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Siliceous dolomites in the Horsethief Creek aureole, S.E. British Columbia: fluid infiltration and mineral.. Niermann, Maria C. 1998

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SILICEOUS DOLOMITES IN THE HORSETHIEF CREEK AUREOLE, S.E. BRITISH COLUMBIA: FLUID INFILTRATION AND MINERAL REACTIONS by MARIA C. NIERMANN A THESIS SUBMITTED IN PARTIAL FULFILLMENT OF THE REQUIREMENTS FOR THE DEGREE OF MASTER OF SCIENCE in THE FACULTY OF GRADUATE STUDIES (Department of Earth and Ocean Sciences) We accept this thesis as conforming to the required standard . THE UNIVERSITY OF BRITISH COLUMBIA November, 1998 © Maria C. Niermann, 1998 In presenting this thesis in partial fulfilment of the requirements for an advanced degree at the University of British Columbia, I agree that the Library shall make it freely available for reference and study. I further agree that permission for extensive copying of this thesis for scholarly purposes may be granted by the head of my department or by his or her representatives. It is understood that copying or publication of this thesis for financial gain shall not be allowed without my written permission. The University of British Columbia Vancouver, Canada Department Date VrA/. 9, MR DE-6 (2/88) ABSTRACT Potassic siliceous dolomites in the Horsethief Creek aureole, S.E. British Columbia, Canada record peak contact-metamorphic conditions of -2.2 kbars and 650 °C in the inner aureole. The protolithic dolostones have varying but small amounts of disseminated quartz (< 5 volume percent) and K-feldspar (<15 volume percent) but have interbedded numerous chert nodules and layers. Subvertical joints and veins cut across layering. Six isograds can be identified: phlogopite; tremolite+K-feldspar; diopside; diopside+phlogopite; diopside+dolomite; and at highest grade forsterite. The prevalence of isobarically divariant mineral assemblages, paucity of high XCQ2 mineral assemblages, widespread development of hydrous minerals often in high abundance (up to 65 volume percent) and a ~7 per mill (%0) depletion (from =23 %o to =16 %o) in 8180 are all evidence for substantial syn-metamorphic fluid flow. The almost complete de-dolomization of the inner and middle aureole suggests that metamorphism was not isochemical and that silica-metasomatism took place. Heterogeneity in the distribution of index minerals and the potassic phases reflect varying bulk composition as well as variable fluid infiltration. Infiltration-driven near-isothermal reactions are documented in the rocks in form of numerous samples showing a texture interpreted to record prograde reaction of diopside to tremolite. A petrographic break is observed at the tremolite+K-feldspar isograd. While rocks up-grade of this isograd are dominated by isobarically divariant assemblages, rocks down-grade are characterized by isobarically univariant assemblages. This break may record a sudden change in infiltration character. The proposed model of contact metamorphism in the impure siliceous dolomites of the Horsethief Creek aureole is: 1) Intrusion of the Horsethief pluton triggered metamorphic reactions in the siliceous dolomites and the underlying pelitic rocks. 2) Buoyant fluids evolving in the metapelites at temperatures >530 °C rose along subvertical joints in and across the different ii lithologies and infiltrated the dolostone. 3) Infiltration of silica-bearing aqueous fluids triggered pervasive reaction in the siliceous dolomites. Rocks in the outer aureole did not experience significant infiltration because temperatures were too low to trigger the dehydration reaction in the underlying metapelites. Plutonic fluids may only have played a minor role, and bedding parallel flow likely was also only minor. iii TABLE OF CONTENTS ABSTRACT ii TABLE OF CONTENTS iv LIST OF TABLES viLIST OF FIGURES .'. viii LIST OF PLATES ix ACKNOWLEDGEMENTS x 1.0 INTRODUCTORY STATEMENTS 1 1.1 Structure of Thesis1.2 Introduction 1 1.3 The T-XC02 Diagram 7 2.0 GEOLOGICAL SETTING 1Q 2.1 Regional Geology 0 2.2 Local Geology 12 3.0 MINERALOGY AND MINERAL CHEMISTRY 18 3.1 Introduction 13.2 Metamorphic Zones 20 3.2.1 Phlogopite zone3.2.2 Tremolite + K-feldspar zone 28 3.2.3 Diopside zone 30 3.2.4 Diopside + Phlogopite zone 3iv 3.2.5 Diopside + Dolomite zone 32 3.2.6 Forsterite zone 34 3.3 Mineral Chemistry 6 4.0 CONTROLLING FACTORS ON THE DEVELOPMENT OF THE METAMORPHIC MINERAL ASSEMBLAGES 40 4.1 Introduction 44.2 Effects of Bulk-Compositional Variations 40 4.3 Thermal Controls 4 4.4 Fluid Infiltration 51 4.4.1 Nominally retrograde assemblages 52 4.4.2 Reaction of diopside to tremolite - Textural evidence for infiltration-triggered near-isothermal reaction 52 4.4.3 5'80 whole rock carbonate data 4 4.4.4 Silica metasomatism and intensity of infiltration 58 4.5 Discussion of Possible Fluid Flow Direction(s), Fluid Source(s) and Heat Advection 60 4.5.1 Heat flow - conductive or advective? 64.5.2 Discussion of the possibility of horizontal fluid flow in the direction of increasing temperature 62 4.5.3 Discussion of the possibility of horizontal fluid flow in the direction of decreasing temperature 5 4.5.4 Discussion of the possibility of vertical fluid flow 67 Metamorphism of the pelite and its implication for the siliceous dolomite 68 Metamorphism of the siliceous dolomite in the middle aureole 69 Metamorphism of the siliceous dolomite in the tremolite + K-feldspar zone 71 The 7_ - profile 71 v The inner aureole - a special case 72 4.6 The Petrographic Break - A Change in Infiltration Character 73 4.6.1 A model of diachronous heat and (horizontal) fluid flow 75 4.6.2 A model of vertical fluid flow 79 5.0 SUMMARY AND CONCLUSION 82 REFERENCES 85 APPENDICES Appendix A: Electronmicroprobe Data and Geothermometry Results 91 Appendix B: XRF Data 108 Appendix C: Stable Isotope Data 113 Appendix Cl: Stable Isotopes- Methods, Data and 813C - Signature 114 Appendix C2: Silica Weight Percent (whole rock) versus 8180 data - Evidence for Silica Metasomatism 116 Appendix D: Mineral Formulas and Molar Volumes 117 Appendix E: Sample Number Conversion 119 vi LIST OF TABLES Table Page 1. List of peak metamorphic assemblages 22 2. Selected results of quantitative analyses of silicates and carbonates 37 3. List of maximum stability ranges and 45 vii LIST OF FIGURES Figure Page 1. Introductory T-XC02 diagram 8 2. Map of the regional geology of the Horsethief Creek area 11 3. Geological map of the working area 13 4. T-XCQ2 diagram for KCMAS-HC system at P=2.2kb 19 5. Map showing the six isograds 21 6. F-content in phlogopites 39 7. Major element whole rock analyses 41 8. Mineral facies diagrams (projected on Qtz-Dol-Ksp plane) 43 9. Map showing situation after elimination of bulk compositional effects 46 10. T-XC02 diagram illustrating effects of thermal controls 47 11. Geothermometry data 49 12. T-XC02 diagram illustrating isothermal reaction 53 13. Oxygen isotope analyses 56 14. Oxygen isotope ranges of selected geological materials 57 15. Diagrams showing the process of 'annealing' of low-grade rocks 76 viii LIST OF PLATES Plate Page 1. View of the field area 15 2. Typical veining observed in siliceous dolomites 17 3. Phlogopite zone rock with normal low phlogopite abundance 24 4. a and b) High abundance phlogopite in deformed phlogopite zone rock 25 5. Phlogopite concentration in vein in rock from phlogopite zone 26 6. Typical 'protolithic' rock with chert nodules and veining 27 7. Typical tremolite - calcite - K-feldspar assemblage 9 8. a and b) Reaction texture of diopside to tremolite 31 9. Bleaching along chert layer 33 10. a and b) Cross cut of a vein with mineral assemblages indicating local silica enrichment 35 ix ACKNOWLEDGMENTS I like to thank the members of my supervisory committee Greg Dipple, Tom Brown, Kelly Russell and Lori Kennedy for their support and discussions during my time at UBC. Special thanks to my main supervisor Greg Dipple whose door was always open and from whom I learned a lot over the years. Thanks to the MDRU for letting me use their facilities, to UBC-geochron people namely Jim Mortenson and Rich Friedman for their ready help on thesis work, and to Mati Raudsepp for his patience with me during my S.E.M. and microprobe work. Thanks also to Terry Gordon and Dave Pattison who, from the beginning on, shared their enthusiasm for this project with me, the former adding a lot of extra fun to my first field season. Many thanks to my field assistants, especially Anne Bagley who probably was the best field assistant one can have in a rainy, snowy and physically very challenging season. The 'pain' of this hard season, as well as the second season in 1996, was significantly eased by the extra effort of 'Helicopter-John' and the fabulous support by the Madsons out of Invermere. I also like to express my thanks to Nathalie Marchildon, without whose help my start at UBC would have been so much more difficult, and to Peter Alt-Epping, Art Calderwood and Mathias Westphal for the interesting discussions and the phantastic support during the last 'hot' phase of writing this thesis. Many thanks also to my mother, especially for the financial support for my 'round the globe trips'. ' Danke' also to my friends back in / from Germany for their never-ending support and concern, especially to Magdalena & Michael. Very special thanx for the 'personal' support also to Susannah, Art, Shin Yi, Mathias and especially Peter who kept me going - Ta and Good Luck to you! x CHAPTER 1 INTRODUCTORY STATEMENTS 1.1 Structure of Thesis Most of the thesis addresses the petrography of mineral assemblages that developed in siliceous dolomites in a contact aureole adjacent to the Horsethief Creek pluton and the analysis of oxygen isotope data from these rocks. These data are used to develop a model of contact metamorphism in the Horsethief Creek siliceous dolomites. Additional data - carbon stable isotopes and whole rock compositions - are used in support of a model involving local, pervasive silica metasomatism as briefly discussed in Appendices B and C. I first discuss a few basics of the interpretation of contact-metamorphic siliceous dolomites in paragraph 1.3 by explaining some of the principles of the temperature (T) - XC02 diagram, where XC02 is the mole fraction of C02 in a binary H20 - C02 fluid. This diagram is routinely used in the interpretation of the metamorphic evolution of siliceous dolomites. 1.2 Introduction The mineralogy developing in chemically simple siliceous dolomites during metamorphism is sensitive to the amount and composition of fluid they interact with. Mineral abundance and composition commonly allow for determination of the amount of influx, fluid source(s) and possible metasomatic changes (e.g., Bucher-Nurminen 1982, Hover Granath etal. 1983, Ferry 1 1991, Holness 1992, Ferry 1994). Stable isotope analyses in combination with petrographic observations provide an additional powerful tool to decipher the infiltration history of these rocks (Lattanzi et al. 1980; Nabalek et al. 1984, Bowman 1994, Holness and Fallick 1997). The succession of mineral assemblages in siliceous dolomites of different contact aureoles is often observed to be different from each other. Assemblages vary in spatial distribution and/or can be of different variances depending on amounts of fluid infiltration and/or differences in flow direction. Ferry (1994) and Dipple and Ferry (1996) showed in a systematic way which mineral assemblage successions are expected to develop in an aureole depending on (a) whether the fluid flow is in direction of increasing temperature (up-T) or in the direction of decreasing temperature (down-T), and (b) the absolute amount of infiltration. They used a 1-dimensional model of coupled fluid flow and mineral reaction (and heat flow in Dipple and Ferry 1996) in rocks with 10 mol% quartz and 90 mol% dolomite, a bulk composition similar to that observed in many contact-metamorphic siliceous dolomites. In numerous contact aureoles siliceous dolomites with predominantly isobarically univariant assemblages are observed (Rice, 1977a,b; Suzuki, 1977). This makes it possible to trace reaction paths, and in many cases progressive metamorphism is proposed for these aureoles (i.e., high grade assemblages developed through all the lower grade stages that are now found in the aureole). These aureoles are interpreted to have experienced up-T fluid flow (Ferry, 1994). Other aureoles exhibit a change in the prevalent variance of assemblages over distance. Predominantly isobarically univariant or isobarically univariant 'key' ('key' with respect to the determination of fluid flow direction) and divariant assemblages are found throughout the aureole, except for the innermost part (usually <150 m from the contact with the intrusion). Here rocks contain periclase or forsterite in predominantly isobarically divariant assemblages (Holness 1992, Moore and Kerrick 1977, Bowman etal. 1994). The observed mineral assemblages in aureoles like these require down-T flow in the inner aureole and possibly up-T flow in (at least parts of) the rest of the aureole (Ferry 1994, Dipple and Ferry 1996). Interestingly the observed 8180 isotope 2 signature sometimes allows for an interpretation of down-T flow in all of the aureole (Bowman et al. 1994). In contrast, mineral assemblages in the siliceous dolomites of the Elkhorn aureole are dominantly isobarically divariant assemblages throughout the aureole except for the periclase zone (<20 m). Here, generally isobarically univariant periclase assemblages are developed (Bowman and Essene 1982). Fluid flow is interpreted to have been vertical in the periclase zone and down-T in the rest of the aureole (Ferry 1994). This interpretation of vertical fluid flow in the inner aureole is in accordance with observations in massive garnet-pyroxen skarns that formed in reentrants of the pluton-wallrock contact. These were also interpreted to have experienced vertical fluid flow (Bowman et al. 1985). The special character of the zone closest to the contact is underlined by the occasional observation of signs of (silica-) metasomatism in this area. For example, forsterite in the inner part of the Adamallo aureole was formed in significantly higher abundance than metamorphic silicates in the outer aureole. It was suggested that the forsterite formed by reaction of dolomite with Si02 dissolved in fluid emanating from the igneous body (Bucher-Nurminen 1994). In most of these aureoles, independent of flow direction, the fluid flow is interpreted to have been heterogeneous. The scatter in isotope values and in the spatial distribution of mineral assemblages commonly cannot be explained if total influx was the same in all areas (Ferry 1994). However, these heterogeneities can be reproduced by models of fluid flow in rocks with heterogeneous permeability (Gerdes et al, 1995; Marchildon and Dipple, 1998). In summary it can be said that hydrologic models for fluid flow in contact aureoles are diverse and potentially very complex. The complexity of the issue of fluid flow in carbonate rocks is underlined by the problems of understanding infiltration mechanism. Massive, pure carbonate rocks are generally not susceptible to pervasive grain-boundary flow of H20-C02 fluids under normal metamorphic pressure-temperature conditions. Experiments with monomineralic carbonates (under static 3 conditions) showed that pervasive fluid flow along grain boundaries is essentially impossible in dolomites and restricted to low pressure (<1.5 kilobars (kbar)) and fluids with intermediate H20-C02 compositions or strong brines in calcite; the dihedral angle theta is always larger than the critical angle of 60° (Holness and Graham 1995, Hay and Evans 1988). Experiments with monomineralic quartz show similar results. The dihedral angle for grains in equilibrium with a H20 fluid is only smaller than 60° at low temperatures (<350 °C) or at high pressures (>8 kbar) (Holness 1993). In the presence of a C02 fluid theta is always larger than 60°. The experimental results are generally confirmed by field observations. Massive, pure limestone in metamorphic terrains show no or only small isotopic alteration even if they are adjacent to lithologies that experienced large influxes and show significant isotope depletion (e.g., Nabalek et al. 1984, Schliestedt and Mattews 1987, Jamtveit et al. 1992). Quartz or chert layers are occasionally interpreted to be permeable (e.g., Heinrich 1993) but infiltration might have been channeled mostly along the lithologic boundaries (Jamtveit et al. 1992, Walther 1996). Consequently, in rocks similar to the monomineralic samples used in experiments, fluid flow under metamorphic conditions is interpreted to occur only or predominantly along fractures. It is not clear however, what percentage of for example carbonate minerals in a rock is sufficient to give the rock the overall permeability behavior of a pure (dolo)marble. Some authors suggest that minerals that make up only about 30 - 40 percent of a rock determine its overall (grain boundary) permeability behavior (Holness and Graham 1991, and citations therein). These values appear to be low. With up to 70 volume percent other material in the rock the possibility of the formation of a interconnected network of fluid 'channels' along grain boundaries seems possible. In nature, most rocks are not monomineralic but exhibit at least low abundances of other minerals, and many natural (impure) carbonate rocks in metamorphic terrains show signs of fluid infiltration. For example all the above mentioned contact-metamorphic siliceous dolomites have quartz contents of about 5-15 volume percent (vol%) in the protolithic rocks and show (with 4 increasing metamorphic grade) mineral reactions and isotopic depletion that generally require low to moderate amounts of fluid infiltration (Dipple and Ferry 1996). Hence, some processes must enhance permeability of these rocks. Etheridge et al. (1983, p.219) noted that the "main controls on permeability are rock type and structure. The lithological influence will result largely from differences in microstructural behavious and rate of fluid production, rather than of primary porosity." There are at least three very potent mechanisms, which are coupled to some degrees, that can cause permeability enhancement in massive rocks. These are (a) reaction-enhanced permeability, (b) hydrofracturing, and (c) deformation. Reaction often enhances permeability because the volume of the produced minerals is (usually) smaller than the volume of the reactants. This leads to an increase in porosity which, if it (at least partially) can be kept open against the pressure of compaction, leads to an increase in permeability (Etheridge et al. 1983, Ortoleva et al. 1987). Additionally if the reactions are devolatilisation reactions, large volumes of volatiles are liberated. This either is supportive of keeping the increased porosity open (against the pressure of compaction), or, if volume reduction of the solids is small, devolatilisation would lead to hydrofracturing. Fluid pressure increases up to or above lithostatic pressure (or, the least compressive stress rj3) at which point the pores crepitate and the fluids are expulsed into the surrounding area by hydro-fracturing (e.g., Fyfe et al. 1978). This in turn can lead to further reaction in the infiltrated region, if fluid infiltration is able to speed up the rate limiting process of reaction. Rocks in which no (devolatilisation) reaction occurs would experience only minor hydro-fracturing. With a sufficient increase in temperature the very small amounts of pore fluids can reach fluid pressures larger than the lithostatic pressure of the rocks. If these two rock types are interlayered, for example in form of "permeable" pelites and "impermeable" carbonates, the carbonate could be affected by the enhanced infiltration in the more permeable pelites. This can lead to increased advective-diffusive infiltration into the 5 carbonate near the lithologic contact (e.g., Bickle and Baker 1990, Jamtveit et al. 1992, Cartwright and Weaver 1997). Deformation is another process interpreted to be effective in enhancing permeability by crystal plastic mechanisms or some other mass transport mechanism (Etheridge et al. 1983, Holness 1997). In recent years studies have increasingly focused on the complexity in the interplay between heating, structural influences, fluid flow and reaction and their effects on the permeability structure (e.g., Etheridge et al. 1983, Nabalekefa/. 