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Petrophysical and physicochemical controlling parameters on stable isotope depletion patterns in carbonate… Lepore, William Adamas 2013

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PETROPHYSICAL AND PHYSICOCHEMICAL CONTROLLING PARAMETERS ON STA-BLE ISOTOPE DEPLETION PATTERNS IN CARBONATE ROCKS FROM AURIFEROUS HYDROTHERMAL FLUID INFILTRATION AT THE LONG CANYON SEDIMENT-HOSTED GOLD DEPOSIT: NE NEVADAbyWILLIAM ADAMAS LEPOREB.Sc., University of British Columbia, 2006A THESIS SUBMITTED IN PARTIAL FULFILLMENT OFTHE REQUIREMENTS FOR THE DEGREE OFMASTER OF SCIENCEinTHE FACULTY OF GRADUATE STUDIES(Geological Sciences)THE UNIVERSITY OF BRITISH COLUMBIA(Vancouver)August 2013© William Adamas Lepore, 2013iiABSTRACTThe Long Canyon deposit is a sediment-hosted gold deposit located in northeastern Nevada, over 150 km east of the Carlin Trend. The deposit geology, geochemistry and mineralogy suggest that it is a Carlin-type deposit. This project tests the extents of auriferous hydrothermal fluid infiltration through carbonate rocks at Long Canyon using patterns of 18O/16O depletion to define the limits of fluid-rock interaction. Patterns of isotope depletion are used to assess the structural and lithologic controls on fluid flow, and the lateral extent of fluid rock interaction beyond the limits of trace element halos genetically associated with gold miner-alization. Carbon and oxygen isotope ratios were measured using the Mineral Deposit Research Unit’s Mineral Isotope Analyzer (MIA) at the University of British Columbia.Gold mineralization at Long Canyon occurs predominantly within deeply oxidized, limestone-hosted solu-tion breccias, located within necklines of an 80 metre thick boudinaged dolomite unit. Alteration associated with gold mineralization comprises decalcification of limestones>>dolomite, fine-grained pyrite growth within Fe-rich host rocks, argillization and silicification. Clay minerals include dickite within the core of the system, flanked by illite and kaolinite. Assuming typical Carlin-type fluid temperatures of 180° to 240°C, δ18O values of ore fluids are calculated at ≤1.6 to 6.3‰. Calculated δ18OH20 that deposited dickite within the core of the mineralized system was calculated at 2.3‰ for a 200°C fluid.In this study several scales of sampling were tested including: contiguous drill assay pulps over cross sec-tions; surface sampling along traverses; closely spaced hand sample coverage down drillholes; and micro-drilled hand samples. Presented here are the results of over 2,800 unique 18O/16O carbonate sample compo-sitions from across the deposit. Results indicate stable isotope depletion mapping around carbonate-hosted ore bodies with sufficient sample density can provide far field vectors to fluid flow paths beyond what traditional lithogeochemistry alone may show. Additionally, it is evident from the pattern of oxygen iso-tope depletion that hydrothermal fluid flow responsible for mineralization was largely constrained to brittle damage zones within boudin necks, or incipient boudin necks, within massive dolomite units. There was minimal lateral flow within or vertically across stratigraphic units without pre-existing structural damage. iiiPREFACEThis research project was designed by the author W.A. Lepore and Dr. K.A. Hickey, in consultation with Drs. S.L.L Barker and M.T. Smith. The research program benefited from and built upon understanding of light stable isotope fluid:rock exchange in Carlin gold systems developed during previous studies designed by K.A. Hickey and the Carlin Gold Systems research team. The research, sampling and data collection were carried out by the author W.A. Lepore, in consultation with Drs. K.A. Hickey and S.L.L Barker. Parts of Chapters 3 and 4 are intended to be reorganized and submitted for publication to a scientific journal. Fig-ure 4.16 was reformatted and included in an Economic Geology Express Letter (Barker, Dipple et al. 2013). ivTABLE OF CONTENTSABSTRACT   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   . iiPREFACE   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   . iiiTABLE OF CONTENTS   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   . ivLIST OF TABLES  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  . viiiLIST OF FIGURES   .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  . ixACKNOWLEDGEMENTS  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  . xxiDEDICATION  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  . xxiiChapter 1: Thesis Introduction and Context .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .   .  11.1 Introduction  . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11.2 Sediment-Hosted “Carlin-type” Gold Deposit Characteristics  . . . . . . . . . . . . . . . . . . 21.3 Long Canyon Deposit Overview  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  41.4 Thesis Organization . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5Chapter 2: Controls on Stable Isotope Depletion Patterns in Carbonate Rocks   .  .  .  .  .  .  .  .  .  .  . 62.1 Introduction  . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 62.2 Petrophysical Properties of Carbonate Sediments   .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  62.3 Hydrothermal Fluid Flow in Carbonate Rocks  . . . . . . . . . . . . . . . . . . . . . . . . . . 82.4 Dissolution-Precipitation Kinetics of Carbonate Rocks   .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .102.5 Light Stable Isotopes in Carbonates and Natural Waters  . . . . . . . . . . . . . . . . . . . . .152.6 Isotopic Fluid Transport Mechanisms and Modeling in Carbonate Rocks  . . . . . . . . . . . .18Chapter 3: Geology of the Long Canyon Deposit   .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .233.1 Introduction  . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .233.2 Previous Work and Acknowledgment of Source of Concepts   .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .233.3 Sampling and Research Methods  . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .233.4 Regional Geologic Setting and History . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .243.5 Deposit Scale Geology  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .253.6 Stratigraphic Setting of the Long Canyon Deposit  . . . . . . . . . . . . . . . . . . . . . . . .263.6.1 Upper Cambrian Notch Peak Formation   .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .263.6.2 Lower Ordovician Pogonip Group   .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .283.6.3 Origin of Dolomitic Units and Sedimentary Facies Interpretation  .  .  .  .  .  .  .  .  .  .  .  .30v3.6.3.1 Dolomitization Models   .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .303.6.3.2 Long Canyon Dolomite Origin Interpretation . . . . . . . . . . . . . . . . . .313.7 Igneous Rocks of the Long Canyon Deposit  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .333.7.1 Mafic Intrusives  . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .333.7.2 Felsic Intrusives and Volcanics  . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .343.8 Structural Geology of the Long Canyon Deposit . . . . . . . . . . . . . . . . . . . . . . . . .343.8.1 Synmetamorphic Fabrics   .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .343.8.2 Boudinage in the Notch Peak Dolomite . . . . . . . . . . . . . . . . . . . . . . . . . .353.8.3 Brittle Brecciation and Faulting in the Notch Peak Dolomite . . . . . . . . . . . . . . .363.8.4 Genesis of Boudinage in Notch Peak Dolomite . . . . . . . . . . . . . . . . . . . . . .373.9 Alteration and Mineralization of the Long Canyon Deposit  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .393.9.1 Carbonate Dissolution and Brecciation  . . . . . . . . . . . . . . . . . . . . . . . . . .393.9.1.1 Dissolution Breccia   .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .393.9.1.2 Interpretation of Ore Fluid Composition From Relative Dissolution of Calcite and Dolomite   .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .423.9.2 Quartz Grain Growth and Silicification within Ore Zones   .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .423.9.2.1 Non-Ore-Stage Quartz . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .423.9.2.2 Ore-Stage Quartz.  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .433.9.3 Calcite Veining and Breccia Cementation . . . . . . . . . . . . . . . . . . . . . . . . .443.9.3.1 Coarse Calcite-Cemented Breccia . . . . . . . . . . . . . . . . . . . . . . . .443.9.3.2 Calcite Veining . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .453.9.4 Gold Mineralization . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .463.9.5 Clay Alteration and Fluid Sources . . . . . . . . . . . . . . . . . . . . . . . . . . . . .483.9.5.1 Methodology of Clay Studies  . . . . . . . . . . . . . . . . . . . . . . . . . .483.9.5.2 Results of Clay Studies   .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .483.9.5.3 Interpretation of Clay Studies  . . . . . . . . . . . . . . . . . . . . . . . . . .493.9.6 Oxidation of the Long Canyon Deposit . . . . . . . . . . . . . . . . . . . . . . . . . .503.10 Lithogeochemistry of the Long Canyon Deposit  . . . . . . . . . . . . . . . . . . . . . . . . .513.10.1 Reactive Iron Control on Gold Mineralization  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .513.10.2 Gold-Silver Ratio of Mineralization . . . . . . . . . . . . . . . . . . . . . . . . . . . .523.10.3 Ore-Stage Trace Element Zonation  . . . . . . . . . . . . . . . . . . . . . . . . . . . .533.11 Interpretation and Discussion  . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .543.11.1 Fluid Flow Path Development in the Long Canyon Deposit  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .543.11.2 Comparison of the Long Canyon Deposit to Carlin-Type Deposits in Nevada . . . . . .56Chapter 4: Mapping Hydrothermal Fluid Flow With High Density Sampling of Oxygen and Car-bon Isotope Ratios in Carbonate Rocks of the Long Canyon Deposit   .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .97vi4.1 Introduction  . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .974.2 Previous Studies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .984.3 Project Design . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1004.3.2 Light Stable Isotope Ratio Mass Spectrometry  . . . . . . . . . . . . . . . . . . . . . 1014.3.3 Off-Axis Integrated Cavity Output Spectroscopy . . . . . . . . . . . . . . . . . . . . 1014.4 Sample Suites and Sampling Methodology . . . . . . . . . . . . . . . . . . . . . . . . . . . 1014.4.1 Phase I Micro-Drilling . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1014.4.2 Phase II Drill Assay Rock Pulps . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1024.4.3 Phase II Drill Core Hand Samples . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1024.4.4 Phase II Micro-Drilling  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  . 1034.4.5 Surface Hand Samples . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1034.5 Results . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1044.5.1 δ18O and δ13C Values of Carbonate Rocks at the Long Canyon Deposit   .  .  .  .  .  .  .  . 1044.5.2 Drill Assay Rock Pulp δ18O Values and Distribution at the Long Canyon Deposit . . . 1054.5.3 Hand Sample Rock Pulp δ18O and Geochemical Values and Distributions . . . . . . . 1074.5.4 Micro-Drilled δ18O and δ13C Values of Long Canyon Carbonate Rocks  .  .  .  .  .  .  .  . 1094.5.5 Surface Traverse δ18O Distribution . . . . . . . . . . . . . . . . . . . . . . . . . . . 1134.5.6 Lithologic Control on Fluid Flow and Oxygen Isotope Depletion Patterns . . . . . . . 1144.5.7 Mechanism of Isotopic Exchange . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1194.5.8 Ore Fluid Isotopic Composition at Long Canyon  . . . . . . . . . . . . . . . . . . . . 1204.5.9 Carbon Isotope Alteration and its Utility as a Tracer for Hydrothermal Fluid Flow  .  . 1204.5.10 Calcite Vein and Breccia Cement Interpretation . . . . . . . . . . . . . . . . . . . . . 1204.5.11 Geometry of Fluid Flow Paths and Oxygen Isotope Depletion Halos Surrounding Ore-bodies 1214.5.12 Effect of Sampling Scale, Location and Density on Apparent Oxygen Isotope Depletion  . 1234.5.13 Distribution of Oxygen Isotope Depletion Compared to Trace Element Metasomatism 1264.5.14 Mechanism for Oxygen Isotope Depletion Beyond Trace Element Metasomatism . . . 1264.5.15 Isotope Mapping on Surface and Implications for Exploration . . . . . . . . . . . . . 1274.6 Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1304.6.1 Transport Theory, Fluid Flow and 18O depletion at Long Canyon   .  .  .  .  .  .  .  .  .  .  . 1304.6.2 Application of Stable Isotope Studies to Exploration for Carlin-type Deposits . . . . . 1364.6.2.1 Exploration Drilling  . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1364.6.2.2 Ore Deposit Definition Drilling  . . . . . . . . . . . . . . . . . . . . . . . . 1364.6.2.3 Sampling Considerations   .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  . 1374.6.2.4 Trace Element Geochemistry Versus Oxygen Isotopes in Exploration  . . . . 1374.6.3 Assessment of Mineral Potential from Oxygen Isotope Depletion Patterns   .  .  .  .  .  . 138viiChapter 5: Synthesis, Summary and Conclusions  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  . 1735.1 Summary   .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  . 1735.2 Synthesis   .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  . 1735.2.1 Evolution of Framework Fluid Flow Paths at Long Canyon  .  .  .  .  .  .  .  .  .  .  .  .  .  . 1735.2.2 Evolution of Ore Fluid at Long Canyon . . . . . . . . . . . . . . . . . . . . . . . . . 1745.2.3 Controls on Fluid Flow at Long Canyon   .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  . 1745.3 Guidelines for Use of 18O-Depletion Studies in Mineral Exploration for Carlin-type Deposits 1755.4 Recommendations for Further Work  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  . 176REFERENCES  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  . 177APPENDIX A: Analytical Techniques and Methodologies   .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  . 183A.1 Oxygen and Carbon Isotope Analysis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 183A.1.1 Delta Mass Spectrometer   .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  . 183A.1.2 LGR Desktop Analyzer  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  . 183A.2 Clay Analysis  . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 184A.2.1 Clay Separation  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  . 184A.2.2 X-Ray Diffraction (XRD) Analysis  . . . . . . . . . . . . . . . . . . . . . . . . . . . 184A.2.3 Near and Shortwave Infrared Analysis   .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  . 184A.2.4 Clay Isotope Analysis  . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 184A.3 Scanning Electron Microscope (SEM) Analysis  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  . 185APPENDIX B: GEOCHEMICAL AND ISOTOPIC DATA TABLES  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  . 186B.1 Terraspec Results  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  . 186B.2 X-Ray Diffraction Analysis Results . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 187B.3 Data Tables   .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  . 189viiiLIST OF TABLESTable 3.1: Calculated fluid δD and δ18O values for H20 in equilibrium with silicate minerals in clay separates collected in this study from 5 hand samples down drill hole LC556C. Whole rock mineral identification by TerraspecTM analytical spectral device (ASD). Clay separate mineral identification by x-ray diffraction (XRD).  . . . . . . . . . . . . . . . . . . . . . .85Table 3.2: Fe values of handsamples from drillholes LC533C, LC555C and LC556C on section L12800N. Two sample populations are presented: samples with Au assays below detec-tion limit (0.005 ppm Au) (n=171); and samples with both no Au and detectable Au as-says (n=268).   .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .87Table 4.1: Duplicate carbon and oxygen isotope data from drillhole LC424, estimated precision (1 SE). 144Table 4.2: Statistical distribution of δ18O values of drill assay rock pulps and hand sample pulps ana-lyzed from drillholes across sections L12800N and L12750N. Drill assay pulp analyses on the left side of the table, hand sample pulp analyses on the right side of table. . . . . . 147Table 4.3: Estimated average precision of stable isotope and geochemical analysis between 1.5 m drill assay pulps and 15 cm hand samples analyzed within same interval. Source Table B.2 in Appendix 2.  . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 158Table B.1: Drill core hand sample descriptions for both Phase I and II drillhole sampling, as well as δ18O, δ13C and selected geochemistry of Phase II sampling from drillholes LC556C, LC555C and LC533C. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 190Table B.2: Calculation of precision between drill core hand sample and drill assay rock pulps for δ18O, δ13C, Au, As, Hg, Tl and Sb values from drillholes LC556C, LC555C and LC533C.  .  .  . 225Table B.3: Micro-drilled samples from drillcore hand samples analyzed for δ18O and δ13C. . . . . . . . 237Table B.4: Surface sample locations, descriptions, sampled lithology and δ18O and δ13C values from samples collected along traverses LC South and LC 12800.  . . . . . . . . . . . . . . . . 245ixLIST OF FIGURESFigure 1.1: Location of the Long Canyon deposit in northeastern Nevada. (Map after Tosdal et al., and Smith et al., 2012). The Sr 0.706 line represents the edge of the rifted Paleozoic continen-tal margin. Shaded bars represent trends of Carlin-type deposits.   .  .  .  .  .  .  .  .  .  .  .  .  .  .  4Figure 2.1: Schematic diagram of mineral-fluid interface during pervasive fluid flow. Dissolved species can be transported to and from mineral surfaces by either advection or diffusion as shown by the solid and dashed arrows, respectively. The fluid adjacent to the mineral interface is often an immobile boundary layer fluid; requiring diffusion through this boundary layer to transport elements to and from the reactive surface. From Schott et al. (2009).  . . . . . . .12Figure 2.2: Schematic representation of rate controlling mechanisms for calcite dissolution as a func-tion of pH and temperature. From Morse et al. (2002). . . . . . . . . . . . . . . . . . . . .13Figure 2.3: (a) Schematic representation of continuous advective fluid flow from a fluid reservoir into a reactive rock mass. Fluid-mineral reaction along flow path involves kinetically controlled oxygen isotope exchange. (b) Changes in δ18O in rock and fluid represented by isotopic fronts as a function of distance along the transport path. Isotopic front migrates along flow path with time. Modified from Cox (2007).  . . . . . . . . . . . . . . . . . . . . . . .19Figure 2.4: Normalized stable isotope composition of a rock colum infiltrated by a reactive fluid over dimensionless distance. Isotopic fronts represent rock column δ18O at that dimensionless time. Note isotopic fronts progress along rock column with time. (a) Peclet number NPe = 100 implying advective infiltration. Infinite Damkohler number (ND) implying a local equilibrium reaction that is transport controlled instead of kinetically controlled. (b) NPe = 100, ND= 1 implying strongly kinetically-controlled fluid-mineral reaction. At the same fluid transport rate as in (a), the kinetic control on mineral reactions does not allow for signifiant isotopic exchange and development of large isotopic gradients within the fluid-flow path. Modified from Baumgartner &Valley (2001). δ18Or, δ18Of and δ18Oi refer to the 18O/16O ratio of the rock at time (τ) and distance, the rock in exchange equilibrium with the infiltrating fluid, and the initial rock, respectively.  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .21Figure 3.1: Profile of a distally steepened carbonate ramp characteristic of the depositional environ-ment present in the Great Basin during Upper Cambrian-Lower Ordovician.  Long Canyon stratigraphy represent near-Outer to Inner Platform carbonates and Continental Clastics. GS=Grainstone, PS=Packstone, WS=Wackestone, MS=Mudstone.  Modified from Cook and Corboy (2004).  . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .57Figure 3.2: Long Canyon stratigraphic sequence modified from Smith (2010). a) Upper Cambrian-Lower Orodovician stratigraphic sequence, with larger 3rd order cycles and smaller 4th and 5th order shallowing cycles interpreted from sequence stratigraphy. Deposit scale stratigraphy in B and C, determined by Fronteer Gold geologists through drillcore; b) Lower Ordovician Pogonip Group; c) Upper Cambrian Notch Peak Formation.   .  .  .  .  .  .57Figure 3.3: Long Canyon property scale geology consists of Upper Cambrian to Permian carbonate and clastic sediments, as well as Cenozoic volcanics. Mapping by Smith and Thompson (2009). Ore deposit outline in yellow as defined in 2010 Q1 by Fronteer Gold.  . . . . . . .58Figure 3.4: Quartz varieties observable in host rocks at Long Canyon. a) Photomicrograph LCW-AL026: Detrital quartz grains, subrounded to subhedral, irregular and broken within calcareous siliciclastic siltstone interbed of the Pogonip Limestone (Opls). b) Photomi-crograph LCWAL411: Metamorphic quartz aggregates aligned in layers within strongly xfoliated calcareous siltstones (Opls) and as pressure shadows on a euhedral pyrite grain.  