscopic shear fold. B: Small F T shear folds develop as a result of localized inhomogeneous shear at different scales. S x ' is axial planar foliation to F T . With continued shear, F T tightens and its axial plane, as recorded by ST', rotates towards the "regional" foliation, ST. C: With continued shear and rotation, ST ' approaches S T and continued rotation is precluded. Continued shear is accommodated by development and motion along extensional shear bands (ESB), unless or until a new F T forms and a new cycle of transposition is initiated. 52 foliation from an early continuous cleavage (Sc) to different degrees of crenulation of this cleavage to produce a later crenulation or crenulation cleavage (Scr, Scc). They also demonstrated the spatial heterogeneity of evolution through such cycles, as a result of strain partitioning over areas as small as a thin section. The above nomenclature is adopted here to describe components of the composite transposition foliation, ST (Plates 9D, 10). Inclusion trails in porphyroblasts (S;) show variable angular relationship with respect to the external matrix foliation (Se; Plates 9D, 12, 13A,B). Sj in biotite (e.g., Plate 12) is usually more strongly rotated with respect to S e than is S; in other porphyroblastic minerals (e.g., Plates 9C,D, 13 A), although the extent of rotation of Sj in biotite is commonly irregular at the scale of a thin section (Plate 11C). Locally, Sj reflects different stages of the transposition cycle than those recorded by S e (Plate 9D; also see Plate 1A,B). This indicates that the relative timing of development of the dominant foliation with respect to the growth of metamorphic minerals is spatially heterogeneous (see Section 3.4.2.1 and previous chapter). As the prograde growth of metamorphic minerals was probably more or less synchronous, this implies that deformation was diachronous and heterogeneous at the scale of the thin section. This diachroneity and spatial heterogeneity of the deformation are further evidence of transposition due to progressive deformation (Williams, 1967, 1985; Hobbs et al., 1976; Tobisch and Paterson, 1988; Mawer and Williams, 1991). In the more intensely metamorphosed rocks in the northern part of the study area, compositional layering (which defines S T in part) is commonly characterized by mm- to cm-thick biotite-rich laminae interlayered with quartz-rich laminae (e.g., Plates 9A, 14A). In particular, rocks which bear evidence of intense strain usually show extensive differentiation of biotite-rich and quartz-rich laminae (Plate 14A). These observations suggest that progressive deformation associated with the development of S T involved dissolution and fluid-aided material remobilization (e.g., Williams, 1972). 3.3.2.2 "FT", tight to isoclinal folds associated with S T As previously noted, Sj is axial planar to tight to isoclinal, intrafolial folds developed at different scales (Plates 8 A, 9 A, 15). Intrafolial isoclinal folds of concordant quartz veins are particularly abundant (Plate 15). These folds generally have strongly attenuated to discontinuous limbs (Plates 9A, 15C), and in places, are seen to be intrafolially refolded (Plate 9A). Given the progressive nature of the deformation associated with these folds and their axial planar foliation, ST, it is likely that stylistically similar folds in these rocks did not all form at the same time (Park, 1969; Roberts, 1977; Williams and Zwart, 1977; 53 Williams and Campagnoni, 1983; Williams, 1985; Tobisch and Paterson, 1988, 1990; Mawer and Williams, 1991). For this reason, and because they are associated with ST, these folds are referred to herein as " F T " . Generally, F T folds plunge gently E N E to WSW, but some dispersion is observed (Fig. 3.3 A). Most F T axes are sub-parallel to mineral and stretching lineations (Fig. 3.3A). Folds with axes parallel to stretching lineations are common in compressional orogenic settings (e.g., King and Rast, 1956; Bryant and Reed, 1969; Olesen and Sorensen, 1972; Borradaile, 1972; Hobbs et al., 1976; Williams and Zwart, 1977; Williams and Campagnoni, 1983; Holdsworth, 1989; Alsop and Holdsworth, 1993; Alsop et al., 1996, 1998). They are often interpreted as resulting from progressive rotation of folds initially at high angle to the stretching lineation by continued deformation (e.g., Bryant and Reed, 1969; Borradaile, 1972; Sanderson, 1973; Escher and Watterson, 1974; Williams and Campagnoni, 1983; Holdsworth, 1989). In this scheme, dispersion in fold axis orientation is either due to irregular initial orientations, or to continued fold generation during axis rotation (Borradaile, 1972; Sanderson, 1973; Escher and Watterson, 1974; Ramsay, 1979). However, other mechanisms have been proposed whereby folds may form with axes initially close to the main transport direction (or stretching lineation; e.g., Butler, 1982; Rattey and Sanderson, 1982; Coward and Potts, 1983; Northrup and Burchfield, 1996; Grujic and Mancktelow, 1995). Rattey and Sanderson (1982) and Coward and Potts (1983) suggested that orogen-transverse folds may form at the lateral tips of thrust sheets as a result of differential orogen-transverse movement, and Butler (1982) documents orogen-transverse folds in lateral ramps to structural horses formed as a result of impingement by a laterally adjacent, younger horse. On the other hand, Northrup and Burchfield (1996) interpret orogen-transverse folds to result from orogen-parallel flow due to vertical strain partitioning in a transpressional setting. In all of these cases, orogen-transverse folds would be expected to show a preferred sense of vergence (e.g., Rattey and Sanderson, 1982). In contrast, rotation of folds with axes initially perpendicular to the transport direction into transport-parallel orientation by passive rotation during continued shear should result in a more or less symmetric distribution of fold vergence (cf. Coward and Potts, 1983). A few rare northwest-southeast trending (i.e., orogen-parallel), tight to isoclinal folds were encountered in the study area (Plate 8B), usually folding impure limestone layers. These folds are typically west-vergent (Plate 8B) and have been interpreted by previous workers to be related to early, west-vergent, regional recumbent folds on the western flank of the Selkirk Fan structure (e.g., Raeside and Simony, 1983; Brown and Lane, 1988; Colpron et al., 1998). No definite occurrence were encountered of non-cylindrical fold with sheath-like geometry, which would be expected if axis rotation 54 produced the transport-parallel fold axes and the observed dispersion in hinge orientations (Cobbold and Quinquis, 1980; Coward and Potts, 1983). 3.3.2.3 Mineral/intersection lineation, L T and transient foliation, ST' In upper biotite- and higher grade rocks of the study area, cigar- (prolate) to elongate penny-shaped (oblate) biotite porphyroblasts are widespread. The long axis of these biotite grains typically have a preferred orientation which defines a distinct mineral lineation on ST-parallel surfaces (Plate 9B). Where oblate biotite grains dominate, the lineation is defined by the intersection of those grains with S T (Plate 11). Whereas in many cases the platy grains do not show a preferred orientation about the axis of elongation (Plate 11C), in a number of instances, a preferred orientation was observed, defining a distinct foliation (Plate 1 IB). This foliation is variably developed at the scale of the outcrop, and is generally discontinuous from layer to layer. A schematic representation of the observed attitudes of lineation-defining biotite is shown in Figure 3.5. Because the mineral/intersection lineation formed during progressive deformation associated with transposition (see below), it is referred to as L T . Similarly, the foliation defined by aligned platy biotite porphyroblasts represents a transient stage of a transposition cycle involved in progressive ST development (Tobisch and Paterson 1988, 1990), and is labeled ST'. Macroscopically, L T lineations plunge gently ENE to WSW (Fig. 3.3A). As mentioned in the previous section, L T and stretching lineations, as well as most F T fold axes show sub-parallel orientations (Plates 9B, 15B). In the several instances where a well-defined ST' was observed, it dipped consistently to the southeast (i.e., NW-vergence; Fig. 3.3A). In thin section, biotite porphyroblasts are cigar- to fish-shaped, and typically asymmetric (Plates 10, 12). Inclusion trail foliation in biotite porphyroblasts (Sj) is typically continuous with, but inclined relative to S e (e.g., Plates 10A,B, 12). The angle between Sj and S e usually varies from grain to grain (Plates 10A,B, 11C, 12C,D; also see Plate 3D) suggesting that the extent of relative rotation of different biotite porphyroblasts with respect to S e was heterogeneous. Some randomly oriented flakes have S; parallel to S e indicating that they did not rotate relative to Se; these grains are generally blocky and do not show strong asymmetry (Plate 11C). In cuts normal to both S T and the mineral lineation, there is generally a preferred sense of rotation of the porphyroblasts, as indicated by the angle between S e and S; and by the asymmetry of fish-shaped grains (e.g., Plates 10C,D, 11C; Lister and Snoke, 1984), although within the same section, some flakes may show opposing sense of rotation or vergence with respect to S T (Plates 11C,D, 12A,B,D). In sections normal to S T and parallel to L T , long axes of porphyroblasts are 55 Figure 3.5 Schematic representation o f the attitude of lineation-defming biotite porphyroblasts S x is parallel to top surface o f block. A : Biotite flakes are irregularly fanned about an axis parallel to L T . Although there is a dominant sense of vergence with respect to S x , both senses o f vergence are encountered. B : Biotite flakes show co-planarity, defining a foliation, S x ' . The intersection o f this foliation with S T is Lr. 56 generally sub-parallel to S T (Fig. 3.5A) although some asymmetry is also usually still preserved, and Sj, where observed, is at high angle to Se, indicating significant rotation of the grains (Plates IOC, 1 ID). Where the sense of rotation can be determined, it is mostly consistent from grain to grain, although grains showing opposite sense of rotation are also encountered in LT-parallel sections (Plate 12A). These observations lend insight into the formation of the L T lineation. Blocky biotite porphyroblasts which grew at high angle to S T (e.g., Plate 11C) were progressively rotated towards parallelism with it. These grains were also affected by crystal plastic deformation: their common asymmetric fish-shape habit (e.g., Plates 10, 11C,D, 12; e.g., Lister and Snoke, 1984) may have resulted from slip along basal cleavage planes; in other cases, deformation bands are observed in biotite porphyroblasts (Plate 12B). Hence, development of the biotite lineation in rocks of the study area involved non-coaxial internal strain, and rigid-body rotation, as illustrated schematically in Figure 3.6. Where rotation was complete, the biotite grains lie sub-parallel to ST. A mineral lineation (LT) is visible on S T surfaces because the grains have been strained and rotated until their long axes are parallel to the maximum stretching direction (Fig. 3.6C). In some cases, large biotite flakes have their long axis parallel to, and are bounded by a crenulation foliation developed during the early to intermediate stages of a transposition cycle (Plate 10A,B; Tobisch and Paterson, 1988), and this may explain the preferred sense of vergence of biotite flakes observed in some rocks: these biotite grains may have grown with a preferred orientation parallel to the transient crenulation cleaveage. In other cases, however, large biotite flakes appear to have grown in more or less random orientation with respect to S T or associated foliations developed during transposition (e.g., Plate 11C). The fanning of biotite porphyroblasts, and of other matrix minerals with shape-preferred orientation, about an axis parallel to the direction of L T (Plates 10A,B, 11, 12; Fig. 3.5) is interpreted to indicate continued growth and rotation of newly formed biotite grains during progressive deformation and this is supported by Se-S; relationships. The presence of late, undeformed and unrotated biotite co-existing with extensively rotated and deformed biotite (e.g., Plate 11C) suggests that locally at least, biotite growth outlasted progressive deformation associated with L T development. The fish-like morphology of biotite flakes and their rotation towards parallelism with S T suggest that extensional strain was associated with the development of the biotite mineral lineation (e.g., Lister and Snoke, 1984). As discussed below, biotite "fish" are commonly associated with occurrences of extensional crenulation or shear bands (Piatt and Vissers, 1980) in the study area (see Section 3.3.2.6 below), which reinforces this conclusion. The asymmetry and vergence of fish-shaped micas can be used to infer the sense of non-coaxial shear in these rocks during development of the mineral lineation (Lister 57 Figure 3.6 Schematic representation of deformation of lineation-defining biotite flakes; S T plane is parallel to the top of the block diagram. A: Initially undeformed and randomly oriented flake B: Rotation towards the S T (flattening) plane and plastic deformation of the flake to produce fish shape. C: Rotation towards parallelism with L T lineation and plastic deformation. The sequence does not represent a time series, as rotation towards S T and L j are probably in part at least synchronous. 58 and Snoke, 1984). On the other hand, the common occurrence of biotite flakes showing opposite senses of rotation within the same sample (Plates 12A,B,D) indicates that deformation related to the development of the biotite lineation involved a pure shear component normal to S T (i.e., flattening; Piatt and Vissers, 1980). 3.3.2.4 Stretching lineation in quartzo-feldspathic rocks A stretching lineation is commonly developed in mica-poor quartzo-feldspathic rocks of biotite-and higher metamorphic grade in the study area, as recrystallized ribbon quartz and stretched feldspar clasts with high aspect ratios (Plate 16), which locally constitutes the dominant structural element of the rock. This stretching lineation generally plunges ENE to WSW and the majority of the stretching lineations measured show parallelism with F T fold axes and L T lineations (Fig. 3.3A). In thin sections, strained psammite from the M2 staurolite zone is characterized by coarse quartz grains with local polygonal outline, and ribbons of large elongate quartz grains separated by high angle grain boundaries which are approximately orthogonal to ST (Plate 16B). Generally, these grains do not show a preferred crystallographic orientation. These textures are interpreted to represent recrystallization of quartz under conditions of low strain rate (static) or high temperature, or both. Commonly, these grains have lobate and sutured outlines, and preserve evidence of crystal plastic strain (deformation lamellae, sub-grain boundaries) and strings of fine grains are developed at grain boundaries, indicating substantial subsequent dynamic recrystallization (Plate 16B). Feldspar porphyroclasts commonly have a high density of sub-grain boundaries and dynamically recrystallized grains at grain boundaries; fine, dynamically recrystallized grains form long, lineation-parallel strings (Plate 16). These textures indicate that strain involved crystal plastic deformation of both quartz and feldspar. Crystal plasticity in feldspar suggests deformation at temperatures characteristic of middle to upper amphibolite conditions (e.g., Tullis and Yund, 1987). This indicates that at least some of the deformation resulting in the stretching lineation was active at or near metamorphic peak conditions in the area. This, and the parallelism of the stretching lineation with the mineral/intersection lineation indicates that the stretching lineation is also the result of the progressive deformation which produced the transposition-related structures described in the previous sections. Locally, stretched feldspars are boudinaged into angular fragments with intervening gaps at high angle to the dominant foliations (Plates 13D, 14D), suggesting that extensional deformation of feldspar probably continued during cooling from metamorphic peak conditions. 