1992, Dutrow and Norton 1995, Holness 1997, and in a review by Ferry 1994). In this study I describe the mineralogy, mineral chemistry and isotope signature of siliceous dolomites in the Horsethief Creek contact-metamorphic aureole in south-eastern British Columbia, Canada. The siliceous dolomites show a normal prograde sequence of isograds but some features are of special interest and attest to the complexity of the contact-metamorphic process. Throughout the mid to inner aureole (<1250 m from contact) almost all assemblages are isobarically divariant and most rocks are completely de-dolomized. Metamorphic minerals occur in unusually high abundance and pervasive distribution, and index minerals are heterogeneously distributed. An abrupt change from predominantly isobarically divariant assemblages in the inner and middle aureole to isobarically univariant assemblages in the outermost aureole is observed in the Horsethief Creek siliceous dolomites. The peak temperature profile and the isotope signature of the aureole also are unusual. The development of the mineral assemblages in the Horsethief Creek aureole must have been influenced by varying bulk composition, temperature, and spatially and temporally heterogeneous fluid infiltration. An attempt is made to outline a model that satisfactorily explains the distribution of mineral assemblages and most of the features special to this aureole. 6 1.3 The T-XC02 diagram Mineral assemblages in contact-metamorphic siliceous dolomites are commonly used to interpret fundamental aspects of metamorphic fluid flow. Carbonate rocks aire particularly useful for this purpose because they are chemically and mineralogically simple and sensitive to the amount and composition of an infiltrating fluid. This leads to the development of only a small number of characteristic mineral assemblages in the temperature range typical for metamorphism. The developed mineral assemblages commonly are the result of infiltration-triggered reaction. Two reaction paths in a T-XC02 diagram are used to illustrate the interaction of siliceous carbonitic rocks and a H20-C02 fluid. These diagrams are commonly used to depict mineral assemblages and reaction paths for impure carbonate rocks. Fluid pressure (P) is assumed to be constant. Hence, they should be applied with caution in regional metamorphic rocks but are generally considered to be applicable on the small scale of a contact aureole. Figure 1 shows a simple T-XC02 diagram at P=lkbar. The XC02is the XCD2 of a fluid in contact with the rock. If the protolithic dolomite - quartz rock is heated without infiltration, the reaction path will follow reaction curves (path 1). This results in the formation of isobarically univariant assemblages where usually only small amounts of products are found besides the reactants. Little reaction progress of the devolatilisation reactions is sufficient to change the fluid composition of the small amount of fluid present in the pore space. As soon as rock and fluid are in equilibrium reaction ceases for a given temperature. This kind of reaction is referred to as an internally (by the rock) buffered reaction (Greenwood 1975). If on the other hand the rocks are infiltrated by an infinite amount of H20-fluid during heating the C02 produced during reaction is not sufficient to change the fluid-XC02 and the reactions will continue at the XC02 of the infiltrating fluid until at least one of the reactants is completely exhausted (path 2). This is followed by a temperature increase at constant XC02. The resulting assemblages are referred to as isobarically divariant assemblages as they are represented 7 Figure 1: T-XfjQ2 diagram illustrating the influence of H20-C02 fluids on the development of mineral assemblages in siliceous dolomites at Pfluid = const = lkb. Path 1 (red dashed line) shows the reaction path for a rock that is not infiltrated during heating. Path 2 (blue solid line) shows the reaction path for a rock that is infiltrated by an infinite amount of a nearly pure H20 fluid. Mineral abbreviation according to Kretz (1983). 8 by a stability field rather than a reaction line in the T-XC02 diagram. The XC02 in this case is governed by the externally derived fluid. 9 CHAPTER 2 GEOLOGIC SETTING 2.1 Regional Geology The quartz monzonitic Horsethief Creek pluton is located in the Purcell Mountains of southeastern British Columbia, Canada (Figure 2). The Purcell Mountains are part of the 'Purcell Anticlinorum' which is dominated by a thick sequence of Proterozoic to upper Paleozoic sedimentary rocks that were deposited on or near the western edge of the North American craton (Cook and Van der Velden 1995). During the Mesoproterozoic, the area was part of an intracratonic basin into which predominately shallow water sediments were deposited, leading to the formation of the Belt Purcell Supergroup. This sequence is unconformably overlain by the Neoproterozoic Windermere Supergroup. Reesor (1973) attributed the lower Windermere rocks to reflect a rifted margin setting and the upper part to comprise continental margin strata. Warren and Price (1992, p. 15) noted that in more detail the "stratigraphic relationships within and between Upper Proterozoic and Lower Cambrian rocks show that this interval was punctuated by several tectonic events, related to extension and/or rifting of the North American continental margin.". This led "to the emergence of 'the Windermere High' (Reesor, 1973) as a high-standing continental block." During the mid-Mesozoic the area experienced extensive tectonic activity when the allochthonous Intermontane Superterrane was accreted to the western side of the North American craton. This caused widespread deformation as recorded in N-NW trending folds and faults, regional scale metamorphism and the intrusion of late- to synkinematic mid-Jurassic plutons 10 Figure 2: Simplified map of the regional geology (after Reesor 1973) 1 ! (Figure 2) in the suture zone that forms today's Omineca Belt (Archibald et al. 1983). A second episode of plutonism occurred during the mid-Cretaceous leading to the intrusion of several post-kinematic and hence discordant igneous bodies (Reesor, 1973). This plutonism is interpreted to be the result of crustal thickening (Archibald et al, 1984). Geochemical characterization of various of the mid-Cretaceous plutons confirms this hypothesis. Brandon and Lambert (1993) invoke crustal anatexis of Precambrian basement gneisses and/or Proterozoic metapelites to produce these later intrusions. The Horsethief Creek Batholith (HCB) belongs to this later group of discordant mid-Cretaceous intrusions. K-Ar dating (biotite) revealed one age of 108 Ma (Reesor 1973) and Rb-Sr dating an age of 109 Ma (Brandon andLambert 1993) for the HCB. After slow cooling from mid- to late-Cretaceous time the area was thrusted eastwards moving "over a step-like ramp in the basement corresponding to the ancient rifted margin of the continent" (Archibald et al. 1984, p.578). This convergence during Late Cretaceous to Paleocene time is attributed to the accretion of the Insular Belt to the west. Seismic data confirms that the Anticlinorum formed when up to 15 km Belt-Purcell to Paleozoic margin sediments were carried eastward over a prominent basement ramp by means of imbricate thrust faults (Cook and Van der Velden 1995). During the Eocene the area was affected by crustal extension and tectonic denudation (Archibald et al. 1984). 2.2 Local Geology This study focuses on the contact metamorphism of siliceous-potassic dolomites from the Mount Nelson Formation of the Belt-Purcell Supergroup along the northern margin of the Horsethief Creek pluton (Figure 3). 12 Figure 3: Geological map of the northern flank of the Horsethief Batholith showing main units, topography (altitudes in meters), sample locations and isograds (for sample location numbers see Figure 5). Isograds as defined within the siliceous dolomite system (CaO-MgO-SiO^-^O-CO^) considering index minerals that can be observed in the field. 13 The intrusion is =15 km in diameter and is one of several discordant mid-Cretaceous igneous bodies that intruded the Proterozoic to early Paleozoic clastic and carbonate sediments of the Belt-Purcell and Windermere Supergroups in mid-Cretaceous time (Archibald et al. 1984). Close to the intrusion, within a 1 - 2 km wide contact aureole, low-grade regionally metamorphosed rocks are overprinted with high-grade mineral assemblages. The regional metamorphism is coeval with Jurassic deformation and is of sub-chlorite grade in the study area (Reesor 1973). The development of systems of steeply dipping joints in all lithologies of the working area is believed to be associated with the folding (Twiss and Moores 1992). The Mount Nelson Formation in the working area consists of a sequence of cliff-forming (siliceous and potassic) dolomitic marbles underlain by calc-silicates and pelites (Figure 3) and with numerous chert interlayers of various thickness (centimeter - decimeters). The thickness of the underlying pelitic layer can only be estimated as >300 m because the contacts are outside the mapped area. Pelitic interlayers in the siliceous dolomites are rare and very thin (cm). The whole sequence dips =35° ENE (Figure 3), and is part of the west flank of a NNW-trending anticline (Reesor 1973). Most of the inner two-thirds of the study-area contains good exposures along a west-facing cliff that radiates northward from the batholith (Plate 1). Exposure deteriorates considerably at the perimeter of the aureole. The irregular map pattern of parts of the inner aureole is due to a combination of irregular topography within a NW draining cirque (Figure 3) and block faulting resulting in poor exposure. The rocks closest to the contact are exposed along an approximately E-W trending ridge that parallels the pluton-wallrock contact. Siliceous and potassic dolomite dominates a coherent stratigraphic package =160 m in thickness that can be mapped from the inner to outer aureole. Individual beds within this package range in thickness from =1 - 50 cm but cannot be traced with confidence through the aureole because of mm to m offset on E-W trending high angle faults. These faults appear be post-contact-metamorphic as no (infiltration-driven) development of metamorphic minerals along 14 15 them is observed. Thin (cm) chert interlayers are common, and in some localities dolomite is characterized by dark, discontinuous elongated chert lenses or nodules that are arranged parallel to bedding. Sets of closely (cm - dm) spaced joints that are roughly perpendicular to the bedding are found in the low-grade rocks of all lithologies in the working area and beyond the extent of the aureole further north. At higher grades, in rocks where abundant metamorphic minerals were produced, the rocks are generally massive. In the siliceous dolomites veins roughly perpendicular to bedding are common (Plate 2). They are thought to be related to the pre-contact metamorphic joint system. The length of the veins is in most cases determined by the width of the sedimentary layers in which they are contained. The width of the veins is usually a few millimeters (mm) and they are spaced on a cm - dm scale. This study focuses on iron-poor, silicic-potassic dolomites which form the bulk of the mappable unit. Scapolite-bearing or iron-rich dolomitic marble is present in =5 % of the samples examined. These samples are not considered further in this study. Sample localities are shown in Figure 3. 16 Plate 2: Photograph of an outcrop near sample location 19/3 (looking South). The rocks show typical veins trending vertical and parallel to bedding. The amount and form of veining differs between the individual layers; sledgehammer for scale. Parts of the inner aureole can be seen in the background. 17 CHAPTER 3 MINERALOGY AND MINERAL CHEMISTRY 3.1 Introduction Mineral assemblages were examined in thin section by optical microscopy and back-scattered electron imaging. Carbonate minerals were identified by staining with the method of Dickson as described in Hutchison (1974). Very fine grained carbonate minerals were identified with scanning electron microscopy (SEM), using a Philips XL30 with back-scattered electron (BSE) imaging. All dolomite occurrences were confirmed with SEM. A T-XC02 phase diagram for potassic siliceous dolomites was constructed using endmember compositions of the minerals: quartz (Qtz), dolomite (Doi), calcite (Cal), talc (Tic), tremolite (Tr), diopside (Di), forsterite (Fo), brucite (Br), periclase (Per), wollastonite (Wo), K-feldspar (Ksp), and phlogopite (Phi) (Figure 4). The PTAX program by Perkins et al. (1980) which is based on the thermodynamic data set by Berman (1988) and H20-C02 mixing properties of Kerrick and Jacobs (1984) was used for the calculation and plotting of the mineral reactions. This use of endmember compositions is justified by the general occurrence of colorless phlogopite which is indicative for very iron poor conditions. The iron poor character of the rocks was qualitatively confirmed by SEM and quantified by microprobe analyses (see paragraph 3.3 below). Fluid pressure was assumed to be =2.2 kbar in accordance with PT estimates derived from mineral assemblages observed in metapelites. (Floriet 1996). Peak metamorphic conditions in the inner aureole are estimated to be =2.1 - 2.5 kbar and =610 °C based on the observation of the breakdown of muscovite + quartz to andalusite + K-feldspar and the occurrence of fibrolite in a 18 Figure 4: T-Xco2diagram depicting phase equilibria in the potassic sili ceous dolomite system [KCMAS-HC (K20-CaO-MgO-Al203-SiC>2-H20-CO2)] at P = 2.2kb. Calculated for endmembers with the PTAX program (Perkins et al. 1980). Fluorine-substitution was not considered. It would cause some changes in the low-temperature area. The thermal stability of (OH-F) minerals increases at the expense of anhydrous breakdown products with increasing F-content (e.g. Abercombie et al. 1987). Reactions are num bered according to their order of appearance in the text. Mineral abbrevia tions after Kretz (1983). 19 few of the highest grade samples, with no partial melting of the metapelites. In the siliceous dolomites, six isograds are identified based on the first appearance of index minerals: phlogopite; tremolite + K-feldspar; diopside; diopside + phlogopite; diopside+dolomite and forsterite (Figure 5). The identified assemblages and isograds are listed in order of increasing proximity to the pluton in Table 1. Mineralogical variations on the sub-thin section scale are frequently found. Samples with two distinct layers or large lenses with assemblages different from those in the matrix are indicated by the suffices 'A' and 'B' to the respective sample number in Table 1. 3.2 The Six Metamorphic Zones Following is a description of the 6 metamorphic zones in prograde sequence. Typical mineral assemblages and textures are described and possible reactions listed. The assemblages are listed in Table 1. Almost all rocks outside the phlogopite zone show static recrystallisation and mineral formation. Few samples showed a preferred mineral orientation. 3.2.1 The Phlogopite Zone This zone cannot be identified in the field because phlogopite is found only in very low abundance - typically 1-3 volume percent (vol%). Furthermore, the lack of outcrop in the outermost aureole prohibits the collection of unmetamorphosed rocks and therefore, although there is a phlogopite zone, the phlogopite isograd cannot be traced. The outermost outcrop lies 1515 m from the pluton and contains phlogopite. The phlogopite isograd must occur at a distance equal or greater than 1515 m away from the pluton. The phlogopite zone rocks are composed of 20 Figure 5: Map showing all isograds and assemblages as defined by the KCMAS-HC system. Note the very heterogeneous distribution of assemblages within the metamorphic zones. The exact outcrop location is at the symbol nearest to the outcrop number. 21 Tablel: Peak metamorphic assemblages in increasing proximity to pluton. Note: X=present; (X)=traces; C=isolated core, XC=cores with conatct to matrix minerals; n/d=not detected (only used in phlogopite zone rocks, where calcite is expected if phlogopite is of metamorphic origin). Doi Qtz Tr Cal Di Fo KSD Phi Doi Otz Tr Cal Di Fo KSD Phi 18/2 - 1 X X n/d X CX) 27/3 - 1 X X X 18/1 - 1 X (X) (X) X (X) 27/3 - 2 X X X 18/1 - 2 X X n/d X (X) 3/4- 1 X X X X 13/2 - 1 X X X 3/3 - 1 X X X X 13/2 - 2 X X X 9/1 - 2 X X X 13/2 - 3 X X (X) X CX) 23/3 - 1 X X X X 13/2 - 4 X X (X) X CX) 23/3 - 2 X X X X 28/1 - 1 X X (X) X (X) 3/1 - 1 X X X 28/1 - 2 X X CX) X (X) 8/3 - 1 X X Di + Dnl X X 28/1 - 3 X X CX) X CX) 8/3 - 2 X X X XC X 28/1 - 4 X X n/d x cx) 8/3 - 3A X X X X 28/1 - 5 X (X) X Tr + Ksn X CX) 8/3 - 3B X X X X 15/1 - 3 X X X 8/4 - 1A X X 15/1 -4 X X X X 8/4 - IB X X CX) 15/1 - 5 X X X 8/2 - 1 X X Fn 15/1 - 6 X X X X 8/1 - 1 X X X X 15/1 - 7 X X X 25/2 - 1 X X X 12/3 - 1 X X X 25/2 - 2 X X X X 12/3 - 2 X X X 25/2 - 3A X X X X 21/3 - 1 X X X 25/2 - 3B X X X 21/3 - 2 X X Di X X 25/2 - 4 X X X X 12/4 - 1 X X CX) X X 25/1 - 1 X X X 12/4 - 2 X X CX) 25/1 - 2 X X X X 12/4 - 3 X X c CX) 6/3 - 1 X X X 21/2 - 1 X X X 6/3 - 2 CX) X X 21/2 - 2 X X c X 6/2- 1 X X X 21/1 - 1 X X X Di + Phi X 4/1 - 1A X X X 21/1 - 3 X X XC X 4/1 - IB X X X X 12/2 - 1 X X X 4/1 - 2 X X 12/1 - 1 X X X 23/1 - 1 X X X X 12/1 - 3 X X XC X 23/1 - 2 X X X 12/1 - 4 X X X (X) 23/1 - 4 X X X X 12/1 - 5 X X X 24/1 - 1 X X X X 22/1 - 1 X X X X 24/1 - 2 X X X 22/1 - 2 X X X X 24/1 - 3 X X X X 22/1 - 3A X X X 24/1 - 4 X X X 22/1 - 3B X X X X 7/1 - 1A X X X 22/2 - 1 X X cx) 7/1 - IB X X X 22/2 - 2 X CX) X CX) 24/3 - 1 X X X X 22/2 - 3 X X X X 24/3 - 2 X X X X X 22/3 - 1 X X X 24/3 - 3 X X 22/3 - 2A X X X 4/3 - 2 X X X X 22/3 - 2B X X X 23/2 - 1 X X X X 22/3 - 3 X X X X 23/2 - 2 X X X X 22/3 - 4 X X X 23/2 - 3 X X X X 17/1 - 1 X X X X 24/2 - 1 X X X X 22/4 - 1 X X X 24/2 - 2 X X X X 22/4 - 2 X X X X 2/5 - 1 X X 11/1-1 X X X 2/5-2 X X X X 10/2 - 1 X X 2/5 - 3 X X X X 10/2 - 2 X X 2/5 - 4 X X X 10/2 - 3 X X 2/5 - 5 X X X X 10/2 - 4 X X CX) 2/5 - 6 X X X 10/2 - 5 X X X 2/4 - 1A X X X 19/3 - 1 X X X 2/4 - IB X X X 19/3 - 2A X X X 2/4 - 2 X X X 19/3 - 2B X X X X 2/4 - 3 X X X X 14/2- 1 X X X 2/2 - 1 X X 22 dolomite - quartz - K-feldspar ±calcite and varying but generally small amounts of phlogopite. In very fine grained, homogeneous rocks containing only traces (<l-3 vol%) of randomly orientated phlogopite are found (Plate 3). In small discrete zones (mm - cm scale) phlogopite is present in greater abundance (as high as =8 vol%) and with preferred mineral orientation (Plate 4 a and b). In one thin section phlogopite is found in the center of a quartz vein (Plate 5). If the phlogopite is of contact-metamorphic origin, it most likely formed by the reaction: 3 dolomite + K-feldspar + H20 = phlogopite + 3 calcite + 3 C02 (1) at 0.00 > XCQ2 < =0.86 (cf. Figure 4). The fact that there are usually only small amounts of calcite found in the matrix and that the amount of calcite increases proportionally with the amount of phlogopite supports the assumption that the phlogopite formed by reaction (1). Calcite may be present but not identified in these samples because in BSE imaging the calcite shows a gray value very close to that of K-feldspar. It is easily possible that calcite which is expected in the presence of phlogopite was simply not detected. Based on these observations the protolith is presumed to be dolomite - quartz - K-feldspar and ± calcite. Chert lenses or nodules (Plate 6) are characteristic of rocks in the phlogopite zone. They are sometimes interconnected by bedding-parallel chert layers and/or sub-vertical quartz veins. Because they are more resistant to weathering they stand out of the more strongly weathered quartz-poor dolostones. The quartz content varies from just a few percent in a fine grained mainly dolomitic matrix up to =50 vol% in layers containing chert lenses. The potassic feldspar content ranges from trace up to =15 vol%. Calcite only occurs in trace amounts in the matrix but makes up for 10 - 90 vol% of the material in occasional small veins. The effect of iron on mineral distribution is not considered in this study because most samples are Fe-poor. However, above sample locations 28/1 and 13/2 (in the Mg-phlogopite zone) is an outcrop that contains two adjacent layers (both =40 - 50 cm wide) with Fe-amphibole in high abundance (=60 vol%). 