c) Photomicrograph LCWAL037: Siliceous dolomite (Cnpd) with fine grained intrac-rystalline quartz aggregates between dolomite crystals, similar morphology to quartz observed in chert. d) Photomicrograph LCWAL287: Ribbon chert margin in diagenetic dolomite (Cnpd). e) Photomicrograph LCWAL294: Mottled Notch Peak Dolomite (Cnpd) of varied grain size with commonly observed high angle quartz veins crosscutting.   .  .  .  f) 6.5 cm diameter (HQ) drill core photo LC083C 205’-213’: Fault zone within Notch Peak Dolomite (Cnpd) with pervasively silicified matrix/cement of white quartz.   .  .  .  .  .59Figure 3.5: Dolomitization of Notch Peak Limestone a) Slabbed core LCWAL117: Well developed zebra dolomite replacing burrowed, cherty Notch Peak Limestone (Cnplbu). Note open space crystal growth of sparry white dolomite in vug. b) Photomicrograph LCWAL280: Saddle dolomite crystal growth of zebra dolomite. Note sweeping extinction and curved crystal faces and cleavages. c) Slabbed core LCWAL070: Dolomitized Upper Silty unit of Notch Peak Limestone (Cnplus) altered to typical tan colour with sparry white dolomite veins and oxidized sulphide pseudomorphs. d) Photomicrograph LCWAL070: Grainsize difference between dolomitized siltstone (lower) and dolomite vein (upper). Iron oxide within vein indicates contemporaneous sulphide growth. e) Slabbed core LCWAL278: Zebra dolomite alteration of oncolitic Notch Peak Limestone (Cnplonc) with late quartz growth in open space vugs. f) Photomicrograph LCWAL278: Paragenetically late quartz growth in dolomite vug with slightly curved saddle dolomite crystal faces with compo-sitional zoning. g) Slabbed core LCWAL119: Sulphide growth (oxidized) syn or post dolomitization as coarse clots within dolomitized Cnplonc. Sulphide growth generally occurs as isolated interstitial grains or on dolomite growth lamellae in crystals. This coarse variety is rarely observed. No significant elevation in base metal content over the 5 ft interval to indicate it was anything other than pyrite. h) Photomicrograph LCWAL278: Sulphide (oxidized) along growth lamellae on dolomite rhomb PREVIOUS PAGE.   .  .  .  .60Figure 3.6: Breccia distribution mapped from drill core photos from all diamond drill holes across sec-tions L12800N (top) and L12750N (bottom). Contoured breccia outlines include multi-ple breccia types including cataclastic, carbonate dissolution collapse breccias and late calcite-cemented breccias. Some breccia outlines are drawn with a prefered orientation of a moderately to steeply northwest dipping fabric intersecting the plane of the W40N - E40S oriented cross-section. NEXT PAGE. See text for details.  . . . . . . . . . . . . . . .62Figure 3.7: a,b) Photomicrographs of LCWAL046 and LCWAL047 showing chlorite altered bio-tite laths and aggregates in unmineralized lamprophyre intrusive. c) Photomicrograph LCWAL073, XPL, oxidized, 4 ppm gold-mineralized lamprophyre intrusive with opaques overprinting biotite laths. d) Photomicrograph, reflected light, same image as figure c, he-matite psuedomorphs overgrown on biotite laths, likely after pyrite, evidence for sulfida-tion. e) Mafic intrusive in drill hole LC555C, weakly gold mineralized, cutting dolomite at a high angle to bedding. Adjacent fractures cutting chert nodule of similar orienation indicate orientation of structural damage exploited by intruding mafic dyke. f) Stereo-graphic projection of dike and sill orientations down drillhole LCG02.  . . . . . . . . . . .63Figure 3.8: a) Thinsection Scan LCWAL315: Well foliated silty limestone with elongated limestone boudins and coarse calcite and pyrite clots. Later cleavage crenulating at high angle to foliation.  b) Coreslab LCWAL274 Folded and foliated silty limestone with stretched limestone boudins.  c) Photomicrograph LCWAL332 Foliated siliciclastic silty limestone layers with elongate boudins of coarse calcite, all crenulated by a later, high angle cleav-age.   d) Thinsection Scan LCWAL406 Foliated silty limestone layering with thin micritic limestone interbeds boudinaged, rotated and overprinted by high angle cleavage. . . . . . .64Figure 3.9: Top) Block model image of Long Canyon lithologies over the resource area. Middle xiLeft) Gold distribution trends in 1 g/t cutoff block model with cross section locations of L12750N, L12800N, L12000N, L11900N and Western Long Section shown. Block model data current as of December 15, 2010 from Fronteer Gold. Middle Right) Local simplified geology in the vicinity of the deposit, modified from M.T. Smith presentation. Anticlines occur in Cnp limestones and synclines in Op limestones in between dolomite blocks due to limestone beds flowing into bouding necklines during extension. Stretching lineation is roughly perpendicular to boudin block orientation consistent with extension in a NW-SE direction. Crenulation lineation overprinting foliation indicates a second NW-SE shorten-ing event. Normal fault on western margin of map area places Upper Ordovician over Lower Ordovician stratigraphy. Lower) Ordovician Pogonip lithologies are removed to show coincident gold block model trend with dolomite boudin necks and fractures. Block models created primarily in Gemcom GEMS by Fronteer Gold geologists. Main mineral-ized zones are displayed for reference, as well as cross section lines 12000N, 127500N and 12800N. PREVIOUS PAGE. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .65Figure 3.10: a) Long Canyon interpreted cross section 12000N from drill core, modified from Rhys (2010), displaying boudin block distribution and limestone bedding/S1. b) Oplsm silt beds display alpha angles that with the drill hole orientation and underlying contact with Cnpd indicate sub-horizontal bedding. c) Cnplus and dolomitized oncolite bedding displaying sub-horizontal bedding at bas of Cnpd boudin block. d) Oplsm bedding with near 0° de-gree alpha angles show bedding/S1 plunging over nose of dolomite boudin. e) Fault zone where Cnpd and most of Cnplbu is missing between Oplw and Cnplonc. f) Opdl clasts within solution breccia, below the level of Cnpd boudin block on SE side of intervening fault zone, indicating clockwise rotation of Cnpd blocks. g) Cnpdl bedding with zebra-texture and high alpha angles displaying near horizontal bedding adjacent to fault to NW which accomodated rotation of dolomite blocks. NEXT PAGE.  . . . . . . . . . . . . . . .67Figure 3.11: Above) Long Canyon interpreted cross section 12000N from drill core, modified from Rhys (2010), displaying different boudin failure modes recognized by Mandal (2000).    Middle) Outcrop scale analogues in Cap de Creus, NE Spain, displaying similar boudin failure modes as above, observed by Druguet (2009) in a marble-metapsammite multi-layer. Normal faulting accompanies boudin failure and boudin neck infill. Similar normal fault orientations hosting mineralization are observed at Long Canyon, although their timing is generally attributed to extension occuring post boudinage deformation. These boudin/fault timing examples show initial damage zones may have been established during earlier deformation. Close examination of extensional boudin neck (left) reveals predominantly inflow of overlying stratigraphy.    Lower Left) Unoriented rill core sample LCWAL274 from LC533C. Foliation development has folded limestone beds as tight isoclinal folds in middle of core. At base of drillcore a limestone bed has been stretched into boudins with an apparent antithetic rotation on boudins. Close examination of boudin necks reveals predominantly inflow of underlying silty limestone. . . . . . . . . . . . . . .68Figure 3.12: Notch Peak Dolomite and dolomitized Limestone fault breccia and overprinting altera-tion.  a) Slabbed core LCWAL196: Subangular matrix supported breccia within dolo-mitized burrowed and cherty limestone (Cnplbu). Matrix is weakly siliceous, oxidized and comminuted. Handsample contains 2.6 ppm Au and interestingly only 24 ppm As.  b) Photomicrograph LCWAL196: Comminuted dolomite breccia matrix with subhedral quartz grain growth, finer grained quartz as cement and iron oxide after sulphide pseu-domorphs.  c) Slabbed core LCWAL380: High angle fault damage zone in Notch Peak Dolomite (Cnpd). Comminuted matrix strongly oxidized and weakly silicified. Hand-sample contains 0.13 ppm Au.  d) Photomicrograph LCWAL380: Dolomite fault brec-cia matrix with sub – euhedral quartz grain growth and finer-grained quartz cementing matrix.  e) Thin Section Scan LCWAL297: Notch Peak Dolomite faulted and milled along xiihigh angle fracture zones. Late high angle stylolites after brecciation.  f) Photomicrograph LCWAL297: Dolomite fault breccia of milled subangular dolomite clasts aligned along fractured damage zones supported in a comminuted matrix. High angle stylolite oxidized and cross cutting breccia.  g) Slabbed core LCWAL298: Notch Peak Dolomite fault brec-cia of angular – subangular clasts aligned in a high angle fabric. Handsample contains 0.13 ppm Au.  h) Photomicrograph LCWAL286: Notch Peak Dolomite fault fabric of comminuted dolomite matrix supporting subangular siliceous dolomite breccia clasts. NEXT PAGE.  . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .70Figure 3.13: Equal Area stereonet plots utilizing GEOrient (Holcombe 2011), depicting selected drill-holes (open dots) from sections 12800N and 12750N with median alpha angle of domi-nant fault/fracture fabric within Cnpd plotted as small circles around drillholes. Alpha an-gles were measured from core photos, with an estimated +/- 10 degree error.  Intersections of small circles represent range of poles to likely fault plane orientations. a) 12800N, possible fault orientations dip 50 degrees towards 273 to 350 degrees. b) 12750N, possi-ble fault orientations dip 65 degrees towards 258 to 002 degrees.  .  .  .  .  .  .  .  .  .  .  .  .  .  .71Figure 3.14: a) Modelled dolomite with drillhole intercepts of Opdl and Cnpdl above and below Cnpd. Where boudin necks occur with wide extensional separation, there is a notable absence of Opdl drill intercepts. Conversely where Opdl intercepts are more frequent, boudinage mode is by shear failure with fracturing of the dolomite without extension. b) Variation of boudinage failure mode with respect to layer-thickness ratio and strength ratio between brittle and ductile layers. c) Variation of boudin aspect ratio with layer-thickness ratio. From Mandal et al, 2000.  . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .72Figure 3.15: Ore stage solution breccia examples and alteration from Pogonip and Notch Peak Lime-stone breccias  a) Coreslab LCWAL021: Matrix supported solution breccia in Oplw limestone with concentration of undissolved detrital, metamorphic and hydrothermal quartz grains shown in inset. Matrix is weakly calcite cemented. Sample interval contains 10.25 ppm Au. b) Photomicrograph LCWAL071: Silicified solution breccia within Cnplw and stratigraphically below dolomitized limestone. Large euhedral and doubly terminated quartz grains in breccia. Breccia margin shows contact with later calcite cement brec-cia. Sample interval contains 0.58 ppm Au. Inset shows thin section source coreslab with early silicified breccia on left and later comminuted, oxidized, calcite cemented breccia on right. c) Coreslab LCWAL150: Cnplw limestone solution breccia, with oxidized ma-trix supporting grey apparently unaltered limestone clasts, all cut by later coarse calcite cement breccia. Sample interval contains 1.96 ppm Au. d) Photomicrograph LCWAL150: Limestone solution breccia exhibiting ‘overpacked’ texture with interpenetration of limestone clasts along stylolite contacts. Iron oxide and clay matrix host both detrital and hydrothermal quartz grains, with calcite cement. Hydrothermal quartz growth occurs out of frame as insitu growth on wall rock and limestone clast margins indicating hydrother-mal quartz grains authigenic. e) Coreslab LCWAL220: Oplsm limestone solution breccia from within high grade mineralized core of drillhole LC556C on 12800N. Calcareous silt bedding fragments in matrix supported breccia with sulphide (now iron oxide pseudo-morphs) deposition predominantly along clast margins and clast fractures indicating sulfi-dation occurred late in solution brecciation event. Thin section shows similar pattern of hydrothermal quartz grain growth, with late pervasive silicification of bedding clasts and matrix. Handsample contains 3.97 ppm Au. f) Coreslab LCWAL221: Oplsm limestone solution breccia above figure e). Stratified insoluble solution breccia residuae and silt bedding clasts. Dark maroon areas are very fine-grained iron oxide pseudomorphs after sulfide clots. Bright white clay is dickite. Thin section study shows abundant hydrother-mal quartz, locally appearing as overgrowths on earlier sulfide. Handsample contains 6.4 ppm Au. g) Coreslab LCWAL224: Oplsm limestone bedding collapse breccia. Iron oxide xiiipseudomorphs after sulfide distribution pattern indicates sulfidizing fluids preferentially exploited fractures and bedding. Sample is pervasively silicified. Handsample contains 10.95 ppm Au.  h) Coreslab LCWAL226: Oplsm limestone pervasively replaced by iron oxide pseudomorphs after sulfide. The extent of pre-sulfidation brecciation is difficult to determine, although surrounding strata relatively intact and hosting similar apparent sulfidation along bedding. Handsample contains 45.10 ppm Au and greater than >1% As, providing a strong geochemical and textural argument for gold deposition through sulfida-tion with arsenical pyrite. NEXT PAGE.  . . . . . . . . . . . . . . . . . . . . . . . . . . .74Figure 3.16: Drillhole LC591C. Commonly observed solution collapse breccia facies. . . . . . . . . . .75Figure 3.17: Extensive fracturing and  small-scale normal faulting occurs in Pogonip limestone overly-ing areas of volume loss through solution brecciation. Occuring ubiquitously above solu-tion breccias, this suprastratal deformation is notably absent where underlying limestones are unbrecciated. Examples from drillhole LC533C on L12800N a) LCWAL349, 224 ft: small scale fracturing and normal faulting of Opls limestone. Silty laminations exhibit carbonated dissolution and flow of siliciclastic material into fracture. Inset shows trash-zone of limestone bedding chips, largely calcite with little quartz, amidst siliciclastic material.   b)LCWAL330, 461 ft: fracturing of Opls limestone with normal slip. Dilatant zones filled with siliciclastic residuum, indicating likely carbonate dissolution along frac-tures, which are filled with calcite veinlets. c) LCWAL339, 330 ft: fracturing within Opls limestone, exhibiting 2-3 distinct fluid flow events, consisting of two oxidation events and calcite deposition. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .76Figure 3.18: Hydrothermal quartz grains and silicification observable at Long Canyon  a) Thin Sec-tion Scan LCWAL021: Ore stage solution breccia at base of Oplw consisting of dissolved limestone and calcareous siltstone residuum with detrital and hydrothermal quartz grains within iron oxide and clay matrix. Sample interval contains 10.25 ppm Au. Inset photomi-crograph of solution breccia residuum with doubly terminated and sub-rounded to sub-euhedral hydrothermal quartz grains amongst irregularly shaped detrital quartz hosted in iron oxide and clay matrix with minor calcite cement. b) Photomicrograph LCWAL220: Hydrothermal quartz, doubly terminated subhedral to euhedral grains grown within Oplw limestone dissolution breccia, overprinted by fine grained-microcrystalline quartz. Handsample contains 4 ppm Au.  c) Thin Section Scan LCWAL059: Bedding collapse solution breccia of Opls calcareous siltstone bedding with oxidation and rotation of silt bedding fragments and concentration of quartz grains.  d) Photomicrograph LCWAL059: Quartz species in iron oxide and clay matrix amidst bedding fragments. Quartz varieties include fine grained aggregates of metamorphic quartz, detrital quartz and doubly termi-nated hydrothermal quartz grains. Sample interval contains 0.13 ppm Au.  e) Coreslab LCWAL037: Several breccia generations within dolomitized Cnplbu, with an earlier light brown coloured, more competent breccia cut by a later maroon, iron oxide rich breccia, all cut by late calcite veins.  f) Photomicrograph LCWAL 037: Dolomitized limestone clast in iron oxide matrix breccia with doubly terminated, euhedral hydrothermal quartz grain. Sample interval contains 0.25 ppm Au. g) Thin Section Scan LCWAL034: Several generations of dolomitized limestone (Cnplbu) breccia. Earlier breccia clast appears to host tectonic fabric of aligned dolomite fragments within comminuted matrix.  h) Photo-micrograph LCWAL034: Early tectonic dolomite fault breccia clast within a later brec-cia hosting hydrothermal quartz grains and possibly a weakly silicified matrix. Sample interval contains 0.40 ppm Au. NEXT PAGE.  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .78Figure 3.19: Quartz and amorphous silica solubility as a function of temperature in a 5 wt percent NaCl solution at 500 bars, modified from Hofstra (2000). Displayed are two possible cooling paths of Long Canyon ore fluids as indicated by quartz textures: 1) a silica-rich xivfluid cools until reaching the quartz-saturation curve, after which it cools along the curve while precipitating euhedral hydrothermal quartz grains; 2) mixing (red path) with a cool-er meteoric fluid occurs and the resulting fluid becomes rapidly quartz-supersaturated; 3) no fluid mixing (blue path) occurs, fluid continues cooling until 180-200 °C, after which quartz precipitation stops (Fournier, 1985; Hofstra, 2000) and fluid continues cooling in a quartz supersaturated state; 4) quartz supersaturation results in precipitation of large volumes of fine-grained, non-texturally destructive silicification within breccias. . . . . . .79Figure 3.20: Post Ore Calcite Cement Breccias within Pogonip and Notch Peak lithologies  . . . . . . a) Thin Section Scan LCWAL033: Dolomitized wispy limestone (Cnplw) clasts with early oxidized and silicified breccia matrix on margins, all brecciated by a later coarse calcite cement breccia. Inset of the thin section source coreslab. Sample interval contains 2.5 ppm Au. b) Photomicrograph LCWAL033: Calcite cement breccia with clasts of an earlier breccia of dolomitized limestone with a comminuted matrix hosting euhedral, dou-bly terminated quartz grains and iron oxides. c) Thin Section Scan LCWAL035: Calcite cement polymictic breccia within dolomitized Cnplbu. Subangular dolomite clasts within an earlier iron oxide rich breccia matrix occur together as large composite clasts within coarse calcite cement. Isolated iron oxide breccia matrix as clasts indicate coarse calcite was not just a cement to the earlier breccia. Sample interval contains 6.79 ppm Au. d) Photomicrograph LCWAL035: Early iron oxide breccia matrix hosting euhedral, doubly terminating quartz grains. Clasts of dolomitized siltstone (Cnplus) identified based on dolomite grainsize indicates early iron oxide rich polymictic breccia was the likely the re-sult of reverse faulting as Cnplus is stratigraphically below Cnplbu. Hydrothermal quartz likely associated with 6.79 ppm Au mineralization in sample interval. All brecciated by later coarse calcite.  e) Thin Section Scan LCWAL060: Early calcareous siltstone bedding collapse breccia within Oplsm with comminuted iron oxide and quartz grain-rich matrix, all brecciated by a later coarse calcite cement. Sample interval contains 0.75 ppm Au.  . . f) Photomicrograph LCWAL060: Coarse, late calcite cement brecciating an earlier quartz, clay and iron oxide-rich matrix supporting calcareous siltstone clasts. Isolated clasts of early breccia matrix indicates coarse calcite was a later event and not a progressive ce-mentation of quartz-bearing breccia.  g) 6.5 cm (HQ) diameter drill core LC173C 153.1’: Coarse, late calcite breccia cement within Cnplw with clasts of silicified early calcareous silt bedding collapse breccia. h) 8.5 cm (PQ) diameter drill core LCBS2A 68’: Coarse, late calcite cement within Oplw brecciating oxidized calcareous siltstone and grainstone interbeds. Liesengang banding which occurs elsewhere nucleating from oxidizing fluid flow along fractures is present here, indicating extension and calcite matrix infill occurred along preexisting fractures. PREVIOUS PAGE.  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .80Figure 3.21: Contoured Au grades >0.01ppm on sections L12800N (top) and L12750N (bottom) over-lain on lithology and contoured breccia distribution. Widths of down hole bar chart dis-playing Au grade is constrained to 5 ppm Au or less to maintain clarity on the figure. Au contours are often drawn with a prefered orientation of a moderately to steeply northwest-dipping fabric intersecting the plane of the W40N - E40S oriented cross-section. This orientation follows that of the breccias, which are the prefered host and first order control on the distribution of Au mineralization. NEXT PAGE.  