59 3.3.2.5 Fractured porphyroblasts and porphyroclasts Split or fractured porphyroblasts are commonly observed in rocks at or above garnet grade. Fractured garnet porphyroblasts are abundant, occurring in the majority of garnet-grade and higher garnet-bearing rocks (see Appendix II for occurrences of split garnet in hand samples and outcrops). Split and/or fractured garnet porphyroblasts are recognizable in the field (Plate 9A,B); split or fractured porphyroblasts of biotite (Plates 12C, 13 A), staurolite (Plate 13B), kyanite, and andalusite (Plate 13C) were also identified in thin section. In general, gaps or fractures are at high angle to S T (Plates 9, 13). In the field, the gaps generally trend N-S, and dip steeply, a trend approximately perpendicular to lineations and F T fold axes (e.g., Plate 9B). Different co-existing minerals commonly show a similar orientation of gaps or fractures (e.g., Plate 13 A) and, except for biotite (see previous section), have not been substantially rotated following splitting or fracturing (Plates 9A,B,C). This suggests that porphyroblast splitting primarily represents a coaxial process (i.e., splits are tensional fractures). A retrograde aggregate consisting of chlorite + quartz + muscovite ± biotite ± plagioclase is commonly developed in gaps of split garnet porphyroblasts (see Plate IB, previous chapter) and of muscovite ± chlorite ± biotite in gaps of split staurolite (Plate 13B) suggesting that hydration during retrogression may have facilitated splitting. In many cases, syn- or post-splitting relative displacement of grain fragments seems to have taken place (e.g., Plate 9C; also see Plate IB). As discussed at length in the previous chapter, early (Ml) garnet porphyroblasts were split prior to, or during a cooling and retrogression event which resulted in the growth of retrograde assemblages in the gaps between split fragments (Plate IB). This retrogression event was followed by M2 prograde recrystallization, during which porphyroblasts developed which were subsequently split (e.g., M2 staurolite and kyanite, and andalusite, Plate 13B,C). This implies that the strain regime responsible for the development of fractures and splits at high angle to S T was active, episodically at least, from prior to the high-pressure peak of M2 until after late-M2 decompression which resulted in crystallization of andalusite. As discussed in the previous section, boudinaged feldspar porphyroclasts were also observed in strongly lineated psammitic rocks (Plates 13D, 14D). 3.3.2.6 Extensional crenulation and shear bands Mesoscopic extensional shear bands and extensional crenulations (e.g., Berthe et al., 1979; Piatt and Vissers, 1980; Lister and Snoke, 1984) are relatively common in rocks of the study area (Plate 14). 60 At the hand sample and microscopic scale, these surfaces are outlined by gentle, approximately symmetric to gently asymmetric microfolds of the dominant foliation with which they merge and anastomose, commonly producing an S-C fabric (Piatt and Vissers, 1980; Lister and Snoke, 1989; Twiss and Moores, 1992, p. 273; Plates 11C, 12C, 14B). Shear bands and extensional crenulation cleavages are commonly closely associated with lineation-defining biotite porphyroblasts (e.g., Plates 11C, 12C). In many cases, strain partitioning resulting from the presence of porphyroblasts (e.g., Lister and Williams, 1983; Bell, 1985) appears to have favored the development of extensional crenulations or shear bands (e.g., Plate 12C,D). As with biotite porphyroblasts, extensional crenulation cleavage or shear bands appear to have undergone variable extents of rotation towards parallelism with ST (the shear plane; e.g., Tobisch and Paterson, 1988), and the angle between these surfaces and S T is usually lower in LT-parallel than in LT-normal cuts (Plate 11C,D). Similar features were interpreted by Tobisch and Paterson (1988) as recording overprinting of different generations of foliations (e.g., layering, crenulation, crenulation cleavage, etc.) during progressive deformation. The fact that extensional shear bands and crenulation cleavage are visible in both LT-parallel and LT-normal cuts perpendicular to ST (Plates 11C,D, 14C) indicates that deformation involved a component of coaxial shear during development of these structures (i.e., shortening perpendicular to ST). Insufficient observations are available to determine whether extensional shear bands systematically show a preferred sense of vergence. However, some notable mesoscopic shear bands in strongly sheared rocks appear to record orogen-parallel, top-to-the northwest displacement in the later stages of ductile progressive deformation (Plate 14A). 3.3.2.7 Southeast-trending crenulation and related folds Locally very intense crenulation of S T and earlier foliations has resulted in the development of a WNW-striking foliation which dips steeply to the northeast and generally intersects ST at high angle (Plate 8C). Intersection of this foliation with ST produces a gently ESE-plunging lineation which is locally very strong (Plate 8D; Fig. 3.3B). This crenulation foliation is axial planar to upright, generally open mesoscopic folds of S T (Plate 8C). The extent of overprint of S T by this foliation is very uneven. Its development is much more advanced in mica-rich rocks, and in general, low-metamorphic grade rocks in the southern part of the study area appears to be more strongly affected than high-grade rocks in the north. This crenulation has been interpreted to be related to large-scale, late NW- to WNW-trending folds of which the broad antiform defined by S T in the study area (Fig. 3.3) would be an example (Van der Leeden, 1976; Brown et al., 1977; Brown and Tippet, 1978). However, an alternative interpretation for the variable attitude of ST at the scale of the study area is discussed in Section 3.4.1 below. 61 3.3.2.8 Northeast-trending crenulation Locally, a crenulation intersection lineation is observed on S T surfaces at high angle to the ESE-trending crenulation lineation (Fig. 3.3B; Plate 8D). No foliation related to this crenulation lineation was measured, nor were any related folds observed, but Van der Leeden (1976; his D 5 structures) observed the same lineations and associated steep, NNE-striking, west-dipping crenulation foliation, which is axial planar to NNE-trending folds. Relative timing of the two crenulation foliations is difficult to establish and appears to vary from instance to instance. Van der Leeden (1976) interpreted them to be contemporaneous and to constitute a conjugate set, akin and of similar orientations to conjugate foliations reported by Simony and Wind (1970) for the Dogtooth range, to the southeast (also see Murphy and Journeay, 1982). 3.4 Discussion The bulk of the field and petrographic observations presented above suggests that rocks in the study area record intense, albeit spatially heterogeneous strain resulting from protracted progressive deformation attended by a complex history of metamorphic recrystallization. High-metamorphic-grade rocks in the northern part of the study area in particular, appear to be the locus of a zone of pervasively intense deformation and transposition, which arguably constitutes a regionally significant shear zone. As discussed below, strain was non-plane, at least episodically, comprising an important component of shortening perpendicular to the dominant, shallow-dipping regional foliation, ST, accommodated at least in part by orogen-parallel extension. The latter part of this discussion presents a tectonic model in which the observed strain features are explained by mechanisms of syn-contraction gravitational spreading ("dynamic spreading"; e.g., Holdsworth and Grant, 1990) of a thrust package during Mesozoic convergence of western North America with accreted terranes to the west. 3.4.1 Intensity of deformation The dominant structural feature of the study area is ST, the transposition foliation. Development of S T resulted from continued cyclical development and rotation of fabric elements (crenulation cleavage, cleavage-parallel mica flakes, extensional shear bands, etc.) towards the local shear plane (e.g., Fig. 3.4; Tobisch and Paterson, 1988, 1990; Mawer and Williams, 1991), which S T is interpreted to parallel. Evidence for progressive rotation of fabric elements towards a regionally consistent S T is pervasive, indicating that this also represents a regional shear plane (e.g., Escher and Watterson, 1974; Williams, 62 1983, Mawer and Williams, 1991). Hence, the progressive and cyclical nature of ST-related transposition is interpreted to approximate a steady-state process (Means, 1981; Williams, 1983). Although this implies that estimating finite strain precisely is impossible, the achievement of a steady-state foliation itself requires that rocks in the study area were affected by intense strain comprising an important non-coaxial component (Means, 198i). Other evidence for high strain in these rocks includes the widespread, albeit heterogeneous occurrence of a number of fabrics usually associated with intense deformation. Although mesoscopic evidence for high strain does include local occurrences of rocks characterized by a mylonitic foliation parallel, or at low angle to layering, and defined in part by regular, shear-induced compositional segregations (Plate 14A) and/or a strong stretching lineation (L-tectonite; Plate 16), most of the features indicative of high strain are more subdued, and suggest distributed, pervasively intense strain at different scales. Some of these features are mesoscopic: isoclinal, locally refolded intrafolial folds with attenuated to sheared-out limbs with axial plane parallel to the dominant foliation (Plates 8A, 9A; e.g., Williams, 1967; Williams and Zwart, 1977; Mawer and Williams, 1991); fold axes parallel to the stretching lineation (Plates 8A, 15B; e.g., Roberts and Sanderson, 1974; Escher and Watterson, 1974; Passchier, 1986; Mawer and Williams, 1991); cross- and anastomosing layering and related extensional shear bands (e.g., Plates 11C, 12, 14A,C; Hobbs, 1965). In addition to evidence for cyclical development of S T , and co-planarity of S T with the regional plane of non-coaxial shear, a number of fabric elements supportive of intense strain are widespread at the hand-sample to microscopic scale. These include obliqueness of newly-crystallized micas with respect to the mesoscopic dominant foliation (e.g., Plates 11, 12; Passchier, 1986; Mawer and Williams, 1991); anastomosing foliation (Plates 10C,D, 11C,D, 12A,C; Piatt and Vissers, 1980; Lister and Snoke, 1984; Passchier, 1986); extensional crenulations, or extensional shear bands, which in many cases produce an S-C fabric (Plates 11C, 12C, 14B,C; e.g., Piatt and Vissers, 1980; Lister and Snoke, 1984; Hanmer and Passchier, 1991); fish-shaped micas (Plates 10C,D, 11, 12; e.g., Lister and Snoke, 1984); fragmented rigid elements (Plates 9, 13, 14D; e.g., Ramsay and Sturt, 1970; Simpson and Schmid, 1983); pressure shadows around rigid grains (Plates 12C, 13 A; e.g., Simpson and Schmid, 1983); delta- and/or sigma-type rigid porphyroclasts (Plate 14D). All these features seem to confirm the interpretation that the study area is a zone of intense distributed ductile strain at all scales. Although the high-strain features just mentioned occur throughout the study area, the extent to which they are developed is spatially variable. The M2-staurolite isograd in the study area approximately 63 coincides with a relatively sharp gradient in the pervasiveness of the different fabric elements; north of the isogradic surface (Figs. 3.2, 3.3 A), rocks show much more consistent co-planarity of fabrics at the hand-sample to microscopic scale than they do south of it. Although high strain features (see above) do occur south of this surface, they are less pervasively developed, suggesting that strain intensity was generally lower than in rocks to the north. The more irregular attitude of foliation elements in the south also indicates that ST-normal shortening was less important than in the north. The macroscopic distribution of S T attitudes is also consistent with this interpretation. Upgrade of the lower part of the garnet zone, S T is much more regular than to the south (Fig. 3.3A), consistent with the interpretation that S T more closely approaches a regional shear plane in the high-grade northern half of the study area, than in the lower-grade southern half. The greater competency of the low-grade rocks in the south may have precluded the development of a steady-state transposition foliation and complete rotation of structural elements towards the shear plane (i.e., the ductility of the rock was exceeded at lower strain than that required for establishment of a steady-state foliation; c.f. Means, 1981). This interpretation for the spatial variability of ST attitudes differs from that of previous authors who attributed it to late (post-ST development) folding (e.g., Van der Leeden, 1976; Brown et al., 1977). Rocks in the Big Fish Creek area therefore, are interpreted to be part of a zone of intense ductile strain, a ductile shear zone, here called the Big Fish Creek shear zone (BFSZ), characterized by spatial gradation in the nature and intensity of strain. As discussed in Chapter 2, development of S T , and hence movement along the BFSZ, was at least in part contemporaneous with M2 metamorphism. The coincidence of the strain gradient with the M2 staurolite isograd suggests that metamorphic recrystallization may have served to facilitate ductile deformation in these rocks. 3.4.2 Nature of deformation 3.4.2.1 Orogen-transverse shear Progressive deformation associated with the development of the transposition foliation in the study area involves an important component of orogen-transverse shear evidenced by the strong ENE-trending stretching lineation developed in quartzo-feldspathic rocks (Section 3.3.2.4; Plate 16). This stretching lineation corresponds to the orientation of maximum stretch (i.e., the X-axis of the finite strain ellipsoid) associated with constrictional strain during progressive deformation. The general orientation of gaps in split and micro-boudinaged garnet and other porphyroblasts at high angle to the stretching lineation (Section 3.3.2.5) is consistent with important constrictional strain. Many split porphyroblasts do not show substantial rotation synchronous with, or following splitting (Plates 9C, 13), which indicates 64 that orogen-transverse shear comprised a component of co-axial stretching. Orogen-transverse shear could also account for parallelism of fold axes, as well as the preferred orientation of the long axis of elongate biotite grains with the stretching lineation, as a result of rotation from an initial orientation at high angle to the maximum stretching direction (e.g., Sanderson, 1973; Escher and Watterson, 1974). However, an alternative explanation for this is discussed below which also accounts for the north-south-directed (i.e., orogen-parallel) shear also evident in rocks of the Big Fish Creek area. 3.4.2.2 Orogen-parallel shear Some fabric elements suggest that deformation associated with progressive development of S T involved a component of orogen-parallel stretching. The most direct evidence are mesoscopic north-dipping ductile extensional shear bands which intersect S T at low-angle and disrupt it, locally producing an anastomosing pattern of foliation (an S - C fabric; Piatt and Vissers, 1980; Lister and Snoke, 1984; Plate 14A). Extensional shear bands and extensional cleavage (Piatt and Vissers, 1980) are also observed at the microscopic scale, and are particularly well-expressed in sections cut normal to both ST and the stretching lineation (Plates 12C, 14C), lending further support for substantial orogen-parallel extensional strain in rocks of the area. The preferred sense of vergence of these and other (see below) fabric elements suggests that non-coaxial, top-to-the-north shear was important, although the locally symmetric nature of shear bands encountered in ST- and LT-normal sections suggests a component of coaxial, orogen-parallel stretching (see below; Piatt and Vissers, 1980). The structures described above are the result of ductile strain, and therefore indicate that at least some of the orogen-parallel deformation occurred prior to substantial cooling of the host rocks following the peak of M2 metamorphism. Furthermore, rotation of these structures towards parallelism with S T implies that orogen-parallel deformation was at least in part responsible for the progressive development of S T itself. Several other ductile fabric elements encountered in the area may in part be explained by orogen-parallel shear. Figure 3.6 illustrates how ST, ST' and F T are interpreted to be related to each other. As mentioned above, it is possible that ST' and F T fold axes were rotated into parallelism with the stretching lineation (e.g., Sanderson, 1973; Escher and Watterson, 1974). However, it is also allowable that ST' and F T resulted from orogen-parallel shear without substantial rotation, such as would be expected at the lateral tip of a spreading nappe (Fig. 3.7). This interpretation is consistent with other evidence for northwest-verging ductile deformation in the form of mesoscopic shear bands and other N-S-directed extensional features described above. This interpretation is also consistent with the observation of the preferred top-to-the-northwest sense of vergence indicated by ST' and F T and other structures (e.g., Plate 65 Figure 3. 