23 Plate 3: Photomicrograph of a mineral assemblage in the phlogopite zone in cross-polarized light (XPL) (sample 18/1-1). Field of view -0.6 mm wide. The assemblage shown is typical for this zone in that it shows very fine grained dolomite (very high interference colors (IC)) - K-feldspar (first order gray)) - and 3 flaky phlogopite crystals (first order gray-to second order purple interference colors (IC)). These crystals are larger than average and where chosen to ease identification. 24 Plate 4: a) Plain-polarized light (PPL) and b) XPL photomicrograph respectively showing the higher abundance of orientated phlogopites (flakes) in deformed rocks of the phlogopite zone (sample 13/1-1). The cause of the small scale, local deformation is not clear. Note the growth of phlogopite grains perpendicular to the foliation in the pressure shadows of a pyrite grain. Field of view is =0.3 mm wide. 25 Plate 5: XPL photomicrograph of a phlogopite zone rock (sample 13/1-1) that shows a vein unusually rich in phlogopite. As K-feldspar veins are not observed the for reaction necessary potassium and aluminum was likely introduced by the fluid. Field of view =0.3 mm wide. 26 Plate 6: Photograph of a typical outcrop in the phlogopite zone (near 13/1) showing bedding parallel chert lenses or nodules and subvertical joints and veins cutting through single layers; hammer for scale. 27 3.2.2 The Tremolite+K-feldspar Zone The tremolite + K-feldspar isograd is located at =1250 m from the pluton - wallrock contact. The first occurrence of tremolite is marked by cm - dm wide layers with 1 - 3 cm long tremolite crystals growing mostly in rosettes and by layers with very fine grained (=1 mm) pervasive tremolite. The latter layers tend to be more abundant and wider. The number and thickness of tremolite-bearing layers increases towards the diopside isograd. The large tremolite crystals are mostly found in layers that contain abundant veins approximately perpendicular or parallel to bedding. In a few samples tremolite growth can be observed to have started in points where veins intersect. A widespread assemblage is tremolite - calcite - K-feldspar, showing homogeneously sized and distributed minerals (Plate 7). The tremolite content can reach up to =65 vol%. Another common assemblage is tremolite - calcite - phlogopite. Both of these assemblages persist to the forsterite isograd. It is not clear which reactions are responsible for the formation of the tremolite in the various assemblages. In potassium free rocks tremolite was likely formed by the reaction 5 dolomite + 8 quartz + H20 = tremolite + 3 calcite + 7 C02 (2) if fluid infiltration occurred during heating. In potassic rocks the situation is more complex. Combinations of reaction (1), (2) and the reaction 6 calcite + 5 phlogopite + 24 quartz = 3 tremolite + 5 K-feldspar + 2 H20 + 6 C02 (3) are possible depending on bulk composition and timing of heat- and fluid flow. This is discussed in more detail in chapter 4. At around 1000 m (±200 m) some samples were found to contain albite grains and Ba-bearing K-feldspars (cf. Appendix B, Figure Bl). 28 Plate 7: XPL photomicrograph of one of the most common assemblages in the Horsethief Creek aureole: tremolite (needles; second order IC) - calcite (very high order IC) - K-feldspar (first order gray IC). Note the even distribution of minerals throughout the rock. Field of view =1.2 mm wide( sample 12/3-2). 29 3.2.3 Diopside zone The diopside isograd is placed =1125 m from the pluton. Only one sample in this narrow zone shows diopside in an equilibrium peak metamorphic assemblage. Samples 12/4-3 and 21/2-2 also contain diopside but the diopside was found only as cores within tremolite with no observed contact with the matrix minerals. In these two samples the diopside is listed with a "C" in the tabulation of peak metamorphic assemblages in Table 1. 3.2.4 Diopside+Phlogopite zone The first appearance of a diopside + phlogopite bearing peak metamorphic equilibrium assemblage is in sample 12/4-1, approximately 1010 m from the pluton - wallrock contact. This sample and 12/1-3 (the next further sample up-grade) show diopside cores within tremolite but these diopsides are observed to be in contact with the matrix minerals. The samples do not show diopside in the fine grained matrix. It is conspicuous that almost all of the diopside bearing samples in the middle to outer aureole show diopside reacting to tremolite. Further up-grade this reaction texture between diopside and tremolite is less commonly observed. It does occur in rocks up-grade to near the forsterite isograd where the unambiguously isobarically univariant assemblage diopside - dolomite - tremolite - calcite (-phlogopite) in sample 8/3-2 shows this texture (Plates 8 a and b). The most common diopside bearing assemblage observed is diopside -tremolite - calcite - phlogopite with diopside often occurring in roundish, 0.3 - 0.7 mm large grains while tremolite forms 1-2 mm long needles. The rocks commonly have a uniform fine grained texture. Original chert nodules are often found completely replaced by diopside and/or (?) tremolite and calcite, and occasionally diopside rosettes are found to grow within chert layers. Tremolite often grows in or nucleates at veins and is also found to generally be coarser directly 30 Plate 8: a) XPL photomicrograph showing a diopside core within a tremolite porphyroblast (sample 8/3-2). The diopside has brownish IC the tremolite first order IC or is extinct, b) Same as a) but in PPL showing the relief contrast between the diopside cores (high) and the tremolite (low). The orange spots are iron oxide staining around pyrites, the pink hue is a result of carbonate staining. Field of view = 1.2 mm wide. 31 adjacent to chert layers. One sample (22/2-1) in this zone was found to have talc cores within tremolite, and another sample (21/1-1) contained a calcite - tremolite - talc - calcite (rim to core) vein. 3.2.5. Diopside+ Dolomite zone Sample 8/3-2 marks the beginning of the diopside + dolomite zone at about 360 m from the contact. This is the only sample recording diopside + dolomite in an isobarically univariant equilibrium assemblage. Sample 8/3-3 also contains diopside and dolomite but they are not found in mutual contact; they are in two different areas of the thin section with different bulk compositions. Close to sample location 8/4 a ~3 cm wide bleached zone was observed below a chert layer. The bleached layer curves (keeping a constant width) around a diopside spray nucleating on the chert - marble boundary (Plate 9). Diopside forming reactions cannot be determined from the observed textures. A possible exception is the case where diopside is found as cores to tremolite. This texture, observed in the isobarically univariant assemblages diopside - dolomite - tremolite - calcite (sample 8/3-2), suggests that the tremolite formed by reaction 4 diopside + dolomite + H20 + C02 = tremolite + 3 calcite (4). The first developed diopside is concluded to have formed by reaction dolomite + 2 quartz = diopside + 2 C02 (5) at XC02 > 0.95. In this case diopside would have formed without any significant amount of H20-fluid infiltrating the system up to that point. 32 Plate 9: Photograph from an rock outcrop close to sample location 8/4. It shows a bleached zone in a carbonate layer underneath a chert layer. Note how the bleached zone follows the diopside spray that nucleates at the chert - carbonate contact. Camera lens cover for scale. 33 3.2.6 The Forsterite Zone The forsterite isograd is placed =270 m from the pluton. In hand sample, forsterite generally occurs as mm large, round, green crystals. Very few samples contain bladed forsterite next to also bladed diopside (e.g., sample 24/2-1). The most common mineral assemblage is forsterite -diopside - calcite ± phlogopite. The relative abundance of forsterite and diopside varies considerably with both extremes of diopside-absent and forsterite-absent assemblages observed. These variations are observed on scales down to thin section size. Veins, or what seem to be more strongly recrystallized areas in rocks from this zone commonly have a higher abundance of diopside than the matrix assemblage - with only up to 30 vol% diopside in the matrix but up to 70 - 80 vol% in the veins (Plate 10 a and b). In a few thin sections (6/2-1, 2/5-5, all samples from outcrop 23/2) simultaneous growth of forsterite and diopside can be inferred from 'mutual' inclusions and small inclusions of tremolite in both kinds of porphyroblasts, suggesting that the reaction 3 tremolite + 5 calcite -> 11 diopside + 2 forsterite + 3C02 + 3 H20 (6) took place at fluid XC02 < =0.54. In most cases no specific reaction is recorded in the rocks. The sample closest to the pluton that contains dolomite is 2/4-1 at a distance of 20 m. No periclase or brucite was detected in this thin section. With assumed maximum temperatures of 600 - 640 °C for that sample the XC02 must have been larger =0.04 to hinder the formation of periclase or brucite (cf. Figure 4). Although there are some assemblages which are quite common in each of the zones as described above it is conspicuous that index minerals as well as the two potassic phases are not uniformly developed within the metamorphic zones and that the geometry of isograds is poorly constrained. The heterogeneity in observed mineral assemblages is evident in the tabulation of peak metamorphic mineral assemblages (Table 1) and also recognizable in Figure 5. 34 mammas • . jmwmmmmm* •>* > Plate 10: a) PPL and b) XPL photomicrograph respectively showing a cross-section of a vein in a forsterite zone rock. Note that the thin section is too thick. Forsterite in PPL with high relief and typical fracturing - diopside with relatively low relief (PPL) and here 1st to 3rd order IC (XPL). Note that the vein is virtually forsterite absent, and how diopside abundance and grain size decrease with increasing distance from the vein. Field of view ~5 mm wide (sample 25/1-2). 35 3.3. Mineral Chemistry Quantitative analyses of mineral compositions were obtained by Wavelength Dispersive Spectrometry (WDS) with a CAMECA SX-50 microprobe. Synthetic olivine and phlogopite and natural diopside and carbonate standards were used. Operation conditions for the silicates were excitation voltage, 15kV; beam current, 20nA; beam size 5pm, and for the carbonates 15 kV, lOnA and lOuxn. Data reduction was done with the "PAP" (j)(pZ) method (Pouchou and Pichoir 1985). Representative results are listed in Table 2 the complete data set is listed in Appendix A. In all minerals the Fe/(Fe+Mg) is very low. Other elements like for example Mn, Ti or Ni are only found in concentrations near or below the detection limit. This justifies the use of Mg-endmember thermodynamic data in the construction of the T-XC02 diagram. The degree of substitution of OH by F or Cl in phlogopites and tremolites varies. The r chlorine content is below or at the detection limit in the Mg-endmember minerals. The fluorine content is higher and ranges from 0.079 - 0.233 (cations per 22 oxygens) in phlogopite and from 0.022 - 0.072 (cations per 23 oxygens) in tremolite. The interaction of F - Cl-bearing fluids and Mg - Fe solid solution minerals is quite complex but the observed trends seem to be in general agreement with the "Mg - Cl avoidance" rule (e.g. Zhu and Sverjensky, 1992). The distribution of fluorine contents in phlogopite (the only hydrous mineral found in all metamorphic zones) is plotted against distance in Figure 6. Throughout the aureole F-contents are relatively homogeneous with only two exceptions: The sample closest to the contact shows a F-content lower than average and the sample =1250 m from the pluton in the upper phlogopite zone shows higher and more scattered F-contents. A description of the rocks with orientated phlogopite grains in the phlogopite zone is given in section 3.2.1. The significance of the data in Figure 6 is not clear. Moore and Kerrick (1976) interpreted the F-depletion near the contact to be the result of influx of igneous fluid while Rice (1977b) and Holness (1997) argue that high F-values in 36 Table2: Microprobe analyses of representative silicate and carbonate minerals a b) Tremolites' a) Phlogopites" Label 19/3-2 22/2-1 12/1-4 Label 2/4-1 25/1 •1 19/3-2 22/1-3 13/1-1 13/1-3 28/1-1 28/2-1 from zone: Di+Phl Di+Phl Di+Phl from zone: Fo Fo Di+Phl Di+Phl Phi Phi Phi Phi K 0.021 0.009 0.024 K 1.788 1.866 1.878 1.769 1.764 1.506 1.823 1.846 Na 0.043 0.017 0.072 Na 0.141 0.055 0.020 0.028 0.027 0.025 0.027 0.024 Mg 4.860 4.926 4.834 Mg 5.271 5.729 5.712 5.653 5.040 5.427 5.106 5.138 Ca 1.946 1.989 1.964 Fe 0.175 0.076 0.141 0.082 0.110 0.183 0.133 ' 0.072 Fe 0.090 0.072 0.116 Oa 0.010 0.022 0.012 0.031 0.017 0.641 0.040 0.029 Mn 0.004 0.005 0.002 Mn 0.006 0.004 0.001 0.000 0.006 0.001 0.001 0.000 Ti 0.006 0.002 0.004 Ti 0.077 0.051 0.022 0.042 0.160 0.073 0.096 0.063 Si 5.462 5.912 5.941 6.129 5.776 5.841 5.747 5.741 Si 7.891 7.948 7.798 Al 2.963 2.054 2.054 1.680 2.436 1.796 2.523 2.715 Al 0.160 0.037 0.270 F 0.079 0.359 0.294 0.335 0.771 0.844 0.553 0.233 F 0.044 0.072 0.047 Cl 0.013 0.042 0.062 0.012 0.033 0.076 0.083 0.054 Cl 0.024 0.008 0.007 Oxide Sum 95.44 95.63 96.09 95.61 96.57 94.39 95.42 95.39 Oxide sum 97.77 97.66 98.033 Fe/(Fe+Mq) 0.032 0.013 0.024 0.014 0.021 0.033 0.025 0.014 Fe/(Fe+Mq) 0.018 0.014 0.023 c) Diopsidesd sample 25/2-3 8/1-1 a All mineral formulas assume Fe as Fe2+, from zone: Fo Fo Oxide sum refers to the sum of metal oxide wt%. d) Forsterites' b Phlogopites: cations per 22 oxygen atoms Na 0.003 0.003 c Tremolites: cations per 23 oxygen atoms Mg 0.999 0.998 sample 2/4-1 25/1-1 25/2-3 8/1-1 d Diopsides: cations per 6 oxygen atoms from zone: Fo Fo Fo Fo e Forsterites: cations per 4 oxygen atoms Fe 0.017 0.010 f Calcites: cation 1 oxygen atom: Ca 0.980 0.992 Mg 1.918 1.935 1.899 1.957 minerals are the one that record maximum temperatures Mn 0.002 0.002 Fe 0.076 0.057 0.096 0.052 e Dolomites cations per 2oxygen atoms Ti 0.001 0.002 Ca 0.002 0.O01 0.001 0.000 minerals with maximum recorded Ca content Al 0.009 0.012 Mn 0.004 0.003 0.004 0.002 Si 1.993 1.987 Si 1.000 1.001 1.000 0.993 Oxide Sum 99.85 99.90 Oxide sum 100.14 99.98 99.90 100.73 Fe/(Fe+Mq) 0.017 0.010 Fe/(Fe+Mq) 0.038 0.029 0.048 0.026 e) Calcites' sample 2/4-1 2/5-1 24/3-2 4/1-1 8/3-2 8/3-3 22/3-3 22/1-1 22/1-3 21/1-1 13/1-3 from zone: Fo Fo Fo Fo Di+Dol Di+Dol Di+Phl Di+Phl Di+Phl Di+Phl Phi Ca 0.910 0.949 0.933 0.934 0.931 0.938 0.936 0.937 0.941 0.945 0.961 Mg 0.086 0.050 0.066 0.064 0.066 0.059 0.062 0.061 0.057 0.054 0.037 Fe 0.003 0.001 0.001 0.002 0.002 0.002 0.002 0.001 0.001 0.001 0.002 Mn 0.000 0.001 0.001 0.001 0.001 0.001 0.001 0.001 0.001 0.001 0.001 oxide sum 99.45 100.44 99.16 99.95 100.13 99.50 99.07 100.05 99.63 100.41 100.35 Fe/(Fe+Mq) 0.035 0.018 0.016 0.025 0.027 0.040 0.028 0.015 0.016 0.017 0.058 f) Dolomites8 sample 2/4-1 24/3-2 4/1-1 8/3-2 8/3-3 22/3-3 22/1-1 22/1-3 21/1-1 13/1-3 from zone: Fo Fo Fo Di+Phl Di+Phl Di+Phl Di+Phl Di+Phl Di+Phl Phi Ca 1.031 1.032 1.034 1.019 1.023 1.021 1.025 1.020 1.020 1.007 Mg 0.950 0.957 0.956 0.960 0.954 0.963 0.962 0.970 0.964 0.968 Fe 0.016 0.008 0.009 0.021 0.020 0.013 0.010 0.010 0.016 0.024 Mn 0.003 0.002 0.001 0.001 0.003 0.003 0.003 0.000 0.001 0.001 oxide sum 99.23 99.41 100.22 99.10 99.40 99.85 99.63 99.59 99.65 99.22 Fe/(Fe+Mq) 0.017 0.008 0.009 0.021 0.020 0.013 0.010 0.011 0.016 0.024 37 minerals in all parts of the aureole are indicative for the involvement of igneous fluid, even in grade rocks. In this case the untypically high values in the outer aureole may reflect localized infiltration of fluid derived from an underlying igneous body of unknown shape or extent. 38 1 .Or 0.9 0.8 0.7 0.6 0.5 0.4 0.3 O 0.2) Q. • F - Phi (randomly orientated grains) AF - Phi (grains with preferred orientation) • • A • A A i • • I 0.1 o .o1-0 200 400 600 800 1000 Distance (m) 1200 1400 Figure 6: Fluorine content of phlogopites over the whole width of the aureole. The distinction between 'randomly orientated grains' and 'grains with preferred orientation' stems from the fact that in the phlogopite zones, phlogopite grains are significantly more abundant in parts of the rocks where the phlogopite grains show a preferred orientation.. 39 CHAPTER 4 CONTROLLING FACTORS ON THE DEVELOPMENT OF THE METAMORPHIC MINERAL ASSEMBLAGES 4.1 Introduction Metamorphic minerals and mineral assemblages commonly reflect peak metamorphic conditions and sometimes they preserve information on the reaction path. To identify the influences of contact metamorphic processes such as heating and fluid flow the effects of varying bulk compositions must be eliminated first. Below the effects of varying bulk composition on the development of the mineral assemblages in the Horsethief Creek aureole are addressed and the competing influences of heating and fluid infiltration discussed. 4.2 Effects of Bulk Compositional Variations on Metamorphic Mineral Assemblages Part of the petrological complexity in the Horsethief Creek aureole as described above can be assumed to reflect variations in bulk composition. The extent of the variations becomes clear when looking at the results of whole rock analyses as shown for major elements only in Figure 7 a-f. MgO content for example varies from =12-22 weight percent (wt%), CaO from =20 - 30 wt% with a distribution pattern very similar to that of the MgO, KfO varies from =0-3.5 wt% and Si02 from =10 - 48 wt%. The complete data set for the whole rock analyses, plots of the distribution of some trace elements and a description of the methodology are listed in appendix B. 40 a) CaO Distance (km) b) MgO e) FeO SiO •5 1 3 • • Distance (km) Distance (km) e) 0.6 Na20 f) K,0 5 0.2 0.1 -f 1 Distance (km) 1.5 Distance (km) Figure 7: a-f) Major element whole rock XRF analyses. 