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .82Figure 3.22: SEM backscatter images of iron oxide pseudomorphs with detectable As. LCWAL040 (top) 35 μm Fe-oxide psuedomorph, likely after pyrite, within quartz-replaced Cnpdl (0.21 ppm Au/ft). Image displays an alteration rim, spot analysis indicates arsenical pyrite with As as an identifiable ore-stage trace element. LCWAL061 (bottom) 5 x 10 μm zoned Fe-oxide psuedomorph after pyrite grain encapsulated within silica overprinting Oplw solution breccia matrix (15.75 ppm Au/5ft). Spot analysis indicates Fe, O and As suggest-ing oxidation of arsenian pyrite. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .83xvFigure 3.23: a) plotted values of calculated fluid δD and δ18O values for H20 in equilibrium with silicate minerals. Error (uncertainty) bars represent possible fluid istopic compositions at temperatures from 180 - 240 °C. XRD analysis of clay separate mineralogy in Table 3.1. Additional y-axis error bars on sample LCWAL260 represent endmember isotopic compositions for 100% illite and 100% kaolinite in sample, as XRD indicates presence of both minerals but not relative abundances. Cenozoic, Cretaceous and Jurassic meteoric water lines (acounting for Paleo-latitudes) taken from Hofstra (1999). Also shown are magmatic and metamorphic fluid isotopic compositions, and the meteoric water line from Taylor (1974). Calculated isotopic composition of fluids in equilibrium with clay minerals fall largely in the magmatic fluids field, while the dickite sample LCWAL221 appears to have a mixed with a component of meteoric fluid. b-g) - handsamples and their location on drillhole LC556C, from which clay separates were prepared and analyzed for δD and δ18O values.  . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .84Figure 3.24: Box and whisker plots of Fe and Au contents of lithologies sampled from drillholes LC533C, LC555C and LC556C on section L12800N.   Boxes represent middle 50% of data. Whiskers are the 5% and 95% values. Black circles are the mean. Black lines are the median. Open circles are top or bottom 5% of data (outliers). Specific values displayed in Table 3.1. a) Fe content range of samples where Au is below detection limit (0.005 ppm Au). b) Fe content range of all samples, Au-mineralized and unmineralized. c) Au content range of all samples. All lithology abbreviations as described in text, and in Table 3.1 footnotes. Of particular note are highest Au contents contained in units Opbx and Cnpbx - breccias within Pogonip Group and Notch Peak Formation rocks, respectively. Lampro-phyres sampled on these sections significantly under-represent typical Au grade observed in mineralized lamprophyres elsewhere in the deposit. . . . . . . . . . . . . . . . . . . . .86Figure 3.25: Cross-sections L127800N (top) and L12750N (bottom) displaying hand contoured As from drillhole sampling at a cutoff of >20 ppm As. Overlain on contoured Au>0.1ppm (red outline), lithologies and mapped breccia.  . . . . . . . . . . . . . . . . . . . . . . . .89Figure 3.26: Cross-sections L127800N (top) and L12750N (bottom) displaying hand contoured Sb from drillhole sampling at a cutoff of >0.15 ppm Sb. Overlain on contoured Au>0.1ppm (red outline), lithologies and mapped breccia.  . . . . . . . . . . . . . . . . . . . . . . . .90Figure 3.27: Cross-sections L127800N (top) and L12750N (bottom) displaying hand contoured Tl from drillhole sampling at a cutoff of >0.1 ppm Tl. Overlain on contoured Au>0.1ppm (red outline), lithologies and mapped breccia.  . . . . . . . . . . . . . . . . . . . . . . . .91Figure 3.28: Cross-sections L127800N (top) and L12750N (bottom) displaying hand contoured Hg from drillhole sampling at a cutoff of >0.25 ppm Hg. Overlain on contoured Au>0.1ppm (red outline), lithologies and mapped breccia.  . . . . . . . . . . . . . . . . . . . . . . . .92Figure 3.29: Brecciated, mineralized and oxidized Ordovician Pogonip Limestone lithologies from 28’ - 499’ from hole LC577C on cross section 12750N. Demonstrated is the extensive brecciation and alteration above solution breccias, scorodite overprint of hematite and li-monitic upper. Arranged as downhole columns from left to right. Lithologies going down stratigraphy include Opls - silty laminated limestone, Oplsm - silty massive limestone, Oplw - wispy limestone and overprinting Opdl - dolomitized grainstone and minor silt interbeds.  . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .93Figure 3.30: Brecciated, mineralized and oxidized Ordovician Pogonip Limestone lithologies from 104’ - 569’ from hole LC556C on cross section 12800N. Demonstrated is the strong solu-tion brecciation adjacent to dolomitized limestone which are not decarbonated. Arranged as downhole columns from left to right. Lithologies going down stratigraphy include Opls xvi- silty laminated limestone, Oplsm - silty massive limestone, Oplw - wispy limestone and overprinting Opdl - dolomitized grainstone and minor silt interbeds.   .  .  .  .  .  .  .  .  .  .  .  .94Figure 3.31: Relatively unaltered Ordovician Pogonip Limestone lithologies from 119’ - 584’ from hole LC452C on cross section 12800N. Demonstrated is the lack of structure within lime-stones and a corresponding lack of decalcification and alteration. Arranged as downhole columns from left to right. Lithologies going down stratigraphy include Opls - silty lami-nated limestone, Oplsm - silty massive limestone, Oplw - wispy limestone and overprint-ing Opdl - dolomitized grainstone and minor silt interbeds.  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .95Figure 3.32: Block model schematic evolution of the Long Canyon deposit.  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .96Figure 4.1: Schematic representation of the scales of sampling employed in this study. See text for details. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 140Figure 4.2: Long Canyon geological mapping from Smith 2008 and Thompson 2009 with author’s iso-topic surface, drillcore handsample and drill assay pulp sampling displayed in green, blue (Phase 1 sampling) and red (Phase 2 sampling), respectively. Long Canyon drill traces and selected section lines also displayed. . . . . . . . . . . . . . . . . . . . . . . . . . . 141Figure 4.3: δ18O values of micro-drilled silty laminae from multiple subunits in a cumulative frequency plot. Inflections in the slope of data distribution represent different sample populations and are selected as: unaltered background δ18O values > 19 ‰; a range from 19 - 17.5 ‰ of uncertainty in the degree of oxygen isotope depletion from hydrothermal fluid infiltra-tion or just the lower end of background values within error; and all values below 17.5 ‰ where all rock is considered depleted from background oxygen isotope values from interaction with an isotopically light hydrothermal fluid. . . . . . . . . . . . . . . . . . . 142Figure 4.4: Plot of δ13C values against δ18O analyzed from hand sample pulps. Median δ13C values of lithologies are represented by shaded boxes and display a secular change in δ13C values with lithological younging from right to left. a) Hand sample pulp δ13C values with > 17.5 ‰ δ18O. b) Hand sample pulp δ13C values with < 17.5 ‰ δ18O. Results show that isotopic exchange between carbonate rocks with a hydrothermal fluid at Long Canyon do not statistically effect δ13C values. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 143Figure 4.5: Line plots of carbon and oxygen isotopes from duplicate analyses of drill assay pulps down drill hole LC424. Samples were analyzed approximately 6 months apart. a) δ18O values. b) δ13C values. Symbols plotted are original analysis, duplicate analysis and the mean of the two analyses. Reproducibility as an estimate of precision (1 SE) is calculated in Table 424Dup.  . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 145Figure 4.6: Plots of δ18O and δ13C values by lithology of all drill assay rock pulps (a to e) and drill hand sample rock pulps (f to j). Rock pulps are from drillholes on sections L12800N and L12750N. Hand samples are from drillholes LC556C, LC555C and LC533C. Global background box displays the global mean δ18O and δ13C values for carbonates of Late Cambrian to Early Ordovician in age, determined from Shields and Veizer (2002). Long Canyon background δ18O and δ13C values determined from handsamples and drill as-say pulps sampled well outside of visual alteration and geochemically displaying no enrichment in pathfinder elements associated with ore-fluids. Plots are differentiated by lithology as follows: (a and f) - all lithologies; (b and g) - Pogonip Limestones; (c and h) - Pogonip and Notch Peak Breccias; (d and i) - Notch Peak Limestones; (e and j) - Pogonip and Notch Peak Dolomites. (Next Page)  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  . 145Figure 4.7: Box and whisker plots of δ18O and δ13C values by lithology of all drill assay rock pulps (a and b) and drill hand sample rock pulps (c and d). Rock pulps are from drillholes on xviisections L12800N and :12750N. Hand samples are from drillholes LC556C, LC555C and LC533C. Boxes represent middle 50% of data from Q1 to Q3. Whiskers are the 5% and 95% values. Black circles are the mean. Black lines are the median. Open circles are top or bottom 5% of data (outliers) further than 1.5*(Q1-Q3) from the box. Triangles are far outliers further than 3*(Q1-Q3) from the box. Plots are ordered from lowest to highest δ18O and δ13C values.   .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  . 147Figure 4.8: Cross-section L12800N displaying δ18O values from drill assay pulps down drillholes as coloured bars of variable width, shorter is more 18O-depleted, against contoured Au>0.1ppm (red outline). (Following page).  . . . . . . . . . . . . . . . . . . . . . . . . 149Figure 4.9: Cross-section L12750N displaying δ18O values from drill assay pulps down drillholes as coloured bars of variable width, shorter is more 18O-depleted, against contoured Au>0.1ppm (red outline). (2 pages forward). . . . . . . . . . . . . . . . . . . . . . . . . 150Figure 4.10: Cross-sections L127800N (top) and L12750N (bottom) displaying hand contoured As from drillhole sampling at a cutoff of >20 ppm As. Overlain on δ18O values from drill as-say pulps down drillholes as coloured bars of variable width, shorter is more 18O-depleted, against contoured Au>0.1ppm (red outline). (3 pages forward).  . . . . . . . . . . . . . . 151Figure 4.11: Cross-sections L127800N (top) and L12750N (bottom) displaying hand contoured Sb from drillhole sampling at a cutoff of >0.15 ppm Sb. Overlain on δ18O values from drill assay pulps down drillholes as coloured bars of variable width, shorter is more 18O-deplet-ed, against contoured Au>0.1ppm (red outline). (4 pages forward).  . . . . . . . . . . . . 152Figure 4.12: Cross-sections L127800N (top) and L12750N (bottom) displaying hand contoured Tl from drillhole sampling at a cutoff of >0.1 ppm Tl. Overlain on δ18O values from drill as-say pulps down drillholes as coloured bars of variable width, shorter is more 18O-depleted, against contoured Au>0.1ppm (red outline). (5 pages forward).  . . . . . . . . . . . . . . 153Figure 4.13: Cross-sections L127800N (top) and L12750N (bottom) displaying hand contoured Hg from drillhole sampling at a cutoff of >0.25 ppm Hg. Overlain on δ18O values from drill assay pulps down drillholes as coloured bars of variable width, shorter is more 18O-deplet-ed, against contoured Au>0.1ppm (red outline). (6 pages forward).  . . . . . . . . . . . . 154Figure 4.14: Trace elements As, Hg, Tl, and Sb plotted against δ18O and coloured by Au grade (a to d) and lithology (e to h). Divided into 4 quadrants, the lower left box represents data points depleted in 18O and below trace element cutoff. See text for details. . . . . . . . . . . . . 155Figure 4.15: Cross-section L12800N displaying δ18O values from 1.5 m drill assay pulps on right and 15 cm hand sample pulps on left, down drillholes LC556C, LC555C and LC533C against contoured Au>0.1ppm (red outline). Previous Page.  . . . . . . . . . . . . . . . . . . . . 156Figure 4.16: (a - c): Plots of δ18O values from drill assay pulps against δ18O values from hand sample pulps taken within same 1.5 m interval. Displayed is line y=x and a linear regression line through the data. (d and e): Plots of δ18O values from hand samples against δ18O values of micro-drill analyses from within the same hand sample. Samples are plotted as a shape representing the rock type from which the sample was drilled. Where multiple samples were drilled from a hand sample, multiple data points are plotted, each representing the difference between δ18O values from hand sample pulp and the micro-drilled carbonate sample.   .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  . 157Figure 4.17: Log-log plots of drill assay pulp Au, As, Hg, Sb and Tl values against hand sample pulps taken within same 1.5 m interval from drill holes LC556C, LC555C and LC533C. Data points are coloured by Hand Sample Au grade. Displayed is line y=x and a linear regres-xviiision line through the data. N = 185. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 158Figure 4.18: Drill core hand samples from drill hole LC556C on L12800N. Drill core is 7.6 cm in diameter. Micro-drilled carbonate spot analyses and their δ18O values are displayed in ‰ (VSMOW). Hand sample bulk δ18O value from assay pulp displayed in lower left corner. Hand sample As and Au content displayed in upper right corner.  . . . . . . . . . . . . . 159Figure 4.19: Drill core hand samples from drill hole LC533C on L12800N. Drill core is 7.6 cm in diameter. Micro-drilled carbonate spot analyses and their δ18O values are displayed in ‰ (VSMOW). Hand sample bulk δ18O value from assay pulp displayed in lower left corner. Hand sample As and Au content displayed in upper right corner.  . . . . . . . . . . . . . 160Figure 4.20:  Pogonip Limestone hand samples from 12.0 - 44.0 m (LCWAL368 - LCWAL357) down LC533C. Drillcore is 7.6 cm in diameter. Shown are bulk δ18O values in bottom left cor-ner and As concentration in upper right.   .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  . 161Figure 4.21: Plots of δ18O values from hand samples against δ18O values of micro-drill analyses from within the same hand sample. Samples are plotted as a shape representing the rock type from which the sample was drilled. Where multiple samples were drilled from a hand sample, multiple data points are plotted, each representing the difference between δ18O values from hand sample pulp and the micro-drilled carbonate sample.  . . . . . . . . . . 162Figure 4.22: Plots of δ18O and δ13C values of micro-drilled carbonate from drill core hand samples. Data points are coloured by stratigraphic unit with sampled rock type represented by shape as either massive micritic limestone, silty limestone laminae, limestone breccia clast or matrix, massive dolomite or dolomitized limestone and dolomite breccia clast or matrix. Global background box displays the global mean δ18O and δ13C values for carbonates of Late Cambrian to Early Ordovician in age, determined from Shields and Veizer (2002). Long Canyon background box determined from range of visually unaltered micro-drilled carbonates. Long Canyon uncertain box represent δ18O values from visu-ally unaltered carbonates which could be background values within analytical error, or the result of oxygen isotope depletion from hydrothermal fluid:rock interaction.  .  .  .  .  .  .  . 163Figure 4.23: Drill core hand samples from drill hole LC287C on Western Dolomite Long Section. Drill core is 7.6 cm in diameter. Micro-drilled carbonate spot analyses and their δ18O values are displayed in ‰ (VSMOW). Samples (a-g) represent an intersection through Pogonip Group carbonates. Samples in (f) and (g) are taken immediately above and below an unmineralized lamprophyre intrusive, as a test for the effect of a lamprophyric intrusion on oxygen isotope ratios in the wall rock. Sample (g) shows the contact between Pogonip Limestone and Notch Peak Dolomite as a conformable contact. Samples (h and i) display variable isotopic alteration within adjacent Cnplus units. Sample (j) displays unaltered Cnplw with background oxygen isotope values.  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  . 164Figure 4.24: Drill core hand samples from drill hole LC533C (a-c) and LC556C (d-f) on section L12800N. Drill core is 7.6 cm in diameter. Micro-drilled carbonate spot analyses and their δ18O values are displayed in ‰ (VSMOW). Hand sample bulk δ18O value from assay pulp displayed in lower left corner. Hand sample As and Au content displayed in upper right corner. Samples represent a) zebra dolomite and b-f) fault breccias within Notch Peak Dolomite.   .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  . 165Figure 4.25: a) micro-drilled carbonate vein δ18O and δ13C analyses, distinguished by lithology and vein type. b) hand sample bulk δ18O plotted versus vein δ18O value.  c-n) micro-drilled carbonate spot analyses and their δ18O values are displayed in ‰ (VSMOW). Hand sam-ple bulk δ18O value from assay pulp displayed in lower left corner. Hand sample As and Au content displayed in upper right corner. drill core is 7.6 cm in diameter.   .  .  .  .  .  .  . 166xixFigure 4.26: Surface isotope and geochemical study south of the Long Canyon deposit in Lower Pogonip Group rocks exhibiting a similar structural environment to the main deposit; spe-cifically brecciation in basal Pogonip units and boudinage within Notch Peak Dolomite. Lithologic units as described in Section 3.3.2 in text. All figures aligned by a common UTM East x-axis. (a): Historic rock sampling from a ridge and spur and general prospect-ing sampling program displaying Au in handsamples.Selected trace element geochemistry displayed as a ranked variable map including (b): As and (c): Sb, of rock samples plotted by UTM Easting and Northing. (d): δ18O values of drilled carbonate material from 50 rock samples collected along a ridgeline traverse. (e): δ18O values plotted against Easting. See text for description. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 167Figure 4.27: Surface isotope and geochemical study near L128000N along an 800 m traverse in Lower Pogonip Group rocks. Lithologic units as described in Section 3.3.2 in text. All figures aligned by a common UTM East x-axis. (a) Historic rock sampling from a ridge and spur sampling program displaying Au in handsamples. Selected trace element geochemistry displayed as a ranked variable map including (b) As, (c) Hg, (d) Tl and (e) Sb, of rock samples plotted by UTM Easting and Northing. (f) δ18O values of drilled carbonate mate-rial from 59 rock samples collected along a ridgeline traverse. (g) δ18O values plotted against Easting. See further description in text.   .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  . 168Figure 4.28: Schematic block diagram of interpreted fluid flow path evolution at Long Canyon. (1) Initial fluid flow occurs along longitudinal cracks within the dolomite. (2) With increasing TIFF, limestone dissolution occurs above and below the dolomite at fracture terminations, generating solution collapse breccia and increasing permeability. (3) Fluid flow continues, enriching Au grade through fluid focusing primarily within limestone solution breccia. Extensive fracture permeability develops within overlying limestone units owing to un-derlying volume loss. Within these secondary fracture flow paths, ore-stage trace element fronts are slightly nested within oxygen isotope depletion fronts.  .  .  .  .  .  .  .  .  .  .  .  .  . 169Figure 4.29: Schematic diagram of an advective fluid flow path along a silty laminated limestone bed within micritic limestone beds. Horizontal arrows represent relative strength of darcy flux. Fluid flow path is fluid-dominated nearest the fluid conduit, becoming progressively more rock-dominated as time-integrated fluid flux (TIFF) decreases to the right. An 18O exchange front is the most distal expression of fluid flow in the rock, with a geochemi-cal front representing the extent of trace element geochemical metasomatism within the fluid flow path of the rock. Rock sample R0 to R4 represent progressive alteration as they are passed by the isotopic and geochemical fronts. Rocks within the same front are more altered by a greater TIFF. Rock samples display bulk hand sample δ18O value in bottom left corner and As content in upper right corner.  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  .  . 170Figure 4.30: Normalized stable isotope composition of a rock colum infiltrated by a reactive fluid over dimensionless distance. Isotopic fronts represent rock column δ18O at that dimensionless time. Note isotopic fronts progress along rock column with time. (a) Peclet number NPe = 100 implying advective infiltration. Infinite Damkohler number (ND) implying a local equilibrium reaction that is transport controlled instead of kinetically controlled. (b) NPe = 100, ND= 1 implying strongly kinetically-controlled fluid-mineral reaction. At the same fluid transport rate as in (a), the kinetic control on mineral reactions does not allow for signifiant isotopic exchange and development of large isotopic gradients within the fluid-flow path. Modified from Baumgartner &Valley (2001) after Bowman and Willett (1994). δ18Or, δ18Of and δ18Oi refer to the 18O/16O ratio of the rock at time (τ) and distance, the rock in exchange equilibrium with the infiltrating fluid, and the initial rock, respectively.  .  .  . 