7 Deformation at the frontal and lateral tips of a spreading nappe. Dashed outline is initial state. A: Cross-sectional view. B: Plan view. C: Different components of strain predicted; arrow points to direction of nappe movement: i. unstrained control volume; ii. transport-parallel (orogen-transverse) simple shear; iii. transport-normal (orogen-parallel) simple shear; iv. flattening parallel to regional shear plane (orogen-parallel and orogen-transverse coaxial extension). 66 14A). In contrast, rotation of these structures from initial orogen-parallel attitudes does not explain this asymmetry (i.e., produces sheath fold geometry; c.f. Coward and Potts, 1983). 3.4.2.3 ST-normal shortening (sub-vertical "flattening") In addition to orogen-parallel and orogen-transverse non-coaxial strain, a number of observations suggests that ductile progressive deformation involved a component of coaxial shortening perpendicular to ST, or the regional shear plane (i.e., a component of flattening strain), especially in the high-grade, northern portion of the study area. Evidence includes split porphyroblasts with resulting gaps at high angle to S T showing little evidence for rotation during or following splitting, and the locally symmetric rotation of planar structural elements such as biotite porphyroblasts and extensional crenulations and shear bands with respect to S T . In particular, the fact that biotite porphyroblasts show symmetric rotation in sections normal to ST and either normal or parallel to the stretching lineation indicates that extension was both orogen-parallel and orogen-normal (i.e., non-plane). Locally striking evidence of compositional segregation (Plates 9A, 14A) suggests that solution-aided mass transfer may have contributed to this component of co-axial deformation. 3.4.2.4 Transport direction The foregoing discussion highlights the fact that ductile strain associated with ST and related structural elements in the study area consisted of orogen-parallel and orogen-normal general non-coaxial strain. In addition, evidence has been presented which indicates that strain was both temporally and spatially heterogeneous. Hence, the present study area is interpreted as part of a long-lived zone of intense general non-coaxial shear of regional significance, the Big Fish Creek shear zone. The preferred facing of transient biotite foliation and of extensional shear bands, and the vergence of folds with orogen-transverse axes indicate that orogen-parallel transport involved top-to-the-north-northwest displacement. Given the prevalence of orogen-transverse structures, and the extent of rotation of structural elements towards L T , or towards S T in LT-parallel sections, it is clear that orogen-transverse shear was more important than orogen-parallel shear in the study area. The sense of displacement associated with the orogen-transverse component of shear is more difficult to ascertain, in large part because of loss of asymmetry in LT-parallel sections, of kinematic indicators such as rotated biotite porphyroblasts, due to nearly complete rotation of these structural elements into parallelism with S T (e.g., Plates 10C,D, 1 ID, 12A). In other areas of the southern Omineca Belt, orogen-transverse displacement 67 associated with similar structures (stretching and mineral lineations, orogen-transverse fold axes) is generally interpreted to reflect top-to-the-east transport (e.g., Read and Brown, 1981; Journeay and Brown, 1986; Journeay, 1986; Lane et al., 1989; Scammell, 1993; McNicholl and Brown, 1995; Johnston, 1996; Schaubs and Carr, 1998), and this may also apply to the present study area. In a limited number of instances, however, observed kinematic indicators appear to indicate top-to-the-west sense of shear (e.g., Plate 14D), which is in agreement with observations by Scammell (1993) and Johnston (1996), for rocks of the Monashee Mountains to the west, and of the Thor-Odin culmination, to the southwest, respectively, of the Big Fish Creek area (Fig. 3.1). The tectonic model proposed in Figure 3.8 implies an upper-plate-up-to-the-east sense of shear for the BFSZ (i.e., thrust motion), with an overlying (and/or partially superimposed), syn-thrusting, west-verging extensional shear zone (here correlated with the Norman Wood fault exposed to the east of the study area; Perkins, 1983; Scammell, 1993). An alternative interpretation would be an upper-plate-down-to-the-west shear zone associated with syn-orogenic extrusion of the underlying rocks, as proposed by Scammell (1993) and Johnston (1996) for other areas of the Selkirk Allochthon. The tectonic model presented in Figure 3.8 is further discussed in Section 3.4.3, below. 3.4.2 Timing of deformation 3.4.2.1 Relative timing of deformation and metamorphic recrystallization Rocks of the study area record a complex and protracted history of prograde metamorphism (see Chapter 2). In pelitic rocks, metamorphic assemblages define a Barrovian sequence characterized by the presence of staurolite and kyanite in the higher grade rocks. Andalusite and pseudomorphs after andalusite (Plate 4B) are widespread in the central portion of the study area and their development post-dates the establishment of the Barrovian field gradient, as evidenced for instance by inclusions of early staurolite in andalusite porphyroblasts (Plate 4C,D). The above observations indicate that substantial decompression occurred between the onset of metamorphism in the area and the late stages of prograde metamorphism highlighted by the growth of andalusite. Although the history of metamorphic recrystallization in the area is complex, the nature of the attending deformation was relatively constant. Extensively recrystallized pelitic rocks in the northern portion of the study area preserve the most complete record of metamorphism and hence, of its relationship to deformation. The features described in the following are characteristic of these rocks. Evidence for the earliest prograde metamorphism (Ml) consists of anhedral cores of composite garnet 68 w E Figure 3.8 Simplified structural cross-section and tectonic model for the Selkirk Allochthon and Selkirk Fan at the latitude o f the present study area. B i g Fish Creek shear zone is a top-to-the-east contractional shear zone at mid-level in the Selkirk Allochthon, in the hangingwall o f the Monashee Decollement. It is active at least as early as the Late Jurassic as east-directed wedging results in westward overthrusting, and thickening o f the western flank o f the Selkirk Fan. B y middle to late Cretaceous time, eastward movement along the B F S Z may still be occurring. B y that time, substantial thinning at mid-levels on the west flank o f the Selkirk Fan has occurred, probably by a combination o f lateral extrusion, and extensional movement along top-down-to-the-west ductile shear zone(s) such as the Norman Wood Creek fault. B F S Z : B i g F ish Creek shear zone; C R F : Columbia River fault; F C F : French Creek Fault; M D : Monashee Decollement; N W F : Norman Wood Creek Fault; PT: Purcell Thrust. Modif ied after Colpron et al., 1998. 69 porphyroblasts which represent early porphyroblasts resorbed and dynamically split (or micro-boudinaged) prior to being overgrown by euhedral, M2 garnet rims (e.g., Plate 9A,B,C; Chapter 2). Development of M2 garnet overgrowth was followed by growth of M2 staurolite. Large andalusite porphyroblasts include, and therefore developed later than, this staurolite. Although M2 garnet tends to be inclusion-poor, abundant inclusions in M l garnet cores typically define an internal foliation (Si) which is typically straight and at low angle to the external foliation, S e (= ST; Plates 9C,D, 13A). In some cases, anhedral garnet cores preserve an internal foliation which appears to represent an earlier stage of S T development (Plate 9D; also see Plate IB). In the less common case where Si is gently curved, it tends to line up with S e near the margin of the core (Plate 13 A). These observations indicate that M l garnet growth was synchronous with, or post-dated the development of ST, the dominant foliation. M l garnet porphyroblasts were split and resorbed prior to being overgrown by M2 euhedral rims. Typically, they are split into two sub-equant fragments with regular intervening gaps (e.g., Plate 2B). As previously described, the gaps generally lie perpendicular or at high angle to ST, and to the direction of maximum finite stretching inferred from stretching lineations. This implies that progressive development of ST and the stretching lineation was at least in part synchronous with garnet splitting. Small euhedral M2 garnet porphyroblasts and M2 euhedral garnet overgrowths on resorbed M l garnet cores are conspicuous. This textural discontinuity between core and overgrowth in composite garnet grains is associated with a sharp compositional discontinuity (Chapter 2, Plate 2). M2 garnet is typically inclusion-poor and spatially restricted to resorption rinds around early garnet, making the timing of its growth with respect to matrix foliation development difficult to establish. However, it is commonly associated with (and in some cases included in) inclusion-rich, euhedral M2 staurolite (Plate 9C) which is locally rotated with respect to, and deflects, foliation, suggesting that both M2 garnet and staurolite developed prior to the cessation of progressive deformation associated with development of ST. The final stage of prograde metamorphism observed in the area is the development of large (up to 15 cm-long; Plate 4B), elongate andalusite porphyroblasts (typically replaced by a muscovite-rich aggregate during subsequent retrogression) which locally include porphyroblasts of M2 garnet and/or staurolite (Plate 4C,D). A number of observations indicate that ductile deformation associated with development of ST attended and/or post-dated growth of andalusite in the area. Andalusite crystals typically lie with their long-axis parallel to ST, either as isolated grains, or as groups of needles defining conspicuous splays ("chickenfeet" texture; Plate 4B). Needles do not have a preferred orientation on ST. Where fresh andalusite is found in thin section, it generally shows evidence of micro-boudinage, with 70 gaps at high angle to S T (Plate 13C). Finally, muscovite in retrograde muscovite-rich aggregates after andalusite is typically aligned parallel to ST (Plate 6D), indicating that ST-related progressive deformation continued until after growth, and during or after retrogression of andalusite in the area. Taken together, the foregoing textural observations indicate that ST-related deformation in the study area, like metamorphic recrystallization, was protracted and attended by important temporal variations in ambient pressure and temperature conditions. Although few absolute time constraints on prograde metamorphism and deformation are available, these important fluctuations in metamorphic conditions suggests that prograde metamorphic recrystallization and the associated deformation, spanned a substantial temporal range. The absolute timing of deformation and metamorphism in the area is discussed below. 3.4.2.2 Absolute t iming of deformation Extensive deformation and metamorphism in the Selkirk Allochthon and adjacent areas of the southern Omineca Belt initiated in Early to mid-Jurassic time as a result of accretion of the Quesnellia island arc terrane to the western edge of North America (Monger et al., 1982; Brown et al., 1986, 1992a; Archibald et al., 1983; Colpron et al., 1996). Although direct supporting evidence is lacking, M l metamorphism in the study area is interpreted to be of that age or younger. The Bigmouth Creek pluton, a magmatic epidote-bearing alkali feldspar-phyric, hornblende granodiorite (Wheeler, 1965) intrudes supracrustal rocks exposed at the northern edge of the study area. Zircon and titanite U-Pb systematics were analyzed as part of this study and results are presented and discussed at length in Chapter 2. Salient points are briefly summarized here. Five discordant zircon separates yielded a lower intercept age of 157 ± 3.3 Ma, and an upper intercept age of 1.75 ± 0.02 Ga. Two titanite fractions were also analyzed which yielded concordant, but not overlapping, ages of 140.5 ± 0.8 Ma and 137.4 ± 1.4 Ma, respectively. As discussed in the previous chapter, these results are interpreted to reflect emplacement and final crystallization of the Bigmouth Creek stock in late Middle Jurassic time, as implied by the zircon lower intercept age. M l metamorphism is interpreted to be broadly synchronous with emplacement of the stock in the Middle Jurassic, based on the fact that mid-Jurassic metamorphism is common in other parts of the Selkirk Allochthon (e.g., Archibald et al., 1983; Brown et al., 1992a; Colpron et al., 1996). U-Pb resetting of titanite in Early Cretaceous time is interpreted to have resulted from a thermal pulse at that 71 time which also produced the M2 metamorphic overprint. The presence of magmatic epidote in the Bigmouth Creek granodiorite indicates minimum final emplacement pressure probably well in excess of 6 kbar (i.e., >20 km depth; Zen and Hammarstrom, 1984; Zen, 1985) in late Middle Jurassic time. The occurrence of M2 kyanite-bearing assemblages indicates that rocks in the northern part of the study area were still deeply buried and that the thickness of that part of the Selkirk Allochthon orogenic wedge was still considerable in Early Cretaceous time. This is in contrast to rocks of the Illecillewaet Synclinorium to the south, in the vicinity of the Fang granodiorite stock where, Colpron et al. (1996) have argued, Barrovian (moderate to high-P) regional metamorphism, and subsequent rapid exhumation and decompression to ca. 3 kbar were mostly over by 168 Ma (mid-Jurassic). In the present study area, the persistence of high pressures to Early Cretaceous time implies that the late M2 decompression associated with development of andalusite, and the associated deformation must be younger (i.e., Early Cretaceous or younger). Thus, ductile deformation in the study area began in the Middle Jurassic, following accretion of the Quesnellia island arc terrane, and continued, episodically at least, until the Early Cretaceous or later. Late Cretaceous (ca. 76 Ma) muscovite and biotite Ar-Ar cooling ages obtained for rocks in the southern part of the study area (Colpron, personal communication, 1996) suggest that final exhumation and cooling did not occur until 50 m.y. later, at least. 3.4.3 Tectonic implications The structural style which resulted from ductile deformation in the Big Fish Creek area is similar to that observed in structurally deeper rocks near the base of the Selkirk Allochthon, throughout the southern Omineca Belt. Features encountered in the Big Fish Creek area which are also widespread in other areas include evidence for long-lived transposition of structures towards a shallow-dipping regional shear plane, a strong ENE to WSW-trending mineral/stretching lineation, orogen-transverse, stretching lineation-parallel fold axes, and evidence for substantial syn-compressional extension parallel to the stretching lineation (e.g., Journeay, 1986; Journeay and Brown, 1986; Lane and Brown, 1988; Scammell; 1993; McNicoll and Brown, 1995; Johnston, 1996; Schaubs and Carr, 1998). Both Scammell (1993) and Johnston (1996) have interpreted extensional strain at the base of the Selkirk Allochthon as resulting from large-scale syn-compressional extension associated with extrusion (e.g., Burchfield and Royden, 1985; Thompson et al., 1997) of a mid-crustal wedge as a result of gravitational collapse. According to these authors, extension took place prior to cessation of eastward 72 displacement of this wedge in the Early Tertiary (Carr, 1992), as recorded by the continued development of the regional ENE-WS W-trending orogen-transverse lineation. Eastward displacement of rocks exposed in the Big Fish Creek area is also evidenced by the widespread occurrence of an ENE-WSW-trending orogen-transverse lineation, illustrating the distributed nature of deformation associated with east-directed transport in the lower part of the Selkirk Allochthon from Late Jurassic to Early Tertiary time. This and the preferred vergence of structures associated with orogen-parallel shear support the interpretation that these rocks were located near the base, and at the lateral edge of a mid-level, east-directed thrust nappe within the Selkirk Allochthon (c.f., Butler, 1982; Coward and Potts, 1983; Fig. 3.7) during mid-Jurassic to Early Cretaceous time at least. Evidence for ST-normal shortening and transport-parallel extension prior to Late Cretaceous cooling, and contemporaneous with apparently rapid decompression (as recorded by the late- M2 development of andalusite) is consistent with an evolution involving extension and crustal thinning (e.g., Jamieson and Beaumont, 1988; Hodges et al., 1993; Thompson et al., 1997) during continued eastward displacement of the Selkirk Allochthon. Syn-compressional extension affected rocks in higher structural levels of the Selkirk Fan as early as the Middle