41 An effective way to illustrate effects of varying bulk composition is by using ternary diagrams. The observed mineral assemblages in the potassic siliceous dolomites can be represented in the Si02 - CaO - MgO - KA102 tetrahedron. As justified in section 3.2.1 it is assumed that prior to metamorphism the siliceous dolomites were composed of dolomite + quartz + K-feldspar ± calcite. All higher grade rocks also contain calcite. This allows the simplification of the compositional phase diagrams by projecting from calcite onto the Qtz - Doi -Ksp plane (Figure 8a). Within the Qtz - Doi - Ksp ternary phase diagram, the effect of minor changes in bulk composition can be visualized. For example, the topology illustrated in figure 8a is for the diopside + phlogopite zone. It demonstrates that the sporadic development of diopside and the complex distribution of phlogopite and K-feldspar-in part reflect bulk compositional variations. The mineral assemblages: Cal+Tr; Cal+Tr+Phl; Cal+Di+Phl; Cal+Di+Ksp; Cal+Di+Tr+Phl; Cal+Di+Phl+Ksp; Cal+Tr+Dol+Phl are all stable on this topology and are all observed in the diopside + phlogopite zone. Diopside-free rocks are not necessarily of tremolite grade but can be assigned to the diopside + phlogopite zone if the bulk composition prevents the formation of diopside. As an example, the assemblage tremolite - calcite - K-feldspar assemblage is stable in fields (4) and (5) of the T-XC02 diagram (see Figure 8b and c). Diopside is produced at the reaction curve between fields (5) and (6). A silica rich sample with the assemblage tremolite - calcite - quartz is stable over fields (3) and (4) and starts producing diopside at lower temperatures along the reaction curve between fields (4) and (5). Hence, the potassic sample still shows tremolite as highest grade mineral in field (5) while the silicic sample already contains diopside in field (5). The stippled field on the ternary marks the approximate range of rock compositions observed in the Horsethief Creek aureole. To systematically depict possible assemblages within and reactions between all the metamorphic zones, the relevant isobarically divariant fields in the T-XC02 diagram were systematically numbered (Figure 8b) and a ternary projection for each of these zones was constructed (Figure 8c). In order to identify changes in mineral assemblages that require a change 42 Figure 8: a) Projection from Cal onto the Qtz-Ksp-Dol plane allowing an easier identification of stable assemblages, and example illustration of the diopside + phlogopite zone in the new simplified system, b) Simplified T-XC02 diagram with the divariant fields systematically numbered for a clear representation of the different metamorphic zones, c) Phase compatibility diagrams for all the identified grades. 43 in temperature or XC02, (and are not an effect of varying bulk composition) the maximum T-XC02 stability range of each mineral assemblage was compiled (boxes in Table 3). New, higher grade metamorphic zones were defined by identifying samples with mineral assemblages that require an increase in metamorphic grade. For example, samples 12/4-2 through 21/1-1 were all assigned to the metamorphic zone (or grade) 5 (Table 3). Their stability fields allow an assignment to any one of the zones 3 through 6 but the down-grade sample 12/4-1 requires a minimum grade 5 and only sample 21/1-3 requires an increase in metamorphic grade above that, i.e. to grade 6. The representative zone for each sample is indicated by the number within the boxes in table 3. Using this procedure we eliminate possible effects of bulk compositional variations, confirm the six earlier identified isograds and also identify what appear to be retrograde assemblages (r#). The upper stability limit of these samples lie below the peak metamorphic grade assigned to the respective areas. Note that these assemblages predominantly occur in samples in the intermediate (to high) temperature range. These "retrograde" assemblages may only appear to be retrograde. They can also be the result of fluid infiltration at or near peak metamorphic conditions as is explained below in section 4.4. See Figure 9 for the spatial distribution of assemblages assigned to zones as described here. 4.3 Thermal Controls If metamorphism is driven by heat flow alone, low-variance mineral assemblages with sharp isograds are expected (e.g., Rice 1977a,b, and section 1.3 this study). The assemblages in the Horsethief Creek aureole are predominantly isobarically divariant and the geometry of the isograds is not very well defined. Additionally, the concept of progressive metamorphism suggests that higher-grade assemblages develop from the lower-grade assemblages observed in an aureole (e.g., Rice 1977a,b). For example, during heating tremolite + calcite should react to 44 Table3: Maximum stability range and representative zone for each assemblage after elimination of bulk compositional effects, resulting in the identification of seemingly retrograde assemblages (r) Note: The numbers in table 3 correspond to the numbers assigned to the stability fields in the T-XQ^ in figure 3b. 12345678 l-> 2 H2 l-)2 1^2 1^ 2 1^ 2 1-»2 in 4<-S JL H HI 6 6 H: 2/4 - 1A 2/4 - IB 2/4 - 2 2/4 - 3 2/2-1 12345678 45 1->2 univariant reactions 6 divariant fields "retrograde" assemblages ELITIC CHERT I-PELITES V % V V "V V V Ni % V V % M V A A MAMA i 100 200 300 m *V S f s rzf \ \ x \ t f f f f \ \ x \ N'VN'V'VHORSETHIEF 'PLUTON S S f * - S \ X * f * f Figure 9: Map showing the distribution of assemblages representing the various metamorphic zones after elimination of bulk compositional effects. The zones are numbered according to the numbering of stability fields in the T-XfjQ2 diagram in figure 8b. Assemblages refer to isobarically univariant (X->Y) or divariant (X) assemblages. "Retrograde" assemblages (X) are apparent retrograde assemblages as described in the text. 46 produce diopside + dolomite. In the next paragraphs it is demonstrated that this may not have been the case in the Horsethief Creek aureole and that heating alone could not have produced the mineral assemblages in the distribution observed in the Horsethief Creek aureole. Other factors also had control on the development of the peak metamorphic assemblages. The sequence of the six identified isograds is consistent with a continuous increase in peaktemperature from =400 to =650 °C at P = 2.2 kbar and at XC02 = 0.6 as schematically indicated in the T-XC02 diagram in Figure 10. 700 650 600 550 o o (0 500 Q. E Q) 450 400 35 tj!0 ' 0'.2 TJ^ TIB 0TET X(C02) r.o Figure 10: T-XC02 diagram showing that the sequence of zones identified could theoretically have been produced by simple heating at XCO2~0.6. To constrain the temperatures attained in the aureole, samples with coexisting calcite and dolomite were analyzed. The MgC03 content of calcite in equilibrium with coexisting dolomite is a function of temperature as defined by the miscibility gap between the two phases (e.g., Goldsmith and Newton (1969)). The Mg-content in calcite can easily be measured and then used to calculate the corresponding temperatures. Here, temperatures were calculated after Anovitz and 47 Essene (1987) with no iron-correction because of the very low iron content observed in the rocks (Table 2). Temperatures calculated using the Fe-correction yielded temperatures within 2 to 3 degrees of the values of the uncorrected calculation. Geothermometry results are plotted in Figure 11a. Each temperature obtained is plotted as a solid circle. A total of 123 spots on 10 thin sections containing coexisting calcite - dolomite were analyzed. These are supplemented by analyses from one forsterite-bearing sample (2/5-1) that does not contain dolomite. This sample was analyzed to constrain the minimum temperature for the forsterite-in reaction assuming that the reaction was a dolomite consuming one. Only the maximum value is plotted for that sample as indicated by the solid triangle at =25 m from the contact. It is found to be within the temperature range observed in dolomite bearing assemblages. The maximum value only is also plotted for the sample from the phlogopite-zone (13/1-1) where a very wide scatter, interpreted to reflect un-equilibrated grains, was observed. Multiple analyses on one grain were possible in some high-grade rocks (see Appendix A2), but in intermediate- to low-grade assemblages small grainsizes precluded multiple analyses on single crystals. In sample 2/4-1, the sample closest to the contact, some calcite grains were analyzed that appear to be inclusions within forsterite. The resulting temperatures of these grains are among -but are not exclusively - the highest recorded. The observed scatter in temperatures for each given sample can be ascribed to re-equilibration during cooling and/or exchange disequilibrium. The Mg-content in some calcite crystals varies significantly. For example, sample 2/4-1 contains calcite crystals with cores that have a higher Mg-content than the rims. In sample 8/3-2 however, calcite crystals with cores Mg-poorer relative to the rims were observed. An additional independent temperature estimation is provided by the peak metamorphic assemblages. The vertical lines and arrows in Figure 11a represent the stability range of the mineral assemblages (as defined in Figure 8b) of each of the samples used for geothermometry. These two independent temperature determinations are in excellent agreement except for sample 48 a) o a> (0 a. E 03 O a> 3 ffi ^_ 0) a E a> 700 600 500 400 300 " Forsterite I' 1 I VV V w 11 Di A Doi V "4"-V V •9 V Diopside + Phlogopite | Di I I I I I I I I 1 I •_! l l I Tr+| !K§PJ Phlogopite i I i i i_ 0 200 400 600 800 1000 1200 1400 1600 Distance (m) b) 700 600 500 400 T T I T I ' I Forsterite h -t W W VW V Di J Do|| Diopside + Phlogopite Di Tr + Ksp 0 200 400 600 800 1000 1200 1400 Distance (m) 1600 Figure 11: a) Plot of geothermometry data. Each solid marker represents a single analysis. The numbers at the top of the diagram refer to the sample numbers. The numbers in parentheses give the number of spots analysed at outcrop(s) at each given distance. The dashed line represents a best fit temperature profile. For description of other symbols refer to the text, b) Maximum possible temperature for the apparent retrograde assemblages compared with the best fit 49 24/3-2 where the highest measured temperature is slightly below what is required for the assemblage to be stable. This internal consistency of the two data sets not only confirms the good quality of the measurements but also the validity of assuming that the fluid pressure was near lithostatic at =2.2 kbar, in accordance with geobarometry data from the pelitic rocks (Floriet 1996). The best estimate of peak metamorphic temperatures (Tmax) throughout the aureole is a smooth line drawn through the maximum calcite - dolomite - geothermometry temperatures. This best fit curve is used as the representative temperature profile below. The open triangles in Figure 11a represent the maximum temperatures at which dolomite free mineral assemblages in other samples are stable. These temperatures are all in agreement with the representative Tmax-profile. In Figure 1 lb the Tmax-profile is compared to maximum temperatures recorded by apparently retrograde assemblages (r# in Table 1). These temperatures are at most only slightly (<20 °) below the Tmax-profile. Hence, nominally retrograde assemblages may have formed at or near the thermal peak and not truly retrograde after substantial cooling. Because the difference in their mineralogy could not be explained by the variable bulk composition either, other factors must have had an important influence. The form of the Tmax-profile is unlike peak temperature profiles predicted by simple models of conductive cooling of igneous intrusions (e.g., Furlong et al. 1991). The plateau may reflect the presence of an igneous body below the middle aureole (cf. Appendix B). However, this is considered unlikely because of the small overall width (1.5 km) of the thermal aureole compared to the =15 km in diameter Horsethief pluton. This problem is discussed in more detail in section 4.5 below. 50 4.4 Fluid Infiltration Because there are several lines of evidence for syn-metamorphic fluid flow, infiltration of a H20 rich fluid must have played a significant role in the evolution of the aureole. The development of hydrous minerals like tremolite and phlogopite, often in high abundance, throughout the inner and middle aureole (Plate 7) from the anhydrous dolomite + quartz + K-feldspar ± calcite protolith requires the introduction of H20. The prevalence of isobarically divariant mineral assemblages and the paucity of high XC02 assemblages (Table 1) also require infiltration, as they are inconsistent with internal buffering of fluid composition by mineral reaction during heating (Greenwood 1975). Infiltration of an H20-rich fluid can trigger mineral reactions in the absence of heating or cooling as outlined in 4.4.1 below. If infiltration is heterogeneous, the development of mineral assemblages with different index minerals can result within a single outcropping. This leads to the conclusion that heterogeneities in the distribution of index minerals which cannot be explained by effects of varying bulk composition or retrograde reaction may have formed as a consequence of heterogeneous fluid infiltration. In section 4.4.2 the diopside forming reaction (5) recorded be mineral assemblages in several samples in the aureole is analyzed in light of the results from paragraph 4.4.1. Having established that fluid infiltration played an important role in the development of the aureole, oxygen and carbon stable isotope data were obtained to place further constraints on possible fluid source(s). The data and its interpretation are presented in paragraph 4.4.3. In paragraph 4.4.4 evidence for silica metasomatism in the aureole is outlined. A comprehensive discussion about possible mechanism fluid infiltration, combining all available data, is given in the next section, section 4.5. Possible explanations for the petrographic break at the tremolite + K-feldspar isograd are also given in that section. 51 4.4.1 Nominally Retrograde Assemblages As outlined above, the mineral assemblages marked "r#" in Table 3 are only nominally retrograde, but their difference in mineralogy cannot be explained by the variable bulk composition. Their maximum possible temperatures are very close to the Tmax - profde indicating their stability at or near peak metamorphic conditions. Infiltration of a fluid with a XC02that is in disequilibrium with the rock triggers mineral reactions as rock and fluid try to re-gain equilibrium. The effects of infiltration of an H20 rich fluid into rocks that were substantially heated prior to infiltration is illustrated in Figure 12. The horizontal arrows in the mid to high temperature part on the T-XC02 phase diagram mark possible isothermal reaction paths that could explain the occurrences of tremolite assemblages in the diopside zone(s) and diopside occurrences in the forsterite zone. Depending on the amount of infiltration and the amount of reaction progress, assemblages with a very different mineralogy can develop. Hence, the "retrograde" assemblages in the Horsethief aureole may record near-isothermal heterogeneous fluid infiltration at the outcrop scale, not reaction during cooling. 4.4.2 Reaction of Diopside to Tremolite - Textural Evidence for Infiltration-triggered Reaction A possible record of a near-isothermal reaction (as described in 4.4.1) is observed in several samples from the aureole in which cores of diopside are preserved within (up to 2 cm long) tremolite porphyroblasts (Plate 8a and b). All of these samples are marked with a "XC" or "C" in the diopside column in table 1 and with a superscript "C" in Figure 9. One of these samples (8/3-2) records the univariant assemblage diopside - dolomite - tremolite - calcite. These observations indicate that diopside reacts to tremolite by reaction (4) (i.e., Di+Dol+H20+C02=Tr+Cal; cf. 52 X (C02) Figure 12: T-Xfjo2 diagram illustrating effects of infiltration of an aqueous fluid into rocks that were heated to temperatures >500 °C prior to infiltration. The blue arrows represent possible reaction paths for isothermal reactions triggered by infiltration. For details see text. 53 Figure 4). This reaction can result from infiltration of a H20-fluid (Figure 7). While the textural observation alone is not unequivocal evidence for prograde reaction, the relatively widespread occurrence of this texture and of other only nominally retrograde assemblages (4.4.1) support the assumption that the diopside -> tremolite reaction is an essentially prograde rather than a retrograde volatilization reaction. Interestingly, the diopside + dolomite = tremolite + calcite reaction curve is the only one with a marked negative slope in the T-XC02 diagram. For this reason the isobarically univariant assemblage diopside - dolomite - tremolite - calcite can persist during down-T fluid flow of low to moderate quantities of an aqueous fluid that infiltrates diopside -dolomite bearing rocks, triggering reaction. Thus, in the case of down-T fluid flow in an aureole, the diopside - dolomite -tremolite - calcite assemblage is the only isobarically univariant assemblage predicted to be stable over a finite distance in the aureole (Ferry (1994), Dipple and Ferry (1996)). Although this assemblage is not the only univariant assemblage in mid- to high-grade Horsethief Creek rocks it is present (sample 8/3-2). Furthermore, numerous dolomite-absent divariant diopside - tremolite - calcite (- phlogopite) samples found in the aureole show the same reaction texture as sample 8/3-2 suggesting that the same reaction proceeded, but went to completion, in these rocks. (Note that isobarically univariant assemblages are generally rare in the Horsethief Creek aureole.) 4.4.3. Whole Rock Carbonate 5'8Q data Isotopic ratios of light elements such as oxygen and carbon are sensitive to many physico-chemical changes in a contact-metamorphic environment. Amongst others, they can be used to determine fluid source, extent of devolatilization during reaction and the nature of fluid-rock interaction (Nabalek 1991). This is possible because "the stable isotope composition of a metamorphic rock is controlled by four factors: (1) the composition of the pre-metamorphic 54 protolith; (2) the effects of volatilization; (3) exchange with infiltrating fluids (with variable composition and fluid/rock ratio); (4) the temperature of the exchange." (Valley 1986, p.445). In this section I briefly discuss the results of carbonate whole rock isotope analyses with respect to the oxygen isotopes. A description of the method, the listing of all raw and converted data and a very brief discussion of the carbon isotope data can be found in Appendix Cl. An excellent overview over the use of stable isotopes in the analyses of fluid flow during contact-metamorphism is given by Nabalek (1991). The whole rock carbonate isotope data obtained for 26 samples over the width of the aureole confirm the infiltration of isotopically distinct fluid(s) and the heterogeneous nature of infiltration. The results are ambiguous as to the source(s) of the fluid(s), but a case of combined vertical influx of pelite-derived fluids and sub-horizontal fluid flow best explains the data. The oxygen isotope data are plotted versus distance from the contact in Figure 13a. There is a clearly discernible 18 O depletion between carbonate samples in the phlogopite zone and those close to the pluton. Values are as high as =23 per mill (%c) in phlogopite zone rocks in the outermost aureole and as low as =16 %c near the contact. Although care was taken to sample unaltered parts of the rocks, the 5180 of =13 %o at =180 m (sample 24-1/4) may reflect influence of late low-temperature alteration. Serpentine is found in altered parts of this rock. The values of =22 - 23 %o in the outer aureole reflect values typical of unaltered dolomites (e.g., Valley 1986). The values of=16 %o in the inner aureole are significantly higher than the values expected for these rocks if they isotopically equilibrated with fluid that emanated from the adjacent granitoid (reference for 'granite'-5180 in Figure 14). However, oxygen isotope depletion by decarbonation reactions generally accounts for a 18 O depletion of no more than =2 %o, independant on whether Rayleigh or batch decarbonation occurs (e.g., Nabalek 1991). The lsO depletion of =7 %c is therefore strong evidence for infiltration of an isotopically lighter fluid. The source(s) of the fluid(s) in this case is (are) not obvious. If the main source of the isotopically distinct fluid was not the pluton, the underlying metapelite is 55 a) 24 22 20 18 16 I 14 4 12 | 10 200 400 600 800 1000 Distance (m) 1200 1400 1600 b) 700 600 500 400 ( -L Figure 13: a) Whole rock carbonate oxygen isotope data. The vertical grey lines mark the locations of the isograds (Fo, Di+Dol, Di+Phl, Di, Tr with increasing distance). The thick, black dashed line outlines the oxygen isotope pattern if the sample at =180 m is assumed to be retrogressed. A 'front' is observed at =1150m (for details see text). The thin solid green line outlines the isotope pattern if all analyses are interpreted to repre sent contact-metamorphic values. A "two-front" pattern emerges with the one front at =1150 m and the other one closeto the contact (=200 m). The effect that the one sample at =180 m has on the overall appearence can be weakend if assuming a more gradual transition as indicated by the fine dashed green lines (for possible explanations refer to text), b) "Best-fit temperature' profile and locations of samples that show the reactions texture of diopside cores within tremolites('C') (as described in paragraph 4.4.2). Note the correlation between isotope 'front', end of thermal plateau and outermost occurrence of the diopside -> tremolite reaction texture. 