171Figure 4.31: (a) Schematic diagram of aqueous advective fluid infiltration fronts of non-dimensional time from t=0 to t=3, occurring through a multi-layered carbonate rock with bimodal xxpermeability. Horizontal arrows represent relative strength of immediate darcy flux. Kinetically-controlled isotope exchange between fluid and mineral occurs behind each front. Shape of infiltration front controlled by rock permeability with greatest initial fluid flux occurring through a highly permeable mm-scale silty limestone layer (slt) interbed-ded with less permeable micritic limestone (lst). (b) Schematic diagram of 18O exchange fronts migrating through the silty limestone layer, illustrated as a function of distance along the fluid flow path. With increasing time the position of the isotopic front migrates along the flow path. Time integrated fluid flux (TIFF) increases along the path from left to right resulting in greater fluid-mineral exchange and increasing 18O depletion in rock. (c) Changes in δ18O of silty limestone at point P1 and micritic limestone at point P2 in (a) as isotopic fronts pass through and TIFF increases. A high fluid flux and rapid isotopic exchange owing to small carbonate grain size in silty limestone precipitates rapid δ18O depletion at P1. A lower fluid flux and more kinetically-inhibited fluid-mineral isotopic exchange results in more gradual δ18O depletion at P2. (d) Silty limestone equivalent to t0. Rock sample displays bulk hand sample δ18O value in top right corner and δ18O spot val-ues from micro-drilling. (e) Silty limestone interlayered with micritic limestone showing 18O depletion only within silty limestone, equivalent to t1. (f) Strongly 18O-depleted silty and micritic limestone layering, equivalent to t3. Modified from Cox (2007).  . . . . . . . 172xxiACKNOWLEDGEMENTSI would like to thank Fronteer Gold for supporting me in going back to school and for funding this thesis. Your work on Long Canyon established the geological framework I needed to successfully carry out this study. Thanks to Newmont for your support. To the Society of Economic Geologists, thank you. Your sup-port of students is exemplary.Ken Hickey, you were an excellent supervisor. You pushed the project in all the right directions and I am grateful to have had the opportunity to work with you. Your input and direction was always insightful and timely. Your critique has made me a better geologist. Thank you, Ken. I can’t thank Shaun Barker and Jer-emy Vaughan enough. My advice to any new student is to hitch your wagon to someone who knows where they are going. Thanks for all the support and guidance, guys. Thank you to my committee members Drs Craig Hart, Moira Smith, Shaun Barker and my external examiner Dr Greg Dipple for your interest and guidance in my project and for not letting me get away with submitting that garbage first draft.To the Hickey Lab. I love you all. Do your dishes, and stay awesome. It has been one of the pleasures of my life to run this gauntlet with you all.Thank you to the MDRU crew and all that you do. Enjoy your new windows and remember all of us who did it in the caves.To summers in Vancouver. You were better than I dreamed you’d be. Obviously as a geologist, we’ll never hang out again.Speaking of seasons, the fall writeup back home on Denman Island with my parents and the woman I love was some of the finest time I’ve spent on that island. Thanks for the company and the cooking.I would like to thank my bicycles. You were worthy companions. I’m sorry I broke one of your frames. I would also like to acknowledge rainy nights on the bike. You were a worthy adversary.Claire Williams, I love you. Your patient support and companionship brings balance to my life. xxiiDEDICATIONTo my family.1Chapter 1: Thesis Introduction and Context1 .1 IntroductionOverview and Motivation for StudyWhen a hydrothermal fluid infiltrates a carbonate rock, oxygen and carbon isotopic exchange processes that occur between carbonate minerals and the fluid can produce changes in the original isotopic ratios of the rock (Criss, Gregory et al. 1987). Analysis of these “exchanged” isotope ratios in the host rock, relative to their original isotopic composition, can help to define the extents of fluid-mineral interaction, as well as provide information on the mode of fluid flow, fluid flux and its chemical evolution through the fluid flow path (Baumgartner and Valley 2001). Previous studies have shown that this isotopic exchange can be detected beyond the extents of ore-stage trace element metasomatism in Carlin-type deposits and can potentially be used as a far-field vector to orebodies (Vaughan, Hickey et al. 2010). This study uses oxygen and carbon isotope ratios of carbonate rocks to determine hydrothermal fluid flow path development around the Long Canyon deposit, a sediment-hosted Au deposit located in northeastern Nevada. The deposit geology, geochemistry and mineralogy suggest that it is a Carlin-type deposit. These deposits typically exhibit narrow alteration zonation patterns compared to other hydrothermal Au deposits (Cline, Hofstra et al. 2005). Exploration for these deposits can be challenging, particularly in blind-drilling, owing to their small alteration footprints. Through oxygen and carbon isotope analysis, this study aimed to improve understanding of hydrothermal fluid flow paths around the Long Canyon deposit, and from this understanding, develop guidelines for use of oxygen and carbon isotope studies in exploration programs at Long Canyon and similar Carlin-type deposits.Previous oxygen and carbon isotope studies around Carlin-type deposits have been unable to resolve individual fluid flow paths owing to a lack of appropriate sample density (Stenger, Kesler et al. 1998; Arehart and Donelick 2006). Historically, this lack of resolution on fluid pathways/oreshoots has made oxygen and carbon isotope studies of limited use in mineral exploration programs. In this project we employ isotopic analysis of contiguous drill assay pulps down drillholes. This high spatial density sampling allows for mapping of meter-scale fluid flow paths. Isotopic exchange patterns present in drill assay pulps are correlated with specific carbonate lithologies and alteration types by high density isotopic analysis of hand samples down select drill-holes. From these hand samples, centimeter-scale control on fluid flow paths can be determined. Presented in this thesis are the results of over 2,800 unique carbonate 2oxygen and carbon isotope sample ratios from across the deposit; making it the largest stable isotope study of a carbonate-hosted hydrothermal ore deposit to-date.Project OutcomesThrough establishing understanding of the geological framework for fluid flow and examining evidence of fluid-mineral isotopic exchange, this study was able to: (1) map fluid pathways over drill cross-sections and surface sampling traverses at the Long Canyon deposit; (2) assess the structural and lithologic controls on fluid flow; (3) determine the extent of fluid-rock interaction and isotopic exchange beyond the limits of trace element halos genetically associated with gold mineralization (antimony, thallium, mercury and arsenic); and (4) provide guidelines for the use of stable isotope studies in mineral exploration programs at Long Canyon and Carlin-type deposits in general. 1 .2 Sediment-Hosted “Carlin-type” Gold Deposit CharacteristicsCarlin-type gold deposits occur predominantly in the northeastern corner of Nevada, concentrated along two parallel, northwest-trending clusters, the Carlin Trend and the Battle Mountain-Eureka Trend ( ). These deposits formed from 42-34 Ma as epigenetic, sediment-hosted, disseminated, auriferous pyrite deposits with the following salient attributes as summarized by Cline et al. (2005):1. Alteration is characterized by carbonate dissolution, argillization of silicates, sulfidation of ferroan minerals and silicification/jasperoidation of typically calcareous sedimentary rocks and lesser intrusive rocks.2. Ore fluids transported Au as a bisulfide complex at temperatures of 250° to 150°C, from moderately acidic (pH≈5), reduced, low salinity fluids containing <4 mole percent CO2, and >0.01 mole percent H2S.3. Gold was deposited from fluid into arsenian pyrite as submicron inclusions during sulfidation reactions where Fe from ferroan host-rocks consumed H2S. Concentrations of As, Hg, Tl and Hg were also precipitated into pyrites (Barker, Hickey et al. 2009). This process is triggered by the availability of Fe2+ to the ore fluid, which is the result of cooling ore fluids becoming increasingly acidic and dissolving Fe-carbonate or increasing permeability through carbonate dissolution for fluids to access detrital ferroan siliciclastic minerals (Hofstra, Leventhal et al. 1991). 4. Mineralization controls vary by individual deposit, but are generally a combination of structural and stratigraphic. Structures utilized by ore fluids are typically high-angle faults to access favourable stratigraphy. Stratigraphic controls include a requirement for ferroan-rich rocks 3for sulfidation, and sufficient permeability for high fluid flux. Often a combination of the two includes secondary fracture permeability development, coupled with dissolution of carbonate for reaction-enhanced permeability. This often creates solution breccia with spectacular Au grades (Emsbo, Hofstra et al. 2003).5. Source of ore fluids has not been resolved unequivocally as calculated δ18OH2O and δDH2O of ore-stage minerals variably suggest a meteoric source, a mixed meteoric-magmatic source or a meteoric-metamorphic source (Cline, Hofstra et al. 2005). It is likely that ore-fluids of whatever source reacted with local meteoric ground waters (Muntean, Cline et al. 2011). Ore-stage orpiment from the Getchell deposit has a calculated δ18OH2O value of 4.2‰, which Cline and Hofstra (2000) interpret as an ore-fluid of mixed magmatic-meteoric origin.6. Host rocks are predominantly Devonian slope facies carbonate sedimentary rocks. The major deposit trends are thought to be spatially associated above old, reactivated basement structures along the edge of the rifted Paleozoic continental margin ( ).45Fig. 1. Location of Long Canyon Property (after Tosdal et al., 2000).  The Sr 0.706 line is used as an indication of the edge of the rifted Paleozoic continental margin.Geological SettingThe Pequop Mountains are underlain primarily by thrust- and low-angle normal fault-imbricated Lower and middle Paleozoic strata (Fig. 2).  Lower Paleozoic strata generally record episodic shallowing of the passive continental margin and migration of the shelf break westward over time.  Slope facies gave way to shelf margin and platformal limestone and dolomite during the Cambrian, with fluctuations of water level from relatively deep to emergent conditions during the Ordovician, then deposition of mainly shelf facies dolomite and limestone during Silurian to Devonian time (Cook and Carboy, 2004). This continental margin assemblage is overlain by a sequence of siliciclastic rocks assigned to the Mississippian foreland basin that advanced eastward over the shelf, ahead of the Antler orogenic highland. Minor Permian limestone is also present locally.  Paleozoic strata in northeastern Nevada and eastern Utah record ductile to ductile-brittle,penetrative deformation possibly related to the mid-Jurassic Elko Orogeny (Thorman, 1970; Thorman et al., 1991).  Other related features include amphibolite to low greenschist-facies metamorphism, low angle thrust faults, low-angle normal faults, northwest-southeast- to east-west-trending stretching lineations, and folding on a regional scale. Thorman et al. (1991) dated this event as pre-165-155 Ma through 40Ar/39Ar dating of undeformed plutons that cut NEVADASr 0.706ArchaeanCrustOceanicCrustUTAHIDAHO100 kmGetchell TrendIndependence TrendLong CanyonARIZONAPaleoproterozoic/ TransitionalCrust118116114 ooo42o40o38o36 o112oCarlin TrendFigure 1 .1: Location of the Long Canyon deposit in northeastern Nevada. (Map after Tosdal et al., and Smith et al., 2012). The Sr 0.706 li e represe ts the dge of the rifted Paleozoic continental margin. Shaded bars represe t trends of Carlin-type deposits.1 .3 Long Canyon Deposit OverviewThe Long Canyon deposit is a sediment-hosted, disseminated gold deposit located 150 km east of the Carlin Trend in northeastern Nevada ( ). It is hosted within a sequence of carbonate rocks of Cambro-Ordovician age and a platform carbonate paleo-environmental setting. Specifically the host stratigraphy consists of laminated silty and micritic limestones overlain by an 80 metre thick dolomite sequence at the top of the Cambrian, overlain unconformably by Ordovician laminated and thin-bedded silty limestones intercalated with micritic limestone beds. The rock sequence records lower greenschist facies metamorphism and strong deformation by a Late Mesozoic orogenic event (Camilleri and Chamberlain 1997) which extended the dolomite layer into a series of elongate boudin blocks. Boudin necks are oriented SW-NE and vary from narrow cracks to gaps > 100m wide. Polyphase brecciation occurs on the margins of boudin necks. Breccia show evidence of 5a cataclastic origin within the dolomite cracks and strong solution brecciation within limestone-hosted breccia located along boudin block margins and cracks above and below the dolomite. These boudin necklines and breccia are a first order control on the distribution of Au mineralization at Long Canyon. Within the structural corridors defined by boudin necks, both mineralized and unmineralized mafic dykes and sills occur, typically <2m wide.The Long Canyon deposit hosts pervasive post-mineralization oxidation, which largely obscures the nature of hypogene mineralization. Fine-grained masses of disseminated Fe-oxide pseudomorphs, presumably after ore-stage pyrite, occur in areas of elevated Au, As, Hg, Tl, and Sb. Accompanying alteration includes decalcification of limestone, silicification and argillization of detrital siliciclastic minerals within silty laminae and dykes. Au mineralization occurs primarily within limestone-hosted breccias along dolomite boudin necklines, and locally within mafic dykes. Orezones form multiple sub-parallel and elongate bodies coincident with boudin necklines and brecciation. Mineralization in the deposit is currently defined for >3km along a northeast trend and contains a combined measured, indicated and inferred resource of 2.2 million ounces at and average grade of ~2.3 g/t Au (Smith, Rhys et al. 2012).1 .4 Thesis OrganizationThis thesis is organized in a classical format, with Chapter 3 and Chapter 4 to be re-worked and re-combined later into 1 to 2 papers for submission to a peer-reviewed journal. Chapter 1 provides an overview and introduction to the thesis subject, objectives and methods utilized. Chapter 2 is a literature review on the petrophysical and physicochemical controlling parameters on isotopic exchange processes between carbonate rocks and an infiltrating hydrothermal fluid. Chapter 3 describes the geology of the Long Canyon deposit, an understanding of which is critical for effectively interpreting patterns of fluid flow. This chapter is based primarily on studies from drillcore carried out during this thesis, synthesis of past work by Fronteer geologists and consultants, and the author’s inherited understanding of the deposit geology from previous employment while working for Fronteer Gold’s exploration team at Long Canyon. Chapter 4 presents the results of carbon and oxygen isotopic studies within carbonate wall rock from drillcore and surface samples. Results are used to interpret the development of fluid pathways, fluid evolution and the implications for using oxygen and carbon isotope studies in a mineral exploration program. Chapter 5 presents the conclusions of this thesis.6Chapter 2: Controls on Stable Isotope Depletion Patterns in Carbonate Rocks 2 .1 IntroductionFluid flow paths of hydrothermal systems infiltrating bedded carbonate rocks are controlled predominantly by heterogeneities in intrinsic rock permeability, fracture-enhanced permeability and reaction-enhanced permeability (Jamtveit and Yardley 1997; Baumgartner and Valley 2001; Cathles and Adams 2005). When studying a carbonate rock-hosted fossil hydrothermal system to characterize the extent, pathways, volume and mechanism of fluid flow, analysis of oxygen isotope ratio variation in carbonate minerals in wall-rocks can be a powerful tool (Jamtveit and Yardley 1997; Baumgartner and Valley 2001). During fluid-rock interaction isotopic exchange occurs, producing isotopic gradients along these pathways (Bowman, Willett et al. 1994; Gerdes, Baumgartner et al. 1995). Well constrained studies of oxygen isotope ratios paired with appropriate models and boundary conditions can help to elucidate the nature of fluid flow and fluid fluxes (Baumgartner and Valley 2001).This study reviews the petrophysical and physicochemical controlling parameters of a hydrothermal fluid infiltrating carbonate stratigraphy and the resulting stable isotope alteration front. Included are reviews on: 1. controls on fluid flow in carbonate sediments; 2. dissolution-precipitation kinetics of carbonate rocks; 3. oxygen isotope systems in carbonate rocks and natural waters and isotopic fractionation between them;4. isotopic transport modeling in hydrothermal systems infiltrating carbonate rocks.Many of these concepts are in themselves highly technical and complex fields of study and warrant much more than the brief review to be provided here. The purpose of this review is to summarize the relationships of the main principles involved in hydrothermal fluid-rock interaction and stable isotope fraction in carbonate rocks to provide the non-specialist reader with a background for interpreting the subject and conclusions of this thesis.  2 .2 Petrophysical Properties of Carbonate SedimentsConsidered the most critical and limiting rock property in controlling fluid flow during water-rock interaction (Jamtveit and Yardley 1997; Lucia 2007), permeability (k) controls the rate at which fluids can 7flow through the connected pore space of a porous medium (Cathles and Adams 2005). It is expressed by Darcy’s Law (Lucia 2007):Q Ak PL=µ∆         (1)where Q is rate of flow, k is permeability, μ is fluid viscosity, (ΔP/L) is the potential drop across a horizontal sample, and A is the cross-sectional area of the sample. Permeability is expressed as an area where, 1 darcy (D) is approximately 1 μm2 and typically used to describe sedimentary permeability as Log Permeability (m2) (Cathles and Adams 2005). Sedimentary permeabilities range from less than 0.01 millidarcy (md) to well over 1 darcy (Lucia 2007). Porosity (Ф), defined as pore volume divided by bulk mineral volume (Lucia 2007) or “effective porosity” as the porosity that is most available for fluid flow (ibid), is in turn the critical variable in describing permeability (Cathles and Adams 2005; Lucia 2007).Permeability is determined by the pore structure of the rock and any process that affects the pore structure affects the permeability (Jamtveit and Yardley 1997). Two critical aspects of pore structure are pore-throat size, and pore-size distribution. Pore-throat size is the limiting factor for fluid flow between interconnected pores, as narrow pore-throat size can impart significantly high adhesive force or surface tension between the fluid and minerals as to impede fluid flow (Lucia 2007). Pore-size distribution refers to the spatial distribution of all pore sizes, including pore-throats, throughout the rock (ibid). An even or uniform pore-size distribution is favourable for continuous permeability throughout the rock (Lonoy 2006). Porosity in sedimentary rocks is commonly interparticle (intergrain and intercrystal) (Choquette and Pray 1970; Lonoy 2006; Lucia 2007), though in carbonate sediments the presence of pore space within peloids and shells can increase the overall porosity (Dunham 1962). While in siliciclastic sediments, improved sorting increases porosity, in carbonate sediments porosity generally increases when sorting decreases (Lucia 2007). Lucia also notes that intergrain porosity decreases with closer packing, thus pore size is related to grain size, sorting and intergrain porosity, and therefore a function of the type of carbonate sediment and its post-depositional history (i.e. mudstone or packstone, diagenesis and/or metamorphism). Porosity and permeability reduction in ancient carbonate sediments results from a combination of processes related to burial diagenesis including: 1) mechanical and chemical compaction; and 2) 8carbonate cementation (Lonoy 2006; Lucia 2007). Compaction processes result in grain interpenetration, grain breaking, grain deformation and stylolitization, all of which reduce pore size (Lucia 2007). Carbonate cementation results in occlusion of pore space, typically leaving either patchy pore-size distribution or intercrystalline microporosity with significantly reduced permeability (Bjorlykke 1997; Lonoy 2006).Carbonate rocks and their fluid flow properties of particular interest in this review and in the broader thesis are ancient mudstones and mud-supported wackestones (Lucia 2007) of variable crystal size (and variably undergone low-grade metamorphism). Carbonate mud-sized particles are generally needle-shaped aragonite crystals around 5 μm which tend to stack openly leading to high initial