Jurassic (Colpron et al., 1996) as the Selkirk Fan structure was developing into a thickened, doubly-vergent orogenic wedge (Malavieille, 1984; Willett et al., 1993; Brown et al., 1993) as a result of eastward underthrusting of a westerly-derived crustal slab below the western edge of the wedge (Figs. 3.8A, 3.9A; Brown et al., 1986, 1993; Price, 1986; Colpron et al., 1998) and east-vergent overthursting of the eastern part of the wedge over the North American continental margin. The lack of evidence for significant post-mid-Jurassic penetrative deformation and metamorphism in rocks of the Illecillewaet Synclinorium (Colpron et al., 1996) suggests that from the Late Jurassic on, rocks at high structural levels in the Selkirk Fan underwent mostly passive eastward translation as a result of movement along shear zones at deeper structural levels (Fig. 3.8). In contrast, Scammell (1993) has argued that rocks near the base of the Selkirk Allochthon to the west of the Big Fish Creek area record deformation associated with eastward displacement of the Allochthon from as early as 135 Ma to as at least as late as 100 Ma. As in the Big Fish Creek area at intermediate structural level, and in the Illecillewaet Synclinorium at high structural level, evidence for syn-compressional extension also exists at the base of the Selkirk Allochthon (Scammell, 1993; Johnston, 1996). Scammell (1993) has suggested the Norman Wood Creek fault, whose trace is exposed several kilometers east of the study area (Perkins, 1983), as a possible candidate for an upper-plate-down-to-the-west, syn-compressional extensional fault in the hangingwall of the Monashee Decollement, and this is 73 A. Early to Middle Jurassic B. Middle to Late Jurassic C. Early Cretaceous ] Continental crust ||^ Ductile crustal root Continental mantle Instantaneous flow line Figure 3.9 Proposed dynamic wedge tectonic model for the Selkirk Allochthon to account for inferred metamorphic and structural evolution. Modif ied after Willett et al. , 1993. A : Early accretion of westerly-derived terrane to the western edge of North America and associated contraction leads to crustal thickening, thermal relaxation o f thickened wedge, and results in M l metamorphic event. B : A stable wedge geometry is established which is maintained by the balance o f far-field contractional stresses and expansion of ductile crustal root due to continued thermal relaxation. Parts o f the orogenic wedge undergo burial while others are uplifted, accounting for limited retrogression between M l and M 2 metamorphic event. Continued growth o f the wedge and associated thermal relaxation locally (at least) produce heating and result in M 2 metamorphism C : Relaxation o f far-field contractional stresses and/or expansion of the ductile crustal root leads to substantial gravitational collapse o f the dynamic wedge, during the late stages o f M 2 metamorphism. Resulting extension produces denudation and decompression as recorded by development o f late-M2 andalusite. 74 tentatively integrated in the model presented in Figure 3.8, to accommodate evidence for ductile, west-vergent deformation encountered in the Big Fish Creek area. Thus it appears that the eastern part of the Selkirk Allochthon, in particular the Selkirk Fan and adjacent areas, represents a long-lived, doubly-vergent dynamic orogenic wedge (Malavieille, 1984; Willett et al., 1993) resulting from far-field compressional stresses induced by continued convergence of accreted terranes to the western margin of North America, and from internal deformation involving both compressional and extensional components, acting to maintain a critically tapered topography (Figs. 3.8, 3.9; e.g., Davis et al., 1983; Dahlen et al., 1984; Willett et al., 1993). A possible model for the evolution of the Selkirk Allochthon and Selkirk Fan which integrates the results of the present study is illustrated in Figure 3.9. Development of a doubly-vergent wedge began in Early to mid-Jurassic time (ca. 173 Ma; Colpron et al., 1996; 1998), but initially may have been limited in size as suggested by the lack of substantial sediment input to the foreland basin until Late Jurassic time (Price and Mountjoy, 1970). By mid- to Late Jurassic time (ca. 160-155 Ma), the eastern part of the wedge had overridden the basement ramp which defined the landward boundary of the thick Proterozoic to Early Paleozoic continental margin sequence (Fig. 3.8A; Colpron et al., 1998), and the resulting gravitational potential in the eastern part of the orogenic wedge was apparently balanced by a substantially thickened western part of the wedge, as evidenced by emplacement of the magmatic epidote-bearing Bigmouth Creek pluton in supra-crustal rocks at depths well in excess of 20 km. Thickening of the western part of the orogenic wedge was the result of underplating of the wedge by westerly-derived material along west-verging thrust faults such as the French Creek Fault (Fig. 3.8), synchronous with eastward displacement along deeper-level, ductile shear zones such as the Big Fish Creek shear zone. Dynamic maintenance, and possibly continued growth of the thickened wedge until the Early Cretaceous (Fig. 3.8B) is evidenced by kyanite-bearing M2 metamorphic assemblages. Eventually, thermal relaxation of the thickened wedge produced an increase in temperature in the mid- to lower levels of the wedge, leading to enhanced ductility, which in turn facilitated continued shear and progressive deformation along such a ductile shear zone, as well as gravitationally-driven ductile extension (Figs. 3.8B, 3.9C), thinning and resulting decompression at these structural levels, in response to the increased overlying load (i.e., "dynamic spreading" and extrusion; e.g., Holdsworth and Grant, 1990). 3.5 Conclusions Rocks in the Big Fish Creek area appear to have preserved a more complete record of deformation and metamorphism in the Selkirk Allochthon than rocks at higher structural levels (due to 75 their "passive" behavior after the Middle Jurassic) and at lower structural levels (due to the continued high-grade metamorphic overprint until the Late Cretaceous, which obliterated the record of earlier evolution). Some key features of this record had previously gone unrecognized. These include: 1) the polycyclic nature of prograde metamorphic recrystallization implying two episodes of heating to amphibolite facies conditions and an intervening episode of cooling; 2) the widespread occurrence, in higher grade rocks of the Big Fish Creek area, of late andalusite which indicates decompression late in the metamorphic recrystallization history; 3) evidence for very intense strain consistent with deformation in a zone of high distributed, general shear; 4) evidence for substantial shortening normal to the dominant shallow-dipping foliation (i.e., flattening) in the area; and 5) a Middle Jurassic emplacement age for the Bigmouth Creek Stock with evidence for themal reworking to Late Jurassic-Early Cretaceous time. These new lines of information suggests a metamorphic and associated tectonic evolution distinct from that proposed in earlier models involving simple, foreland-, and downward-propagation of metamorphic and deformation fronts within the southern Omineca Belt (e.g., Parrish et al., 1995), or cessation of tectonic and metamorphic activity by mid-Jurassic time in the Northern Selkirk Mountains (e.g., Colpron et al., 1996). Hence, the Big Fish Creek area provides information crucial to the unraveling of the tectonic evolution of the southern Omineca Belt in Mesozoic time. Based on the structural and metamorphic features observed in the Big Fish Creek area, the following major conclusions may be synthesized: 1. Rocks in this area record intense, non-coaxial strain which resulted from long-lived ductile progressive deformation, and involved important components of orogen-transverse and orogen-parallel extension (i.e., sub-vertical shortening normal to the dominant transposition foliation ST, which approaches the regional shear plane). 2. These rocks are therefore interpreted to be part of a ductile shear zone of regional significance, herein named the Big Fish Creek shear zone. The presence of this shear zone at intermediate levels in the Selkirk Allochthon illustrates the distributed nature of strain associated with east-directed displacement of the Selkirk Allochthon during Mesozoic to Early Cenozoic time, rather than its being restricted to movement along the sole thrust of the Allochthon, the Monashee Decollement. 3. Structural and metamorphic evidence for diachronous syn-compressional extensional strain and crustal thinning in the Big Fish Creek area, combined with similar interpretations at both higher, and lower structural levels by other workers, suggests that the Selkirk Allochthon, at least in the area of the Selkirk 76 Fan structure, constituted a dynamic orogenic wedge from as early as the Middle Jurassic, to at least as late as the Late Cretaceous. The dynamic behavior of this wedge was controlled by far-field forces associated with convergence of accreted terranes and the North American continental margin, by gravitational forces, and by rheological modification and internal deformation of the wedge in response to these different controls. 4. Based on U-Pb age dating of the Bigmouth Creek granodiorite, the high-pressure, early stage of regional metamorphism recorded in rocks of the area can be assigned to mid-Jurassic time. This implies that syn-metamorphic progressive deformation began at least as early as the Middle Jurassic. Thermal resetting, in the Early Cretaceous, of Bigmouth Creek stock titanite U-Pb systematics constrains the timing of M 2 metamorphism to Early Cretaceous time. Late-M2 decompression was therefore Early Cretaceous or younger, and predates the end of ductile deformation associated with the transposition foliation in the area. This also puts an Early Cretaceous upper boundary on syn-compressional extension and gravitationally-driven thinning of the Selkirk Allochthon orogenic wedge. 77 Chapter 4 IRREGULAR ISOGRADS, REACTION INSTABILITIES, AND THE EVOLUTION OF PERMEABILITY DURING METAMORPHISM 4.1 Introduction Rocks in the Big Fish Creek area show evidence of fluid flow attending recrystallization. Retrogression of peak mineral assemblages after both prograde metamorphic events recorded in these rocks was attended by fluid infiltration, as indicated by the hydrous nature of the retrograde minerals (e.g., chlorite and biotite after M l garnet, muscovite-rich aggregates after M2 staurolite and andalusite). However, the fact that the same prograde minerals show variable degrees of replacement by hydrous phases indicates that fluid-flow-driven retrogression was heterogeneous (e.g., M l garnet, compare Plate IB and 1C; M2 staurolite, compare Plate 5C,D and Plate 6B; late-M2 andalusite, compare Plate 6C and 6D). When attempting to understand the rates and patterns of fluid flow in mid-crustal environments such as the Big Fish Creek area, the nature of permeability is critically important. However, it remains poorly constrained at best. Metamorphic mineral reactions can alter porosity leading to modification of permeability and fluid flow distribution. Reactions that increase porosity will tend to enhance permeability and focus fluid flow (e.g., Rumble et al., 1982). If fluid flow drives mineral reaction, a positive feedback may develop between infiltration and reaction which will result in runaway flow-focusing (e.g., Ortoleva et al., 1987). Fluid flow will become increasingly heterogeneous. Alternatively, reactions that decrease porosity may reduce permeability and divert flow away from the reaction site (e.g., Cook et al., 1997). Either way, mineral reactions alter fluid flow patterns and the resulting distribution of chemical alteration and mineral assemblages. Recently, the heterogeneity of stable isotopic alteration has been used to constrain the spatial variability of permeability during metamorphism of carbonate rocks (Bowman et al., 1994; Cartwright, 1994; Gerdes et al., 1995). However, the mineralogical record of fluid flow has not been similarly evaluated. Because infiltration drives decarbonation reactions in these rocks (e.g., Ferry, 1994b), the distribution of metamorphic mineral assemblages is also a record of permeability distribution. Moreover, if metamorphic modification of permeability was an important control on fluid flow distribution, then this must be reflected in the distribution of mineral assemblages. 78 Here we model reactive transport through heterogeneous permeability to quantify the spatial variability of fluid flow required to produce observed distributions of mineral assemblages in contact metamorphosed siliceous dolomites. We compare these results to fluid flow heterogeneity predicted from the distribution of stable isotopic alteration, and show that in a number of contact aureoles the mineralogic and isotopic records of fluid flow are extremely heterogeneous. We propose that this heterogeneity is the result of permeability enhancement by mineral reaction. 4.2 The record of fluid flow Contact metamorphosed siliceous dolomites of the Alta (Utah) and Ubehebe (California) aureoles show similar distributions of mineral assemblages. At Alta, the minerals talc, tremolite, forsterite, and periclase (inferred from secondary brucite) define metamorphic isograds with increasing proximity to the Alta stock (Fig. 4.1 A; Moore and Kerrick, 1976). At Ubehebe, the progressive sequence of isograds is defined by tremolite, forsterite, and periclase (Roselle, 1997). Applying one-dimensional (ID) models of reactive transport of H2O-CO2 fluid in siliceous dolomite, Ferry (1994b) showed that these mineral assemblage distributions, in particular the presence of periclase and absence of diopside, require pervasive, grain-scale infiltration of aqueous fluid in the direction of decreasing temperature (e.g., Fig. 4.2B). He estimated a time-integrated fluid flux for Alta in excess of 1000 moi fluid/cm2 of rock. Whole-rock carbonate l 8 0 is variably depleted in the inner portions of both the Alta and Ubehebe aureoles (Fig. 4. IB; Bowman et al., 1994; Roselle, 1997). The magnitude of 1 80 depletion matches that predicted to result from infiltration of H20-rich, isotopically light magmatic volatiles. Bowman et al. (1994) estimated a time-integrated fluid flux for Alta of approximately 4000 mol/cm2. They attributed the spatial variability of 1 80 depletion to hydrodynamic dispersion resulting from heterogeneous fluid flow. In their ID reactive transport models, hydrodynamic dispersion results in a broadening of the isotopic alteration front, but does not account for the scatter of values about that front (Fig. 4. IB). Gerdes et al. (1995) used two-dimensional (2D) stochastic models to show that heterogeneous fluid flow could produce this isotopic scatter. Because the distribution of mineral assemblages depends on the amount of fluid flow (Fig. 4.2B), the heterogeneity of flow implied by the stable isotopic scatter must also influence the geometry of mineral isograds if the two processes are coupled. At Alta, periclase-bearing and periclase-absent marbles are interlayered on a meter scale (Cook et al., 1997), and forsterite- and periclase-grade rocks 79 B Metasediments ~2 Mississippian H Cambrian 40° 34' 00" 111° 36' 00" 30 + c£b j . • S & 8 (9 ^ H-span of Per isograd ZONE • Per © Per or Fo ® Fo oTr + Tic 500 1000 Distance from intrusion (m) 1500 Figure 4.1 A : M a p o f southern portion o f Al ta , Utah contact aureole (adapted from Cook, 1992). Talc (Tic), tremolite (Tr), forsterite (Fo), and periclase (Per) isograds mark first appearance o f these minerals as pluton is approached. However, low-grade assemblages persist into higher-grade zones: forsterite-grade rocks occur within periclase zone to 50 m from contact, as reported by Smith (1972; filled triangles) and Cook (1992; open triangles). B : Whole-rock carbonate 8 1 8 0 (standard mean ocean water) o f siliceous dolomites vs. distance from contact (modified after Cook, 1992). Span of periclase isograd is zone over which periclase- and forsterite-grade rocks overlap. Dashed curve denotes alteration predicted by model o f Bowman et al. (1994). Mineral abbrevations as in A . 80 Figure 4.2 A : Isobaric (1 kbar) phase diagram for model siliceous dolomite showing mineral assemblage stability as function o f temperature and mole fraction o f C 0 2 in H 2 0 - C 0 2 fluid. Dashed lines show mineral assemblage sequences and equilibrium fluid compositions for three different time-integrated fluid fluxes, during down-T flow of H 2 0 from the intrusive contact. Thermodynamic data from Berman (1988). D o l : dolomite; Qtz: quartz; C a i : calcite; D i : diopside; Brc : brucite; al l others as in Figure 4.1. 