56 rS180 (SMOW) 40 30 20 10 0 -10 -20 I I I I I I -30 -40 -50 1 1 — Shales? ^™ Siliceous Dolomites3 l-type Granitic rocks2 ^™ S-type Granitic rocks2 1 after Hoefs 1997 2 after Nabalek 1991 3 after Valley 1983 Figure 14: Oxygen isotope signature of various geological materials that are possible sources for the fluid(s) that infiltrated the siliceous dolomites in the Horsethief Creek aureole. 57 another possible source. Metapelitic rocks are presumed to have a 5lsO signature of = 15 - 20 %o prior to contact-metamorphism (Figure 14). These values are in agreement with the 8lsO observed in the Horsethief Creek aureole. Another interesting observation is the wide scatter of 8180 at any given distance, especially in the inner aureole. Values range from as high as 21%o to as low as 17%o at =100 m distance from the contact. This points to strongly heterogeneous fluid flow in this area. On the other hand there is no systematic change in whole rock carbonate 8lsO with distance or style of infiltration. There is no correlation between the amount of veining in the rock and the carbonate whole rock 8180 values (Dipple and Niermann, 1997). Preliminary tests of 'spot-analyses' across veins by means of micro-drilling confirms this. No changes in isotope signature is found across veins into matrix material (Dipple, pers. com.). This suggests that infiltration was homogenous on the scale of individual beds and that vein and pervasive flow were well mixed. This attests to the overall strongly pervasive nature of infiltration in the Horsethief Creek aureole. A more comprehensive discussion about possible fluid sources and fluid flow direction is possible if the isotope values are not discussed by themselves, but rather are analyzed in combination with petrographic observations and heat flow considerations. This is done in section 4.5 below. 4.4.4 Silica metasomatism and intensity of infiltration The pervasive development of high abundant tremolite in most rocks and the general absence of dolomite in the inner and middle aureole suggests that metamorphism was not isochemical and that silica was introduced into the system by the infiltrating fluid. This is also reflected by the observation of often silica richer bulk composition in and around veins, including the chaotic veining in the inner aureole (Plate 10) (cf. Figure 8c, see also Appendix C2.) Possible sources of 58 the silica are the interbedded chert layers, and/or the source(s) of the infiltrating fluid, namely the pelites and the pluton. The strongly pervasive character of infiltration and silica metasomatism in all except the outermost aureole (>1250 m from contact) is unusual for contact metamorphosed siliceous dolomites. Siliceous dolomites in other contact aureoles are often reported to show more isobarically univariant assemblages and/or unreacted dolomite (Moore and Kerrick 1976; Rice 1977a b; Holness 1992). Infiltration coupled with extensive silification is, if at all, only observed to occur relatively close (< =150 m) to the contact (e.g. Bucher-Nurminen, 1982). The introduction of large amounts of silica by the fluid may have a profound influence on the permeability structure of the aureole. For example, the tremolite-in reaction (2) is coupled with a volume decrease of =24 % if all Si02 is provided by quartz. In contrast, the same reaction is coupled with a volume increase of=19 vol%, if all Si02 is provided by Si02{aq), dissolved in the infiltrating fluid. While the overall permeability is not only a function of porosity, these different volume changes of the solids can be expected to influence the infiltration geometry. In the Horsethief Creek aureole both kinds of reactions likely took place but the extent of each reaction and its influence on the permeability structure are not clear. If one assumes that all silica was brought into the system, some first order estimates can be made about the amounts of fluids infiltrating the rocks. However, in the case of the Horsethief Creek aureole this seems futile, because of 'complications' like the interlayered chert layers. Any fluid which loses silica during reaction with the rocks would be "recharged" (become silica saturated again) when it comes in contact with one of these interlayers. As this process cannot easily be quantified, every estimate of the amount of influx would likely overestimate the amount of externally derived fluid. 59 4.5 Discussion of Possible Fluid Flow Direction(s), Fluid Source(s) and Heat Advection In this section the possible interplay of heat and fluid flow and different scenarios of fluid infiltration, with respect to fluid flow direction and source(s), are discussed. In paragraph 4.4.1 the atypical form of the Tmax profile and the possibility of heat advection are briefly discussed. In paragraphs 4.4.2 to 4.4.4 scenarios of up-Temperature, down-Temperature and vertical fluid flow in the aureole are outlined and evaluated. These models do not consider the phlogopite zone, down-grade of the petrographic break at the tremolite + K-feldspar isograd. The outermost aureole may have different reaction history. Possible explanations are outlined in paragraph 4.4.5. Two competing models are introduced to try to explain the sudden change in petrologic and infiltration character at the tremolite + K-feldspar isograd. An attempt is made to tie the implications of this observation into a model consistent with the observations in the rest of the aureole. 4.5.1 Heat Flow - Conductive or Advective? Meaningful discussion of fluid flow direction and fluid sources requires the consideration of the combined petrographic, isotope and heat flow data. The former two are discussed in chapter 3 and section 4.4.3 respectively. Consideration of heat flow recorded in form of maximum temperatures reached in the rocks can support attempts to explain the overall contact-metamorphic history of an aureole. No in depth analysis is attempted here, but some aspects of possible heat flow mechanisms in the Horsethief Creek aureole are worthwhile considering. For an excellent review of thermal modelling see Furlong etal. (1991). The profile of maximum temperatures (Tmax) recorded in siliceous dolomitic rocks in the Horsethief Creek aureole is not in agreement with simple conduction of heat from an intrusion 60 with near vertical contacts to the wallrock. This form of purely conductive heat flow results in up-temperature concave Tmax - profiles (Jaeger 1964). Simple models can provide valuable constraints on the thermal evolution, although in more detail it is observed that, for example, the geometry of the intrusion is a very important factor in controlling the spatial distribution of temperature in an aureole (Furlong et al. 1991, and citations therein). The absolute values of Tmax throughout an aureole are strongly influenced by the intrusive geometry (Bowers et al. 1990). However, Furlong et al. (1991, p.444) pointed out, that in "... a purely conductive regime, the average maximum width of an aureole will be approximated by ... simple models... .". For initial temperature distributions, and temperature increase (of the wallrocks) necessary for reaction similar to those assumed in the Horsethief Creek aureole [AT (intrusion - wallrock) = 500 °C, AT necessary for reaction 100 °C], the intrusion of a cylindrical solid igneous body would lead to the development of an aureole with a width =1.5 a (a = half-width of intrusion [km]). In case of the Horsethief Creek pluton with a = 7, the aureole should be =10 km wide if heat flow was purely conductive. This is a distance much larger than that observed. Furthermore, the possibility of an intrusive body below the aureole, causing the unusual thermal plateau, may also be excluded. While such an additional heat source can lead to the development of a plateau in the temperature profile, it also is expected to result in an increase in the overall width of the aureole. Hence, the contact between the Horsethief Creek intrusion and the siliceous dolomites is assumed to be near vertical. Therefore, the significant deviations from the theoretically predicted conductive heat flow require the consideration of effects of advective heat transport (Furlong et al. 1991). Systematic relatively closely spaced fractures in rocks can lead to a permeability high enough to allow for advection of heat (Furlong et al. 1991). Sets of narrowly spaced joints or veins are observed in all lithologies in the Horsethief Creek aureole. An additional requirement is that high permeability is maintained over (some) time to have significant effects on fluid- and heat flow. In the Horsethief Creek aureole, fluid infiltration and production can be expected to have been high. 61 The influx of external H20-rich fluids into the siliceous dolomites from the large pluton and / or the underlying pelitic rocks, in which dehydration reactions proceed during contact metamorphism, can be expected to have been high. In addition, internal H20 -C02-fluid production in the pervasively reacting siliceous dolomites themselves was also high. This results in fluid pressures likely high enough to maintain a relatively large porosity and permeability against pressure of compaction and and/or sealing processes (e.g. Fyfe et al 1978). Possible implications of advective heat flow will be addressed in the discussions about fluid flow direction and sources below. 4.5.2 Discussion of the possibility of horizontal fluid flow in the direction of increasing temperature The scenario of up-Temperature (up-T) fluid flow in the siliceous dolomites in the context of a large hydrothermal convection cell is unlikely. If the convection cell operated in an open system, significant amounts of meteoric water should be involved. Meteoric waters have a very low positive to high negative 8'80 signature (8180 range in Figure 14) and therefore should lower the 8I80 values below those recorded in Horsethief Creek rocks (e.g., Valley 1986 and citations therein, Holness and Fallick 1997). Small amounts of meteoric waters may have been involved, especially if their isotopic signature was altered by rocks they passed before infiltrating the siliceous dolomites. Hydrogen isotope analyses of tremolites could possibly clarify whether meteoric waters were involved. The hydrogen isotope signature of meteoric water is very distinct from that of other geological materials and should be reflected in the minerals even if the oxygen isotope signature is ambiguous. Tremolite seperation was hindered by the often very fine grained and interlocking minerals and by widespread weathering in the strongly cleaved amphiboles. 62 If small(er) scale convection occurred without involvement of significant amounts of meteoric waters the system was likely 'sealed' by a relatively impermeable cap above the intrusion (Etheridge et al. 1983, Furlong et al 1991). The recorded 18 O depletion seems in accordance with up-T flow of fluids with an 8180 of =18 %o at the time of influx into the dolomites in the outer aureole. This value is in accordance with what can be expected for a fluid that equilibrated with the surrounding mainly pelitic sediments. Flow up a temperature gradient causes lsO depletion in the rocks because the fractionation factor (a) for fractionation of 18O between rock and fluid changes with temperature. In the case of carbonate, a > 1 at low temperatures, and approaches 1 with increasing temperatures. This leads to a lsO depletion of marble towards the intrusion (Dipple and Ferry, 1992). Although the temperature gradient observed in the Horsethief Creek aureole is distinct from those utilized by Dipple and Ferry, the overall depletion (including depletion by devolatilisation reactions) is comparable to that predicted for the given temperature range. The interpretation of the 6180 isotope pattern observed in the inner aureole alone (in the scenario of up-T flow) is more uncertain. If the outlier is representative the 18 O depletion seems to be too high to be the result of temperature induced 18 O depletion only . If the outlier is unrepresentative, the 8180 values in the inner aureole seem to actually increase towards the pluton, or at least stay constant. This is unlikely to be the case with up-T fluid flow where the temperature gradient is expected to be steepest close to the contact. This should lead to some 180 depletion. Hence, up-T fluid flow in the inner aureole seems unlikely. However, as pointed out in the introduction (paragraph 1.2) the inner aureole often seems to have a special position. Fluid flow here might have been sub-vertical. The isotope values in the inner aureole are in a range expected with involvement of pelitic fluids emanating from the underlying pelites. Whether the sequence of mineral assemblages observed in the aureole are in agreement with up-T flow is somewhat difficult to assess. If the one diopside - dolomite - tremolite - calcite assemblage observed is not representative, the assemblages could have formed with moderate 63 amounts of fluid infiltration (very tentative value of=100 moles / cm2 time integrated flux (q°) cf. Figure 2d in Dipple and Ferry 1996). Much higher amounts of fluid flux seem to result in the development of finite areas in which isobarically univariant assemblages are stable (esp. in the inner aureole) (Ferry 1994, Ferry and Dipple 1996). However, if the aforementioned isobarically univariant is representative, up-T flow is very unlikely. The T^ distribution in the aureole cannot satisfactorily be explained with up-T fluid flow. As pointed out above purely conductive heat transport seems unlikely in the Horsethief Creek aureole. Heat advection by fluid in context of a convection cell can lead to a decrease in the width of an aureole in the up-T fluid flow region (decrease with respect to an aureole which experiences purely conductive heat flow) ( Dipple and Ferry 1996; Furlong et al. 1991). However, the extent of shortening observed in the Horsethief Creek aureole seems too high to be the result of effects of up-T flowing 'cool' fluids. Furthermore the unproportionally wide metamorphic zone(s) of mid to high grade (diopside + phlogopite plus diopside zone(s)) (Figure 13b) seem difficult to explain with this kind of model. If a decrease in the overall width of the aureole occurs, caused by heat-advective fluid flow up-T, the width of the different metamorphic zones seem to decrease proportionally and high-grade zones are in first approximation narrower than low-grade zones (Furlong et al. 1991). Even with a complex geometry of (steeply dipping) intrusion-wallrock contacts, where unusually wide higher-grade zones can develop in and near reentrants of the contact, these higher grade zones are not wider than the mid-to low grade metamorphic zone (Bowers et al., 1990). In summary it can be said that although up-T flow with a strong vertical component in the inner aureole cannot entirely be dismissed, it seems unlikely. Many uncertainties are connected with this hypothesis. 64 4.5.3 Discussion of the possibility of horizontal fluid flow in the direction of decreasing temperature Down-Temperature (down-T) flow in context of a convection cell seems unlikely in the Horsethief Creek aureole. Those parts of a convection cell in which down-T flow occurs are commonly situated above the top of the pluton (Furlong et al. 1991 and citations therein). Present day erosion levels in the working area show that, while it is not clear where exactly the rocks under consideration are located with respect to the original vertical extend of the intrusive body, they were not above the pluton. Elevations in the intrusive rocks reach higher than those in the working area. In the case of likely lithologically controlled down-T fluid flow the crystallizing pluton can be expected and is often observed to be the main source of infiltrating fluid (e.g., Valley 1983), and in theory, a sharp front is expected to develop in the infiltrated rocks. Upstream from such a front the rocks should be fully equilibrated with the infiltrating igneous fluid while downstream from it the isotopic composition of the carbonate rocks are unchanged. No igneous values with 5lsO as low as =10 %o are observed in Horsethief Creek carbonates. One explanation could be disequilibrium in isotope exchange between the magma derived fluid and the infiltrated rocks. But the high temperatures recorded in contact aureoles are generally considered to be sufficiently high for a fast equilibration, and in many other aureoles, where the fluid is considered to be of igneous origin, values of 8I80 =10 %o are recorded (e.g. Nabalek et al. 1984, Cook et al. 1997, and examples in Valley 1986). Furthermore, the amount of fluid emanating from an intrusion and the distance of 18 O depletion in an aureole with down-T flow can be approximated by some simple calculations (Dipple and Ferry, 1992). Fluid emanating horizontally from a crystallizing magma with 10 wt% H20 is expected to drive isotopic depletion to a distance of around 0.087 r ( with r = radius of a (cylindrical) pluton [km], 0.08 mole oxygen/cm3 rock, density of intrusion = 2.5 g / cm3 and 65 complete exsolution and expulsion of the H20). In the Horsethief Creek case r ~1, and 18 O depletion should reach no further than 600 m . This distance can be expected to be even smaller with a higher mole oxygen per cm3 rock value of the originally very dolomite rich lithology (0.125 mole oxygen / cm3 rock). This distance of 18O depletion is significantly smaller than observed. If the pluton was the main source of the fluids these must have been channeled into the siliceous dolomites to provide the required larger amounts of fluids. Such a channeling was proposed for the siliceous dolomites in the Alta aureole (Cook et al, 1997) where the Grizzly thrust just above the pluton is considered to be the low permeability 'cap'. This likely caused layer parallel fluid flow at its bottom. In case of the Horsethief Creek rocks no likely candidate for such a barrier is known. However, if significant amounts of igneous fluids infiltrated the aureole their isotopic signature must have been altered immediately after, or before, flowing across the pluton -wallrock contact to produce the relatively high 8'80 values in the inner aureole . This could result from mixing with another fluid. A possible source for additional fluid are the underlying pelites. At the time of intrusion, low-grade meta-pelite are presumed to have a 8'80 signature of =15 - 20 %o (Figure 14). In the case of (sub-)vertical flow isotopic exchange across the pluton-wallrock contact may have changed the isotopic values of fluid on either side of that contact. The mineral assemblages in the siliceous dolomites are generally in agreement with down-T flow of moderate to high total fluid fluxes, depending on whether the diopside - dolomite -tremolite - calcite sample is representative or not (Dipple and Ferry, 1996). The isobarically univariant assemblages in the phlogopite zone may be an exception. This is further discussed below in section 4.6. The Tmax profile is likely not in agreement with down-T flow in all of the aureole (if the unusual Tmax profile is the result of heat advection). With heat advection by down-T fluid flow the width of the aureole is not expected to decrease. 