porosity (Lucia 2007). However, when subjected to compaction, the aragonite crystals are highly unstable and tend to dissolve, contributing to further compaction and calcite cementation, and resulting in loss of porosity, pore-size and permeability (Lucia 2007). Studies have shown, summarized in (Lonoy 2006), that cemented carbonate rocks, particularly mud-supported/dominated, will typically have intercrystal microporosity and correspondingly low permeabilities. The preceding sections have summarized briefly petrophysical aspects of carbonate rocks which contribute to their intrinsic permeabiliy. However a rock’s bulk permeability reflects fluid flow rates resulting not only from fluid flow through rock matrix, as mentioned, but also fluid-focusing along fractures and faults, as well as dynamic permeability (Lonoy 2006) resulting from secondary dissolution and corrosion of pore-throats to increase permeability or mineral precipitation to reduce permeability during interaction with a hydrothermal fluid. The following sections will address these processes.2 .3 Hydrothermal Fluid Flow in Carbonate RocksFluid flow pathways and sources for aqueous fluids infiltrating carbonate sediments vary greatly depending on the age and geologic history of the sediments. In the case of Carlin-type deposits in Nevada, the rocks hosting most of the gold form part of an ancient carbonate platform/ramp facies package. At the time of mineralization these rocks say within the brittle upper crust at a depth of 2-5 kilometres, having previously undergone burial and low-grade metamorphism from orogeny-induced crustal thickening, subsequent uplift and now at the onset of crustal extension (Cline, Hofstra et al. 2005; Muntean, Cline et al. 2011). Shallow emplacement of magmas and circulating meteoric waters produced a mixed magmatic-meteoric and moderately CO2-rich fluid (2-4 mol %), rising under thermal buoyancy along 9extensional structures into the carbonate sediments. These fluids at depth were likely hydrostatically to suprahydrostatically pressured (Hofstra and Cline 2000). As has been discussed, fluid flow through sedimentary carbonates occurs in one of the following two ways (Frimmel 1992; Bjorlykke 1997): (1) porous or pervasive media flow (flow in the pore network of the rock matrix between sedimentary particles; or (2) fracture or channeled flow (flow along faults and fractures in rocks). Mass transport through pervasive media flow is controlled by advection, dispersion/diffusion, sorption on mineral surfaces and fluid-mineral reactions (Jamtveit and Yardley 1997). Fluid-mineral reactions are typically controlled by reaction kinetics, rather than the transport mechanism (Morse and Arvidson 2002), and will be dealt with in the following section. In order to form large carbonate-hosted hydrothermal ore bodies, a very large fluid flux is required (Hofstra and Cline 2000; Cathles and Adams 2005; Cox 2005). To invoke pervasive fluid flow as the primary fluid flow mechanism, fluids would need to pass through the pore space of the rock with sufficient fluid velocity to accommodate the fluid fluxes required, something that is unlikely without significant channelized flow for vertical fluid flux across stratigraphy (Jamtveit and Yardley 1997; Cathles and Adams 2005; Cox 2005). In the ancient carbonate sediment packages in Nevada, owing to their apparently low permeability and the amount of ore fluid required, it is likely that fluids travelled through the crust by fracture or fault flow (Frimmel 1992; Cline and Hofstra 2000; Hofstra and Cline 2000; Cox 2005), with pervasive media flow processes occurring on a more localized deposit scale when rock permeability is favourable. In ancient brittle carbonate rocks where pore-networks permeability is too low for continuous fluid flow, faults and fractures provide high permeability pathways for fluid (Jamtveit and Yardley 1997; Pedersen, Wangen et al. 1997). Fractures hosting fluid flow can be a variety of sizes and morphologies from: (1) interconnected networks of grain-scale fractures or cracks (Jamtveit and Yardley 1997; Pedersen, Wangen et al. 1997; Cox 2005);  (2) millimetre-scale fracture with permeabilities many orders of magnitude higher than surrounding sediments (Pedersen, Wangen et al. 1997); and (3) wide fault damage zones of fracturing with extremely high transmissivity capabilities for fluids. Fracture-width is obviously a controlling factor for fluid flow with wider aperture corresponding to increased fluid flow volumes and velocities (Pedersen, Wangen et al. 1997). High fluid velocity can be critical to continued flow through fractures. If fluid velocities are slow enough to allow significant heat loss to occur through conduction to the surrounding sediments, dropping silica solubility of hydrothermal fluids may cause precipitation and blockage of fractures (Deloule and Turcotte 1989; Cathles and Adams 2005). There is however a 10competing process a where a cooling, CO2-bearing fluid will produce H+ ions that attack and dissolve wall rock carbonate when quartz is precipitated (Simmons and Christenson 1994). Carbonated dissolution will typically occurs at a much faster rate, maintaining permeability (Bjorlykke 1997). This reaction-enhanced permeability will be addressed in the following section on the dissolution-precipitation kinetics of calcite – a critical mechanism for hydrothermal fluid infiltration of carbonate rocks.2 .4 Dissolution-Precipitation Kinetics of Carbonate RocksThe dissolution-precipitation kinetics of carbonate rocks, with a primary focus on calcite (limestone), are examined in this study for several reasons: (1) the dissolution and reprecipitation of CaCO3 or MgCO3 at the crystal-fluid interface is a primary control on enabling isotopic fractionation of an infiltrating fluid and the development of isotopic alteration gradients in host carbonate rocks at temperatures where diffusion is inhibited (Bowman, Willett et al. 1994; Simmons and Christenson 1994; Baumgartner and Valley 2001; DePaolo 2011); and (2) the degree of dissolution or precipitation of carbonate as well as associated isotopic gradients within fluid pathways can allow for inference on fluid composition and temperature history (Deloule and Turcotte 1989; Sanford and Konikow 1989; Bjorlykke and Egeberg 1993; Simmons and Christenson 1994; Pedersen, Wangen et al. 1997; Morse and Arvidson 2002). While there are numerous studies exploring the dissolution rates of calcite and other carbonate minerals, the majority of these do not consider natural conditions of porous or fractured media fluid flow (Arakaki and Mucci 1995) and there are surprisingly few studies in the literature which directly address permeability-porosity enhancement from mineral dissolution during acidic fluid flow in carbonate rocks (Baumgartner, Gerdes et al. 1997). Key factors influencing these phenomena include: (1) the carbonic acid system and the influence of CO2 concentration within the infiltrating fluid; (2) concepts of kinetic equilibrium, disequilibrium and saturation state; (3) rate-controlling mechanisms on the diffusion of reactants and products through solution and surface reactions; and (4) effective pH and temperature ranges influencing dissolution and precipitation.Calcite dissolution and precipitation reactions and their kinetics are described extensively in the literature (Weyl 1958; Lund, Fogler et al. 1975; Plummer, Wigley et al. 1978; Sjöberg and Rickard 1984; Morse and Mackenzie 1990; Lasaga 1998; Morse and Arvidson 2002) among many others and are summarized here. Natural waters have a wide range of pH, PCO2 s and temperatures; all parameters which control saturation state (Ω), or the degree of saturation with respect to carbonate minerals, defined as the ratio of 11the ion activity product to the solubility product (for dissociation of CaCO3 to Ca+ and CO32-, ksp = [Ca+][CO32-]) (Morse and Mackenzie 1990; Morse and Arvidson 2002). ΩcalciteCa COspa ak=+ −232         (2)Simply put, if Ω = 1, the solid and solution are in equilibrium; if Ω < 1, the solution is undersaturated and mineral dissolution can occur; and if  Ω > 1, the solution is supersaturated and mineral precipitation should occur (Morse and Mackenzie 1990). The following set of equations (3, 4, and 5) represent the heterogeneous multi-step CO2-carbonic acid system (Morse and Mackenzie 1990; Arakaki and Mucci 1995):CO H O H COaq l aq2 2 2 3( ) ( ) ( )+ ⇔        (3)H CO HCO Haq aq aq2 3 3( ) ( ) ( )⇔ +− +       (4)HCO CO Haq aq aq3 32− − +⇔ +( ) ( ) ( )        (5)The importance of dissolved CO2 concentration on fluid pH and aCO32- is apparent from these equations (Arakaki and Mucci 1995). There are several elementary reaction mechanisms described by equations (6), (7), and (8), attributed to calcite dissolution in a wide range of pH andPCO2 (Plummer, Wigley et al. 1978; Chou, Garrels et al. 1989), though in a moderate-weakly acidic environment under consideration the equation representing the reversible dissolution of calcite via protonation by hydrogen ions (eq 6) is most representative (Lund, Fogler et al. 1975; Arakaki and Mucci 1995)CaCO H Ca HCOkI323+ ++ + −⇔          (6)CaCO H CO Ca HCOkII3 2 30 232+ +⇔ + −        (7)CaCO Ca COkkIIIIII3232−⇔ + −+         (8)with the net rate of dissolution given by:R k a k a k k a aI H II H CO III III Ca HCO= + + ++ + −−2 30 2 3       (9)where kI, kII, and kIII, are first order and temperature dependent rate constants and k-III is a second order rate constant and a function of PCO2 (Plummer, Wigley et al. 1978; Arakaki and Mucci 1995). Plummer 12(1978) observed that the backward rate k-III dominated eq (9) during calcite precipitation reactions, hence the importance of the activity of calcium and carbonate ions on saturation state and its impact on calcite precipitation. These general equations (6, 7, 8 and 9) represent chemical kinetic models describing the characteristics of a solid interacting with a bulk solution. However, calcite-solution interaction is more complex with intermediate chemical species forming as complexes on mineral surfaces, and a diffusion boundary layer (DBL) of <5 - >30 μm thick present between the mineral surface and bulk solution (depending on hydrodynamic conditions this may be thinner and more turbulent rather than stagnant) where products and reactants reside often in near-equilibrium conditions and chemical species concentrations can be orders of magnitude different from the bulk solution (e.g., Figure 2.1) (Plummer, Wigley et al. 1978; Sjöberg and Rickard 1984; Chou, Garrels et al. 1989; Arakaki and Mucci 1995; Morse and Arvidson 2002). FluidMineralD i f f u s i o nB o u n d a r y  L a y e r  ( D B L )Figure 2 .1: Schematic diagram of mineral-fluid interface during pervasive fluid flow. Dissolved species can be transported to and from mineral surfaces by either advection or diffusion as shown by the solid and dashed arrows, respectively. The fluid adjacent to the mineral interface is often an immobile boundary layer fluid; requiring diffusion through this boundary layer to transport elements to and from the reactive surface. From Schott et al. (2009). Sjoberg’s (1984) study provided the framework for subsequent work by describing three pH regimes of low pH (<4), moderate pH (4-5.5) and higher pH (>5.5) as displaying transport controlled kinetics, mixed kinetics and chemically controlled kinetics, respectively, for calcite dissolution (Figure 2.2).13Figure 2 .2: Schematic representation of rate controlling mechanisms for calcite dissolution as a function of pH and temperature. From Morse et al. (2002).At low pH (<4), the surface reaction is so fast that the diffusion of  hydrogen ions through the DBL to the surface becomes the rate-limiting step (Lund, Fogler et al. 1975) and the dissolution rate is almost directly proportional to the H+ concentration, thus kI becomes the dominant term in eq (9) (e.g. transport controlled). At higher pH (>5.5) the dissolution rate is independent of H+ concentration and calcite dissolution/precipitation rate is controlled by mineral surface chemical complexation reactions forming on Ca+, CO32- sites and interaction with H2O (Sjöberg and Rickard 1984; Arakaki and Mucci 1995; Morse and Arvidson 2002). Between these pH ranges there exists a transition kinetic regime of mixed H+-dependence and H+-independence (Figure 2.2). In most natural environments of hydrothermal fluid-carbonate rock interaction, particularly in weak-moderately acidic (4.5<pH<6) and porous and fractured media flow, the kinetics of carbonate mineral dissolution-precipitation reactions are most likely the controlling step, i.e. surface controlled (Plummer, Wigley et al. 1978; Arakaki and Mucci 1995).The effects of temperature and pressure on calcite dissolution and precipitation have predominantly been studied in low pH conditions in systems far from equilibrium e.g. “transport controlled/ H+ dependent” and seawater. Temperature effects in relation to Sjoberg’s (1984) pH regimes (Figure 2.2) is to shift their bound-aries to lower pH with increasing temperature, seemingly an extremely significant effect when considering the pH boundary for H+ dependence/independence of reaction kinetics and that natural systems would in-creasingly exhibit surface controlled kinetics. Morse in his (2002) review paper cites a limited knowledge of temperature influence on calcite dissolution kinetics in the surface-controlled near equilibrium region common in natural systems and a subsequent substantially limited ability to predict calcite dissolution kinetics in the “real world”. However, there are some fundamental temperature and pressure effects which 14can change the bulk fluid composition, leading to dissolution-precipitation reactions. The most important temperature effect for CO2 – H2O low-temperature hydrothermal systems such as we are considering is the solubility of carbonate from 100-300°C as determined by (Ellis 1959).  Ellis first determined through exper-iments of varying temperature and PCO2 that the solubility of calcite decreases with increasing temperature, as does the solubility of CO2 gas in water. These findings were later refined by (Plummer and Busenberg 1982) for the temperature range 0 - 90°C by determining the solubilities of chemical species involved in the carbonic acid system. Plummer found that from 0 to 50-75°C solubilities increased, after which they began to decline rapidly with higher temperatures in agreement with Ellis’ findings. The following reaction:2 323 2 2HCO Ca CaCO H O CO− ++ ⇔ + +                (10)describes heating of a CO2-rich fluid near calcite saturation, where decreasing solubilities of carbon dioxide and calcite drive equation (10) to the right, precipitating calcite and releasing gaseous carbon dioxide. Incidentally, this same reaction describes precipitation of calcite in a boiling zone within a hydrothermal system, as decreasing hydrostatic pressure on a rising CO2-rich fluid decreases the solubility of carbon dioxide in fluid, releasing gas and precipitating platy calcite (Simmons and Christenson 1994). Conversely, cooling of a CO2-bearing fluid increases the solubility of carbon dioxide, causing eqs (3 and 4) to proceed to the right, H2CO3 dissociates producing H+ ions, the fluid pH lowers and acidic fluids will now either dissolve carbonate wallrock, or attack feldspars through hydrolysis reactions to produce clay alteration (Fournier ; Simmons and Christenson 1994). Paragenetically-late calcite veins might reflect cooler meteoric waters, or a collapsing hydrothermal system, already near calcite saturation and migrating down recently heated fluid pathways to become supersaturated and precipitate calcite, as described in the Broadlands-Ohaaki geothermal system (Simmons and Christenson 1994). This study reviews the controlling mechanism(s) involved in a carbonate exchanging its isotopically-heavy oxygen 18O for isotopically-light oxygen 16O with an infiltrating fluid via this reversible reaction (O’Neil, Clayton et al. 1969):13163 218 13183 216CaC O H O CaC O H O+ ⇔ + ,                (11)where the net observable result is varying amounts of carbonate dissolution and a 18O-depleted rock. Kinetic equilibrium theory describes a reaction between a solid and a solution where exchange of mass 15occurs simultaneously in both via a forward dissolution reaction and a reverse precipitation reaction. At equilibrium the rates of the opposing reaction are equal and thus no change is observable even though dissolution and precipitation of carbonate is continually occurring (Morse and Arvidson 2002). For an observable net result of dissolution, the system must be in chemical disequilibrium to some degree, where the forward reaction proceeds at a faster rate, consuming carbonate reactants (ibid), yet throughout, the reverse reaction is precipitating 16O into new carbonate.2 .5 Light Stable Isotopes in Carbonates and Natural WatersOxygen has three stable isotopes (atoms whose nuclei contain the same number of protons but a different number of neutrons) in the following abundances: 16O – 99.763%; 17O – 0.0375%; 18O – 0.1995%. Carbon has two stable isotopes: 12C – 98.89%; 13O – 1.11%. Hydrogen has two stable isotopes: 1H – 99.9844%; 2D – 0.0156% (Hoefs 2009). The isotopic composition of a substance is expressed as the ratio R of the abundance of heavy isotopes over light isotopes (e.g. 18O/16O) (Faure and Mensing 2005). Isotopic compositions are measured by mass or laser spectroscopy (Barker, Dipple et al. 2011) and reported relative to an internationally accepted standard, which for oxygen is Vienna Standard Mean Ocean Water (VSMOW) and carbon (and sometimes oxygen) is Vienna Peedee Belemnite (VPDB) (Coplen 1994; Hoefs 2009). R values are expressed in terms of a parameter delta (δ), which is expressed in eq (12) for oxygen:δ 18 310OR RRspl stdstd=−× 0                  (12)where Rspl = 18O/16O ratio of the sample, Rstd = 18O/16O ratio of the standard (VSMOW), expressed in terms of per mill (Faure and Mensing 2005). δ18O values (X) reported relative to VSMOW and VPDB are related by eq (13) (Morse and Mackenzie 1990):δ δ18 181 03086 30 86O x OVSMOW VPDB( ) ( ). .= +                (13)When undergoing phase changes through chemical reactions, the difference in molecular weight imparts a higher velocity to lighter isotopes, making them more reactive in chemical reactions dependent on temperature. When it comes to precipitation reactions from aqueous solutions, heavy oxygen isotopes will preferentially precipitate into the solid phase (Criss and Taylor 1986; DePaolo 2011). This partitioning 16of isotopes between phases resulting in different isotope ratios is called “isotopic fractionation” (Hoefs 2009). The main mechanisms by which isotopic fractionation can occur are: 1. equilibrium isotopic exchange reaction – e.g. a calcite crystal and calcite-saturated solution in equilibrium where dissolution and precipitation reactions are constantly occurring at an equal rate, progressively increasing the ratio of light oxygen isotopes in the calcite relative to solution, until the original isotopic ratios have “equilibrated”2. kinetic isotope effects occur as a result of differences in reaction rates of isotopic components where a chemical reaction is incomplete or unidirectional owing to a variety of possible reasons including temperature and/or chemical compositions – e.g. a reaction between calcite and fluid where dissolution and precipitation reactions occur to some degree until fluid becomes buffered by the rock and unable to isotopically exchange, resulting in incomplete exchange of isotopic components, relative to an equilibrium state where rock and fluid are exchanged to a state of isotopic equilibrium. The extent of fractionation between two stable isotopes of the same element is controlled by the difference in mass and the temperature of the environment in which the chemical reaction or phase change is occurring (Faure and Mensing 2005). Fractionation under equilibrium conditions is quantitatively described by the fractionation factor (α), which is defined as the ratio of the isotopic ratios of solid to liquid involved in fluid-mineral exchange (ibid). Fractionation factors for many key mineral-fluid systems have been calculated by researchers (O’Neil, Clayton et al. 1969; O’Neil 1986) at equilibrium conditions and STP. Fractionation factor values decrease with increasing temperature (ibid). Bulk oxygen isotope ratios within ancient carbonate sediments are the result of what is often a complex geological history, involving interaction with numerous oxygen sources with which isotopic exchange occurred. These can include primary precipitation from seawater; diagenetic processes including pressure and temperature driven recrystallization through microsolution-precipitation (Morse and Mackenzie 1990); interaction with pore waters causing dissolution or carbonate cementation (Lucia 2007); recrystallization through metamorphic fluids (Baumgartner and Valley 2001); circulating meteoric waters causing dissolution and/or new carbonate deposition/veining (Sanford and Konikow 1989); and hydrothermal fluids sourced from meteoric waters equilibrating with magmas at depth (Muntean, Cline et al. 2011) infiltrating and isotopically exchanging with host carbonate rocks (Radtke, Rye et al. 1980). Carbonate rocks analyzed today derive their primary oxygen isotope ratios