81 Figure 4.2 B: Mineral assemblage stability as function of distance from intrusive contact and time-integrated fluid flux. Thin dashed lines denote assemblage sequences for same fluxes as in A. Displacement of l s O alteration front is shown as heavy dashed line. Assemblage fields are: Dol-Qtz: dolomite + quartz protolith; Tic: Tlc+Dol+Cal; Tr: Tr+Dol+Cal; Di: Di+Dol; Fo: Fo+Dol+Cal; Brc: Brc+Cal; Per: Per+Cal; Di-Tr: Di+Tr+Dol+Cal. 82 overlap over =250 m (Fig. 4. IB; metamorphic grades are defined in Fig. 4.2). Roselle (1997) observed a similar overlap in the Ubehebe aureole, and Ferry and Rumble (1997) documented a heterogeneous distribution of periclase in siliceous dolomites of the Silver Star and Beinn an Dubhaich aureoles. Thus, mineral assemblages in contact-metamorphosed carbonate rocks also appear to record heterogeneous fluid flow. 4.3 Models We used finite-difference methods to model 2D heterogeneous fluid flow through siliceous dolomite initially composed of 90% dolomite and 10% quartz by volume, a composition representative of massive siliceous dolomites at Alta (e.g., Moore and Kerrick, 1976). The stability of mineral assemblages in this system varies with temperature and fluid composition (Fig. 4.2A). Each rock in the model aureole was heated and equilibrated to peak metamorphic temperature prior to infiltration. Further mineral reaction was driven by infiltration of magmatic H 2 0 down the steady-state peak thermal gradient (down-T flow, Figs. 4.2B and 4.3). We did not account for heat flow because doing so adds considerable numerical complexity but does not appreciably affect the sequence or general displacement of mineral isograds with infiltration (Dipple and Ferry, 1996). However, without accounting for heat flow, primary and secondary brucite cannot be distinguished. The presence of either brucite or periclase is used to define the periclase zone at Alta (Moore and Kerrick, 1976) and in our analysis. Stable isotopic alteration is driven by down-T flow of pure magmatic H 2 0 with a 8 1 8 0 of 7%o. Fluid-rock exchange is approximated by the calcite-H20 system (e.g., Bowman et al., 1994; Gerdes et al., 1995); initial 5 I 8 0 of calcite is 25%o. Steady-state heterogeneous permeability was represented by stochastic distributions generated with geostatistical software (Deutsch and Joumel, 1992), by fixing the mean, variance, and spatial correlation of the natural logarithm of permeability (ln[k]). The permeability maps had a geometric mean of 10"15 m 2. We changed the spatial variance (a2) of ln(k) between simulations (e.g., Gerdes et al., 1995) to assess how heterogeneous permeability affects mineral assemblage and 5 l 8 0 distributions. Spatial correlation was anisotropic, with characteristic correlation distances (i.e., the variogram ranges; de Marsily, 1986) of 300 m longitudinal and 30 m transverse to flow (Gelhar, 1993; Gerdes et al., 1995). Steady-state distributions of flow velocities were obtained by imposing high fluid pressure at the inflow boundary (intrusive contact), low fluid pressure at the outflow boundary (unmetamorphosed country rock) and impermeable flow-parallel boundaries. Fluid flow, although heterogeneous, was generally away from the intrusive contact. Fluid production by metamorphic mineral reactions was neglected. 83 650°C 550°C 450°C 350°C B Per Per q m = 3500; o 2 = 2.0 Brc Fo q m = 3500; a 2 = 0.2 Brc Fo i i 500 1000 Distance from contact (m) Tic Tic 3 1 1500 Figure 4.3 Mineral assemblage maps for down-T flow of initially pure H 2 0 fluid into siliceous dolomites for different spatially averaged time-integrated fluid fluxes (qm) and variances of ln(k). Mineral fields are as defined in Figure 4.2. Note that irregularity of mineral isograds is more pronounced where permeability is very heterogeneous. 84 At each timestep, fluid infiltrated the rock and reaction proceeded until chemical equilibrium was reached. This was done for each of 30 000 nodes until the prescribed final time-integrated fluid flux (qm) was attained. Results are presented as mineral assemblage maps and 8 I 8 0 distributions for different values of q m and 0"2(ln k) (Figs. 4.3 and 4.4). Unless otherwise mentioned, results are for a q m of 3500 mol/cm2, which yields a good fit to observed isograd and 1 8 0 front displacements at Alta. 4.4 Results Figure 4.3 illustrates the effect of heterogeneous permeability on the geometry of mineral zones and isograds. For a given q m , the sequence of mineral assemblages and displacement of mineral isograds predicted for 2D heterogeneous flow are similar to the ID solution (cf, Figs. 4.2B and 4.3). However, heterogeneous fluid flow produces irregular isograds; in more permeable zones, isograds are displaced towards lower temperature (Fig. 4.3). Our results indicate that even moderately heterogeneous permeability fields yield irregular mineral isograds, as illustrated in Figure 4.3C for a permeability field with o 2(ln k) = 0.2, a value typical of relatively homogeneous near-surface aquifers (e.g., Gelhar, 1993). However, the preservation of forsterite-grade assemblages in the inner 50 m of the Alta (Fig. 4.1) and Ubehebe (Roselle, 1997) aureoles requires a more heterogeneous distribution of periclase such as that in Figure 4.3B, which records flow through a highly heterogeneous permeability field (a 2(ln k) = 2.0). The pattern of l 8 0 depletion in the inner aureole at Alta also requires highly heterogeneous permeability (a2(ln k) = 2.0; Fig. 4.4). In general agreement with Gerdes et al. (1995), our results show that flow through a more homogeneous permeability field would not reproduce the observed 8 1 8 0 distribution (Fig. 4.4). The observed distribution of mineral assemblages and l 8 0 alteration at Alta is consistent with down-T flow of 3500 mol/cm2 through a very heterogeneous permeability field. 4.5 Discussion Our results highlight two important features of reactive transport during contact metamorphism at Alta: (1) the geometry of the periclase isograd and the scatter in 8 1 8 0 values record similar fluid-flow heterogeneity; (2) this heterogeneity was extreme. Similar isotopic and mineral assemblage distributions in other aureoles (e.g., Ubehebe, Silver Star, Beinn an Dubhaich) indicate that highly heterogeneous fluid flow may be a common feature of contact metamorphism. 8 5 350°C 304 span of Per isograd at Alta 500 1000 Distance from intrusion (m) 1500 Figure 4.4 Comparison of model results with Alta observations. Top frame is as in Figure 3B, with Per and Brc fields combined. Bottom frame shows model 5 1 8 0 distribution for qm = 3500 mol/cm2 through permeability fields with a2(ln k) = 0.2 (dashed) and o2(ln k) = 2.0 (solid). Note agreement between model 8 1 8 0 distribution for o2(ln k)= 2.0 and 8 1 8 0 distribution at Alta. Also shown are span of periclase isograd (arrow) and tremolite zone (oblique rule) at Alta. 86 Periclase and 8 I 8 0 distributions at Alta are both consistent with highly heterogeneous permeability (a2(ln k) = 2.0; Fig. 4.4). Only the most heterogeneous near-surface aquifers show such high variability (e.g., Gelhar, 1993). In fact, the range of permeability (nearly 5 orders of magnitude) suggested by comparing Alta data with model results exceeds by nearly two orders of magnitude that estimated by Ferry (1994a) for 266 samples of diverse metamorphic rock types from across northern New England. Given the limited geographic extent of the Alta aureole (Fig. 4.1 A), fluid flow was very heterogeneous there, and perhaps also at Ubehebe, where periclase isograd geometry and 8 l s O scatter are similar to those at Alta (Roselle, 1997). We interpret the extreme heterogeneity of fluid flow in the Alta aureole to be the result of flow-focusing during metamorphism rather than a direct reflection of extreme heterogeneity in the protolith. Flow-focusing is an expected consequence of infiltration-driven reaction involving a decrease in solid volume (i.e., porosity enhancement) due to the development of positive feedback between flow, reaction and permeability enhancement (e.g., Ortoleva et al., 1987). Because prograde isograd reactions in siliceous dolomites involve a significant decrease in solid volume (e.g., 25% for the periclase isograd reaction), reaction enhancement of permeability is a likely cause of the high heterogeneity of fluid flow. Thus, although the synmetamorphic permeability structure might mimic protolith heterogeneity, the heterogeneity in recorded time-integrated fluid flux in Alta dolomites was probably amplified by metamorphic reaction. Attainment of flow focusing depends on the relative rates of fluid flow, mineral reaction, and compaction. Sluggish reaction kinetics will tend to homogenize the chemical record of fluid flow and permeability because areas of fast fluid flow will be furthest from fluid-rock equilibrium. Also, slow reaction rates will lead to kinetic dispersion which will tend to inhibit flow-focusing (Steefel and Lasaga, 1990). We therefore suggest that the highly heterogeneous distribution of periclase-grade rocks at Alta reflects a close approach to equilibrium for the periclase-forming reaction. Because 8 l s O scatter and periclase isograd geometry record a similar flow heterogeneity, oxygen exchange probably also closely approached equilibrium, as argued by Bowman et al. (1994). Flow focusing will be inhibited if compaction is more efficient at destroying porosity than flow and reaction are at creating it. The highly heterogeneous fluid flow recorded at Alta and Ubehebe suggests that metamorphic fluid flow was fast compared to compaction. The observation that regionally metamorphosed rocks record smaller spatial variations in permeability (e.g., Ferry, 1994a) may reflect 87 fundamental differences in the relative rates of fluid flow and compaction in contact and regional metamorphic environments. 88 Chapter 5 CONCLUSIONS This thesis presents the results of two separate studies: a field-based study of the tectono-thermal evolution of the Selkirk Allochthon, a metamorphosed and deformed welt of supracrustal rocks in the southern Omineca crystalline belt of British Columbia, Canada (Chapters 2 and 3); and a study of the nature of multi-dimensional reactive fluid flow during metamorphism of carbonate rocks, using a comparison of numerically simulated distributions, and natural distributions of mineral assemblages resulting from metamorphic fluid flow (Chapter 4). The salient conclusions of these two studies are summarized here. Detailed characterization of the distribution of metamorphic mineral assemblages, of the metamorphic reaction history, and of the nature of structural elements related with broadly syn-metamorphic ductile deformation in the Big Fish Creek area of the Selkirk Allochthon leads to the following conclusions: 1. Two distinct metamorphic events affected rocks in study area: - M l , the earlier event, is recorded by relict porphyroblastic minerals which were subjected to cooling and retrogression prior to development of M2 prograde mineral assemblages. M l was probably regional in extent. - M2 is characterized by a Barrovian sequence of metamorphic zones in the study area. From south to north, this field gradient includes the following metamorphic zones in rocks of pelitic composition: chlorite, biotite, garnet + chlorite, staurolite + biotite, kyanite + staurolite. Late-M2 andalusite is widespread in rocks of the garnet + chlorite, and staurolite + biotite zones, suggesting substantial late-M2 decompression. 2. The Bigmouth Creek stock was emplaced in late Middle Jurassic as indicated by magmatic zircon U-Pb ages (157 ± 3.3 Ma). Magmatic epidote in Bigmouth Creek granodiorite suggests that emplacement depth was substantially greater than 20 km. Earliest Cretaceous U-Pb ages (140.7 ± 0.8 Ma, and 137.4 ± 1.4 Ma) from titanite from the granodiorite are interpreted to be due to thermal resetting at that time. 89 3. Based on evidence for widespread mid-Jurassic metamorphism in other areas of the Selkirk Allochthon, and for deep emplacement of the Bigmouth Creek pluton at that time, M l metamorphism is interpreted to be of mid-Jurassic age. M2 is interpreted to be the result of the thermal pulse which reset U-Pb systematics of titanite from the Bigmouth Creek pluton in Early Cretaceous time Thus, the timing of M2 metamorphism is Early Cretaceous. 4. Rocks in the area record intense, non-coaxial strain which resulted from long-lived ductile progressive deformation, and involved important components of orogen-transverse and orogen-parallel extension (i.e., shortening normal to the shallow-dipping dominant transposition foliation ST, which approaches the regional shear plane). These rocks are therefore interpreted to be part of a ductile shear zone of regional significance, herein named the Big Fish Creek shear zone. The presence of this shear zone at intermediate levels in the Selkirk Allochthon illustrates the distributed nature of ductile strain in the Selkirk Allochthon in Mesozoic time. 5. A mid-Jurassic age for M l metamorphism implies that pre- to syn-metamorphic progressive ductile deformation began at least as early as the Middle Jurassic. Late-M2 decompression at least as late as the Early Cretaceous predates the end of ductile deformation in the area. This puts an Early Cretaceous upper boundary on syn-compressional extension and gravitationally-driven thinning of the Selkirk Allochthon orogenic wedge. 6. Structural and metamorphic evidence for protracted compressional strain, and synchronous, episodically at least, extension and thinning in the present study area, combined with similar interpretations at both higher, and lower structural levels by other workers, suggests that the Selkirk Allochthon constituted a dynamic orogenic wedge from as early as the Middle Jurassic, to at least as late as the Late Cretaceous. The dynamic behavior of this wedge was controlled by far-field forces associated with convergence of accreted terranes and the North American continental margin, by gravitational forces, and by rheological modification and internal deformation of the wedge in response to these different external controls. Chapter 4 deals with the nature of permeability during metamorphism at mid-crustal levels. A good example of the impact of heterogeneous, and probably temporally evolving fluid flow and permeability in a metamorphic system is offered by the spatially variable extent of retrogression of prograde mineral phases to hydrous assemblages in rocks of the Big Fish Creek area. Work presented in Chapter 4 addresses the question of the spatial variability and the temporal evolution of fluid flow and 90 permeability in metamorphic terranes and leads to conclusions which may have applicability to the nature of retrogression in the Big Fish Creek area. Major conclusions of the study presented in Chapter 4, of two-dimensional reactive flow in siliceous dolomites are as follows: 1. Numerical models of two-dimensional reactive fluid flow can reproduce the patterns of mineral assemblage distributions observed in natural metamorphic terranes. 2. The spatial heterogeneity of fluid flow predicted from comparison of numerical simulations with natural distributions of mineral assemblages is consistent with that predicted from comparison of natural distributions and numerical simulations of stable isotopic alteration, using the same permeability distribution. This lends support to the applicability of the numerical simulations. 3. Comparison of natural and numerically simulated distributions of mineral assemblages and stable isotopic alteration suggests that the heterogeneity of permeability during metamorphic infiltration of siliceous dolomites in the Alta, Utah contact aureole, was several orders of magnitude larger than the inferred heterogeneity of permeability of the protolith. This suggests that the heterogeneity of hydraulic properties is amplified by fluid flow and reaction during metamorphism. 4. 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Magmatic epidote and its petrologic significance. Geology 12: 515-518. Zen, E. 1985. Implications of magmatic epidote-bearing plutons on crustal evolution in the accreted terranes of northwestern North America. Geology 13: 266-269. 