66 In summary, lithologically controlled horizontal fluid flow in the direction of decreasing temperature throughout all of the Horsethief Creek aureole is unlikely. The isotopic signature in the inner aureole is too high, and 18 O depletion is observed at greater distances than can likely be expected in the case of horizontal flow of plutonic fluid only. Some combination of down-T flow in the middle and outer aureole and some other process in the inner aureole seems possible. Hence, although both up- and especially down-temperature fluid flow scenarios cannot strictly be excluded, uncertainties are linked with both scenarios. Additional uncertainty in the determination of fluid flow direction and fluid source(s) is introduced by the fact that the above used models are fairly simple. Many of the complex and interacting processes in aureoles, like for example changes in permeability and flow pattern with reaction, are not considered and could potentially alter the model results significantly (e.g., Furlong et al. 1991, Cartwright and Buick 1996). However, interpretation of Horsethief Creek data with the methods currently available suggests that the pelitic rocks underlying the siliceous dolomites may have played a significant role in the infiltration history of the aureole. 4.5.4 Discussion of the possibility of vertical fluid flow Further evidence that vertical fluid flow of fluids liberated by dehydration reactions in the pelites may have played a significant role in the infiltration history of siliceous dolomites comes from a more detailed comparison of petrographic observations and geothermometry results with the isotope data. If one had to point out an 'isotopic front' in the system it could be placed at =1150 m as indicated by the dashed line in Figure 13. Interestingly, the thermal plateau at temperatures of =560 °C as shown in Figure 13b ends at =950 m, and the temperature of =525 °C necessary for the diopside-forming reaction under internally buffered conditions (reaction (5)) is reached at 67 =1180 m. The outermost sample showing the diopside to tremolite reaction texture analyzed in section 4.4.2 was found at =1050 m. This indicates that infiltration up to that distance from the pluton occurred in rocks that were significantly pre-heated. Diachronous heat- and fluid flow is thought to be not uncommon in contact metamorphic environments. A model of rapid heating of the inner aureole followed by infiltration of an aqueous fluid was invoked to explain the mineral assemblages observed in argillitic rocks in the Notch Peak aureole (Nabalek et al. 1984). The formation of forsterite, in rocks of the rapidly heated rocks of the inner aureole, directly from the protolith (and possibly with silica in the fluid) was proposed for other aureoles (Bucher-Nurminen 1981, Holness 1992). What is remarkable in case of the Horsethief Creek aureole is, that this process of heating preceding infiltration is not restricted to the inner aureole but seems to have been operative to distances >1 km. Metamorphism of the pelite and its implications for the siliceous dolomite In the pelites, fluids are expected to be liberated at temperatures of =520 - 540 °C by means of the reaction chlorite + muscovite + 2 quartz = cordierite + biotite + 4 H20 (7) (Bucher and Frey 1994, Pattison and Tracy 1991, Fig. 47), a reaction observed in the contact-metamorphic pelitic rocks by Floriet (1996). This match of data supports the hypothesis that the pelitic rocks were an important source of fluid(s) that infiltrated the siliceous dolomites in the Horsethief Creek aureole. The fluids evolving from the pelites would be buoyant and have the tendency to escape upwards. Relatively pure massive carbonate rocks are considered to be comparably impermeable and were found to be barriers for uprising fluids, leading to layer-parallel flow along lithologic boundaries (Hay and Evans 1988, Holness and Graham 1995, Nabalek et al. 1984, Ferry 1989, 68 Cook 1997). However, in case of the Horsethief Creek aureole sub-vertical, pre-contact-metamorphic joints which developed in all lithologies and trend across bedding could have played a major role in facilitating infiltration across layering. Metamorphism of the siliceous dolomite in the middle aureole At temperatures >525 °C, similar to those necessary to drive reaction in the metapelites, the thermally driven diopside producing reaction (5) proceeds in the siliceous dolomites. This reaction is coupled with a negative reaction volume of the solids of 39 %, which is large, especially when compared with the volume reduction of 'only' 23 % during the tremolite-in (2) and of 13 % during the phlogopite-in (1) reactions (for molar volumes see Appendix D). Hence, this diopside forming reaction (5) leads to a local reaction -enhanced porosity and possibly permeability. It is likely restricted to areas where both reactants are readily available namely the contact between chert and siliceous dolomites, as transport of material has to occur by diffusion. This is a relatively sluggish and small scale process (e.g. Bickle and McKenzie 1983). A possible example of a rock that may have experienced such a reaction is shown in plate 9. Once fluid is released by the underlying pelites, it rises and infiltrates into the siliceous dolomites predominately by means of the pre-contact metamorphic joints, and possibly along the chert - carbonate contact (Heinrich 1993, Nabalek et. al. 1992, Jamtveit et al. 1992, Plate 7 ). This leads to a combined upward and sub-horizontal fluid flow. The fluid channeling along fractures is reflected by the abundant systematic veins in the siliceous dolomite. The veins are now observed to be filled with tremolite or to have been nucleation 'points' for tremolite sprays. Signs for sub-horizontal flow are less obvious. Observations of textures like that shown in Plate 7 are rare. Possible further evidence for sub-horizontal fluid flow may come from the overall development of the aureole outbound the diopside isograd. This is discussed in detail below. 69 In parts of the rocks in the middle aureole, where diopside was formed prior to infiltration, subsequent infiltration would trigger tremolite formation by reaction (4). In the parts of the rocks where no diopside was formed prior to infiltration tremolite may have formed directly from the protolith by reaction (2) at temperatures of infiltration and reaction >520 °C, which are the temperatures necessary to liberate the fluid in the metapelites. This hypothesis is supported by the fact that (a) although the diopside -> tremolite reaction texture is found throughout most of the aureole it is not very common; (b) one of the most widespread assemblages in the aureole is tremolite - calcite - K-feldspar which can form by reaction (2) in the presence of abundant silica, likely brought into the system by SiOybearing fluids (4.4.5). With massive infiltration but no silica metasomatism the tremolite + K-feldspar + calcite assemblage could only have formed in quartz rich areas (i.e., near the chert layers by the combined or successive reactions (1) (2) and (4) (cf. Figure 4, esp. reaction stoichiometry of reaction (4)). In areas with less silica available some dolomite and/or phlogopite should remain preserved in the rocks. Such assemblages are present, but especially dolomite-bearing ones are rare. If temperatures continued to rise after the onset of infiltration, and if infiltration of an aqueous- (Si02aq) -fluid persisted, diopside could form from tremolite during peak thermal, conditions by reactions: tremolite + 2 quartz + 3 calcite = 5 diopside + H20 + 3 C02 (8) or 3 tremolite + K-feldspar + 6 calcite =12 diopside + phlogopite + 6 C02 + H20 (9). The different reaction paths for the formation of tremolite and diopside may have let to the development of different textures. Diopside that is interpreted to have formed from the protolith is found to be large and bladed. Other diopside, which may have formed from tremolite along reaction (8) or (9), occurs in relatively small (mm) roundish crystals rather than in long blades, while the tremolite occurs in smaller needles. No reaction texture (tremolite -> diopside) was found that could support this assumption. 70 Metamorphism of the siliceous dolomite in the tremolite + K-feldspar zone In the aureole outbound of the diopside isograd, i.e., in the tremolite + K-feldspar zone, tremolite had to form from the protolith because maximum temperatures reached in this zone are too low for the diopside producing reaction (5). These temperatures also are below the =520 -540 °C at which the pelitic rocks dehydrate. Hence, the influx of fluid into the tremolite zone rocks was likely from higher grade, sub-horizontally channeled along the chert - carbonate contacts. The Tmax-profde Possible further support of the hypothesis of vertical fluid flow comes from the analysis of the Tmax-profile in light of possibly vertical fluid flow. As pointed out above (4.5.1) the overall width of the aureole is relatively small. Aureoles of around the same width (1-2 km) are observed to have developed around plutons that are considerably smaller (2-5 km) than the =15 km in diameter Horsethief Creek pluton (e.g., Bowman 1994; Rice 1977a). In contrast to these examples, isograds in the Elkhorn aureole, where the existence of strongly vertical flow during contact metamorphism is well established (Bowman and Essene 1982, Bowman et al. 1985), are found to be telescoped to =100 m only. The (highest grade) periclase (brucite after periclase) zone is observed to be up to 20 m wide. The Black Butte stock around which this aureole developed is about 1 km wide and an aureole of several hundred meters in width could be expected if heat flow was purely conductive. The mineral assemblages in the Horsethief Creek aureole are different so that the interpretations from Elkhorn are not directly transferable. It is interesting to note however, that in the Horsethief Creek aureole, not only the overall width of the aureole is relatively small, but also that a very wide (=640 m) relatively high-grade diopside + phlogopite 71 zone developed, which is succeeded down-grade, i.e. outside the zone of proposed vertical flow, by a narrow -125 m wide tremolite + K-feldspar zone. With strong vertical fluid flow the conductive heat profile might become significantly distorted by vertical heat advection. 'Heat' is taken out of the system before it can reach the outer parts of the aureole, so that the zone of heating is considerably shortened, similar to what is observed in the Horsethief Creek aureole. No attempt to quantify the effects of heat advection by vertical fluid flow was made. The inner aureole - a special case Despite strong evidence for possible involvement of large amounts of pelitic fluid, some involvement of plutonic fluid is likely, especially in the inner aureole where chaotic veining was observed. This kind of fracturing is commonly developed in host rocks near intrusions and is interpreted to be the result of overpressuring of fluids expulsed from the cooling pluton. The wide scatter of 8180 values at any given distance in the aureole can be interpreted to reflect interaction of the rocks with different amounts of the fluids caused by heterogeneities in infiltration pattern. However, the scatter is especially well marked in the stoped and chaotically fractured rocks close to the contact where differences in permeability are expected to be large. In addition, the involvement of igneous fluid was likely largest in the zone near the pluton wallrock contact. Hence, the variability in 8180 values may not (only) reflect differences in the total amount of fluids the rocks interacted with but (also) different proportions of igneous and metamorphic fluid that equilibrated with the rock. In the latter case the sample at =180 m in figure 13a which was originally assumed to be retrograded may represent a rock that equilibrated with a relatively large proportion of igneous fluid instead. As of the timing of peak metamorphism in the aureole in the case of predominantly vertical fluid influx, it can be said that reaction / metamorphism likely progressed continuously with the 72 thermal peak 'wandering out' from pluton. Details of the relative timing of metamorphism in various parts of the aureole are not clear. In summary it can be said that a likely model of fluid infiltration into siliceous dolomites in the Horsethief Creek aureole is one of strong vertical flow of fluid, which was released from the underlying pelites with the onset of contact metamorphism, probably mixed with sub-horizontal down-T fluid flow and some involvement of plutonic fluid. However, so far all attempts to decipher the contact-metamorphic evolution of the rocks ignored the phlogopite zone. This zone seems to be distinct from the rest of the aureole. A sharp transition from isobarically divariant assemblages in most of the aureole to isobarically univariant assemblages in the phlogopite zone occurs. Implications of this observation are discussed below. Furthermore, the nature of metamorphism affecting the calc-silicate layer that lies between the pelite and the siliceous dolomite is not known. Outcrop is sparse as the unit lies in the topographic lows. The rocks are made up of mm - cm wide biotite - diopside - quartz rich layers. Joints were observed but veins were rarely found. 4.6 The Petrologic Break - A Change in Infiltration Character Another striking feature of the aureole is the significant break in the petrologic character at the tremolite + K-feldspar isograd, at an inferred peak temperature of =500 °C. Up-grade of the tremolite + K-feldspar isograd the aureole is dominated by isobarically divariant mineral assemblages while the phlogopite zone down-grade of the tremolite + K-feldspar isograd is characterized by the isobarically univariant assemblage dolomite - K-feldspar + minor phlogopite ± calcite (Table 3, Figure 9). A talc zone is predicted by the models of Ferry (1994), except for very low fluid flow (time integrated flux <100 mole / cm2) during down-T flow. No talc zone is observed in the Horsethief 73 Creek aureole. This could result if the infiltrating fluid had a XCG2> =0.54 because infiltration of C02-rich fluids would not trigger any talc producing reaction (Figure 4). The absence of a conspicuous phlogopite zone in the outer aureole however is unlikely to be the consequence of a fluid-XC02 too high to trigger reaction. This would require the infiltrating fluids to have a XC02 > =0.82. The possibility of such high XC02 values to develop is remote when significant amounts of H20-bearing fluid infiltrate a rock. A change in infiltration regime at the tremolite + K-feldspar isograd is likely required. While the inner and middle aureole has to have experienced substantial and pervasive infiltration of an externally derived H20-bearing fluid at or near the peak of metamorphism, the mineral assemblages in the phlogopite zone must have been heated in the absence of significant infiltration. Reaction here was governed by internally buffered fluid composition. This apparent lack of large amounts of fluid infiltration into phlogopite zone rocks can be explained as either reflecting a) a real lack of significant amounts of infiltration, b) conditions that did not allow for reaction despite infiltration, or c) both. In the first case, if the main fluid source were the underlying pelites, vertical influx would cease approximately at the diopside isograd and the horizontal flow component might not have been strong enough to transport fluids out into the phlogopite zone rocks. This, however, is not very likely because layers with abundant Fe-amphibole were found in the (Mg-) phlogopite zone. This requires that significant amounts of fluid were present at a time where temperatures were high enough to drive reactions. Hence, cases a) and c) can basically be excluded. Hence, fluid infiltration likely occurred but for some reason no reaction was triggered in the phlogopite zone (Mg-) rocks. In this case fluid infiltration could have been restricted to a time in the thermal development of the aureole when the outer aureole had not been heated up enough to allow for reaction with the infiltrating H20-C02 fluid. The thermal peak in the outermost aureole would only be reached after infiltration ceased. This idea of diachronous heat and fluid flow is similar to that proposed for the argillitic rocks in the Notch Peak aureole by Nabalek et al. (1984). 74 Alternatively, if fluid infiltration and heating were coexisting, some other process must have hindered reaction. 4.6.1. A Model of diachronous heat and fluid flow The first model evoked to explain the petrologic break at the tremolite + K-feldspar isograd presumes diachronous heat- and fluid flow, whereby the intrusive body is assumed to be the heat source as well as a main source for the fluid(s) that infiltrated the siliceous dolomites. This model is built on results of computer simulations by Dipple and Ferry (1996). They simulated the contact-metamorphic development of a simple model aureole (initial rock composition 10 mol% quartz and 90 mol% dolomite) by coupling mineral reaction, heat- and fluid flow. The results showed that rocks in the middle to outer aureole can experience early infiltration and reaction at temperatures below peak metamorphic ones, leading to the development of low-grade mineral assemblages. With continued heating after the cessation of infiltration the low-grade assemblages react to higher grade isobarically divariant assemblages along shallow reaction curves with thermal maxima (Figure 15a). Traces of a possible early talc zone in the outer aureole are erased during this "annealing" process (Figure 15 b). Rocks even further away from the pluton (in the outermost aureole) do not reach sufficiently high temperatures during infiltration to drive any mineral reaction. Only during the continued, post-infiltration heating internally buffered isobarically univariant assemblages are produced (Figure 15b). In the Horsethief Creek aureole the missing talc zone and the observation of generally only small amounts of phlogopite in isobarically univariant assemblages can be interpreted to reflect such an annealing process where fluid flow ceased early in the thermal evolution of the outer aureole. 