from the parent seawater, but are likely altered from these ratios during their diagenetic and metamorphic history (Shields and 17Veizer 2002). Several compilation studies (Veizer, Ala et al. 1999; Shields and Veizer 2002) have shown that the Phanerozoic δ18O values for carbonate rocks have been steadily increasing, from about +24‰ (VSMOW) at the onset of the Cambrian to about +30‰ (VSMOW) at present. This mean trend line of a large globally-derived database is generally considered to reflect oxygen isotope evolution of seawater throughout the Phanerozoic. Oxygen isotope values in carbonate rocks can be exposed to numerous events resulting in isotopic exchange during their geologic history, thus, when determining background (primary) isotopic values for host carbonate rocks to a hydrothermal system, background values must be established (assumed to be as near to primary as possible) locally and/or regionally from the most visually or geochemically unaltered sediments as possible. In addition, globally determined values can assist in providing a useful baseline for comparison. The 13C values in marine carbonate rocks closely resemble the dissolved bicarbonate from which it precipitated (Hoefs 2009). The δ13C values of limestones for the last 3.5Ga have fluctuated with regular secular variation within 0±3‰ (VPDB) (Veizer, Ala et al. 1999). This is directly related to carbonate sedimentation and burial rates with 12C preferentially removed from seawater into carbonate rocks. Meteoric waters are depleted in δ18O relative to SMOW; the result of a multi-step process: (1) evaporation from surface waters concentrates 16O in the vapor phase; (2) successive precipitation events occur from water vapor (clouds) as moist air cools and 18O-enriched rain falls (still depleted in 18O relative to SMOW); (3) the resulting water vapor with increasingly lower δ18O values relative to SMOW, remains to fall as rain further inland (Hoefs 2009). The δ18O composition of mean worldwide precipitation is estimated to be -4‰ (Craig and Gordon 1965). In meteoric water processes, hydrogen isotopes are fractionated in proportion to oxygen isotopes defining the following relationship:δ δD 8= +18 10O                   (14) which is known as the “Global Meteoric Water Line” (Hoefs 2009). Analysis of clays and hydrous minerals (micas) in the upper crust for D/H and 18O/16O ratios can in some cases be used to infer fluid source (magmatic or meteoric/hydrothermal) as well as differentiating between magmatic hydrothermal alteration and oxidizing meteoric fluids when ratios are compared to established ranges for natural waters 18(Hofstra, Snee et al. 1999; Faure and Mensing 2005). Magmatic waters are commonly estimated to be 6 to 10‰ for δ18O values and -50 to -80‰ for δD values (Hoefs 2009). In Carlin-type deposits models of  ore fluids derived from magmatic waters mixing with Cenozoic circulating meteoric waters, δ18O values are calculated between ~ 0 to 5‰ (Cline and Hofstra 2000).2 .6 Isotopic Fluid Transport Mechanisms and Modeling in Carbonate RocksWhen hydrothermal fluids infiltrate a rock mass, fluid-mineral isotopic and chemical disequilibrium typically exists, and exchange reactions will occur to varying degrees of partial to total equilibration between them along the fluid flow path (Blattner and Lassey 1989; Bowman, Willett et al. 1994; Gerdes, Baumgartner et al. 1995; Baumgartner and Valley 2001). Stable isotope fronts develop within the flow path, separating rocks upstream that contain mineral constituents in isotopic equilibrium with infiltrating fluids, from rocks downstream which host infiltrating fluids which have equilibrated with the isotopic composition of the rock (e.g. Figure 2.3). These upstream and downstream regions separated by an isotopic ratio gradient or front are referred to as fluid dominated and rock dominated regions, respectively (Baumgartner and Valley 2001). The shape of the isotopic front depends on (1) the magnitude of fluid flux/infiltration; (2) reaction kinetics for isotopic exchange between fluids and minerals; (3) the diffusive or dispersive characteristics of the rock matrix and structure; and (4) the relative isotopic compositions of the fluid and minerals in the rock (degree of initial disequilibrium) (Frimmel 1992; Bowman, Willett et al. 1994; Baumgartner and Valley 2001). Several important mechanisms of stable isotope transport by an infiltrating fluid in rock are diffusion, dispersion and advection, described in depth by (Baumgartner and Valley 2001). Diffusion is typically associated with stagnant fluid environments or when fluids are absent in high temperature metamorphic environments and involves isotope migration along or through grain boundaries where chemical potential gradients of isotopes drive exchange. However for hydrothermal fluid infiltration, particularly in lower temperature regimes and those with higher fluid velocities, the bulk of isotope and other dissolved solutes transport occurs through advection, i.e., mass transport in a moving fluid phase. Dispersion is always a part of fluid infiltration, as fluid follows tortuous paths of different length as it flows through cracks or along grain boundaries. Different path lengths and variable isotopic exchange with grain boundaries by diffusion or solution-precipitation reactions results in fluids remixing at intersections with varied isotopic compositions (Baumgartner and Valley 2001). This effect of hydrodynamic and chemical dispersion (Dipple 1998) from heterogeneities in permeability and mineral chemistry is attributed to significant isotopic ratio variation in oxygen depletion patterns of fossil 19hydrothermal systems (Bowman, Willett et al. 1994; Gerdes, Baumgartner et al. 1995). Channeled fluid flow at shallow crustal levels is basically a largely dispersive process (Frimmel 1992).distanceδ18Orockδ18Ouid isotopic frontt3t2t1t0baFigure 2 .3: (a) Schematic representation of continuous advective fluid flow from a fluid reservoir into a reactive rock mass. Fluid-mineral reaction along flow path involves kinetically controlled oxygen isotope exchange. (b) Changes in δ18O in rock and fluid represented by isotopic fronts as a function of distance along the transport path. Isotopic front migrates along flow path with time. Modified from Cox (2007).The primary parameters of interest used in transport modeling of isotope exchange front development (continuous flow path instead of closed system) that relate the above-mentioned characteristics of fluid flow are the Peclet number and the Damkohler number (Lassey and Blattner 1988). The Peclet number NPe gives the relative importance of advective mass transport to diffusive transport in the system and is defined as:NqLnDPe=                    (15)where q is the specific fluid flux, n is the porosity, D is the dispersion coefficient (which contains a factor for diffusion) and L is an arbitrary scaling distance (Bowman, Willett et al. 1994). The Damkohler number gives the relative importance of mass exchange by fluid-rock reaction to mass transported by advection and measures the relative importance of kinetic (non-equilibrium) effects (Blattner and Lassey 1989; Bowman, Willett et al. 1994) and is defined as:20Nk nLqkD k=                    (16)where kk is the reverse reaction rate for an equation like (11) describing oxygen exchange between H16O18O and CaC18O 16O2. A large Damkohler number results from high reaction rates relative to transport rates, implying near-equilibrium exchange. A very small Damkohler number, on the other hand, results from low reaction rates relative to transport rates and thus implying little exchange, or significant kinetic effects (Bowman, Willett et al. 1994). These equations along with definitions for dimensionless distance and time are used by (Bowman and Willett 1991) to calculate solutions to differential equations describing one dimensional transport models of an idealized fluid flow system. Modeling of fluid-rock interaction through one dimensional fluid-transport models in terms of these parameters generally considers two endmember regimes of non-reactive and near-equilibrium exchange with Damkohler numbers of near-zero and infinity respectively, and an intermediate (and geologically realistic) model of non-equilibrium, kinetically-controlled water-rock interaction (Bowman, Willett et al. 1994). Equilibrium exchange modeling is presented in Figure 2.4a and a kinetically-controlled model in Figure 2.4b. These models describe the shape of isotopic fronts developed along an arbitrary distance in the fluid flow system (Lassey and Blattner 1988; Bowman and Willett 1991; Frimmel 1992; Bowman, Willett et al. 1994; Gerdes, Baumgartner et al. 1995). High Damkohler numbers (fast reaction) produce steep alteration fronts and slow reaction rates resulting in broader, more diffuse isotope fronts with rock oxygen isotopic signatures far from equilibrium with infiltrating fluids. Earlier attempts to quantify total fluid fluxes during the life of a hydrothermal system as water-rock ratios (W/R) based on the observed isotopic oxygen shift in the rocks (Taylor 1974) assumed complete equilibration between fluid and rock within a closed system, with any non-equilibrium exchange resulting in underestimation of W/R ratio (Bowman, Willett et al. 1994; Hoefs 2009). Discussions surrounding interpretation of paleohydrothermal systems based on W/R ratios should therefore be approached cautiously, as isotopic ratios are often the result of kinetically-controlled exchange, temperature gradients, or simply within a rock-dominated portion of the flow path (Bowman, Willett et al. 1994; Baumgartner and Valley 2001). W/R ratio is also highly dependent on the size of the defined system (the theoretical volume of rock). Additionally, describing fluid volumes in relation to unit mass of rock is inappropriate for transport equations, which use Darcy flux units of moles or m3 water/m2/sec.21446 Baumgartner & Valley relative to fluid infiltration rates, and with increasing values local equilibrium is approached.  The general kinetic-dispersion-diffusion-infiltration equation. Equations (47) and (48) can be inserted into Equation (46), along with the definitions of dimensionless distance and time: w c_f18OwW  1NPew2 c_f18Owz2 w c_f18Owz  (1 )) NDk KDk c_f18O  c_k18O§©¨·¹¸k  1n¦ 1NPew2 c_ f18Owz2 wc_ f18Owz  (1))ND KD c_f18O  c_k18O§©¨·¹¸  (51)                      The summation included for completeness is over all mineral species k. It reduces to the last term for the case of a single fluid and mineral species containing oxygen. Figure 13 shows the solutions of the equation for different Peclet and Damköhler I numbers calculated for a fluid composed of water and a one-dimensional rock column. Bowman et al. (1994) presented many more examples. The curves shown in Figure 13a are solutions for an infinite Damköhler I number at a Peclet number of 100. In this case of rapid Figure 13. Normalized stable isotope composition of a rock column infiltratedby a reactive fluid. (a) The solution assumes local equilibrium, with a Pecletnumber (NPe) of 100 and infinite Damkohler I number. (b) solution calculatedfor the case of a Peclet number of 100 and a Damköhler I number (ND) of 1. The dimensionless parameters are given by: normalized concentration c_  (G r18O G f18 O ) / (G i18 O  G f18 O ) ; distance z  x / L ; dimensionless time W  t) / (X f L )  (after Bowman and Willet 1991—note that captions for their Figures 1 and 3 are switched). (δ18Or-δ18Of)(δ18Oi-δ18Of)distancedistance(δ18Or-δ18Of)(δ18Oi-δ18Of)Figure 2 .4: Normalized stabl isotope composition of a rock colum infiltrated by a reactive fluid over di-mensionless distance. Isotopic fronts represent rock c mn δ18O at that dim nsionless time. Note isotopic fronts progress along rock column with time. (a) Peclet number NPe = 100 implying advective infiltration. Infinite Damkohler number (ND) implying a local equilibrium re ction that is transport controlled instead of kinetically controlled. (b) NPe = 100, ND= 1 implying strongly kinetically-co trolled fluid-min ral reaction. At the same fluid transport rate as in (a), the kinetic control on mineral reactions does not allow for signifi-ant isotopic exchange and development of large isotopic gradients within the fluid-flow path. Modified from Baumgartner &Valley (2001). δ18Or, δ18Of and δ18Oi refer to the 18O/16O ratio of the rock at time (τ) and distance, the rock in exchange equilibrium with the infiltrating fluid, and the initial rock, respectively.Transport modeling theory has largely been developed and successfully applied to describe natural environments in the context of contact metamorphic environments wher  distance from initial fluid inlet to points of interest in an isotopic profile is easily quantifiable and a particular rock’s isotopic composition can be defined in relation to the fluid flow path and isotopic front (Dipple and Ferry 1992; Bowman, Willett et al. 1994; Gerdes, Baumgartner et al. 1995). In many natural systems however, particularly lower temperature hydrothermal systems where the magmatic heat source is deeply buried and the scale of the system is unknown, boundary conditions such as flow-path length and direction can be difficult to define. Often the only study section available of a paleohydrothermal system will not encompass the entire flow path (i.e., lack of resolution when only viewing fluid-dominated portions), may be oriented obliquely to the flow direction rather than parallel, or may only provide a view to fluid flow within the sides of the flow path. These sides, termed ‘dangling elements’ (Cox 2007), are fluid pathways (cracks 22or permeable rock layers) which are connected to the main fluid channelway or ‘backbone’, but consist of lower overall permeability and therefore lower fluid flux. Application of transport models to such systems should be used with caution while understanding the limitations, although certain parameters of fluid flow conditions within the system may still be discernable from investigating isotopic and/or geochemical alteration patterns within these environments. Basic transport modeling equations typically contain a single variable for porosity, thus describing the system’s porosity as homogenous, though in natural systems this is usually not the case. Due to heterogeneous permeabilities of infiltrated sedimentary layers, there may often be multiple main channelways within a system, each with their own fluid flow characteristics along their sides. Combined with the effects of hydrodynamic dispersion on isotopic ratios (Bowman, Willett et al. 1994), this can produce lobate and irregular patterns of isotopic alteration (Gerdes, Baumgartner et al. 1995) which may confound understanding of the fluid flow system. Measurements of isotope ratios at hand sample scale allow for determination of dominant fluid flow type (Frimmel 1992) and estimates of reaction kinetics by measuring infiltration distance from fracture surface into boundary of fluid filled fracture. While idealized theoretical models may have unrealistic or impossible parameters for natural systems (Frimmel 1992), they may still be used effectively to make educated inferences about characteristic fluid flow parameters of a fossil hydrothermal system from its preserved isotopic profile in the surrounding rocks. In turn, this insight into fluid flow pathways observed at the hand sample and deposit scale can be applied to regional isotopic ‘hot spots’ as a means of vectoring towards potential hydrothermal ore bodies.23Chapter 3: Geology of the Long Canyon Deposit3 .1 IntroductionThis chapter presents the geology of the Long Canyon deposit through studies of the host lithologies, structure, mineralization and alteration. The ultimate aim of this thesis is to map fluid flow paths through the deposit using oxygen isotope depletion and trace element metatsomatism in the host rocks. By establishing the geological framework of permeable pathways for fluid flow and alteration, patterns of oxygen isotope depletion and trace element metasomatism can be more easily interpreted in Chapter 4.3 .2 Previous Work and Acknowledgment of Source of ConceptsStratigraphic sections around the deposit were measured by consulting geologist J. Thorson for Fronteer Gold (Thorson 2007; Thorson 2008) following work by Thorman (1970). Mapping was carried out Smith (2008) and Smith and Thompson (2009). Lithostratigraphic subunit nomenclature used in this study was developed by Fronteer Gold geologists (including this author). Mapping over the Long Canyon deposit by Smith (2008) first identified boudinage in carbonate rock units, which led to the structural model that would ultimately be used for drill targeting to expand the deposit, and is used as the basis of this study. Long Canyon was identified as a solution collapse breccia-hosted Au deposit occurred early on, but the relative timing of breccias to Au mineralization was unclear. A consultant report on the structural setting of mineralization (Rhys and Ross 2010) built upon Fronteer Gold geologist’s and this author’s current understanding of the deposit at the time. Their work helped demonstrate the complexity of brecciation history at Long Canyon and only a comparatively brief review is presented here. Additional source material includes a paper by Smith et al. (2010), to which this author contributed geological interpretation while employed by Fronteer Gold and internal technical reports by Smith (2011). Geological solids used in this study were created by Fronteer Gold staff and consultants, an effort to which the author contributed while employed by Fronteer. 3 .3 Sampling and Research MethodsThe geological information presented within this chapter represents a synthesis of the author’s geological understanding developed while employed by Fronteer Gold on the Long Canyon project from September 2008 to September 2009, as well as further understanding developed during this study from drill core logging and sampling, thin section and geochemical analysis. Data for this study is largely generated from the study and interpretation of drill core, with minor surface work. In this study 440 hand samples 24were taken from 16 drillholes across the deposit for macroscopic description, preparation of 118 thin and polished sections, clay identification by ASD and XRF, clay δD and δ18O isotope analysis and SEM examination of ore-stage pyrites, all presented in this chapter; and δ18O and δ13C carbonate isotope analysis, presented in Chapter 4. Fronteer drillhole core photos were used extensively for examination of stratigraphic, fault, vein, breccia and alteration and mineralization relationships. Several detailed cross-sections on local section lines L12750N and L12800N, displaying rock unit and breccia distribution, were created from core photo analysis and export of drillholes from Fronteer’s Gemcom GEMS drill database. These cross-sections are used throughout the thesis as the underpinning for geochemical and isotopic data interpretation. Fronteer’s drillhole geochemical database was used in conjunction with geochemical analyses of 273 drillcore hand samples selected by the author in this study. Fire assay Au and Ag and multi-element ICP lithogeochemical analyses were carried out by ALS Chemex in Vancouver, BC, for both datasets. Geological descriptions are based largely upon macroscopic and microscopic observations by the author from drillcore except as stated in Section 3.2. Observations and interpretations generally corroborate observations from past technical reports by Moira Smith of Fronteer Gold and Dave Rhys and Kika Ross of Panterra Geosciences. Research into the genesis of dolomitization, boudinage and mineralization and alteration types were carried out by the author at UBC. This study did not set out to determine a detailed paragenetic sequence for generations of calcite veining present at Long Canyon, and sampling protocol generally sought to avoid calcite veining in order to maintain consistency with bulk geochemical and isotopic sampling. Observations and interpretations are the author’s own, except where stated otherwise. 