102 Appendix I: Photographic Plates 103 Plate 1 Garnet morphology and evidence for two prograde garnet growth episodes separated by retrogression. A. Sample 94-NM-75B, M2 staurolite + biotite zone. Garnet porphyroblast showing a dominant, anhedral inclusion-rich core, and thin, euhedral rim. Note rind of more or less randomly oriented biotite flakes around bottom part of grain, which is overgrown by inclusion-poor garnet. Plane polarized light; width of field is 6 mm. B. Sample 94-NM-l 1 IB, M2 staurolite + biotite zone. The two garnet grains are within the same strain shadow and probably represent two clasts from an original grain. The core portions of these grains are anhedral and relatively inclusion-rich whereas the rims are inclusion-poor and euhedral to subhedral. Euhedral rims overgrow biotite + chlorite + quartz + plagioclase mass developed as a result of retrogression of early inclusion-rich garnet represented by the anhedral cores. Also note that where small staurolite porphyroblasts are developed around the garnet, the inclusion-poor garnet rim is less euhedral. Elemental composition maps and a quantitative compositional profile of the left grain are presented in Plate 2A, and Figure 2.3, respectively. Plane polarized light; width of field is 1 cm. C. Sample 94-NM-85C, M2 garnet zone. Partially chloritized inclusion-rich garnet porphyroblast. The chlorite ± quartz ± plagioclase ± biotite rind around the porphyroblast outlines its original shape. Note that biotite seems to be developed preferentially outboard of the chlorite-rich rind. Some of this biotite may have grown during M2 and be prograde. No euhedral rim is present on garnet. Also note that matrix foliation deflects around the original garnet outline, and that inclusion trails in the porphyroblast are inclined with respect to this foliation. Plane polarized light; width of field is 6 mm. D. Sample 94-NM-71A, M2 garnet zone. Garnet porphyroblast completely pseudomorphed by chlorite + biotite + quartz ± plagioclase aggregate. Some of the biotite may have grown during M2 and be prograde. Outline of original porphyroblast is still visible and matrix foliation gently deflects around it. Plane polarized light; width of field are 6 mm. Note: anomalous colours are an artifact of printing. 104 105 Plate 2 Elemental composition maps of composite garnet. A. Sample 94-NM-l 1 IB, M2 staurolite biotite zone. This is the left grain shown on Plate IB, and a quantitative compositional profile is shown in Figure 2.3. B . Sample 94-NM-71B, M2 staurolite + biotite zone. Scale bars are 50 um. Note: anomalous colours are an artifact of printing. 106 107 Plate 3 Chlorite-zone, and biotite-zone textures. A. Sample 94-NM-45B, chlorite zone. Dominant foliation is defined by chlorite + muscovite-rich laminae and quartzose microlithons. Note waviness resulting from gentle crenulation of dominant foliation. Width of field is 3 mm. B. Sample 94-NM-2ME, chlorite zone. Syn-crenulation chlorite porphyroblast. Inclusion trails in the porphyroblasts define a foliation, Sj, which is more gently crenulated than the dominant matrix foliation, S e (dominant foliation is sub-horizontal, and the crenulation dips to the left). Width of field is 6 mm. C. Sample 94-NM-48G, low-grade portion of M2 biotite zone. Three foliations are present: a dominant crenulation cleavage (Scc); an earlier, extensively transposed continuous cleavage (Sc) preserved as recrystallized, folded muscovite in microlithons; and a late crenulation (Scr) which is well developed only in some laminae. Biotite occurs primarily as fine-grained flakes within chlorite + biotite clots (arrow). Width of field is 6 mm. D. Sample 94-NM-l 18B, high-grade portion of M2 biotite zone. Large, irregularly oriented biotite flakes. Note that the foliation, Sj, defined by inclusion trails in these porphyroblasts is not regular from grain to grain, and does not parallel either of two external foliations, a continuous cleavage (Sc), and a crenulation (Scr). These spatial relationships are explored in detail in Chapter 3. Width of field is 6 mm. Note: anomalous colours are an artifact of printing. Note: anomalous colours are an artifact of printing. 108 109 Plate 4 M2 mineral textures, staurolite + biotite zone. A. Outcrop 96-NM-226. Fresh staurolite locally forming crosses, occurring in staurolite + biotite intergrowth over garnet porphyroblasts. Also present are discrete staurolite porphyroblasts in groundmass away from staurolite + biotite clots after garnet. Coin is 2.3 cm in diameter. B . Outcrop 94-NM-101. Muscovite ± chlorite pseudomorphs after coarse andalusite porphyroblasts defining a "chickenfeet" texture. Lens cap is 6 cm in diameter. C, D . Sample 96-NM-214K. Euhedral staurolite porphyroblast partially included in poikiloblastic andalusite; C. plane-polarized light; D. cross-polarized light. Width of field is 6 mm. Note: anomalous colours are an artifact of printing. 110 I l l Plate 5 M2 mineral textures, staurolite + biotite zone and kyanite + staurolite zone. A. Sample 94-N M - l 19D, staurolite + biotite zone. M l garnet porphyroblast partially retrogressed to a chlorite + biotite + quartz + plagioclase mass which define its original shape. Inclusion-rich relict core of M l garnet is rimmed by euhedral, inclusion-poor M2 overgrowth. Small, euhedral M2 garnet porphyroblasts are present, some of which overgrow matrix foliation (e.g., small grain in upper right corner). M2 staurolite is developed around composite garnet and includes small M2 garnet grains. Width of field is 6 mm. B. Sample 94-NM-l35D, kyanite + staurolite zone. Close spatial association of kyanite, staurolite and garnet. Only relict garnet remains. Note sets of parallel fractures in staurolite and kyanite porphyroblasts, at high angle to matrix foliation. Width of field is 6 mm. C. Sample 94-NM-l 29E, high-grade portion of the staurolite + biotite zone. Large staurolite porphyroblast clearly deflects matrix foliation, and Si is strongly inclined to Se. Note the close association of garnet porphyroblast partially visible on right hand side of photo, and staurolite. Width of field is 6 mm. D. Sample 94-NM-129C, from same outcrop as C. Staurolite porphyroblasts do not noticeably deflect matrix foliation and show Si parallel to Se. Width of field is 6 mm. Note: anomalous colours are an artifact of printing. 112 113 Plate 6 Retrogression and deformation of prograde M2 minerals. A. Sample 95-NM-171B, staurolite + biotite zone. Partially to completely retrogressed staurolite porphyroblasts. Composite M1-M2 garnet porphyroblast surrounded by muscovite-rich aggregates which are pseudomorphs after large staurolite porphyroblasts. On the right hand side of the garnet, a small relict staurolite grain is preserved in the core of a pseudomorph. Note that pseudomorph outlines deflect matrix foliation. Also note post-Mi fracture in garnet filled by inclusion-poor M2 garnet, and inclusion-poor subhedral rim. Width of field is 6 mm. B. Sample 95-NM-175A, staurolite + biotite zone. Completely retrogressed staurolite (?) porphyroblast. Muscovite ± chlorite pseudomorph (ps) after staurolite closely associated with partially chloritized and fractured garnet porphyroblast. Note that shape of original staurolite is still discernable, and that it gently deflects matrix foliation. Width of field is 6 mm. C. Sample 96-NM-214J, staurolite + biotite (+ and) zone. Fractured and boudinaged andalusite porphyroblast. Biotite is seen to have developed in a boudin neck in andalusite, just to right of center. Width of field is 6 mm. D. Sample 94-NM-103D, staurolite + biotite (+ and) zone. Muscovite ± chlorite pseudomorph after andalusite, with relict andalusite preserved in core. Note alignment of muscovite and chlorite flakes after andalusite with dominant matrix foliation. Width of field is 3 mm. Note: anomalous colours are an artifact of printing. 114 115 Plate 7 Sillimanite-bearing rocks from western reaches of the study area (see Fig. 2.6). A. Sample 95-NM-156, sillimanite + staurolite zone. Fibrolitic sillimanite + biotite intergrowth developed on what is interpreted to have been a chlorite-rich pseudomorph after garnet, in close association with staurolite. Note local presence of undeformed chlorite which may be a product of early garnet retrogression. Width of field is 6 mm. B. Sample 95-NM-158C, sillimanite + garnet + biotite zone. Late, euhedral garnet developed over fibrolitic sillimanite + biotite intergrowth on early garnet pseudomorph. Early garnet may have been split, as suggested by the bilobate outline of the pseudomorph, and is interpreted top be of M l timing. Width of field is 6 mm. C. Sample 95-NM-158C, sillimanite + garnet + biotite zone. Late, euhedral overgrowth on partially resorbed, inclusion-rich garnet core. Garnet core is interpreted to be of M l timing, and garnet overgrowth was the result of staurolite-breakdown reaction during M2 (see text). Width of field is 6 mm. D. Sample 95-Big-9, sillimanite + garnet + biotite zone. Matrix sillimanite is fibrolitic to finely crystalline, aligned in the dominant foliation, and folded along with other matrix minerals. Width of field is 6 mm. Note: anomalous colours are an artifact of printing. 116 117 P l a t e 8 Macroscopic structural elements. A : Folded and transposed ST. View looking east (outcrop 94-NM-54) at tight fold in psammitic sequence. Note thickened hinge, and attenuated to pinched-out upper limb of most prominent fold in the lower half of the outcrop. B: N-S trending folds, outcrop 95-NM-150, looking north. Isoclinal intrafolial fold (arrow) in semi-pelitic sequence. C : View looking east at semi-pelite outcrop on main N-S ridge in the central portion of the study area (outcrop 94-NM-30). S T is defined by the layering and schistosity, and dips moderately to the ESE. Undulations are the result of late, steep, SE-trending crenulation. Note pinched-out layers in the left half of the outcrop (arrow). D: Two intersection (crenulation) lineations, outcrop 94-NM-l 1, southern part of the study area. The lineation running from upper left to lower right is the more conspicuous of the two, and is the result of the intersection of steeply-dipping, ESE-striking crenulation cleavage with the dominant foliation ST- A less well developed lineation is visible which is at high angle to it. This is the NNE-trending lineation. Crenulation cleavage associated with this lineation is very poorly developed. Note: anomalous colours are an artifact of printing. 118 119 Plate 9 Split or fractured garnet porphyroblasts. A : Hand sample from outcrop 94-NM-l 19, staurolite + biotite zone, cut parallel to L T lineation. Note asymmetric fold of concordant quartz vein in lower part of sample, and intrafolially refolded leucocractic layer near the center (large arrow). Split garnet porphyroblasts (small arrows) show gaps at high angle to ST, the dominant foliation highlighted by compositional layering consisting of biotite-rich, and quartz-rich laminae. B: Mineral/intersection lineation, L T , defined by alignment of elongate biotite flakes on ST surface, outcrop 94-NM-l 11. Note split garnet porphyroblasts with split at high angle to L T . C: Sample 96-NM-165C, ST-normal, LT-parallel cut. Split garnet with gap at high angle to S r. Si in garnet fragments is approximately parallel to the external foliation ST, indicating that no post-splitting rotation took place. Note that many fractures in staurolite porphyroblasts are sub-parallel to garnet split. Also note local, thin, inclusion-poor overgrowth on split garnet fragments. Width of field is 10 mm. D: Sample 94-NM-75B, composite ST. Differing extents of transposition in matrix and in garnet porphyroblast. Dominant foliation in large garnet is a crenulation cleavage (Scr) which overprints earlier continuous cleavage preserved in microlithons (Sc). In matrix, a later-cycle continuous cleavage (Sc) is parallel to S c r in garnet. No clear trace of earlier foliation is found. Late shear bands (SSb) overprint dominant foliation in both garnet and matrix. Width of field is 6 mm. Note: anomalous colours are an artifact of printing. 120 1 2 1 Plate 10 Different stages of progressive transposition of the dominant foliation in the same outcrop, staurolite + biotite zone. A, B: Sample 94-NM-215A, showing Scr, crenulation of an earlier foliation (Sc) defined by the preferred orientation of muscovite flakes. The orientation of the crenulation foliation (Scr) is irregular and some later shear bands (SSb) appear to cross-cut crenulations (e.g., in lower left quadrant of picture, see B). Note that many large biotite flakes lie parallel to S c r, but have Sj approximately parallel to Sc. Some biotite flakes appear to be randomly oriented. This sample shows a less advanced stage of a transposition cycle than sample in C,D. C, D: Sample 94-NM-215B, showing strong crenulation cleavage (Scc) parallel to layering, marked by muscovite-rich laminae. Note irregular, anastomosing nature of this strong crenulation cleavage, and compositional layering/segregation associated with it. Earlier foliation Sc is extensively transposed, but preserved in quartz-rich microlithons as recrystallized, isoclinally folded thin muscovite-rich laminae. Note large, foliation-parallel fish-shaped biotite flakes lying approximately parallel, and locally showing Si at high angle, to S c c (arrow). Width of field is 6 mm for all photomicrographs. A and C are in plane polarized light; B and D are in cross-polarized light. Note: anomalous colours are an artifact of printing. 122 123 Plate 11 Lineation-defining biotite porphyroblasts in LT-normal and LT-parallel sections. A, B: Sample 94-NM-l66B, staurolite + biotite zone. A shows LT-parallel cut in which most biotite flakes lie parallel to ST. B shows LT-normal cut in which biotite flakes lie at shallow to moderate angle to ST , and show a preferred vergence with respect to it, defining a foliation, ST'. Coin is approximately 17 mm in diameter. C, D: Sample 96-NM-205B, staurolite + biotite zone. C shows LT-normal cut. An S-C-type fabric is defined by two shallow-dipping (in this view), discontinuous, cross-cutting foliations (thick lines), forming eye-shaped domains (dashed). A left-dipping foliation is defined by the preferred orientation of most matrix muscovite flakes, and is at shallow angle to earlier compositional layering highlighted by a quartz-rich band running from lower left to center right (large arrow). Some medium-sized biotite flakes parallel this foliation. A right-dipping foliation is highlighted by the preferred orientation of thin, medium-sized biotite flakes. Note that some biotite flakes lie sub-horizontal in the figure, at angle to both left-dipping and right dipping foliations. These flakes tend to be fish-shaped, and elongate inclusions (Sj) are at high angle to both foliations. Also note other biotite flakes which are blocky to rectangular, do not lie parallel to a foliation, have Sj parallel to local external foliation, and appear to overgrow both left- and right-dipping foliations (e.g., 2-3 flakes lumped together in the upper right corner of the figure, small arrow). D shows L T -parallel cut. One strong preferred orientation of platy minerals (sub-horizontal) is visible, and foliation is much more regular in this cut than in the LT-normal cut. Most biotite flakes are elongate to fish-shaped. The majority of flakes lies sub-parallel to the foliation, but a few flakes lie at a shallow angle to it (arrows). Width of field for C, D is 6 mm. Note: anomalous colours are an artifact of printing. 124 125 Plate 12 Rotated and deformed biotite porphyroblasts. A: Sample 96-NM-201B, staurolite + biotite zone. Si-normal, LT-parallel cut showing large biotite fish with opposite senses of asymmetry (two flakes left and right of center, inferred local sense of shear shown by arrows). Note Sj in flake on the right is rotated relative to ST. Also note pinch-and-swell aspect of lamination in quartz vein at bottom. B: Sample 96-NM-208A, staurolite + biotite zone. ST-normal, LT-parallel cut showing large twinned biotite flakes. Also note that different porphyroblasts appear to show different sense of rotation. C: Sample 96-NM-214K, staurolite + biotite zone. Sr-normal, LT-normal cut. Extensional crenulation (Sec, lower left to upper right) and earlier, dominant foliation (Sc, sub-horizontal) anastomose to form an S-C fabric. Note close association of crenulation with large, rotated and locally split biotite porphyroblasts D: Sample 96-NM-225F, staurolite + biotite zone. Two large biotite flakes in the center have opposite shape asymmetry, and opposite apparent sense of rotation of Si relative to Se. Width of field is 6 mm for all photomicrographs. Note: anomalous colours are an artifact of printing. 126 127 Plate 13 Split or fractured porphyroblasts. A: Sample 96-NM-201 A, staurolite + biotite zone, ST-normal, LT-parallel cut. Fractures in garnet (lower right) and gap in biotite (upper left) have similar orientation approximately perpendicular to dominant foliation ST. Width of field is 6 mm. B: Sample 95-NM-166B, staurolite + biotite zone, ST-normal, LT-parallel cut. Post-kinematic staurolite porphyroblast is split, but relative displacement of resulting fragments is difficult to estimate. Width of field is 6 mm. C: Sample 96-NM-214J, boudinaged andalusite porphyroblast, showing split at high angle to ST, and development of biotite in boudin neck just right of center. Width of field is 6 mm. D: Photomicrograph of sample 94-NM-l09, staurolite + biotite zone, showing boudinaged, elongate feldspar grain. S T is approximately horizontal and splits are at high angle to it. Width of field is 10 mm. Note: anomalous colours are an artifact of printing. 