75 a) 650 Per 600 u 550 450 400 350 0.0 —1—1 1— #"* bol P = 1 kbar Fo + Do! + Cal -1 Brc + Ca] am Dot C ^^Di + Dof Tr + Doi + Cal Doi + Qtz 1 Talc zone anneals to 1 Tremolite during •post-fluid flow beating Jrr + Dol A ^TTIc + E ol + Cal I" i I Di 0.2 0.4 0.6 0.8 1.0 Figure 15: a) T-Xco2 diagram for pure siliceous dolomites (CMS-HC) at P=lkb (after Dipple and Ferry 1996). The thick red arrow indicates the reaction path for rocks that developed talc assemblages during infiltration early on in the thermal development of the aureole and 'anneal' during heating after the termination of infiltration (for more details refer to text), b) (I): Mineral abundances over the width of the aureole resulting from a simulation of contact metamorphism with diachronous ^ heat- and fluid flow in pure siliceous dolomites. A total of 400 moles H20/cm2 flowed down-temperature with a Darcy flux of lxlO-8 cm/s for 29,000 years. (II) Modes after an additional 171,000 years of heat flow without infiltration. The two figures show the effect of 'annealing': low-grade divariant assemblages (high percentage Tic area) react to divariant mid-grade assemblages (high percentage Tr area) and rocks in a greater distance from the contact only develop univariant assemblages (low percentage [=0%]) assemblages (after Dipple and Ferry 1996 - their Figures 3 a-b). 0 4 ^—Fo 1 1 (1) Mineral abundance at time of termination Tr T| of fluid flow — Ti (29'000 yrs) < is "^Fo 1 i ! 1 Tr n (II) Mineral abundance after additional heat flow (+171,000 yrs) Tic Distance from Pluton 76 There are however some problems connected with this model of diachronous heat and fluid flow. First, if an early (during infiltration) developed phlogopite zone annealed to a divariant tremolite bearing assemblage it could only be by reaction 6 calcite + 5 phlogopite + 24 quartz = 3 tremolite + 5 K-feldspar + 2 Ff20 + 6 C02 (10). This reaction requires a very large amount of Si02. Although it can be argued that silica is highly abundant in form of chert layers and nodules it is questionable whether that silica was available for reaction in all parts of the rock. With no infiltration of an external fluid, diffusion alone might not be sufficient to transport for reaction necessary components through all of the rocks (e.g., Bickle and McKenzie, 1987). Annealing of rocks in the during infiltration developed phlogopite and tremolite + phlogopite zones would not occur in areas that were outside the regions reached by diffusive silica. The assemblages tremolite - phlogopite - calcite and possibly also phlogopite -dolomite - calcite in the rocks that that did not experience reaction (2) during infiltration would be preserved because both of the reactions that could possibly proceed at higher temperatures, i.e. reactions (2) and (3), require (the not available) Si02 as a reactant. In the Horsethief Creek aureole none of the rocks in the lower tremolite + K-feldspar zone exhibit an isobarically divariant phlogopite - dolomite - calcite assemblage, and no dolomite was found preserved in rocks of this grade. The assemblage tremolite - calcite - phlogopite which is expected to be widespread in the upper tremolite + K-feldspar zone, is observed, but only in rocks in the lower tremolite + K-feldspar zone and not as widespread as would be expected. A second line of evidence that refutes the hypothesis of annealing of the outer Horsethief Creek aureole is that a very steep and hence unlikely temperature gradient of =480 7km would have been required during the time of infiltration. High-grade rocks down to the diopside isograd at =1125 m are interpreted to have experienced infiltration only after reaching temperatures of >520 °C (paragraph 4.4.2). In the annealing scenario the tremolite + K-feldspar-bearing assemblages in rocks in the current (lower) tremolite + K-feldspar zone are presumed to have formed by annealing (i.e., heating to T > 500 °C after the cessation of fluid infiltration) of 77 phlogopite-assemblages which had formed during the earlier fluid infiltration event. The current phlogopite zone in >1250 m is presumed to have formed at about the same time at which the 'old' phlogopite zone annealed, i.e., only after cessation of fluid infiltration. The 'old' phlogopite isograd is presumed to have been approximately where now the tremolite + K-feldspar isograd is positioned (i.e., 1250 m). Hence, with a presumed intermediate XC02 of the infiltrating fluid (e.g., no talc zone) temperatures must have been around 460 °C at 1250 m. Only if rocks in >1250 m had temperatures <460 °C infiltration of that fluid would not have triggered reaction (1). Hence, there must have been a temperature difference of about 60 degrees over 125 m. This very steep temperature gradient seems not very likely. Third, a possible contradiction are the occasionally found higher abundance of phlogopite (up to 8 vol%) in samples where the minerals formed in a preferred orientation (Plates 4a and b) or in veins (Plate 5). This requires that locally fluid was available at temperatures necessary for reaction but that it generally did not infiltrate (most of) the rocks, leading to the widespread development of internally buffered isobarically univariant phlogopite assemblages. If the fluid is not the same as that of the main infiltration event it has to be linked to some late but not retrograde (T >450 °C required, cf.Fig 13b) minor infiltration event(s). This is possible but at this point it seems philosophical to speculate about the likelihood of the coincidence between late infiltration and existence of temperatures necessary for reaction. In summary it can be said that although the process of diachronous heat- and fluid flow is capable of producing a break in petrologic character, with divariant assemblages up-temperature and isobarically univariant assemblages down-temperature from it, this is not likely to be the case in the Horsethief Creek aureole. 78 4.6.2 A model of vertical fluid flow The second model evoked to explain the petrologic break at the tremolite + K-feldspar isograd in the Horsethief Creek aureole implies that the underlying pelites are the main source for the fluids that infiltrated the siliceous dolomites, and that the temperature for the critical dehydration reaction (7) is =525 - 540 °C. Heating to critical temperature(s) and infiltration are basically concurrent events. These 'critical' temperatures are recorded only in the rocks inbound the diopside isograd in =1125 m. Hence, all rocks down-grade from the diopside isograd would not experience vertical infiltration of pelitic fluid. However, there are layers with tremolite in high abundance in >1125 m. Hence, some fluid infiltration must have occurred in these rocks and was likely subhorizontal in nature. As pointed out earlier permeability is likely higher at the lithologic contact between the carbonate and chert layers and fluids in all parts of the aureole may have had a subhorizontal component in fluid flow direction. If fluids enter rocks in the tremolite + K-feldspar zone mostly channeled along these contacts (possibly spreading out along veins), reaction (2) can take place at and near these contacts and around the veins, as far as Si02aq is advective-diffusively transported into the dolostones. This observation is in accordance with field observations. The width of tremolite bearing layers generally decreases towards the lower tremolite + K-feldspar zone (less reactive fluid available), but the abundance of tremolite in these layers is usually high. Further down-temperature, in the phlogopite zone, reaction (1) would not take place if no K-feldspar was available near the fluid conduits, i.e. near the chert layers and/or veins. Relatively pure dolomite is considered relatively impermeable to grain-edge flow (Holness and Graham 1995, Hay and Evans 1988). The feldspar is usually relatively low in abundance and evenly distributed throughout the rock and therefore not readily available for reaction with fluid. Hence, with no pervasive infiltration no pervasive reaction (1) can occur in the phlogopite zone, and 79 without pervasive reaction no pervasive infiltration can occur. This is in agreement with observations in these rocks where mostly isobarically univariant assemblages developed. The hypothesis of initial channeling of fluids and development of pervasive infiltration only by means of reaction-enhanced permeability is further supported by the occurrence of the thick actinolite layers with highly abundant amphibole in the (Mg-) phlogopite zone. Iron endmembers of solid solution phases react at lower temperatures than Mg-endmembers (e.g., Philpotts, Chap. 17, 1990). This explains the possibility of the Fe-amphibole forming reaction to occur in a zone where temperatures for Mg-amphibole reactions are already too low. (No quantification of temperature lowering as a function of iron content was done for this case.) Unlike the phlogopite reaction, where K-feldspar, or, the components potassium and aluminum, are needed, actinolite formation (analogous to the tremolite forming reaction (2)) only requires the addition of silica (and H20) to the Fe-dolomites. Fluids that infiltrate along the chert - carbonate contacts are likely silica saturated. Reaction is not only possible at the chert - carbonate contact but also along the veins and as far as silica - bearing fluid can infiltrate the rock. Potassium and aluminum necessary for reaction (1) in the phlogopite zone in contrast had to be brought into solution in the pelites. It is unlikely that any significant amount of these component escaped filtering-out by reaction in the higher-grade rocks they were in contact with before entering the phlogopite zone. Furthermore unlike the silica content, no "recharge" can occur once the fluid left the 'source-rock'. In summary it can be said that this model of fluids which were predominantly liberated by dehydration reaction in the metapelite and infiltrated the overlying siliceous dolomites, presents a possible explanation for the observations made in siliceous dolomites of the Horsethief Creek aureole. Some plutonic fluid likely was involved especially in the inner aureole. Fluid flow direction had a strong vertical but also some sub-horizontal component; the latter likely the result of changes in lithology. Thin chert layers interlayered with the carbonate rock may have caused preferred flow along their contacts with the carbonate while also keeping fluid silica saturated. 80 However, the model is based on numerous assumptions that were not quantified. More detailed work would be needed to further constrain the hypotheses that are presented here. Furthermore, the situation in detail can be expected to have been even more complex. This is indicated for examples by the occurrence of talc cores within tremolite and tremolite - talc veins in rocks in the middle aureole. This attests to localized infiltration early and late in the thermal evolution of this part of the aureole. 81 CHAPTER 5 SUMMARY AND CONCLUSION Complex interactions between a protolith of variable composition, and heat- and fluid flow led to the development of a =1.5 km wide distinct aureole in potassic siliceous dolomites. Peak metamorphic conditions of =2.2 kb and 650 - 450 °C for the inner and outer aureole respectively are recorded in the rocks. The prograde isograds recorded are: Phi, Tr+Ksp, Di, Di+Phl, Di+Dol and Fo with increasing proximity to the contact. 'Irregularities' in the distribution of these index minerals can be explained in part by varying bulk composition and in part by spatial and temporal heterogeneity in fluid infiltration. Infiltration of a significant amount of a Ff20-bearing fluid is recorded by the widespread occurrence of hydrous minerals often in high abundance (tremolite up to 65 vol%) and the prevalence of isobarically divariant mineral assemblages throughout nearly all of the aureole (<1250 m). The almost complete disappearance of dolomite in the inner and middle aureole attests to the widespread availability of silica for reaction. This requires at least local metasomatism. The interbedded chert layers and nodules may have been a main source for that silica. In addition, the chert - carbonate contacts likely acted as fluid channels along strike. Abundant methodical sets of subvertical veins filled with metamorphic minerals a record of significant sub-vertical fluid flow and attest to the importance of structural controls in facilitating infiltration. The siliceous dolomites in the Horsethief Creek aureole are interpreted to have experienced fluid infiltration with mixed vertical and down-T fluid flow components. The distribution and variance of the mineral assemblages suggest down-T fluid flow, but the isotope signature excludes the idea of the pluton being the (only) source of the infiltrating fluid. The involvement of 82 significant amounts of (isotopically unaltered) meteoric fluids can also be excluded. Other additional sources are necessary to explain the unusually high 5180 values in the inner aureole. The underlying pelitic rocks which dehydrate with the onset of contact metamorphism are the most likely source. The evolving fluid would be buoyant and rise and have an isotope signature that can explain the relatively high 8180 values in the inner aureole. Fluid flow across lithological boundaries was facilitated by a pre-contact metamorphic subvertical joint system. Once silica-saturated aqueous fluid has entered the dolostone, reactions will be triggered that further enhance permeability leading to pervasive infiltration and reaction. Furthermore the overall small width of the aureole and the wide plateau in the maximum temperatures recorded in the aureole hint at possible advective heat flow out of the (siliceous dolomite) system - possible with strong vertical fluid flow. A different fluid regime governed the outermost aureole (phlogopite zone) where only isobarically univariant phlogopite-bearing metamorphic assemblages are observed. Temperatures in this zone were below that necessary to trigger the critical dehydration reaction in the pelites. This resulted in less fluid being available. Fluid that did find its way into the phlogopite zone was presumably brought into the siliceous dolomites by way of sub-horizontal channeling along the chert - carbonate contacts. Unlike the tremolite (+ K-feldspar) - zone rocks in which the addition of only silica and H20 to the dolostones is required for reaction (both of which are available in the fluid), dolostones in the phlogopite zone require the presence (or addition) of potassium and aluminum rather than silica for reaction. These components were not in the fluid (not in significant amounts) and the matrix - K-feldspar is present only in variable and relatively small amounts and relatively evenly distributed in the matrix. Hence, in most rocks it was not available in large quantity for reaction near the fluid channels. Without reaction - enhancement of permeability the bulk of the dolostone was relatively impermeable. 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0.002 100.84 241-5 0.002 54.276 0.010 42.255 0.032 0.011 0.047 0.232 4.037 0.002 100.91 241-6 0.014 54.069 0.003 42.012 0.065 0.002 0.002 0.184 3.790 0.002 100.14 241-7 0.002 54.310 0.003 42.307 0.146 0.002 0.002 0.233 4.060 0.017 101.08 241-8 0.002 54.377 0.003 41.924 0.039 0.005 0.002 0.166 3.977 0.028 100.52 241-9 0.005 54.159 0.008 42.031 0.054 0.004 0.002 0.205 3.943 0.012 100.42 241-10 0.012 54.187 0.005 42.367 0.078 0.011 0.002 0.220 4.193 0.002 101.08 241-1 1 0.010 53.633 0.007 42.037 0.063 0.015 0.002 0.201 4.630 0.051 100.65 241-12 0.002 53.384 0.003 42.220 0.036 0.002 0.002 0.209 5.126 0.010 100.99 251 1-13 0.002 54.428 0.003 41.829 0.040 0.006 0.002 0.179 3.179 0.012 99.68 2511-14 0.002 54.017 0.003 41.915 0.045 0.002 0.013 0.158 3.049 0.005 99.21 251 1-15 0.002 54.435 0.016 41.787 0.018 0.002 0.037 0.158 3.143 0.005 99.60 251 1-16 0.009 .54.649 0.004 41 .999 0.041 0.002 0.019 0.172 2.607 0.017 99.52 251 1-17 0.009 54.942 0.003 41.934 0.014 0.011 0.052 0.147 2.559 0.009 99.68 2511-18 0.002 54.693 0.003 42.183 0.039 0.002 0.019 0.156 2.877 0.002 99.98 241-19 0.003 54.549 0.009 42.018 0.025 0.019 0.002 0.140 2.743 0.002 99.51 811-1 0.022 55.736 0.003 42.168 0.014 0.002 0.039 0.121 2.626 0.002 100.73 811-2 0.002 54.270 0.003 39.443 0.034 0.002 0.002 0.149 2.438 0.007 96.35 811-3 0.002 54.241 0.005 41 .486 0.097 0.013 0.011 0.140 2.554 0.002 98.55 811-4 0.002 54.641 0.003 41.787 0.067 0.002 0.002 0.122 2.595 0.002 99.22 2523-1 0.002 52.097 0.007 39.375 0.073 0.003 0.002 0.176 3.619 0.010 95.36 2523-2 0.006 51.882 0.005 39.919 0.051 0.028 0.034 0.161 3.718 0.037 95.84 2523-3 0.009 42.080 0.004 30.102 0.018 0.054 0.002 0.157 3.324 0.042 75.79 2523-4 0.002 53.637 0.014 41.750 0.044 0.031 0.002 0.187 4.528 0.019 100.21 2523-5 0.002 53.153 0.003 41.743 0.042 0.005 0.013 0.172 4.767 0.002 99.90 2523-6 0.002 53.049 0.009 41.919 0.002 0.002 0.004 0.190 4.656 0.002 99.84 2523-7 0.002 54.920 0.003 41.921 0.058 0.015 0.002 0.156 2.895 0.002 99.97 2523-8 0.002 53.491 0.093 42.153 0.075 0.031 0.002 0.133 3.904 0.024 99.91 2523-9 0.007 53.841 0.009 42.030 0.049 0.004 0.002 0.148 3.71 1 0.002 99.80 Forsterite (cations per 4 oxygens) Label NbCat(Na) NbCat(Mq) NbCat(AI) NbCat(Si) NbCat(Ca) NbCat(Ti) NbCat(Cr) NbCat(Mn) NbCat(Fe) NbCat(Ni) Oxide sum 241-1 0.000 1.917 0.000 0.998 0.003 0.000 0.000 0.004 0.079 0.001 100.99 241-2 0.000 1.922 0.000 0.998 0.001 0.000 0.000 0.004 0.076 0.001 100.98 241-3 0.000 1.916 0.000 1.001 0.002 0.000 0.000 0.004 0.077 0.000 100.61 241-4 0.000 1.918 0.000 1.001 0.001 0.000 0.000 0.003 0.076 0.000 100.84 241-5 0.000 1.914 0.000 1.000 0.001 0.000 0.001 0.005 0.080 0.000 100.91 241-6 0.001 1.918 0.000 1.000 0.002 0.000 0.000 0.004 0.076 0.000 100.14 241-7 0.000 1.912 0.000 0.999 0.004 0.000 0.000 0.005 0.080 0.000 101.08 241-8 0.000 1.925 0.000 0.996 0.001 0.000 0.000 0.003 0.079 0.001 100.52 241-9 0.000 1.918 0.000 0.999 0.001 0.000 0.000 0.004 0.078 0.000 100.42 241-10 0.001 1.908 0.000 1.001 0.002 0.000 0.000 0.004 0.083 0.000 101.08 241-11 0.000 1.901 0.000 1.000 0.002 0.000 0.000 0.004 0.092 0.001 100.65 241-12 0.000 1.889 0.000 1.002 0.001 0.000 0.000 0.004 0.102 0.000 100.99 2511-13 0.000 1.935 0.000 0.998 0.001 0.000 0.000 0.004 0.063 0.000 99.68 2511-14 0.000 1.927 0.000 1.003 0.001 0.000 0.000 0.003 0.061 0.000 99.21 2511-15 0.000 1.937 0.001 0.998 0.001 0.000 0.001 0.003 0.063 0.000 99.60 251 1-16 0.000 1.941 0.000 1.001 0.001 0.000 0.000 0.004 0.052 0.000 99.52 2511-17 0.000 1.948 0.000 0.998 0.000 0.000 0.001 0.003 0.051 0.000 99.68 2511-18 0.000 1.935 0.000 1.001 0.001 0.000 0.000 0.003 0.057 0.000 99.98 2511-19 0.000 1.938 0.000 1.001 0.001 0.000 0.000 0.003 0.055 0.000 99.51 811-2 0.000 1.997 0.000 0.974 0.001 0.000 0.000 0.003 0.050 0.000 96.35 811-3 0.000 1.946 0.000 0.998 0.003 0.000 0.000 0.003 0.051 0.000 98.55 811-4 0.000 1.946 0.000 0.999 0.002 0.000 0.000 0.003 0.052 0.000 99.22 2523-1 0.000 1.945 0.000 0.986 0.002 0.000 0.000 0.004 0.076 0.000 95.36 2523-2 0.000 1.926 0.000 0.994 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core (c) restandardised Mg+Ca=1 T calc: Anovitz & Essene midway (m) Label inclusion (incl) oxide sum NbCat st Mg NbCat st Ca T/K T/C 241-1 100.39 0.052 0.948 807.00 533.85 241-2 100.27 0.055 0.945 820.55 547.40 241-3 99.96 0.065 0.935 854.75 581.60 241-4 100.49 0.059 0.941 834.58 561.43 241-5 100.54 0.053 0.947 810.89 537.74 241-6 100.23 0.057 0.943 827.50 554.35 241-7 r 101.17 241-8 c 100.78 0.051 0.949 803.94 530.79 241-9 (c) 100.04 0.063 0.937 846.90 573.75 241-10 r 100.15 0.063 0.937 847.99 574.84 241-11 r 100.22 0.058 0.942 829.58 556.43 241-12 101.01 241-13 100.14 0.056 0.944 823.97 550.82 241-14 100.71 0.041 0.959 761.55 488.40 241-15 r 98.72 241-16 (c) 99.45 0.087 0.913 915.21 642.06 241-17 -Oi 98.82 241-18 99.01 0.074 0.926 881.65 608.50 241-19 100.28 0.067 0.933 859.80 586.65 241-20 (r) 100.41 0.051 0.949 805.64 532.49 241-21 c 99.88 0.053 0.947 811.07 537.92 241-22 100.59 0.078 0.922 890.74 617.59 241-23 100.31 0.085 0.915 910.31 637.16 241-24 incl? 100.24 0.071 0.929 871.48 598.33 241-25 99.21 0.069 0.931 866.33 593.18 241-26 c 100.38 0.064 0.936 849.66 576.51 241-27 m 98.93 241-28 r 100.17 0.061 0.939 840.95 567.80 241-29 99.89 0.085 0.915 911.71 638.56 241-30 99.09 0.081 0.919 900.98 627.83 241-31 "incl" with doi 98.35 241-32 98.57 241-33 99.33 0.084 0.916 908.68 635.53 241-34 "incl" 99.43 0.085 0.915 909.54 636.39 241-35 99.34 0.083 0.917 905.94 632.79 241-36 100.21 0.081 0.919 899.72 626.57 2432-1 99.63 0.063 0.937 845.40 572.25 2432-2 99.86 0.065 0.935 854.90 581.75 2432-3 100.34 0.041 0.959 761.44 488.29 2432-4 98.77 2432-5 99.16 0.066 0.934 856.00 582.85 2432-6 98.83 2432-10 98.60 2432-11 98.70 2432-12 98.77 2432-13 99.78 0.043 0.957 772.65 499.50 2432-14 beam into a doi 100.39 2432-15 cracks 99.23 2432-16 99.51 0.065 0.935 854.28 581.13 2432-17 99.23 0.056 0.944 823.50 550.35 2432-18 99.22 0.043 0.957 772.30 499.1 5 411-1 99.95 0.064 0.936 850.04 576.89 411-2 99.51 0.062 0.938 844.59 571.44 411-3 100.07 0.059 0.941 833.36 560.21 411-4 101.04 411-5 100.30 0.062 0.938 844.37 571.22 411-6 99.54 0.046 0.954 785.43 512.28 411-7 99.89 0.059 0.941 833.62 560.47 104 rim (r); core (c) midway (m) Label inclusion (incl) 411-8 41 1 -9 411-10 411-11 411-12 411-13 411-14 411-15 411-16 411-17 411-18 411-19 832-1 832-2 832-3 832-4 832-5 832-6 832- 7 832-10 832-1 1 832-12 832-13 832-14 832-16 832-17 832- 1 8 833- 1 833-2 833-3 833-4 833-5 833-6 833-7 833-9 833-8 833- 10 833-1 1 833-12 833-13 833-14 2233-1 2233-2 2233-3 2233-4 2233-5 2233-6 2233-7 2233-8 2233-10 2233-1 1 along vein 2233-12 (Tr-Cal/no doi contact) 2233-13 2233-15 2233-17 2233-19 2233-20 2233-21 " 2233-22 2233-23 2233-24 restandardised Mg+Ca=1 T calc: Anovitz & Essene oxide sum NbCat st Mg NbCat st Ca T/K T/C 100.73 0.060 0.940 836. 