3 .4 Regional Geologic Setting and HistoryThe Long Canyon deposit is situated in northeast Nevada on the eastern flank of the Pequop Range. The Pequops are underlain by Lower – Middle Paleozoic strata, consisting of shelf carbonate rocks formed on the episodically shallowing – emergent passive margin of continental North America, with contributions of terrigenous siliciclastic sediments derived from the craton (Coats 1987; Cook and Corboy 2004). Carbonate sediments were formed on the middle to inner carbonate shelf, which extended from Western Utah to Eastern Nevada, with deepening stratigraphy westward to the platform margin, slope and basin facies in Central Nevada (Figure 3.1)(Cook and Corboy 2004). The rock sequence records lower greenschist facies metamorphism and structures from one or more contractional deformation events, which Camilleri and Chamberlain (1997) attributed to crustal 25thickening and tectonic burial during southeast-directed thrusting occurring in the hinterland of the Late Mesozoic Sevier orogenic belt. Camilleri and Chamberlain (1997) note metamorphic grade increases to upper amphibolite facies westward from the Pequop Range to the Wood Hills, indicating deeper burial to the west. The rocks were exhumed by normal faulting along the west-rooted Pequop fault during the Cenozoic, likely related to the decompression-related detachment complex forming in the Ruby Mountains and East Humboldt Range to the west of Long Canyon (Mueller and Snoke 1993; Camilleri and Chamberlain 1997). Eocene volcanics were deposited from 41-39 Ma from magmatic centers across much of northern Nevada (Brooks, Thorman et al. 1995), and are present in the Pequops north of Long Canyon. This age of volcanism is coincident with the formation of gold mineralization in the Carlin Trend located ~150 km to the west (Cline, Hofstra et al. 2005). Further exhumation of Paleozic strata occurred along the west-rooted Mary’s River fault system after Eocene volcanism (Mueller and Snoke 1993). Subsequent Mid-Miocene – Holocene extension resulting in development of the characteristic basin and range topography of Nevada and Utah; mountain ranges separated by alluvial valleys bounded by normal slip ‘range bounding’ faults (Snow and Wernicke 2000; McQuarrie and Wernicke 2005). The Long Canyon deposit is located in the footwall of a range bounding fault, approximately 1,500 – 500 metres west of the fault (Smith 2011).3 .5 Deposit Scale GeologyThe Late Mesozoic orogenic event extended an 80 metre-thick dolomite layer into a series of elongate boudin blocks. Boudin necks are oriented SW-NE and vary from narrow cracks to gaps > 100m wide. Polyphase brecciation occurs on the margins of boudin necks. Breccia show evidence of a cataclastic origin within the dolomite cracks and strong solution brecciation within limestone-hosted breccia located along boudin block margins and cracks above and below the dolomite. These boudin necklines and breccia are a first order control on the distribution of Au mineralization at Long Canyon. There are three main types of breccia present across the Long Canyon deposit. In decreasing relative age determined from crosscutting relationships they are: 1) dolomite-hosted breccias, most often occurring with a planar, high-angle fabric of angular clasts within dolomite blocks and as masses of chaotic breccias along dolomite boudin block noses. 2) Multi-generational solution collapse breccias within limestones, often mineralized and spatially associated with boudin block margins and sites where dolomite hosted cataclastic breccias terminate in limestone above and below. 3) Coarse calcite cement breccias which overprint all other breccia species and exhibit open-space calcite growth.26Faulting across the Long Canyon property likely reflects the multiple contraction and extensional events of its history with both reverse and normal faulting of stratigraphy. Several faults accommodate hundreds of metres of normal stratigraphic offset around the property (Figure 3.3). Thrust faults are mapped in Pogonip rocks as top to the southeast (Smith 2008). At the deposit scale, normal faulting does not significantly offset stratigraphy or mineralization. Specific timing is generally poorly constrained.3 .6 Stratigraphic Setting of the Long Canyon DepositThe Long Canyon auriferous hydrothermal system is hosted within Paleozoic carbonate middle and inner platform facies sediments deposited on a distally steepened carbonate ramp during the Upper Cambrian-Lower Ordovician (Figure 3.1)(Cook and Corboy 2004). Stratgraphic sequences comprise, from oldest to youngest (Figure 3.2 and Figure 3.3), (1) Upper Cambrian Notch Peak Formation (Cnpl), a variably bioturbated, oncolitic and calcareous silt-laminated limestone capped by an 80 metre thick package of massive dolomite (Cnpd), unconformably overlain by; (2) Lower Ordovician Pogonip Formation (Opl) carbonate rocks of massive and calcareous silt-laminated limestone, calcareous siltstone, quartzite and minor dolomite, overlain by; (3) Middle Orovician Pogonip Formation Eureka quartzite. The stratigraphic sequence is upright and generally trending northwest and gently dipping shallowly to the northeast. 3 .6 .1 Upper Cambrian Notch Peak FormationThe Notch Peak Formation is a product of carbonate deposition during shallowing of the North American continental shelf over at least four 3rd order cycles through the Upper Cambrian followed by inundation of the sequence. Massive and bioclastic limestones were deposited through the first three cycles, with the fourth relative sea level shallowing resulting in an emergent carbonate platform and local development of the Notch Peak Dolomite (Cnpd) as the top of the Cambrian sequence (Cook, Taylor et al. 1989; Cook and Corboy 2004; Thorson 2007; Thorson 2008; Smith 2010). Detailed descriptions of stratigraphic units within the Cnpl at the deposit scale are provided below, bottom to top, and can be seen in the detailed stratigraphic section in Figure 3.2C. Cnplw – the “Notch Peak wispy limestone’ is a light-medium grey micritic mudstone to bioclastic grainstone with 10-30% dark grey, thin wispy silt laminations. Oolitic beds are common lower in stratigraphic unit. Thickness of this unit is poorly defined as it is the lowest unit in the stratigraphic sequence that was consistently intersected in drill core. 27Cnplus – the ‘Notch Peak upper silty limestone’ unit is a sequence of 2 – 4, half metre to one metre thick, massive silty limestone beds, thinning upwards and intercalated with massive micritic or commonly oncolitic limestone beds. Silt laminations are dark grey and comprise fine-grained detrital silicate minerals, detrital quartz grains and carbonate cements. Brassy euhedral pyrite grains occur within silt beds and often within contained carbonate fenestrae. Cnplonc – the ‘Notch Peak oncolitic limestone’ is a massive, light – medium grey oncolitic mudstone-grainstone containing scattered 0.5 cm oncoids, in concentrations from 1 – 20%. The unit can often contain gradational intercalations of massive, calcareous silt laminations on the lower contact. The unit is approximately 10 metres thick and gradational upwards into the burrowed limestone unit of Cnplbu.Cnplbu – the ‘Notch Peak burrowed limestone’ is a massive, light to medium grey mudstone – grainstone unit of bioturbated/burrowed limestone with a distinctive mottled appearance. Beds of oncoids, ooids and light grey chert nodules are locally observed. The unit is 15-20 metres thick with the lower contact often poorly defined where the oncolitic limestone begins. Cnpdl – ‘the Notch Peak dolomitized limestone’ is a dolostone comprising planar-s to nonplanar-a, sensu Sibley (1987), dolomite-replaced Notch Peak Limestone units Cnplbu, Cnplonc, Cnplus and (less frequently) Cnplw at the base of the Notch Peak Dolomite (Cnpd, described below). Dolomite occurs as: 1) massive, dark, generally fabric-retentive replacive dolomite of limestone; 2) white, sparry, non-ferroan dolomite as late veins and fracture fill within earlier dark replacive dolomite and overlying Cnpd; and 3) bedding concordant zebra dolomite, which is best developed in burrowed limestone (Cnplbu) (Figure 3.5a). Nonplanar saddle dolomite occurs locally, particularly within sparry, white coarse-grained zebra dolomite (Figure 3.5b). The term ‘zebra dolomite’ is applied to dolostone with alternating thin, irregular beds of dark grey and white dolomite. Burrows are typically replaced by dark dolomite within a lighter grey dolomite replacement of surrounding micritic limestone. White, sparry dolomite parallels these burrows and contains voids with open-space dolomite crystal growth and local geopetal features. The prevalence of zebra dolomite within Cnplbu likely reflects initial flow of dolomitizing fluids preferentially through more permeable bioturbated limestone. Voids within the white zebra dolomite (Figure 3.5a,b) are locally filled with one of several late mineral growths: (1) a clear, very fine-grained to amorphous silica (Figure 3.5e,f); and (2) an Fe sulphide, observable only as an oxidized hematite pseudomorph (Figure 283.5g). Fe oxide is also observed intergrown with dolomite crystals as compositional banding on crystal growth planes (Figure 3.5h). Oncoids within Cnplonc are often preserved when dolomitized (Figure 3.5e). Relict limestone bedding observable within the dolostone does not display any notable differences to limestone beds at similar stratigraphic levels that were not dolomitized.The stratigraphic extent of dolomitization within the Notch Peak Limestone is not uniform across the property. The thicknesses of replaced limestone stratigraphy varies from 1-2 m to 10-40 m. Along cross sections L12750N and L12800N, the thickness of Cnpdl is near uniform, extending consistently from the base of the Cnpd into the middle of the Cnplus (Figure 3.6). Cnpdl is generally thicker along the western margin of the study area.Cnpd – the ‘Notch Peak Dolomite’ is an 80 metre thick, medium to dark grey dolostone. Although typically massive and medium to coarse-grained with little remnant limestone texture or bedding, this dolostone can also exhibit thin sections of fabric-retentive dolomitization that preserves features of the original limestone including thin oolite and/or oncolite beds, laminated silt wisps, bioturbated and nodular limestone and chert ribbons. There are typically several 10-20 metre thick sections of pale, coarsely crystalline and non-fabric retentive dolomite within the unit. The upper contact with the Pogonip Limestone is very abrupt (Figure 3.2B), and is thought to represent an unconformity surface as in drillcore it can be seen locally truncating dolostone bedding and is overlain by laminar carbonate rocks. This unconformity has also been observed in surface outcrops as a paleokarst (Thorson 2007; Thorson 2008; Smith 2010). The lower contact is most commonly overprinted by the later dolomitization of Notch Peak limestone (see Cnpdl above). 3 .6 .2 Lower Ordovician Pogonip GroupSediments of the Lower Ordovician Pogonip Group at Long Canyon comprise six main stratigraphic units as defined by the mapping of Smith (2008), following (Thorman 1970) (see Figure 3.3 and Figure 3.2A). These stratigraphic units, from the base up, are: 1) Op1, a basal unit of thin to thick bedded grey limestone (mostly fine to coarse grainstone with rare chert nodules) intercalated with massive and laminated calcareous siltstone beds. Silty beds contain irregularly shaped, subangular to angular quartz grains 25-200 μm in size Several metre-thick dolomite lenses occur within the basal sediments of the unit; 2) Op2, a massive silty, sandy and burrowed grainstone with an upper thin dolomitic horizon; 3) Op3, a well indurated, 15 metre thick quartzite unit, the ‘Kanosh Quartzite’; 4) Op4 a massive limestone unit 29similar to the unit underlying the Kanosh quartzite below; 5) Op5, a recessive thin bedded, non-calcareous shale unit, the ‘Kanosh Shale’; 6) Op6, a massive limestone with abundant silt wisps, often with a thin, fossiliferous dolomite horizon at the top of the sequence. Overlying this sequence is the massive, regionally extensive, well-indurated and cross-bedded Eureka Quartzite. The lowermost unit which hosts the majority of gold mineralization at Long Canyon above the Cnpd is informally broken down into 3 main subunits plus a dolomitized-equivalent subunit (Figure 3.2B).Oplw – the “Lower Pogonip wispy limestone’ is a medium to dark grey, thin-bedded/laminated silty limestone with massive grainstone interbeds that contain chert ribbons and are locally replaced by dolomite. This dolomitization often obscures the lower unconformable contact with the Notch Peak Dolomite which has been described as an unconformity surface. Distinction between the two units is typically derived from the presence of silt wisps in light – medium grey dolomitized Oplw, whereas the medium to dark grey Notch Peak Dolomite is more massive and coarser grained. Oplsm – the “Lower Pogonip silty massive limestone’ is a light to dark grey limestone consisting of laminated silt beds 0.5 – 1.5 metres thick intercalated with massive grainstone beds of similar thickness containing sporadic thin silt laminations. The upper transition of laminated siltstone beds to grainstone beds is a knife sharp undulating surface, while the lower bed contact is transitional from massive grainstone to progressively increasing silt deposition. There are typically 4 to 5 such cycles, and while no attempt has been made to correlate individual beds across drillholes and the deposit, it is likely possible to do so. Pervasive, massive dark grey dolomitization occurs locally within massive limestone interbeds both here and within the Oplw unit below. Euhedral brassy pyrite is present within silt laminations. Massive limestone beds in Oplsm are strongly stylolitic denoting significant compaction.Opdl – the “Lower Pogonip dolomitized limestone” is a medium to dark grey dolomite which occurs in lenses 0.5 to 5 metres thick within primarily micritic/grainstone limestone interbeds of Oplw and Oplsm units. The distribution of Opdl is uneven across the property. In places dolostone bodies comprise several stacked lenses (Figure 3.6) or, if at all, only present as a <5 metre lens directly overlying the Notch Peak Dolomite. Dolomite is generally planar-s and polymodal with coarser grain-sizes occurring in replaced micrites and finer grain-size occurring in silty carbonate rocks. Locally, individual fine-grained pyrite grains occur intergrown with dolomite crystals.30Opls – the “Lower Pogonip silty laminated limestone’ is a light to dark grey, thinly laminated silty limestone. Silt laminations are typically 0.1 cm to 2 cm in width, intercalated with light grey grainstone – micritic limestone of similar widths, and undulated where unstrained. Brassy, euhedral pyrite 0.1 – 0.4 cm in diameter is present within silt laminations; often located within fenestrae or coarse-grained white calcite boudins. Flat pebble conglomerate beds are common. 3 .6 .3 Origin of Dolomitic Units and Sedimentary Facies Interpretation3 .6 .3 .1 Dolomitization ModelsDolomitization models that describe the environment and method of formation for massive dolostones are varied and remain subject to considerable controversy despite years of research (Machel 2004). Several widely accepted models of formation of stratiform replacive dolomite bodies within platform carbonate rocks and secondary saddle dolomite formation are presented below.Evaporative Reflux Model – A widely accepted and popular model often invoked for dolomitizing large bodies of platform carbonate rocks, the reflux model, involves evaporation of seawater in a restricted lagoonal and shallow marine setting on a carbonate platform behind a barrier (Adams and Rhodes 1960; Machel 2004). Seawater becomes hypersaline, and the Mg2+ to Ca2+ ratio increases owing to loss of Ca2+ to the precipitation of gypsum and anhydrite (Morse and Mackenzie 1990; Diehl, Hofstra et al. 2010). The resulting dense and Mg2+-rich brine permeates downwards and seawards through carbonate sediments, producing dolomitized platform carbonate bodies (Machel 2004).Penecontemporaneous/Organogenic Sabkha Model – In both peritidal and hemipelagic, anoxic, organic-rich carbonate platform environments, penecontemporaneous dolomites appear to be able to form through organogenesis as a byproduct of bacterial sulphate reduction (Mazzullo 2000; Machel 2004). Dolomite forms in this environment within a few years to tens of years after deposition and is typically non-ferroan owing to concurrent pyrite precipitation (Mazzullo 2000). According to the organogenesis model, only a small volume percent (5-70%) of the total near surface carbonate sequence is dolomitized. However, when occuring in a sabkha environment where storm tides episodically flood an evaporative supratidal environment, organogenesis pairs with seepage reflux mechanisms of the dense hypersaline brines and dolomite forms thin lenses and layers up to 100 volume percent within the upper 1-2 meters of sediment (Machel 2004). 31Saddle Dolomite/“Hydrothermal” Dolomitization Model – Saddle dolomite forms as both a primary void-filling cement and as a replacement mineral above 60-80°C; it has a distinctively distorted crystal structure which manifests as warped crystal faces and cleavage planes (Radke and Mathis 1980). While it often occurs as an epigenetic gangue mineral in MVT metal sulphide deposits (Radke and Mathis 1980; Sass-Gustkiewicz, Dzulynski et al. 1982), it is more significant, volumetrically, as the dominant constituent in replacement dolostone bodies along faults and bedding (Radke and Mathis 1980; Diehl, Hofstra et al. 2010). The method of formation is controversial. In several models for regional burial dolomitization, Mg ions are transported within connate formational brines above 80°C through topographic recharge, tectonic compaction or thermal convection (Morrow 1998). In another model, infiltrating hydrothermal fluids scavenge Mg ions locally by dissolution of earlier dolomite. Evidence of dissolution is observed in zebra dolomitized rocks where earlier dolomite contains vugs, partially infilled with white saddle dolomite (Machel 2004). Saddle dolomite is often associated with a ‘hydrothermal’ origin in the literature owing to its formation above 80°C (Machel and Lonnee 2002; Machel 2004). Machel (2002; 2004) is critical of the widespread misuse of applying a genetic interpretation of a dolomite as hydrothermal; 1) without any independent temperature of dolomite formation constraints relative to the host rock temperature, which in deep burial environments can be isothermal with such dolomitizing brines; and 2) based solely on the observations that the dolomite is saddle dolomite, or that the dolomite is texturally and spatially associated with base metal minerals. Machel (2002) does acknowledge that despite the above-mentioned misinterpretations, there are plenty of environments where saddle dolomite has formed from a hydrothermal fluid, i.e., an aqueous solution that is warm or hot relative to its surrounding environment, sensu White (1957). 3 .6 .3 .2 Long Canyon Dolomite Origin InterpretationThere are three stratigraphic levels were dolomite occurs at Long Canyon: 1) within Notch Peak Limestone units Cnplw, Cnplus, Cnplonc and Cnplbu at the base of the massive Cnpd (Figure 3.2C); 2) at the top of the Notch Peak Formation as the massive dolostone (Cnpd) (Figure 3.2B,C); and 3) as discrete dolomite horizons throughout the Lower Pogonip Group units.Cnpd Origin – Thorson (2007) describes meteoric karsting and erosion in the massive dolostone at the top of the Notch Peak Limestone. Dolomitization in this case would have occurred under emergent 32conditions, possibly through extensive reflux dolomitization of the carbonate platform at the end of the Cambrian. Opdl Origin – The basal units of the Lower Pogonip Limestone host dolomitized lenses that have been proposed by Fronteer geologists to be of late burial/hydrothermal origin owing to: 1) their proximity to structures including faults and breccia zones in basal lenses; 2) the limited distribution of dolomite bodies across the deposit; and 3) the occurrence of intracrystalline pyrite. Dolomitization is fabric retentive and superimposed primarily on thick bedded grainstones of the Oplw and Opsm limestones in stacked 0.25 – 1.5 metre thick lenses that pinch and swell along bedding. Little dolomitization occurs in massive and laminated silty limestones. Why the more permeable massive and laminated silty limestone beds were not preferentially dolomitized over the relatively low permeability massive grainstones is not explained by a hydrothermal fluid infiltration origin.Sequence stratigraphy is a useful tool in determining the environment of formation for the dolomite lenses within the basal Lower Pogonip Group sediments and the two dolomitic units occurring beneath Kanosh and Eureka Quartzite units higher in the sequence. Mutti and Simo (1994), following Wilson (1975), observe evidence of repeating fluctuations of relative sea level in the cyclicity of upward-shallowing carbonate platform strata, as a mechanism to expose carbonate shelves to either sabkha or reflux dolomitization. Such sequences occur in Early Ordivician Upper Knox Group carbonate rocks in the Appalachians where Montanez and Read (1992) and Pope and Read (1998) observe metre scale 4th order cycles of sea level rise and fall (104 – 105 years) within peritidal sequences that resulted in stacked lenses of dolomite intercalated with tidal flat laminated mudstones. The metre scale cycling of micritic limestones/grainstones sharply overlying massive silt laminations, and dolomite development within grainstones of Oplsm, indicates rapid sea level fluctuations over a peritidal mudflat to lagoonal carbonate depositional environment. The Opdl dolomite bodies developed within these cycles as discrete lenses replacing primarily micritic carbonate rocks and minor calcareous silt laminations likely developed as a result of restricted, evaporating lagoonal brines refluxing through carbonate muds, with pyrite development owing to organogenic dolomitization processes.Cnpdl Origin – The replacive dolomite occuring within Notch Peak limestone units Cnplw, Cnplus, Cnplonc and Cnplbu is interpreted as a later dolomite phase than the massive Cnpd based on the 33following observations: 1) Cnpdl hosts zebra-textured saddle dolomite indicating dolomitization created void space through volume reduction as dolomite replaced calcite. Saddle dolomite crystals grew within these voids. Similar voids and zebra texture dolomite are not observed in the massive, dark, coarse-grained and largely fabric-destructive dolomitization in the overlying Cnpd; 2) Saddle dolomite forms in environments above 80°C. As the overlying Cnpd is interpreted to have formed under near surface conditions from reflux dolomitization or a similar process, the saddle dolomite clearly formed later under intermediate to deep burial conditions; 3) Hematite pseudomorphs after pyrite occur both intergrown with dolomite crystals as compositional zoning in the early replacive dolomite (Figure 3.5h), as well as after zebra dolomite development as coarse clots of hematite (likely after pyrite) Figure 3.5h). Together these textures suggest Cnpdl may have developed during a protracted fluid event which dolomitized host limestone along with limited Fe sulphide deposition, developed void space subsequently filled by saddle dolomite from same fluid, and precipitated pyrite and quartz in vugs.Cnpdl is structurally modified by regional metamorphism that is interpreted to be Jurassic to Cretaceous in age. This provides a maximum age for dolomitization. Based on similarities to other occurrences of zebra dolomite in the Great Basin, this dolomite likely formed as burial dolomitization from basinal brines ‘ponding’ beneath the massive Notch Peak Dolomite, replacing Notch Peak limestones, developing void space, with further influx of Mg-rich fluids depositing white saddle dolomite within voids creating distinctive zebra-textured dolomite.3 .7 Igneous Rocks of the Long Canyon Deposit3 .7 .1 Mafic Intrusives Mafic intrusive dikes and sills typically 0.5 – 2 metres wide intrude carbonate strata across the Long Canyon property. The rocks are variably fine to medium-grained, biotite porphyritic with minor alkali feldspar, carbonate and quartz in the groundmass (Figure 3.7). Unmineralized and unoxidized intrusives are ubiquitously propylitically altered. Geochemically the rocks are low in silica, 42-44% SiO2, and ultrapotassic, 4-5.5 % K2O. Their chemistry and mineralogy suggests they are lamprophyres. It is notably difficult to classify ultramafic intrusives as positively lamprophyres (Woolley, Bergman et al. 1996) and these intrusives K2O/ Na2O ratios are generally >>3 indicating they may actually be lamproites rather than lamprophyres (Woolley, Bergman et al. 1996). However, for the purposes of this study these dikes will be informally referred to as lamprophyres (sensu lato). 