128 129 Plate 14 Extensional structural elements. A: Biotite-rich high-strain zone in semi-pelitic sequence, outcrop 96-NM-235, staurolite + biotite zone, looking west, LT-normal cut. Note pinched-out psammite layers and quartz veins. An outcrop-scale shear band is highlighted by passages of dense biotite-rich layers which cut through S T layering at shallow angle, from top left to bottom right of picture. Inferred sense of shear is fop-down-to-the-north. B: Hand sample 94-NM-71A, staurolite + biotite zone, showing extensional crenulation cleavage/shear bands (SSb) at moderate angle to dominant foliation (Sc), producing an S-C'-type fabric. Width of field is 8 cm. C: Sample 96-NM-200B, staurolite + biotite zone, ST-normal, LT-parallel cut. Large extensional shear band runs from upper left to lower right. Note counterclockwise apparent rotation of large chlorite flakes and matrix white micas immediately to right of shear band. Width of field is 6 mm. D: Hand sample 96-NM-223A, staurolite + biotite zone, Si-normal, stretching lineation-parallel cut. View is to south. A feldspar grain on the lower left corner is boudinaged, and the resulting split surfaces are at a <90° angle to ST. Several winged and sigmoidal feldspar porphyroclasts indicate a top-to-the-west sense of motion, consistent with boudinaged clasts of feldspar. Coin is 17 mm in diameter. Note: anomalous colours are an artifact of printing. 130 131 Plate 15 Folds with LT-paralleI axes. A: Folded quartz veins, outcrop 94-NM-82, garnet zone. Dismembered hinges of folded quartz veins form ridges on the dominant foliation surface (which mostly parallels the outcrop surface). Compass points north. B: Outcrop 94-NM-l 13, staurolite + biotite zone. Dismembered hinges of folded quartz veins on S T surface approximately parallel L T , highlighted by elongate biotite grains (both trend from lower left to upper right on figure). Coin is 24 mm in diameter. C: Same outcrop as B, looking approximately normal to L T and fold axes, and showing asymmetric, tight fold of thin, sub-concordant quartz vein. Note thin biotite segregation highlighting compositional layering-parallel dominant foliation. Coin is approximately 1.5 mm thick. Note: anomalous colours are an artifact of printing. 132 133 Plate 16 Stretching lineation. A: Outcrop 94-NM-l27, staurolite + biotite zone, looking north. Strained gritty psammite with stretched quartz and feldspar (white) porphyroclasts. B: Sample 94-NM-101, staurolite + biotite zone. Large, elongate quartz grains with straight grain boundaries at high angle to foliation define ribbon quartz. These grains show deformation twins and sub-grain boundaries (open arrow), and small, dynamically recrystallized grains at their boundaries (closed arrow). Cross-polarized light. Width of field is 6 mm. C, D: Sx-normal, LT-normal and LT-parallel cuts, respectively, sample 94-NM-l20Q, staurolite + biotite zone, showing ribbon quartz and strongly stretched feldspar porphyroclasts in C. Coin is 24 mm in diameter. Note: anomalous colours are an artifact of printing. 134 135 Appendix II: Outcrop locations and list of samples 136 Figure A-II.1 Locations of outcrops and samples discussed in text. 137 CD 00 CO 73 a CD (-1 o £ * # >s * * . g a 'S o > — 13b ea s hi-an 3 o .(-i hi-an ren cn >^  O CO CO D 1 to o >N H x> "ca "O x i c*-« CD CD o CD SO c w> o a CO •c co av & 1% C« i i CO cS sposi CD CD JO sposi 73 ri d o\ 73 ri ayei ddi a CD S3 CO u o <2 u CO [MENTS e 73 e o 1 73 [MENTS lopmen ure-sol' isoclin ure-sol' lotchy" vein w COIV devel press early press folde w/"b gt on H W CO >»' Z Z Z Z Z Z Z z Z Z z Z z Z z z Z Z z z z Z z Z Z Z z Z CO CO z a, CO a >-> Y(2 Y(2 ne ne ne CO CO *f] s o -t—» -4—» o o c f 1 cO CO his his his his his CO his CO D CD CO CD 1 CO 1S-SI CO -t—» CO CO CO CO CD CO -8" s o o o o o o g ea § 1S-SI XI XI XI X . " ^ CD i—1 CO <t> CO CO CO CO CO CO B o o o CO o O o CD ,"t3 •—' arb chl- co CO o o CO 1 ^ CO CO e CO CO CO CO -t—' emi-pel TYPE mpure hl-sch ;uartzii arb ch arb ch arb chl-hl-sch arb ch arb ch ms-sch t ms-s arb ch fu imp. alc-sil npure arb ve hl-ms-arb ch hl-ms-hl-ms-chist chist mesto arb sci hl-ms-hl-ms-hl-ms-ms-sch] t psan: -t—* CO o f schis f schis samm emi-pel TYPE o a- o ti o o o ms-sch XI o fu imp. o o o o o o CO CO — O O o o ms-sch] CO o o a, CO [PLE < PQ U P LE ME an < PQ u L2 s <; PQ U Q hJ 5A 5B V9 6B 6C 6D [PLE <N CN tN CN CN CN CN CN i n oo oo oo 00 oo oo 00 Os Os Os OS ON ON ON [PLE s s s 5j s s s 5j 2 < Z z z Z z z z Z z z z z 5? z '3 z z z z z Z 4-N z z z z z z z Z <0" 4 4 4 4 -er 4 4 4 T •<a- 4 4 4 T 4 4-N 4 4 t/3 Os ON ON os Os ON Os OS OS ON ON ON Ov ON ON ON ON ON ON ON ON ON ON ON ON OS Os ON ON Os ON ON ON ON 138 COMMENTS early ST-// isoclinal folds overprinted by hi-angle late cren cleavage big Ab porphyroblasts GARNET Y.split OH CO G , CO o, CO CO Y,split Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z oo H Y(l) >< TYPE gt schist psammite strained psammite schist psammite psammite impure limestone gritty psammite bt schist? ms schist schist schist carb schist qz-fs ms-schist carb schist JS bt ms-schist limestone SAMPLE 94-NM-l 6L 94-NM-l 6M 94-NM-l 6U 94-NM-l 6W 94-NM-l 6X 94-NM-l 6Y 94-NM-l 6Z 94-NM-l 6AA 94-NM-l 7 94-NM-l 8B 94-NM-l 8C 94-NM-18S 94-NM-l 8T 94-NM-l 9L 94-NM-l 9M 94-NM-l 9U 94-NM-20L 94-NM-20M 94-NM-20U ON 94-NM-21U 94-NM-22 94-NM-23A 94-NM-23B 94-NM-23C 94-NM-23D 94-NM-24 94-NM-26L1 94-NM-26M 94-NM-26U 94-NM-27A 94-NM-27B 94-NM-29A 94-NM-29B 139 «4H COMMENTS float good solution cleavage P-sol cleavage, similar to 117; early fold overprinted by cren GARNET Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z on >-> >* >H >* TYPE chl-schist chl-schist bt chl-schist chl-schist carb schist schist gt-sill schist fu qz-vein carb schist? carb chl-schist carb chl-schist carb chl-schist calc-silicate calc-silicate fu calc-silicate ms-chl-schist ms-chl-schist ms-chl-schist ms-chl-schist w/lm SAMPLE 94-NM-30L 94-NM-30U 94-NM-31L 94-NM-31U 94-NM-33L 94-NM-33U(?) 94-NM-33A 94-NM-34 94-NM-36A 94-NM-36B 94-NM-37A 94-NM-37U 94-NM-38L 94-NM-38M 94-NM-38U 94-NM-39 94-NM-42A 94-NM-42B 94-NM-44A 94-NM-44B 94-NM-44C 94-NM-44D 94-NM-45A 94-NM-45B 94-NM-45D 94-NM-46A 94-NM-46B 94-NM-47A 94-NM-47B 94-NM-48A 94-NM-48B 94-NM-48C 94-NM-48D 94-NM-48E 140 COMMENTS bt // foliation bt // foliation bt-porph (l-2mm); post-k late bt-foliation syn-k gt P-sol cleavage post-crenulations? chl-prophyroblasts solution cleavage sub-// layering strong compositional segregation post-k garnet; Ab porphyroblasts GARNET Z Z Z Z Z Z Z Z Z Z Z Z Z Z Y.split Y,split Y,split Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Y.split c« H >< >< >* Y(2) > H >< >< TYPE psammite psammite gray layered marble gray layered marble fine-gr. psammite ab-porph chl-sch. semi-pelite ms-chl-schist fine-gr psammite calc-pelite SAMPLE 94-NM-48F 94-NM-48G 94-NM-48H 94-NM-49A 94-NM-49B 94-NM-49C 94-NM-50A 94-NM-50B 94-NM-51 94-NM-52 94-NM-53 94-NM-54A 94-NM-55 94-NM-56 94-NM-59 94-NM-60A 94-NM-60B 94-NM-60C 94-NM-60D 94-NM-60E 94-NM-61B 94-NM-61U 94-NM-62L 94-NM-62U 94-NM-63L 94-NM-63U 94-NM-64 94-NM-64B 94-NM-65L 94-NM-65L2 94-NM-65U 94-NM-67A 94-NM-67B 94-NM-68A 141 COMMENTS "blotchy-chl"+gt rock gt chloritized gt chloritized GARNET Y,split Y,split Z Y,split Z » • Z Y,split? Y,split Z Y,split Y,split? Z Z Z Z Z Y,split Y.fract'd Y,split Y.fract'd Y,fract'd OO H > > 00 00 S(2) • S(2) S(2) TYPE calc-pelite calc-pelite marble blotchy amphibolite psammite fine-gr gritty psam gray marble blotchy amphibolite amphibolite SAMPLE 94-NM-68B 94-NM-68C 94-NM-69 94-NM-69A 94-NM-69B 94-NM-69C 94-NM-69D 94-NM-70A 94-NM-70B 94-NM-70C 94-NM-70D 94-NM-71A 94-NM-71B 94-NM-72A 94-NM-72B 94-NM-72C 94-NM-73 94-NM-73S 94-NM-73T 94-NM-74A 94-NM-74B 94-NM-74C 94-NM-74D 94-NM-75A 94-NM-75B 94-NM-75C 94-NM-76Q 94-NM-76R 94-NM-78A 94-NM-78Q 94-NM-78R 94-NM-78S 94-NM-78T 94-NM-78U 142 COMMENTS gt chloritized gt chloritized strong compositional segregation gt chloritized strong compositional segregation GARNET Y.split? Z Y,split? Y.split Y,split Z Z Z >-> Z Z Y.split Y.split Y,split Y,split CO Z Z Z oo 00 00 TYPE semi-pelite pelite pelite psammite pelite/psammite blotchy psammite semi-pelite psammite SAMPLE 94-NM-79A 94-NM-79B 94-NM-79C 94-NM-80A 94-NM-80B 94-NM-80C 94-NM-80L 94-NM-80U 94-NM-81A 94-NM-81B 94-NM-81C 94-NM-82A 94-NM-82B 94-NM-82C 94-NM-82D 94-NM-84 94-NM-85A 94-NM-85B 94-NM-85C 94-NM-86A 94-NM-86B 94-NM-86S 94-NM-86Q 94-NM-87A 94-NM-87B 94-NM-88A 94-NM-88B 94-NM-88C 94-NM-89A 94-NM-89B 94-NM-89C 94-NM-90A 94-NM-90B 94-NM-90C 143 CD COMMENTS gt chloritized core of c.f preserved andalusite from loose block gt chloritized GARNET Z Z Z Z Z Z Z Z Z Z Z Z Z Z Y.split Y,split Y,split Y,split Y,split Y,split Y,split Y,split Y,split Y,split Y,split z ' Y.split Z Y.split Y.split Y,split H Y(4) >-> Y(2) Y(2) Y(2) TYPE bt-rich psammite sheared gritty psam. SAMPLE 94-NM-90D 94-NM-91A 94-NM-91B 94-NM-91C 94-NM-92A 94-NM-92B 94-NM-92C 94-NM-92D 94-NM-93A 94-NM-93B 94-NM-93C 94-NM-93D 94-NM-94A 94-NM-94B 94-NM-97A 94-NM-97B 94-NM-97C 94-NM-98A 94-NM-98B 94-NM-98C 94-NM-98F 94-NM-l 03 A 94-NM-l 03B 94-NM-103C 94-NM-l 03D 94-NM-l 05 94-NM-l 07 A 94-NM-l 08 A 94-NM-l 08B 94-NM-l 08C 94-NM-l 09 94-NM-l 1 OA 94-NM-l 1 OB 94-NM-l IOC 144 of samples COMMENTS pseudomorph after actinolite (garbenschieffen) extensional shear bands strong crenulation extensional shear bands P-sol cleavage P-sol cleavage P-sol cleavage bt-foliation/gneissosity 3 foliations List i GARNET Y,split Y,split Y,split Y,split Y.split >- >H >-Y, fract'd 2 oo hi Y(2) S(2) Y(2) >* TYPE ab-porph ms-sch. ab-porph ms-sch. psammite SAMPLE 94-NM-l 10D 94-NM-l 10E 94-NM-l 11A 94-NM-l 1 IB 94-NM-l HQ 94-NM-l 12A 94-NM-l 12B 94-NM-l 12C 94-NM-l 12D 94-NM-l 12E 94-NM-l 13 A 94-NM-l 13B 94-NM-l 13C 94-NM-l 13D 94-NM-l 13E 94-NM-l 13J 94-NM-l 13K 94-NM-l 13L 94-NM-l 14A 94-NM-l 15A 94-NM-l 15B 94-NM-l 15C 94-NM-l 16A 94-NM-l 16B 94-NM-l 16C 94-NM-l 16D 94-NM-l 17A 94-NM-l 17B 94-NM-l 17C 94-NM-l 18A 94-NM-l 18B 94-NM-l 18C 94-NM-l 18D 94-NM-l 18E 145 COMMENTS graded bedding late bt associated w/ P-sol cren. Cleavage oriented; looking S on thin section rotated grt talc-trem talc-trem syn- to late-k stt and ky oriented oriented, highly strained H ._ GAPvNE Z "o. CO el cn Z Y,spl Z Z Y,spl CO H CO Z Z Z "3. CO CO '3. CO CO Z Z GAPvNE OO H >< >< oo P(2) >-> > >- > >* TYPE psammite psammite pl-amphibolite stt-bearing pelite talc schist stt-ky (ms) schist stt-ky (ms) schist gritty psammite chl-schst/phyllonite W 119A 119B 119C 119D 120Q CN 125A 125B 126 A 126B 126C 126D 126E 126F 127A 127B 129 A 129B 129C 129D 129E 129Q viei 131B 135A 135B 135C 135D 135E 135G? 141A 141B CN 142B SAMPL 94-NM-94-NM-94-NM-94-NM-94-NM-94-NM-94-NM-94-NM-94-NM-94-NM-94-NM-94-NM-94-NM-94-NM-94-NM-94-NM-94-NM-94-NM-94-NM-94-NM-94-NM-94-NM-94-NM-94-NM-94-NM-94-NM-94-NM-94-NM-94-NM-94-NM-95-NM-95-NM-95-NM-95-NM-146 COMMENTS chlorite schist oriented sheared psammite oriented oriented gt-> sill-rich mats oriented GARNET Y,split Y,split Z Z Z Z Z Z Z Z Y,split Y.split Y.split Y.split Y.split Y.split 00 H >-> C N C N C N > >< C N C N C N C N > TYPE gritty psammite gritty psammite med-gray pelitic schist semi-pelite pale-green pelite semi-pelite psammite SAMPLE 95-NM-143 95-NM-145 95-NM-145B 95-NM-147 95-NM-148 95-NM-149 95-NM-150 95-NM-151 95-NM-153A 95-NM-153B 95-NM-153C 95-NM-154B 95-NM-155A 95-NM-155B 95-NM-156 95-NM-157A 95-NM-157B 95-NM-158 95-NM-158B 95-NM-158C 95-NM-159A 95-NM-159B 95-NM-159C 95-NM-160 95-NM-162 95-NM-164A 95-NM-164B 95-NM-165A 95-NM-165C 95-NM-165D 95-NM-166A 95-NM-166B 95-NM-167 95-NM-167B 147 COMMENTS oriented gt chloritized large actinolite splays GARNET Y.split Y.split Y.fract'd Y.split Y.split Z Y.fract'd Y.split Y.fract'd > Z Y.split Y.split Y.split Y.split Y.split Y.split Y.split Y.split? Y.split 00 H Y(2) Y(2) Y(2) TYPE gritty psammite calc-pelite calc-pelite calc-pelite calc-pelite calc-pelite amphibolite calc-silicate fu? calc-silicate SAMPLE 95-NM-168 95-NM-168B 95-NM-169 95-NM-170A 95-NM-170B 95-NM-171A 95-NM-171B 95-NM-172 95-NM-173A 95-NM-173B 95-NM-175 95-NM-175A 95-NM-175B 95-NM-176A 95-NM-176B 95-NM-176C 95-NM-176D 95-NM-176E 95-NM-177 95-NM-178 95-NM-184A 96-NM-200A 96-NM-200B 96-NM-201A 96-NM-201B 96-NM-202A 96-NM-203A 96-NM-203B 96-NM-204A 96-NM-205A 96-NM-205B 96-NM-207A 96-NM-207B 96-NM-208A 148 COMMENTS oriented; looking N on thin section GARNET Y.split Y,split Y.split Y,split Y.split Y.split Y.split Y.split oo H Y(2) > > >< I* > >< >* TYPE amphibolite psammite SAMPLE 96-NM-208B 96-NM-209 96-NM-210A 96-NM-214A 96-NM-214B 96-NM-214C 96-NM-214D 96-NM-214E? 96-NM-214F 96-NM-214G 96-NM-214H 96-NM-214J 96-NM-214K 96-NM-214L 96-NM-215A 96-NM-215B 96-NM-217A 96-NM-217B 96-NM-219A 96-NM-219B 96-NM-220A 96-NM-220B 96-NM-221A 96-NM-221B 96-NM-223 96-NM-225A 96-NM-225B 96-NM-225C 96-NM-225D 96-NM-225E? 96-NM-225F 96-NM-226A 96-NM-226B 96-NM-226D 149 COMMENTS extreme compositional segregation fold GARNET Z Y.split Z Y.split Z Z Z Z Z Z Z Z Z oo H >- >< >- >- >* CN >-TYPE marble mylonite marble marble marble marble, coarse-gr marble marble marble marble SAMPLE 96-NM-229A 96-NM-230A 96-NM-230B 96-NM-230C 96-NM-231A 96-NM-231B 96-NM-231C 96-NM-232A 96-NM-232B 96-NM-233A 95-NM-235A 96-NM-235B 96-NM-235C 96-NM-236A 96-NM-237A 96-NM-237B 96-NM-238A 96-NM-238B 96-NM-239A 96-NM-240A 96-NM-240B 96-NM-242A 96-NM-242B 96-NM-242C 96-NM-242D 96-NM-242E 96-NM-242F 96-NM-242G 96-NM-243A 95-Big-2 95-Big-5A 95-Big-5B 95-Big-6 150 cu C+H O '3 COMMENTS w a T/S CN Y(2) TYPE TYPE floal float float MPLE < m o < 00 PQ oo ON © -12A -12B r~H •F-IA •F-IB MPLE •Big-Big--Big-Big-Big-Big--Big-PQ -Big-Big- 1 PH U •F-IA •F-IB < 00 in ON in ON in ON IT) os in ON m ON in ON in ON m ON m Os NO ON NO ON NO ON 151 Appendix III: Mineral Analy! 152 CD ~ 2 2 CQ os CQ °° CQ "> CQ CQ ^ CQ •* CQ ^ CQ <^> CQ "3. 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B CN o « B 00 < 00 O N r - O N i n © vo T f © CN © © T f 00 T f vo © © © CN © i n © r o CN T f d d d d d vd oo vo r» vq © © T f d r o CN CN r o © © © T f O N CN © © © CN © r o © O N oq © T f CN T f © © © © © vd oo VO O N © VO © . V O T f © CN CN CN r o © r--oo O N T f CN © © © r o © © r o CN T f © © © © — i VO CN VO © Os T f CN C N C N © © r o CN CO © © © © © CN © VO © T f CN T f d d d © © vd oo r o r o ° ° O CN CN O »-i o < U PL, % 9 o S o ^ 5 tf s r i 153 PQ os PQ ° ° PQ <> PQ ^> PQ ^ PQ ^ 0 3 c o CQ o" PQ — ro i n T T . ro <n O O oo 00 m o o i n o ro T t O T t m •n <-< od © oo o o d d © OS © ro — i CN i n ro O cs i n o T t — I ro —< «n ro vi -i oti oq (-• od o s oo o o o 00 o 00 CN OS ro © © d d OS d p p oq d o s o s T t ro o 00 *—I o r—i ro CN d d d d OS d ro — i O ro vo ro O >n d os O vo m CN i n —; CN d —<° o s ro ro CN CN CN o p VO o o <n T t T t CN d d d d T t d o os o o o VO o VO o >n vo CN CN CN d d d d T t d — i CN — i — i T t co o O T t r ~ O ro oo oo o os os o 1—1 o o 00 o o ro VO o T t d d d d OS d ro VO O O CN VO I— T t vo >n o oo oo m >—i oo r o <—i O T t CN — ; T t vo r o —< od CO T t r o co o s O Os o s 00 o o CN m Os r o d d d d OS d r o CN O O d —<' ^ " CN >—i o o o T t »—H m p OS r o OS i n d d d d d r o OS —I f~ i n vo OS OO CN r o —I vo r o —i O m r o p o p d OS CN O CN OS d od o s >n CN OS o i n r o r o i n CN T t d d d d od d CN T t CN 00 o VO oo T t r o d d d d 00 d CN m i n CN CN Os >n o o s T t r o r- o oo r-i T t OS VO OS , , CN f—H r o CN T t d d d d 00 d r o r o CN tN O O w h <, U 2 O CN O cs cs cs CN U ffl Zi « ft - H CN PQ o t3 CN Os r o r o _ r o m Os OO r o OS o o r o o o VO O T t T t p o •—i T t 1—1 OS d o d d T t CN d d d d CO d PQ ° ° PQ ^ CQ ^ PQ 5 PQ n PQ £ j "& CN o v^ 00 O O OO CN <-< 00 CO CN vo r o — i o o d r o T t r o o T t CN o •—' i n o CN d d d d d i n vq vo vq r o O oo p m m o m oo o r o »—1 00 T t r o r o T t r o 1 - 1 r~ d I - H CN os d d d d 00 d oo r o r o r o vo T t r~ o s o vo o T t r o <-i 00 o oo o T t o VO VO CN OO o p *—< CN CN r o d d d d OS d Os vo 00 p i n T t i n —'< od co O r o o o OS OS oo o o o >n 00 m o o T t OS d d d d OS d vo T t , , Os VO o •n . - H Os r o 00 o o o VO o i n CN T t CN i n T t d d d d OS d OS _ _ >n T t i n < od r o OS 00 00 T t i n T t - H r o p CN CN © o s © O OO CN od d o s d OS p o p CN o CN t--o 00 CO d d d d OS d r o © o T t OS r--CN CO d d d d od d T t — I T t C~~ o o O OS oo o oo o oo o o o T f VO CN o OS CO d d d d OS d T t i n —< OO O 00 o CN o p OS p r o r o O m T t d d d d OS d r o r o ° ° O CN CN O < U PH o o o o § q CS CS CS CN U CQ Z W PH 154 Appendix IV: Structural Measurements 155 o o. •5 Gt split trend plunge o FT trend m NO CN Stretching lineation plunge Stretching lineation trend Mineral lineation plunge NO Mineral lineation trend OO dip H CZ) strike dip o in CN CN r--T f m o CO 00 T f m T f ro NO CN o CO T f T f NO CN ro NO T f o CN o CN ON T f CN ON CN 00 CN T f CN oo m ro m T f in ro T f T f T f T f T f CO T f ro ST strike in NO in NO o ON m ON ON T f T f r-- r s 00 NO 00 ON r -T f NO CN ON o NO ro o ro ON o oo O O T f in NO o oo ro m ON T f oo oo <N C -T f oo o ON ON CO ro ro r-ro Grid position > 1300| 1275| 1000J m CN oo m CN ON o in ON 1025| oou 1325| 1325| 1375| 1575| 1625| 39251 42001 3550| 3450| O m r -m CN r -o o ON m r -o in o in CN T f m r~ r -o o ro m CN CO o m T f in CN r -o o NO in <N NO O O NO Grid position X 1400 1500 2050 2075 2020 2250 2350 2825 28751 29001 28751 30251 30531 3100| 29251 32251 30251 17751 1775| 19751 1650| 1700| 24001 25001 24001 26001 2675 30251 30501 31501 3275 j 1 Outcrop # 1 J94-NM-2 1 J94-NM-3 I94-NM-4 [94-NM-5 1 [94-NM-6 1 I94-NM-7 1 