63 563. 48 100.53 0.060 0. 940 837. 33 564. 1 8 100.11 0.058 0.942 829. 14 555. 99 100.21 0.042 0. 958 767. 49 494. 34 99.33 0.061 0. 939 841. 03 567. 88 99.92 0.062 0.938 844. 02 570. 87 100.28 0.062 0. 938 842. 23 569 . 08 99.96 0.059 0. 941 834.45 56 1. 3 0 99.56 0.060 0. 940 836 04 562. 89 99.71 0.064 0. 936 848 94 575 79 99.16 0.061 0. 939 841 83 56 8 6 8 100.03 0.064 0.936 849 70 576 55 100.13 0.066 0 934 857 36 5 8 4 2 1 100.30 0.059 0.941 834 88 561 73 100.52 0.060 0 940 837 58 564 43 100.59 0.063 0 937 848 37 575 22 100.82 0.066 0 934 855 50 582 35 100.41 0.056 0 944 824 34 55 1 1 9 101.18 101.12 100.73 0.047 0 953 787 93 5 1 4 78 100.26 0.045 0 955 780 21 507 0 6 101.50 100.00 0.063 0 937 847 26 574 1 1 100.61 0.039 0 961 752 22 479 07 99.12 0.043 0 957 770.05 496 90 101.48 100.82 0.049 0 951 798 24 525 0 9 99.67 0.055 0 945 820 25 547 1 0 100.14 0.053 0 947 810 67 537 .52 99.43 0.055 0 945 817 74 544 .59 100.63 0.053 0 947 812 83 53 9 .68 101.33 100.04 0.050 0 950 800 .98 527 . 8 3 99.50 0.059 0 941 834 .65 56 1 .50 100.19 0.055 0 945 817 .87 544 .72 99.29 0.043 0 957 769 .99 49 6 .84 99.80 0.038 0 962 749 .09 475 .94 100.27 0.052 0 948 808.41 53 5 .26 100.70 0.058 0.942 828.71 555 .56 100.16 0.059 0 941 833 .41 560 .26 99.33 0.057 0 .943 826 .55 553 .40 98.91 99.24 0.056 0 .944 822 .43 549 .28 99.22 0.044 0 .956 775 .10 50 1 .95 98.51 99.56 0.037 0 .963 741 .21 468 .06 98.77 98.85 99.07 0.062 0 .938 842 .97 56 9 .82 99.59 98.08 98.73 99.14 98.54 98.97 99.36 99.33 98.64 98.71 98.99 105 rim (r); core (c) midway (m) Label inclusion (incl) oxide sum 2233-25 98.91 2233-27 98.84 2233-28 98.83 2233-29 98.80 2233-31 98.61 2233-32 98.51 221 1 -1 99.67 221 1-2 100.14 221 1-3 99.21 221 1 -4 100.25 221 1 -5 100.33 221 1 -6 99.75 221 1-7 99.85 221 1-8 99.85 221 1-9 99.99 2211 -1 0 100.00 2211 -1 1 100.05 2211-12 100.47 2211-13 99.56 221 1-14 100.94 2213-37 98.61 2213-38 100.45 2213-39 99.02 2213-40 99.88 2213-41 99.45 2213-42 99.63 2213-43 99.43 2213-44 99.12 2213-45 100.10 2213-46 100.06 2111-1 99.75 2111-2 99.78 2111-3 99.68 2111-4 100.02 2111-5 100.07 2111-6 100.07 2111-7 100.73 2111-8 in vein (cal-tlc-tr-cal) 101.16 2111-9 in vein (cal-tlc-tr-cal) 99.96 2111-10 in vein (cal-tlc-tr-cal) 101.35 2111-12 100.73 2111-11 100.36 2111-13 100.97 2111-14 area beside vein, no Dio 100.24 2111-15 100.66 2111-16 100.61 2111-17 100.41 2111-18 100.36 2111-19 99.78 21 11 -20 100.63 2111-21 99.79 1313-2 vein 101.41 1313-5 vein 100.33 1313-6 100.44 1313-7 Ksp-Cal intergrowth 101.04 1313-8 Kso-Cal interqrowth 100.70 1313-10 100.35 1313-13 veinlet (Cal) 101.40 1313-14 " 101.33 1313-17 vein (Dol-Cal) 101.39 1313-18 101.40 restandardised Mg+Ca=1 T calc: Anovitz & Essene NbCat stMg NbCat st Ca T/K T/C 0 058 0 942 831 38 558 23 0 058 0 942 831 33 558 1 8 0 056 0 944 823 06 549 9 1 0 054 0 946 816 55 543 40 0 059 0 941 831 98 5 5 8 83 0 054 0 946 814 36 54 1 2 1 0 056 0 944 821 50 548 35 0 058 0 942 829 82 5 5 6 6 7 0 061 0 939 840 24 56 7 0 9 0 057 0 943 827 26 554 1 1 0 058 0 942 830 39 557 2 4 0 057 0 943 826 83 5 5 3 6 8 0 056 0 944 823 97 550 8 2 0 048 0 952 792 88 5 1 9 73 0 050 0 950 800 97 527 8 2 0 057 0 943 825 69 552 5 4 0 049 0 951 795 84 52 2 6 9 0 041 0 959 762 98 48 9 8 3 0 041 0 959 761 65 48 8 50 0 042 0 958 765 91 • 492 76 0 017 0 983 585 14 3 1 1 99 0 038 0 962 748 43 475 28 0 021 0 979 633 78 360 63 0 046 0 954 785 85 5 1 2 70 0 037 0 963 743 22 470 07 0 052 0 948 807 78 534 6 3 0 052 0 948 809 97 536 8 2 0 050 0 950 798 91 525 76 0 016 0 984 576 14 302 9 9 0 021 0 979 636 85 363 70 0 052 0 948 806 34 533 1 9 0 053 0 947 810 73 537 5 8 0 054 0 946 815 19 542 04 0 035 0 965 729 94 456 79 0 036 0 964 . 736 96 46 3 8 1 0 046 0 954 784 77 5 1 1 6 2 0 053 0 947 810 06 536 9 1 0 007 0 993 21 1 82 -6 1. 33 0 012 0 988 500 01 22 6 8 6 0 009 0 991 401 66 1 2 8 5 1 0 037 0 963 740 57 46 7 42 106 dolomite free sample rim (r); core (c) midway (m) restandardised Mg+Ca=1 T calc: Anovitz & Essene inclusion (incl) Label oxide sum NbCat st Mg NbCat st Ca T/K T/C 251-1 100 45 0.049 0 951 796 20 523.05 251-2 101 10 251-3 101 43 251-4 100 35 0 047 0 953 789 88 516.73 251-5 100 98 0 045 0 955 778 93 505.78 251-6 100 44 0 050 0 950 799 62 526.47 251-7 101 31 251-8 102 65 251-9 100 81 0 046 0 954 783 91 510.76 251 -1 0 101 35 251 -1 1 101 03 251-12 101 11 251-13 101 03 251-14 100 42 0 046 0 954 784 99 511.84 251 -1 5 101 20 251-16 100 84 0 046 0 954 783 26 5 10.11 251-17 101 73 251-18 101 33 251-19 101 03 107 APPENDIX B XRF DATA 102 Whole rock analyses - Method and Raw Data Rock chips of about 4.5 x 2.5 x 2.0 cm were cut, crushed and ground to powder in a ring-mill. The powders were used for whole rock analyses as well as stable isotope analyses. Powders were thoroughly mixed to make sure that grains were homogeneously distributed. The powders were sent to the Geochemical Laboratories at McGill University, Montreal where they were analyzed by X-ray fluorescence. Estimates of mineral abundance in the analyzed samples are given below. They are listed in decreasing grade. Table Bl: Rough volume percentage of carbonates versus hydrous versus non-hydrous silicates in rocks analyzed carbonates hydrous silicates non-hydrous silicates calcite and/or tremolite phlogopite Forsterite/Diopside/ dolomite K-feldspar 95 Ho 2/5-1 45% 55% 96 Ho 24/2-1 55% 10% 35% 96 Ho 23/2-1 30% 5% 65% 96 Ho 23/2-2 25% 5% 70% 96 Ho 23/2-3 35% 5% 60% 96 Ho 24/3-2 40% 15% 45% 96 Ho 24/1-4 70% 10% 20% 95 Ho 6/2-1 40% 5% 55% 96 Ho 25/1-2 30% 15% 55% 96 Ho 25/2-2 50% 5% 45% 95 Ho 3/1-1 45% 40% 15% 96 Ho 3/1-2 40% 10% 50% 95 Ho 3/3-1 45% 40% 15% 96 Ho 27/3-2 60% 35% 5% 95 Ho 19/3-1 55% 5% 50% 96 Ho 10/2-5 25% 75% 96 Ho 22/4-1 65% 25% 10% 96 Ho 22/3-1 50% 50% 96 Ho 22/1-2 65% 30% 5% . 95 Ho 12/1-4 55% 40% 5% 95 Ho 12/2-1 20% 80% 96 Ho 21/3-2 30% 70% 96 Ho 28/1-2 mostly dolomite with little Ksp -Qtz -Phi 95 Ho 13/2-2 mostly dolomite with little Ksp -Qtz -Phi 95 Ho 13/2-1 mostly dolomite with little Ksp -Qtz -Phi 96 Ho 18/1-1 mostly dolomite with little Ksp -Qtz -Phi 109 03 H fi N! 3 u o a CO O c Cu O o r o 03 u o o (U c O c < c O H c O f 03 C/3 oo T-H IT) O T-H co CN o r-H in CN CO vo 00 CO CN r-H Tf ON Tt Tt Tt r-H CO CO oo o O oo vo in ON q in Tt in in CN oo r-H Tt o t> o r-; CO oq ON t-- d d ON ON ON t> od ON ON od od ON ON ON t> d ON d 00 d Os ON rS ON ON o o ON ON ON ON ON ON ON ON ON ON ON ON ON o ON o ON o OS ON Os ON i—i i—i r-H i—i r-H Tt Tt CN r- oo vo CN Tt oo o r-H t> m t> CO ON Tt o r-H oo co oo CN O o CO (N co <N ON CO '—' r- Tt CO CO o OO o CO ON CO VO CO in in CN m Tt d CO d oo CO 00 in T-H rS d os vd d d d ON VD CN Tt CO co CO CN d ON ON <N •-H --H i-H CN CN T-H r-H ^H CN CN CN CN i-H CN --H "—1 r-H CO Tt CO CO VO t> CN ON VO in ON t> co r> Tt m O Tt co 00 T-H T-H 00 co CN co CO Tt VO l-H Tf CN i-H •-H CN VO CO CN i-H Tt VD Tt CN co "-H m CN c-- CO CN co vo CN r- t> r- r- Tt CN T-H o vo vo r-~ VO r» T-H r-H ri H CO ON o r-H CN V CN *-H r-H V 00 VO VO ON CN T-H VO Tt VO co CN vo o r-H 00 co m Tt ON Tt ON t> ON o Tt ON CO Tt CO T-H CO CO Tt CO Tt CO CO in Tt Tt VO CN CO CN CN Tt CN CN CO CO CN CN to Tf oo m >n m ON oo ON CN ON CO oo Tt Tt O in Tt O VO ON r- oo o m CO r> oo Tt o 00 co ON r-H r~ m CO m -r^ CN Tt co CN m CO CN oo Tt t> <N Tt Tt m CN vo Tt in T-H CO Tt CO T-H \/ CN oo o r- co < .V CN CN ON t> ON c-» t-- r- in CN m r- oo vo o\ oo r-H m m vo CN r-H VO Tt VO r-H co 'co co CN CO co co CN CN m <N Tt CO CO CN Tt CN Tf co CO m CN Tt CO T-H O o o o O o o o O o o O O O o O o o O o o o O o o o o d d d d d d d d d d d d d d d d d d d d d d d d d 'TT t-- CO Tt T-H VO o ON o ON m VD ON Tt CN VD in in CN in m VO c-» CN ON 00 q Tt oo i-H ON in Tt Tf Tt 1—1 r-H in CO in o Tt CN oo CO ON d fN d d d CN d d d d CO CN d d d 1 (N co m VO Tt vo oo _ o 00 CO in o Tt m o Tt vo co ON m ON m ON o o i-H o o o T-H T-H l-H O O o o CO o CN r-H CN o Tf r-H o r-H r-H d d d d d d d d d d d d d d d d d d d d d d d d d d in CN CO in o T-H m o CN o r- Os ON r-H r- CN VO oo in CN CN ON CN in Tf vq q vq ON ON r-H r-H co Tt i> CN 00 CN in r-H CO Tt Tt r-' in ON r-- ON 00 r-- oo vd vd 00 t> i> d d CN t> CN co r-i r-H vd od r~ CN CN CN CN CN CN CN CN CN CN CN CN CN CN CN CO CN CN CN CN CN CN CN CN CN CO ON vo r- T-H T-H CN ON CN in co ON CN co r- r-H ON O O VD m VD co co r- vo oo T-H ON oo oo CO r- 1—J T-H CN i—l t> CO vq m in oo ON in CN CO r» r-H co c-oo ON od od VD ON T1 ON 00 l> in 00 Tt rS in ON Tt od vd in od CN Tt 00 ON oo CN CN co VO m O ON o C^ _l Tt co CN 00 Tt ,_ ,_ vo CO o r- co Tf CN r-H co VO vo in Tt Tf m CN Tt m VD r-H CN VD in Tt Tt in CO Tt Tt m co CO Tt in Tt O o o o o o CN o o o r-H i—i o o o o o o o o o O O O o o d i d d d d d d d d d d d d d d d d d d d d d d d d d iin CO o 00 Tt o Tt in VD _ vo ON Tt 00 o VO 00 m Tt VO CN CN Tt r- in ON oo CN ON q CN r-H Tt ON r-H ON q oo 00 00 o t> ON ON 00 d i r-H d d d T-H d H CN d i—1 d d d d d d d d d d CO Tt m CN CO CN O VO ON CN r-H o ON co co Tt co ON in m m 'in CN Tf Tt in CN vq 00 ON 00 co i—j ON ON VD i—j OS Tt CN vq o\ r-H Tt CN co •-H i—1 T-H CN i-H i-H CN d CN* CN CO r-H CN l-H CN r-H CN co d in CN r-H d l-H vo T-H 00 Tf co o in CN CN t> ON T—H T_( O oo Tt O Tf in co CO co CN vo CN VO >n VD ON VO vo o Tt CN Os CN VO CN ON Tf in ON o ON m CN in o r-H o o O O o o 1—1 O O o o o o o CN o o O o d d d d d d d d d d d d d d d d d d d d d d d d d d |VO 00 co co t> VO Tt in m VO Tt Tt in oo r- o CN CN in Tt ON —JL. oo ON oo CO oo p t> *-H ON Tt r-H Tt m m 00 00 vq ON I-; CN oo CN co o V") Cs ON ON d CO vd 00 od Tt od d r-H T-H ON d ON vd r-H co Tt r-H CN r-H in T-H CO T-H Tf Tt CO CN CN CN Tt Tt co CN co CN Tt CN Tt CN Tf Tt Tt CN r-H 00 00 T-H i—i CN co CN Tt CN CN CN r-H in r-H ri CN Tt r-H CN CN CN r-H r-H 1 <N CN CN CO T-H i r-H CN 1 CN CO CO CN Tf co r-H CN co r-H CN CN T-H I/-) CN r-H co Tt CO CO CO Tt Tt in in - r- ON cB CN CN CN CN CN i—( 00 CO CO 00 CN CN CN CN CN CN CN vo* CN CN co CO co CN i—1 CN CN CN r-H CN CN T-H o o CN CD r-H CN l> m" co^ in CN <n in m ON o" CO o" CO CD* CN r-H in co o" vo PH OH 1) ri O O .O : p : o : !P 110 Ill Zinc 0.5 1 Distance (km) Barium 3000 2500 2000 ? a. B 1500 ra m 1000 500 0 1——m-0.5 1 Distance (km) Copper Vanadium 20 ? 15 a a. > 10 • • 0.5 1 Distance (km) 0.5 1 Distance (km) Figure Bl: Plots of distribution of trace elements Zn, Ba, Cu and V which show a generally homogeneous distribution. A noteworthy exception are two Barium values at =1000m distance which are three times the average values. These results coincide with the petrographic observation of Ba-rich feldspars in the same area. Source of the barium might be (late) magmatic fluids from underlying igneous bodies of unknown form or extent. Such a presence of igneous bodies below the contact-metamorphosed rocks may also explain the unusual =550°C plateau in the temperature profile recorded in the Horsethief Creek siliceous dolomites. 112 APPENDIX C STABLE ISOTOPE DATA IIB Cl Stable Isotopes - Methods - Data - Interpretation of 513C signature Powders (whole rock carbonate - see appendix B) were sent out to the Stable Isotope Laboratory at Texas A&M university. Analyses are done with a Finnigan Kiel II analysis system using a Finnigan MAT 251 isotope mass spectrometer. The isotopes were measured against standard NBS-19 and converted into Vienna-PDB (VPDB) values by using 8180 NBS imPDB = -2.2%c and 513CNBS19/VPDB = +1.95%o. The oxygen data was subsequently converted into VSMOW scale by the relation: 8180 =8'80 +8'80 + 10~3 8180 8lsO u '"a/VSMOW u ^vJb T u '"'b/VSMOW u wa/b u wb/VSMOW as recommended by Coplen (1994). Table Cl-1: Stable isotope analyses data, samples sorted in increasing grade. sample # 8180 8,3C 8,80 8,3C (NBS 19) (NBS 19) (VSMOW) (VPDB) 2/5-1 -12.39 -3.45 18.17 -5.40 23/2-3 -11.45 -4.73 19.13 -6.68 23/2-2 -12.02 -5.08 18.55 -7.03 23/2-1 -11.43 -4.53 19.16 -6.48 24/2-1 -9.96 -0.91 20.67 -2.86 24/3-2 -9.69 -1.33 20.95 -3.28 24/1-4 -17.04 -4.66 13.38 -6.61 25/1-2 -13.16 -4.14 17.38 -6.09 25/2-2 -12.42 -5.02 18.14 -6.97 6/2-1 -13.60 -2.33 16.92 -4.28 3/1-2 -13.92 -5.78 16.59 -7.73 3/1-1 -13.19 -5.74 17.34 -7.69 3/3-1 -13.52 -4.89 17.01 -6.84 10/2-5 -12.40 -2.61 18.16 -4.56 27/3-2 -11.93 -2.33 18.64 -4.28 19/3-1 -12.27 -3.62 18.29 -5.57 22/4-1 -12.10 -1.55 18.47 -3.51 22/3-1 -10.07 -2.26 20.55 -4.21 22/1-2 -10.64 -1.93 19.97 -3.88 12/1-4 -10.19 -1.25 20.43 -3.20 12/2-1 -12.90 -3.45 17.64 -5.40 21/3-2 -10.55 -1.23 20.06 -3.18 13/2-2 -8.32 -0.18 22.35 -2.13 13/2-1 -8.17 0.18 22.51 -1.77 28/1-2 -8.59 0.66 22.07 -1.29 18/1-1 -8.43 0.21 22.24 -1.74 114 5,80 (VSMOW) Figure Cl: This plot shows that not only 8180 but also 813C values were considerably lowered during contact metamorphism. Possible explanations for the 813C depletion of ~6%o from ~-2%o to -8%o are Rayleigh fractionation with or without infiltration of an igneous fluid (813C ~ -4 to -6) (see review by Valley 1986). Although no quantification of the amount of carbon remaining in the system (F) was done, a rough estimate of these values is considered to be F=0.3-0.5 (for 25-55vol% calcite remaining in the system). For these values of F, Rayleigh fractionation seems sufficient to explain the lowering in 813C values. The outlier could reflect either interaction with an igneous fluid which would simultaneously lower both isotope values or it could reflect a lowering of the carbon isotope signature by Rayleigh fractionation analogous to the other samples and subsequent lowering of the oxygen isotope values by low-temperature alteration with involvement of surface waters (see text for more details). The physical meaning of Rayleigh fractionation is one of a fairly open system where evolving fluids are basically constantly removed from the system. This agrees with petrographic observations of closely spaced veins that act as conduits. The linear correlation between 8lsO and 813C is in contrast to the basically missing correlation of oxygen or carbon with distance (cf. Figure 13 and Table Cl-1). The reason for this is not clear. 115 C2 Silica Weight Percent (whole rock) versus 8180 - Evidence for Silica Metasomatism 50 45 T 40 35 u !r 30 w 25 * > CN 20 O (0 15 10 + 5 + • Si02 wt % • Si02 wt % in Phi zone rocks —f-14 •+- •+- —I 24 12 16 18 5180( VSMOW) 20 22 Figure C2-1: This Si02 weight percent versus whole rock carbonate oxygen isotope plot shows that there is basically no correlation between the oxygen data and the whole rock XRF data except for a step function at 8180 -22. This step coincides with the step in silica content over distance at ~1250m (see figure 7) which also is the location of the petrographic break. Hence, there is no correlation between silica content and oxygen isotope data in rocks that record pervasive infiltration (<1250m) but because all of these rocks have significantly higher silica contents than the average phlogopite zone rock (>1250m) it shows that the fluid influx is somehow coupled with silica increase. The random correlation within infiltrated rocks reflects the complex infiltration pattern suggested for the Horsethief Creek aureole. 116 APPENDIX D MINERAL FORMULAS AND MOLAR VOLUMES 117 Mineral formulas and molar at standard state lbar, 298.15 K (after Berman, 1988): dolomite CaMg(C03)2 64.32 cm3mol" calcite CaC03 36.90 cm3mof tremolite Ca2Mg5Si8022(OH)2 272.68 cm3mol" diopside CaMgSi206 66.20 cm3mor] talc Mg3Si4O10(OH)2 136.10 cm3mof phlogopite KMg3(AlSi3O10)(OH)2 149.77 cm3mol" K-feldspar KAlSi3Os 108.69 cm3mol" forsterite MgSi03 43.66 cm3mol"1 periclase MgO 11.25 cm'mor' brucite Mg(OH)2 24.68 cm3mol"1 quartz SiQ2 22.69 cm3mof 118 APPENDIX E SAMPLE NUMBER CONVERSION II? Table El: List of original and modified sample num! ers. inclu line location distances from contact originally used modified distance from originally used modified distance from sample numbeis sample numbeis pluton (m) sample numbeis sample numbeis pluton (m) 95Hol8-2 18/2-1 1520 96H 0 27-3/2 27/3-2 610 95Ho 18-1/1 18/1-1 1440 95Ho34 3/4-1 540 95H 0 18-1/2 18/1-2 1415 95Ho3-3 3/3-1 540 95H 0 13-2/1 13/2-1 1355 95H o9-l/l 9/1-1 530 95H 0 13-2/2 13/2-2 1355 95H 0 9-1/2 9/1-2 530 96H 0 28-2/1 13/2-3 1355 96H 0 23-3/1 23/3-1 505 96H 0 28-2/2 13/24 1355 96H 0 23-3/2 23/3-2 505 96H 028-1/1 28/1-1 1350 95Ho3-l 3/1-1 365 96H 0 28-1/2 28/1-2 1350 95H o8-3/l 8/3-1 360 96H 028-1/3 28/1-3 1350 95H 0 8-3/2 8/3-2 360 95Ho 13-1/2 28/14 1350 95H 0 8-3/3 8/3-3A 360 95H 0 13-1/3 28/1-5 1350 dito 8/3-3B 360 95Ho 15-1/3 15/1-3 1250 95Ho84/l 8/4-1A 320 95Ho 15-1/4 15/14 1250 dito 8/4-1B 320 95Hol5-l/2a 15/1-5 1250 95Ho8-2 8/2-1 305 95Hol5-l/2b 15/1-6 1250 95H 08-I/I 8/1-1 270 95Hol5-l/2c 15/1-7 1250 96H 0 25-2/1 25/2-1 260 95H 0 12-3/1 12/3-1 1240 96H 0 25-2/2 25/2-2 260 95H 0 12-3/2 12/3-2 1240 96H 0 25-2/3 25/2-3A 260 96Ho21-3/l 21/3-1 1150 dito 25/2-3B 260 96H 0 21-3/2 21/3-2A 1150 96H 0 25-2/4 25/24 260 dito 21/3-2B 1150 96H 0 25-1/1 25/1-1 240 95H 0 124/1 12/4-1 1125 96Ho25-l/2 25/1-2 240 95Hol24/2b 12/4-2 1125 95H 0 6-3/1 6/3-1 220 95Hol2+4'l 12/4-3 1125 95Ho6-2 6/2-1 220 96Ho21-2/l 21/2-1 1075 95H 0 4-1/1 4/1-1A 215 96H 021-2/2 21/2-2A 1075 dito 4/1-1B 215 dito 21/2-2B 1075 95H 0 4-1/2 4/1-2 215 96Ho21-l/l 21/1-1 1010 95H 04-1/3 4/1-3A 215 96Ho21-l/2 21/1-2 1010 dito 4/1-3B 215 96Ho21-l/3 21/1-3 1010 96H 0 23-1/1 23/1-1 165 95H 0 12-2/1 12/2-1 990 96Ho23-l/2 23/1-2 165 95H 0 12-1/1 12/1-1 920 96H 0 23-1/4 23/14 165 95Hol2-l/3a 12/1-3 920 96H 0 24-1/1 24/1-1 160 95Ho 12-1/4 12/14 920 96H 0 24-1/2 24/1-2 160 95Ho 12-1/5 12/1-5 920 96Ho24-l/3 24/1-3 160 96H 0 22-1/1 22/1-1 865 96Ho24-l/4 24/14 160 96H 0 22-1/2 22/1-2 865 95Ho7-l 7/1-1A 140 96Ho22-l/3 22/1-3 A 865 dito 7/1-IB 140 dito 22/1-3B 865 96H 0 24-3/1 24/3-1 115 96H 0 22-2/1 22/2-1 815 96H 0 24-3/2 24/3-2 115 96H 0 22-2/2 22/2-2 815 96H 0 24-3/3 24/3-3 115 96Ho22-2/3 22/2-3 815 95Ho4-3 4/3-1 115 96H 0 22-3/1 22/3-1 800 95Ho4-3b 4/3-2 115 96H 0 22-3/2 22/3-2A 800 96H 0 23-2/1 23/2-1 110 dito 22/3-2B 800 96H 0 23-2/2 23/2-2 110 96Ho22-3/3 22/3-3 800 96Ho23-2/3 23/2-3 110 96H 0 22-3/4 22/34 800 96H 0 24-2/1 24/2-1 100 95H o 17-1 17/1-1 760 96H 0 24-2/2 24/2-2 100 96H 0 224/1 22/4-1 700 96Ho2'-5/l 2/5-1A 25 96H 0 224/2 22/4-2 700 dito 2/5-1B 25 95Holl-l 11/1-1 680 96Ho2'-5/2 2/5-2 25 95Ho 10-2/1 10/2-1 670 96Ho2'-5/3 2/5-3 25 95Ho 10-2/2 10/2-2 670 96Ho2'-5/4 2/54 25 96HolO+l/l 10/2-3 670 96Ho2'-5/5 2/5-5 25 96HolO+l/2 10/24 670 96Ho2-5 2/5-6 25 96HolO+ 1/3 10/2-5 670 96Ho2'4/l 2/4-1 20 95H 0 19-3/1 19/3-1 655 96Ho24 2/4-2 20 95H 0 19-3/2 19/3-2 655 96Ho24+ 2/4-3A 20 95Hol4-2/lb 14/2-1 615 dito 2/4-3B 20 96H 0 27-3/1 27/3-1 610 96Ho2-2 2/2-1 20 120 


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