I94-NM-8 1 I94-NM-9 1 94-NM-lO 1 94-NM-l 1 1 94-NM-l 2 1 94-NM-l 3 1 94-NM-l 4 1 94-NM-l 5 | 94-NM-l 6 | 94-NM-l 7 1 94-NM-l 8 1 94-NM-l 9 1 94-NM-20 1 94-NM-21 1 94-NM-22 1 94-NM-23 1 94-NM-24 1 94-NM-25 1 94-NM-26 1 94-NM-27 1 94-NM-28 1 94-NM-29 1 94-NM-30 1 94-NM-31 1 94-NM-32 | 156 •3 Gt split trend plunge ON FT trend NO NO CN Stretching lineation plunge Stretching lineation trend Mineral lineation plunge Mineral lineation trend P. •3 H <Z) strike dip VO ZZ I ON m CN <N CN OO <N T i -e n r o -cr T T CN c o c o CO CN o T T NO T T ON CN 00 CN o CN ON CN m m T t NO CO o r o T T CN o r o CM r o NO CN CN CN T I -CS r o <N <N c n c n T T </") CN ST strike r-r o t - -o r o O r - ON TT CN T T o ON ON NO O Tt-c~ T T OO T T CN o CN r -ON ON CN ON c n NO CN NO o NO ON r- NO > n r o NO r -r o O i n T f NO i n r o NO T t NO O CN CN CO NO r - r -<n Grid position > 1625| O o ON m t-o o o o 00 O c n r -c n CN r -< n CN NO O > n CO « n CN r o 1025| 1150| o o ON O « n o o > n r - -r - 1750| 1575| 1475| 1400| 1225| m ON 1150| 24251 29751 49751 Grid position X 3250 2525 2350 2350j 22501 2325 2175| 2250) 2225 j 27001 29001 30251 3050| 31751 30001 3025) 2950| 29001 28751 2625| 2650| 29501 28751 3200| | Outcrop # |94-NM-33 |94-NM-34 |94-NM-35 VO m • s z. •4 O N |94-NM-37 | |94-NM-38 | |94-NM-39 | |94-NM-40 | |94-NM-41 | J94-NM-42 | 94-NM-43 | 94-NM-44 | 94-NM-45 | 94-NM-46 | 94-NM-47 | 94-NM-48 | 94-NM-49 | 94-NM-50 | 94-NM-51 | 94-NM-52 | 94-NM-53 | 94-NM-54 | 94-NM-55 | 94-NM-56 | 94-NM-57 | 94-NM-58 | 157 CO O OH '•3 Gt split trend plunge r - -C N VD C N r o m o o C N T f C N i n C N FT trend T f r -00 OO ON C O T f C - -T f r~-C O t - -C O VO C N Stretching lineation plunge o Stretching lineation trend o 00 Mineral lineation plunge o o C N T f C N T f C N o T f r o i n r o r o C N o o C N T f C N ON Mineral lineation trend o o ON ON o o VO VO r o VO r -m C N r -o o C N 00 VO i n VD r -00 i n dip H o o strike dip I n C N r o r o r~ C N o o r-C N r o C N > n T f o m ON C O T f C O VO T f r~ C N r -C N VO ON VO T f o C N r o r o r - C O r o T f C N 00 i n C N i n r o i n C N o o C N o C O 00 C N m C N i n C N m C N o o m i n IS| strike r -C N c o r-~ i n r -T T o o r ~ T t r ~ i n ON T f C O m c o C O o o C O C N r- » n T f C O O m VO C N VO o r o r~ ON C N C N i n r o r o r o c o o C O r ~ T f r o ON C N r o T f C O m r o C O O m v o r o C N C N C N i n r o T f o C O NO o r o T f o r o i n o o C N Grid position > 5025 5100 1900 2050 2150| 2225 2275 23251 37001 3650J 36001 3650| 47251 47251 5175| 5150| 5150| 52001 52751 5300| 37751 3800] 3750| 37751 3850| 3800| 51251 5050| 50251 47001 Grid position X 3250 3350 2925 2900 2900 2900 2925 2925 4775 4675 4550| 43001 2150 28751 34001 34751 3550J 36251 31751 3850| 40501 39751 3750 [ 34751 34251 33751 3300J 3225 j 32251 2950| 1 Outcrop # |94-NM-59 |94-NM-60 |94-NM-61 | |94-NM-62 |94-NM-63 | |94-NM-64 | |94-NM-65 |94-NM-66 J94-NM-67 j [94-NM-68 | |94-NM-69 | |94-NM-70 | |94-NM-71 | 94-NM-72 | 94-NM-73 | 94-NM-74 | 94-NM-75 | 94-NM-76 | 94-NM-77 | 94-NM-78 | 94-NM-79 | 94-NM-80 | 94-NM-81 | 94-NM-82 | (all isoclinal | folds on same| surface) | 94-NM-83 | 94-NM-84 | 94-NM-85 | 94-NM-86 | 94-NM-87 | 94-NM-88 | 94-NM-89 | 158 cn O dip | Gt split trend plunge o o m o C O c n C N C N C N o C O C N C N C O C N C N C N T T C N C N FT trend o o o o Os r -o o r -o t--T f r-Os V O > n V O T t " O r-Os C N Stretching lineation plunge C O C N Stretching lineation trend V O Mineral lineation plunge Tt- Os o C N C N C N ZZ 1 o C O C N v o C N T t -C N c n C S c n o o r o Mineral lineation trend OO V O os V O o o o c o v o V O V O i n r~-i n V O OO r -T f o o T f r - r -i n r - V O V O g. •3 C N C O c n *cr V O H strike o C O v o C O V O C N dip c n C N r - r o C N T I -C N m C N C O C N v o C N m C N m C N < n C N T T C N v o C N C N C N C O C N 00 C N O C N c n c n C O T i -c s f N C O C N C N v o C N T f C N v o c s C O C N o C S r -C N m o r o ST strike 00 V O r o cn C N V O C O o C N 00 c n c o Os o T f c n v o r o C O T * - o r o o o c n Tt-C N C N OS C N C N Os C N •cr Os C N c n r -C N v o m V O Tt-C O r o r -C N C O i n r o C O o v o C O C O c n i n C N r o C N r s C O O C O Grid position 1 > 24251 2525| 26751 29001 32251 33751 3750| 39251 45001 47751 4750| 47001 47751 46001 4050| 42001 49001 48751 49001 49001 51501 5275| 54251 Grid position 1 X 2950| 2975 2950 2875 29501 30001 31751 30751 26251 25001 2675 i 2850! 28501 29501 3925| 39001 36751 3600) 3525| 34001 35501 3650| 3675| Outcrop # | 94-NM-90 | 94-NM-91 1 94-NM-92 1 94-NM-93 1 94-NM-94 | 94-NM-95' 1 94-NM-96 1 94-NM-97 1 94-NM-98 1 94-NM-99 | 94-NM-100 | 94-NM-l 01 1 94-NM-102 I 94-NM-103 1 94-NM-104 1 94-NM-105 1 94-NM-106 1 94-NM-107 1 94-NM-108 1 94-NM-109 1 94-NM-l 10 | 94-NM-l 11 1 94-NM-l 12 1 94-NM-l 13 | 94-NM-l 14 | 159 1 dip | |Gt split | trend 1 plunge in o C O oo CN O CO oo o |FT 1 trend oo r- o ON o NO CN o CN in oo CN | Stretching lineation 1 plunge CO NO CN o CO o | Stretching lineation | trend NO ON OO NO ON NO CO r -| Mineral lineation u M C J 3 "E CN 140 T f in ro r -CN ON CN o T f •n ro | Mineral lineation 1 trend r -T f r-CN <--ON NO o r~ O C- OO NO in NO 1 dip H 0 0 | strike 1 dip •n CN CN CN in CN in CN o CN o CO o O o CN in T f NO ro r -CN ON CN in CN o ro in ro CN CO T f o T f i - -CN i n C O in ro in in C O o CN C O NO in ON in 0 0 oo CN m rN CN |ST | strike CN T f ON C O G O CN T f oo C O no in ro ON C O ro oo ON CN ro oo CN CO ON CN C O in ro ro O CN ro O CN ro O ro ro O o T f C O •n CN ro o CN ro ON in ro n m in ro ro CO 0 0 CN C O oo CN 0 0 o ro r -CN m ON CN O ro in NO rsi ON oo o ro o ON m i n o in | Grid position 1 I* 2625 26001 25751 25751 49501 49001 4925! 49251 4500! 4975! 50251 54751 54251 52501 60251 5950| 57751 5525| 49001 66501 65001 48251 48251 5000| 50751 20251 18751 | Grid position 1 X 1 35251 1 3475) 1 3275! [ 3200! 1 3175 | 32501 32501 33251 35001 3875! 3875! 41001 42251 39251 46001 45501 44251 4325| 38001 45501 45001 14251 20251 2000| 19001 2725] 26751 1 Outcrop # 1 194-NM-l 15 1 194-NM-l 16 1 I94-NM-117 1 I94-NM-118 1 194-NM-l 19 1 194-NM-l 20 1 I94-NM-121 1 194-NM-l 22 1 194-NM-l 23 1 194-NM-l 24 1 194-NM-l 25 1 194-NM-l 26 1 194-NM-l 27 1 194-NM-l 28 1 194-NM-l 29 1 194-NM-l 30 1 I94-NM-131 1 194-NM-l 32 1 194-NM-l 33 1 194-NM-l 34 1 I94-NM-135 1 I95-NM-140 1 I95-NM-141 1 I95-NM-142 1 I95-NM-143 1 I95-NM-144 1 I95-NM-145 1 I95-NM-146 1 I95-NM-147 1 I95-NM-148 1 I95-NM-149 1 160 CO 1 dip | o ON o ON |Gt split | trend O r~ cn r- o O N plunge N O ON cn T f ON 00 O |FT | trend o 00 cn cn O C N O O t - -C N C N O N cn N O C N o N O C N | Stretching lineation 1 plunge ro T f T f C N | Stretching lineation | trend O O O o r-N O C N | Mineral lineation 1 plunge C O C S 00 cn C N C N c- T f | Mineral lineation | trend o ON >n cs ro oo cn N O O N t ~ O O O r— 00 cs cn cn C N dip | o T f r-H CZ) | strike cn N O cn 1 dip T f C N oo C O «n N O C O N O C O C N cn C N o C N cn cn •<t C O C N T i - N O TT C S C N C N cn C N in cs ro C S O O cn C N C N C N C N cn C N T f TI-C S C N C N N O C N C O C N cn C N m N O C N C N N O C N C O < N C O C N |ST | strike N O N O C O o o C O O C N C N O O N O 00 N O o r- cn cn C O ro ro O ro ro ro cs C N C O cn o C O cn ro o ro O N r-C N ON ON C N T t o ro N O o ro C N O O C N 00 o C O r-ON C N N O TI-C S C N o ro cn r~ C N C O r~-C N cn r-cs o o C O 00 N O C N cn T f C N m T f cs ON N O C N T f T f C N 1 Grid position 1 17001 16501 1575| 26501 26001 24251 25001 22001 18501 51001 56751 55001 56501 57251 5750| 54001 53251 5300| 52751 52251 1 Grid position 1 X 1 2625 1 -2550| 1 2500| 1 2600J 1 2500| 1 2375 1 2525 1 2525 1 2525 1 2600 j 3075 1 3000 1 2875 1 2875 1 28251 o o N O in f-in o o N O cn cs N O o cn N O 1 Outcrop # 1 195-NM-150 1 I95-NM-151 1 195-NM-152 1 195-NM-153 1 195-NM-154 1 195-NM-155 1 195-NM-156 1 195-NM-157 1 195-NM-158 1 195-NM-159 1 195-NM-160 1 95-NM-161 195-NM-162 1 195-NM-163 1 195-NM-164 1 195-NM-165 1 195-NM-166 1 195-NM-167 1 195-NM-168 1 195-NM-169 1 195-NM-170 1 I95-NM-171 1 195-NM-172 1 195-NM-173 1 195-NM-174 1 195-NM-175 1 161 2 CO O OH o un Gt split trend in NO plunge oo ZZ FT trend cn o T f r-Stretching lineation plunge oo Stretching lineation trend o m Mineral lineation plunge CN T f CN Mineral lineation trend CN NO CN ON NO dip VI strike dip T f T f CN T t CN T f o CN CN cn m T f o >n o cn NO CN NO o O ST strike VO CN in cn o CN CO cn o cn ON o oo CN oo CN m ON ON CN o ON CN in NO CN o CN cn oo T f Grid position >- 5325J 5175J 5250| 36251 3550| 48001 48001 49751 5075) 50001 49501 49001 54251 54251 5600| 4525| 46251 45751 44001 44751 49001 4800| 48001 45001 4250] 42751 4250| 4300| 4300| Grid position X 2100| 2700| 2775 4850 4925 30501 33251 35251 34751 35751 36001 36001 3800| 39251 4350| 25251 2500J 2400 j 2375 j 23251 3150| 3100| 30501 30001 28251 2800| 27751 27501 2650| (Outcrop # |95-NM-176 j (95-NM-177 | |95-NM-178 | (95-NM-180 | |95-NM-181 | |95-NM-182 | |95-NM-183 | |95-NM-184 | |95-NM-185 | J95-NM-186 | ! 95-NM-187 | 95-NM-188 | 95-NM-189 1 95-NM-190 | 96-NM-200 | 96-NM-201 | 96-NM-202 1 96-NM-203I 96-NM-204 | 96-NM-205 | 96-NM-206 | 96-NM-207 | 96-NM-208 | 96-NM-209 | 96-NM-210 | 96-NM-211 | 96-NM-212 1 96-NM-213 | 96-NM-214 | 96-NM-215 | 96-NM-216 | 96-NM-217 | 96-NM-218 1 96-NM-219 | 96-NM-220 | 96-NM-221 | 96-NM-222 | 96-NM-223 | 162 o P. •3 Gt split trend plunge | o FT trend cn r--Stretching lineation | plunge | Stretching lineation | trend Mineral lineation | plunge | VO VO Mineral lineation | trend | o o 00 CN TT dip | H co strike dip | oo ro VO CN CN ON m VO r--CN ON ST strike o ON tN m ON tN O O N tN >n ON CN 00 r-CN o oo CN O ON O oo CN Grid position 1 > 41251 56751 58251 58001 58001 58001 58501 59001 59751 60751 61501 62001 62751 63501 0099 66501 67001 68001 6825| Grid position 1 X 27001 lO IN VO o m V O m CN i/i m r~ cn o o m m o o >n CN o o o o o o in ON o <n oo >n CN oo in CN 00 o in 00 in 00 m CN ON o o oo Outcrop # 96-NM-224 96-NM-225 96-NM-226 1 96-NM-227 96-NM-228 1 96-NM-229 1 96-NM-230 1 96-NM-231 1 96-NM-232 96-NM-233 1 96-NM-234 1 96-NM-235 1 96-NM-236 96-NM-237 96-NM-238 96-NM-239 1 96-NM-240 1 96-NM-241 I 96-NM-242 163 Quartz veins 1 dip 1 o os o Os o Os Quartz veins 1 strike cs •<f o ro rs ro plunge | oo rs rs Late folds trend \ r~-o o Intersection lin. | plunge | o cs 00 cs m cs o o cs o o cs T f rs Intersection lin. | trend | m o T f m vo r-Os T f T f ov T f 00 o cs Late crenulation | dip I T f i—• Late crenulation | strike | OO CS T f 00 cs Grid position 1 13001 12751 10001 «n cs oo m cs OS o in OS in cs o o o m cs ro m cs ro m ro m m rs VO 39251 42001 35501 34501 o m r~ m cs o o o\ m O m o m cs T f m t- o o ro m cs ro o T f »n cs r~ o o vo m cs VO o o VO m cs VO o o Os m r-00 o o 00 o m m cs r-m cs VO o ro in (N m 10251 11501 o o Os Grid position 1 X 1400| 15001 2050! 2075! 2020| 22501 23501 2825| 28751 2900| 2875| 3025| 3053| 31001 2925| 3225| 30251 17751 1775| 19751 1650| 1700| 2400| 25001 2400 2600] 26751 30251 3050] 3150| 3275 32501 2525 23501 2350| 2250| 2325 2175 2250! 2225 2700 2900 3025 Outcrop # 94-NM-2 1 94-NM-3 94-NM-4 94-NM-5 1 94-NM-6 94-NM-7 94-NM-8 94-NM-9 1 94-NM-10 94-NM-l 1 1 94-NM-12 1 94-NM-13 1 94-NM-14 94-NM-15 94-NM-16 94-NM-17 94-NM-18 94-NM-19 1 94-NM-20 1 94-NM-21 94-NM-22 94-NM-23 94-NM-24 94-NM-25 94-NM-26 94-NM-27 00 (S 1 5 z 1 T f Os 94-NM-29 94-NM-30 94-NM-31 94-NM-32 94-NM-33 94-NM-34 94-NM-35 I94-NM-36 I94-NM-37 :94-NM-38 94-NM-39 1 I94-NM-40 I94-NM-41 94-NM-42 94-NM-43 94-NM-44 I94-NM-45 164 o Quartz veins | dip I o ON Quartz veins | strike in CO plunge o CN NO CN m ON Late folds trend m CO m ON ON Intersection lin. | plunge | oo CN CN o CO NO CN 00 m CN T T CN m CN 00 CN m CO o T T T T o T T ON CN Intersection lin. | trend o O o CN CN o o CO m m o o ro NO o 00 00 o ON O Late crenulation | dip | oo NO Late crenulation | strike oo ro Grid position 1 o m oo m r~ r - 1750| 15751 1475| 1400| 1225| in ON 1150| 2425| 29751 49751 50251 5100| 1900| 20501 21501 22251 22751 2325| 3700| 36501 36001 36501 4725| 4725| 5175] 51501 5150| 52001 5275| 53001 37751 3800| 3750| 3775| Grid position 1 X 30501 3175 30001 3025 2950| 2900| 2875 2625 2650| 2950| 2875 32001 32501 33501 2925 2900 29001 29001 2925 2925 4775 4675 4550 43001 2150 2875 3400| 3475 35501 3625 3175| 38501 4050| 3975 3750| 3475 Outcrop # 1 94-NM-46 1 94-NM-47 1 94-NM-48 1 94-NM-49 1 94-NM-50 1 94-NM-51 1 94-NM-52 1 94-NM-53 1 94-NM-54 1 94-NM-55 1 94-NM-56 1 94-NM-57 1 94-NM-58 1 94-NM-59 1 94-NM-60 1 94-NM-61 1 94-NM-62 1 94-NM-63 1 94-NM-64 1 94-NM-65 1 94-NM-66 1 94-NM-67 1 94-NM-68 1 94-NM-69 1 94-NM-70 1 94-NM-71 I 94-NM-72 1 94-NM-73 1 94-NM-74 1 94-NM-75 1 94-NM-76 1 94-NM-77 1 94-NM-78 1 94-NM-79 1 94-NM-80 1 94-NM-81 1 94-NM-82 1 94-NM-83 | 165 Quartz veins | Quartz veins | . strike plunge Late folds trend Intersection lin. plunge | oo NO in CN NO CN m CN o o CN CN CN ON oo T f CN r— CN cn NO T f CN Intersection lin. trend NO «n m T f 8 T f ON cn O o 00 00 cn ON 00 oo ON 00 ON ON CN oo T f o o Late crenulation o t ~ -CN CN CN cn Late crenulation strike NO cn 00 o CO Grid position 1 > 3850| 3800| 5125| 5050| 5025| T f CN 25251 26751 2900| 32251 3375| 3750| 3925| 45001 4775| 4750| 4700| 47751 4600| 4050| 4200| 4900| 4875| 4900| 4900| 5150| 5275| 5425| 26251 2600| 2575 2575| 4950| 4900| 4925| 4925| 4500| 4975| Grid position 1 X 3425 33751 3300| 3225 3225| 29501 2950 2975| 2950| 2875| 2950 j 3000 i 3175| 3075| 26251 2500| 2675| 2850| 2850| 2950| 3925| 3900| 3675| 3600| 3525| 3400| 3550| 3650| 3675| 35251 3475| 3275| 3200! 3175| 32501 3250| 3325| 3500! 3875| Outcrop # 94-NM-84 1 94-NM-85 1 94-NM-86 | 94-NM-87 1 94-NM-88 1 94-NM-89 1 94-NM-90 | 94-NM-91 | 94-NM-92 | 94-NM-93 | 94-NM-94 | 94-NM-95 | 94-NM-96 | 94-NM-97 ! 94-NM-98 | 94-NM-99 | 94-NM-100 | 94-NM-101 | 94-NM-102 | 94-NM-103 1 94-NM-104 1 94-NM-105 1 94-NM-106 | 94-NM-107 | 94-NM-108 | 94-NM-109 | 94-NM-l 10 | 94-NM-l 11 | 94-NM-l 12 | 94-NM-l 13 | 94-NM-l 14 | 94-NM-l 15 1 94-NM-l 16 | 94-NM-l 17 1 94-NM-l 18 1 94-NM-l 19 1 94-NM-120 | 94-NM-121 1 94-NM-122 1 94-NM-123 | 94-NM-124 | 166 1-1 a O Quartz veins | dip 1 Quartz veins | strike plunge | Late folds trend Intersection lin. | plunge | oo ro CN CN r - oo Intersection lin. | trend | un co m cn CN CO Late crenulation | dip 1 ro vO NO r~ ON NO Late crenulation | strike o co oo cn o m Grid position 1 > 50251 54751 54251 52501 60251 59501 57751 55251 49001 66501 65001 48251 48251 50001 50751 20251 18751 17001 16501 15751 26501 26001 24251 25001 2200| 18501 5100| 56751 55001 56501 57251 5750| 5400| 5325| Grid position 1 X 38751 . 41001 42251 39251 46001 4550 44251 43251 3800! 4550( 4500| 1425 2025 20001 19001 . 2725) 2675 26251 25501 25001 26001 25001 2375 ' 2525 2525 2525 26001 3075 3000| 2875 2875 2825 1600! 1575 Outcrop # 94-NM-l 25 1 94-NM-l 26 1 94-NM-l 27 1 94-NM-l 28 1 94-NM-l 29 1 94-NM-l 30 1 94-NM-131 1 94-NM-l 32 1 94-NM-l 33 1 94-NM-l 34 1 94-NM-l 35 1 95-NM-140 1 95-NM-141 1 95-NM-142 1 95-NM-143 1 95-NM-144 1 95-NM-145 1 95-NM-146 1 95-NM-147 1 95-NM-148 1 95-NM-149 1 95-NM-150 1 95-NM-151 1 95-NM-152 1 95-NM-153 1 95-NM-154 1 95-NM-155 1 95-NM-156 1 95-NM-157 1 95-NM-158 1 95-NM-159 1 95-NM-160 1 95-NM-161 1 95-NM-162 1 95-NM-163 1 95-NM-164 1 95-NM-165 1 95-NM-166 1 95-NM-167 | 95-NM-168 | 95-NM-169 | 95-NM-170 | 95-NM-171 | 95-NM-172 | 167 o Quartz veins | dip 1 Quartz veins | strike plunge | Late folds trend Intersection lin. | plunge | Os CN o Intersection lin. | trend in O o Late crenulation | dip ] Late crenulation | strike Grid position | 53001 52751 5225| 5325| 51751 52501 36251 35501 48001 48001 49751 50751 50001 4950| 4900| 5425| 54251 56001 45251 •46251 4575| 4400| 44751 4900| 4800| 4800| 45001 4250| 4275| 4250| 4300| 43001 4125| 56751 5825| Grid position | X 0091 1625! 16501 2100! 27001 2775| 4850| 4925! 3050) 3325J 3525| 3475| 3575J 3600] 3600| 3800| 3925J 4350! 2525) 2500| 24001 2375 2325 3150| 3100] 30501 30001 28251 28001 2775 2750| 26501 2700 1625 1650! Outcrop # 1 95-NM-173 1 95-NM-174 1 95-NM-175 1 95-NM-176 I 95-NM-177 1 95-NM-178 | 95-NM-180. | 95-NM-181 | 95-NM-182 | 95-NM-183 | 95-NM-184 | 95-NM-185 | 95-NM-186 | 95-NM-187 | 95-NM-188 | 95-NM-189 | 95-NM-190 | 96-NM-200 | 96-NM-201 | 96-NM-202 ] 96-NM-203 | 96-NM-204 | 96-NM-205 | 96-NM-206 | 96-NM-207 | 96-NM-208 | 96-NM-209 | 96-NM-210 | 96-NM-211 | 96-NM-212 | 96-NM-213 | 96-NM-214 | 96-NM-215 | 96-NM-216 | 96-NM-217 | 96-NM-218 | 96-NM-219 | 96-NM-220 | 96-NM-221 | 96-NM-222 1 96-NM-223 1 96-NM-224 ;96-NM-225 J96-NM-226 168 ( to Quartz veins | g. •3 Quartz veins | strike plunge | Late folds trend Intersection lin. | plunge | o Intersection lin. | trend o CN O CN Late crenulation | dip | Late crenulation | strike Grid position 1 > 58001 58001 5800| 58501 59001 59751 60751 61501 62001 62751 63501 0099 1 6650J 6700| 68001 68251 Grid position 1 X cn CN cn cn r-ro o o m cn t— o o cn (N o o o o o o «n ON o «n 00 >n CN 00 cn CN 00 o cn oo cn r-00 >n CN as o o 00 Outcrop # 1 96-NM-227 1 96-NM-228 1 96-NM-229 1 96-NM-230 1 96-NM-231 1 96-NM-232 1 96-NM-233 1 96-NM-234 1 96-NM-235 1 96-NM-236 1 96-NM-237 1 96-NM-238 1 96-NM-239 1 96-NM-240 1 96-NM-241 1 96-NM-242 | 169 

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