34Lamprophyre dike contacts cut bedding and earlier metamorphic fabrics at high angles (Figure 3.7e). Dikes often occur as NE trending swarms within moderate to high angle brittle damage zones. An optical televiewer downhole survey by Golder Associates in 2009 down drillhole LCG02 recorded orientations of lamprophyres. LCGO2 is located on the western margin of the deposit and cuts a thick sequence Cnpd, within which lamprophyre orientations were measured. Results are displayed as a stereographic projection in Figure 3.7f. There are three populations of dyke orientations; sub-horizontal sills (2-3° dip); low angle dikes (20-32° dip); and moderate angle dikes (46-53° dip). The median dip direction of intrusives is 308°.3 .7 .2 Felsic Intrusives and Volcanics Felsic igneous rocks are not spatially associated with the Long Canyon deposit, but can be observed on the property to district scale at several locations, see Figure 3.3. The Nanny Creek Volcanics are present north of the deposit in the northern Pequop Mountains, described by (Brooks, Thorman et al. 1995) as rhyolite ash flow tuffs overlain by intercalated andesitic to dacitic flows and flow breccias and rhyolite ash flow tuffs collectively yielding 40AR/39Ar dates between 41 and 39 Ma. On the west side of the Pequop Mountains, Bedell (2010) presents recent U/Pb data for two devitrified rhyolite dikes yielding ages of 41.0 +/- 0.8 Ma and 39.1 +/- 0.7 Ma. 3 .8 Structural Geology of the Long Canyon DepositStructure was a critical control in localizing auriferous fluids at Long Canyon. Structures within the rocks at Long Canyon rock comprise metamorphic fabrics, large scale boudinage and brittle faulting and brecciation. Boudin geometries and spatially associated brittle brecciation appear to be first order controls on the loci of Au mineralization. Boudin development in host rocks at Long Canyon is explored in detail.3 .8 .1 Synmetamorphic FabricsThere are two main metamorphic fabrics present at Long Canyon. The earliest and most prominent, S1, is defined by: 1) Near bedding-parallel foliation, best developed as aligned phyllosilicate minerals within silty laminated Pogonip and Notch Peak limestones (Figure 3.8). Quartz grains occur as very fine-grained aggregates within silty Pogonip limestones. Quartz masses are aligned with foliation and located within what appear to be pressure shadows of euhedral pyrite grains (e.g., Figure 3.4b). 2) Local isoclinal folding of Pogonip and Notch Peak limestone bedding into the plane of foliation (Figure 3.8b). 3) A southeast-plunging lineation defined by stretched bioturbations (Smith 2010) in the plane of foliation, which is consistent with NW-SE elongation. 4) Small scale boudinage of micritic limestone interbeds within laminated siliciclastic sediments (Figure 3.8b). 35The second fabric, S2, occurs as a crenulation cleavage within limestones (Figure 3.8a,c). It is moderately northwest-dipping and has a mod-high angle to the enveloping S1 orientation (Smith 2010) and (Rhys and Ross 2010). Crenulations appear best developed within siliciclastic beds as opposed to massive micritic limestone and dolomite. Folding of earlier foliation as chevron style kink folds occurs within limestones and appears to occur more frequently above boudin necks. Folds are generally southeast vergent with S2 as an axial planar fabric (Rhys and Ross 2010). The metamorphic fabrics are consistent with the Mesozoic metamorphic history of the Pequops as described by Camilleri and Chamberlain (1997), where several phases of contraction accommodated along top-to-the-southeast thrust faults active during the Late Jurassic to Late Cretaceous. Generation of S2 crenulation cleaveg is attributed to this phase of contraction.3 .8 .2 Boudinage in the Notch Peak DolomiteThe presence of a large scale boudinage of the Notch Peak Dolomite was first identified in surface outcrops by Smith (2009). Planar dolomite units, 80 – 100 metres in thickness, pinch out abruptly along linear edges; forming elongate blocks with gently north plunging axes typically oriented approximately SW-NE. Boudin necks vary from narrow cracks to gaps >100 m wide. S1 foliated silty limestone units above and below the dolomite that are typically parallel to dolomite bedding, steepen over the nose of abruptly thinned-out dolomite blocks, ‘flowing’ into the space between dolomite blocks, suggesting S1 foliation and boudinage occurred during the same deformation event. Limestone units that should be stratigraphically separated by a thick dolomite bed, now lie in direct contact with each other (Figure 3.10). Brittle damage zones developed within dolomite units where little separation has occurred, are spatially associated with breccias in Pogonip and Notch Peak limestone units. These limestone-hosted breccias display characteristics of carbonate dissolution and solution collapse breccia (eg. Figure 3.10a,e).Boudinaged dolomite blocks display variable separation along SW-NE oriented boudin necks and high angle brittle damage zones between rotated boudin blocks (Figure 3.11). In the southern half and western margin of the deposit there is wide separation of dolomite boudin blocks. In the northern end and eastern margin of the deposit dolomite boudin blocks are broken along brittle high-angle fault planes, defined by brittle damage zones, which extend consistently for > 1 km along a SSW-NNE trend within the dolomite. Although they exhibit little separation, they commonly have a slight clockwise rotation across them.363 .8 .3 Brittle Brecciation and Faulting in the Notch Peak DolomiteDolomite-hosted breccias are present at Long Canyon principally hosted within dolomite blocks along fault and brittle damage zones (Figure 3.12 and Figure 3.6), along dolomite block margins in contact with limestones and particularly along dolomite boudin noses. Breccias are typically clast-supported with a high-angle fabric of aligned clasts relative to horizontal to sub-horizontal dolomite bedding (Figure 3.12). Clasts are typically cobble to pebble sized and finer and range from angular to subrounded. Matrix is typically highly comminuted dolomite grains and minor quartz grains (Figure 3.4c-e). Breccias are often only semi-lithified; can be pervasively silicified with milky white, coarse quartz (Figure 3.4f); or host calcite cement which typically occurs with minor amounts of Au mineralization and/or weakly anomalous concentrations of Sb, Tl, Hg and As pathfinder elements. During the course of this study, faults were not specifically examined on surface across the property and those observed in drill core were unoriented. Dolomite-hosted breccias often define planar, high-angle damage zones, or faults, from drill core intercepts, Faults most relevant to the scope of this study are those hosted in dolomite, defining incipient boudin necks, and spatially associated with gold mineralization in the northern part of the deposit in the ‘RC Crack Zone’ Figure 3.9. Examples of dolomite-hosted fault breccia fabrics are seen in Figure 3.12. The angles of dominant fault or fault breccia fabrics relative to the core axis (alpha angle) were measured from drillcore across sections 12800N and 12750N (Figure 3.10). There is an estimated error in measurements of +/- 10°. These alpha angles were estimated from core photos where the core was most orthogonal to the camera. Median alpha angles of breccia fabrics were determined for each drillhole and plotted on stereographic projections (Figure 3.13). The poles to the breccia fabric or fault orientation plot as small circles around the drillholes, giving a range of possible orientations of the poles to the fault plane. On both stereonets in Figure 3.13 for data from sections 12750N and 12800N, these small circles intersect at two areas. These areas represent a range of possible fault and fracture plane apparent dips in the plane of the section oriented between 258 – 002 dipping 50-65°. Small changes in plotted median alpha angle orientations of 5-10° (within error) shift small circles such that a variety of solutions from stereographic analysis are possible within a west to north dip-direction. From this analysis, breccia fabric and fault orientations can be described as moderately to steeply dipping in westerly to northerly direction, sensu lato.373 .8 .4 Genesis of Boudinage in Notch Peak DolomiteBoudinage is a commonly observed structure in multilayered rocks deformed at mid-crustal levels. Occurring in an overall transpressional deformation environment, strain becomes highly partitioned (R. Jones and Geoff Tanner 1995) within less competent layers enclosing more competent layers. Layer parallel tension results, with localized brittle extension failure or shear fractures in the strong layer and ductile folding and layer flow in the rheologically weaker layer (Ramberg 1955; Mandal, Chakraborty et al. 2000). Observations from both natural systems (Peacock and Sanderson 1992; R. Jones and Geoff Tanner 1995; Druguet, Alsop et al. 2009) and clay modeling simulating hypothesized stress conditions (Ramberg 1955; Mandal, Chakraborty et al. 2000) display similar modes of boudinage, providing a semi-quantitative assessment of factors controlling the mode of deformation. Lithologic factors controlling the mode of deformation and the resulting boudin geometries, are examined below, specifically the effect of changing brittle layer thickness on resulting boudin geometries, failure mode and fracture spacing.Boudins display varying modes of fracturing within tranpressional or extensional zones, classified by (Mandal, Chakraborty et al. 2000) through experimental determination into three types: tensile boudinage, shear fracture boudinage and extensional shear fracture boudinage. The two most important factors controlling fracture mode are: 1) competency contrast between brittle and ductile layers, or strength ratio (F) – defined as tensile strength in brittle layer/flow strength in ductile layer (Peacock and Sanderson 1992); and 2) the thickness ratio (Tr) between them (Mandal, Chakraborty et al. 2000). Through their physical models, Mandal et al. determined thickness ratios less than 4.5 produced tensile Mode I failure in brittle boudins with high boudin aspect ratios (Ar), significant separation between boudins and rectangular geometries. When the Tr is high or F is low, failure mode changes to Mode II shear failure with boudin geometries more rhombohedral with Ar closer to 1 or <1, i.e. tighter fracture spacing. These relationships are displayed in Figure 3.14b and c and examples of both seen at Long Canyon in Figure 3.11.Notch Peak Dolomite boudin blocks at Long Canyon display both tensile and shear failure modes (Figure 3.11). Shear failure and development of closely spaced faults within dolomite, without significant separation of boudin blocks, occurs primarily in the northeastern domain of the deposit north of the Discovery Zone (see examples in Figure 3.9, Figure 3.14 and Figure 3.6). South and West of the central Discovery Zone, boudin blocks display significant separation within their necks, characteristic of an extensional failure mode (Figure 3.14). Boudin blocks occurring by extensional failure tend to contain less fracturing and brittle damage zones. These observations correspond with an increased thickness and 38occurrence of Pogonip dolomitized limestone lenses (Opdl) in the northeast or Shear Fracture Boudinage domain of the deposit; and a paucity of Opdl sequences within limestones occupying boudin necks in the southwest or Tensile Fracture Boudinage domain (Figure 3.14). This is observed in particular along cross sections 12800N and 12750N (Figure 3.6) where a high density of brittle fractures and faults within dolomite is coincident with thick lenses of dolomitized limestone in the overlying limestones which thicken the overall brittle dolomite layer by up to 30%.Dolomitization of limestone stratigraphy at the base of the Notch Peak Dolomite occurred pre-boudinage. Evidence for this is largely from bedding relationships and structure, and is observed in roadcuts and outcrop where flat-lying zebra dolomite is abutted by steeply dipping limestone units, S1, as commonly seen on dolomite boudin neck margins throughout the deposit. Additional evidence can be observed from bedding orientation patterns in drill cores, particularly at the base of dolomite boudin block margins where dolomitized limestone bedding is flat-lying and brittley deformed adjacent to high angle limestone bedding. Additionally, brittle shear faulting and damage zones associated with brittle deformation during boudinage of dolomite blocks penetrate primary Notch Peak Dolomite as well as underlying hydrothermal dolomitization of Notch Peak Limestone stratigraphy.Review of dolomite distribution within the multilayer carbonate sequence at Long Canyon as well as distribution of observed boudin failure modes leads to several conclusions: 1. Thin dolomite sequences, i.e. less secondary dolomite contributes to extensional fracture failure and separation between boudin blocks and wider fracture spacing or boudin neck distribution. 2. Thicker Opdl and Cnpdl dolomite alteration increase the thickness ratio, inhibiting extensional failure and boudin separation. 3. Thin dolomite replacement sequences within Opl limestone increases the bulk shear strength of an otherwise ductiley deforming limestone layer buttressed against a brittle dolomite, serving to lower the competency contrast and strength ratio between dolomite and limestone domains and enhance shear fracture boudinage failure. 4. The prominent NNW deflection of an extended dolomite boudin neckline from the otherwise consistent NNE orientation of boudin necks and fractures within the center of the modeled dolomite Figure 3.14 may be in response to deflection of strain around this domain of lower competency contrast and increased thickness ratio where shear fracture boudinage dominates.393 .9 Alteration and Mineralization of the Long Canyon DepositOre-stage and later fluid flow events throughout the geologic history of rocks hosting the Long Canyon deposit produced numerous alteration types, summarized tentatively in chronological order with some overlap comprise: 1) carbonate dissolution and solution collapse brecciation; 2) sulfidation with subsequent gold and trace element deposition; 3) partial to complete silicification; 4) clay alteration of siliciclastic minerals and intrusives (argillization); 5) calcite veining and breccia cementation; 6) oxidation within the deposit. Decarbonation, mineralization, silicification and clay alteration are pre to late ore-stage, with evidence for protracted/multiple fluid flow events. Calcite veining likely has variants of late ore-stage to post ore-stage. Oxidation occurred pervasively throughout the deposit, the timing of which is uncertain.3 .9 .1 Carbonate Dissolution and BrecciationCarbonate dissolution (decarbonation) is a prevalent alteration type at Long Canyon spatially associated with gold mineralization. The most visible sign of carbonate dissolution is volume loss resulting in collapse breccias and thinning of silty limestone units with coincident small-scale crackle breccias and concentration of insoluble quartz and siliciclastic minerals (e.g., Figure 3.15). These solution collapse breccias are primarily developed within limestone, without volumetrically equivalent dissolution of proximal dolomite, if any, during the same fluid event(s). Within breccias along highly permeable boudin noses, contained dolomite clasts are locally rounded and embayed implying dissolution, although earlier fault-related communition and mechanical abrasion could have generated the texture. Within dolomite-hosted fault breccias where mineralogical evidence indicates hydrothermal fluid flow (authigenic quartz growth (e.g., Figure 3.12, discussed in Section 3.9.2 below) there is no observed dolomite dissolution. While in many Carlin-type deposits dolomite dissolution by ore fluids is common, and sometimes a requisite for subsequent gold mineralization (e.g. Meikle) (Hofstra, Leventhal et al. 1991; Emsbo, Hofstra et al. 2003; Cline, Hofstra et al. 2005), at Long Canyon, dolomite dissolution does not typically accompany calcite dissolution. 3 .9 .1 .1 Dissolution BrecciaThe majority of Au mineralization at the Long Canyon deposit is hosted within silty limestone-hosted breccias which display characteristics consistent with varying degrees of dissolution of carbonate minerals and rock collapse. This style of brecciation is classified here as solution brecciation. Solution breccia development also occurs without significant Au mineralization. Silty limestone and massive 40limestone-interbedded host rocks which have undergone solution brecciation display a down-stratigraphy variation in development of breccia facies, similar to those defined by Loucks (2007) for classification of paleokarsts. Figure 3.16 displays selected intervals from vertical drillhole LC591C through a representative solution breccia facies sequence developed within Pogonip limestones above a dolomite hosted fault and fracture zone (e.g., Figure 3.9) and described here:1) Collapse breccia hanging walls consist of contiguous, relatively undisturbed bedding. High angle fracturing exploited by thin calcite veins is typically observed in strata above breccia, and is conspicuously absent in stratigraphy above which no solution brecciation has occurred. This fracturing is termed suprastratal damage (Figure 3.17), relative to underlying brecciation. Displacement across veining and fracturing is commonly observed.2) Silty laminated limestone beds become fractured, with rounded bedding fragments collapsing downards into jigsaw breccias. Intercalated micritic limestone layers display rounding and cavity embayment. Concentrations of insoluble siliciclastic, clay and iron oxide minerals occur along seams and rounded cavities. Brecciation and limestone volume loss occurs along fractures. Bedding across thin, high angle faults displays normal displacement.3)  Coarse clast to matrix-supported chaotic breccia of highly disturbed and down-dropped clasts from overlying beds into clay-rich, mud to sand matrix of insoluble siliciclastic minerals. Limestone clasts are sub-angular to sub-rounded and locally embayed. Dolomite clasts are typically angular to subangular. Clasts of earlier breccias are common. Breccia cavities were likely fluid saturated during development as evidenced by: (1) pervasive mud matrix support, (2) infiltration of matrix and fine-grained clasts along fractures, mimicking pebble dikes, and (3) lack of shelter cavities beneath clasts.4) Fine-clast chaotic breccia and basal sediments, dominantly mud matrix and insoluble sediment and bedding chips along base of breccia. Basal sediment thicknesses range from 0.5 metres to as much as 5-7 metres. Breccia base typically transitions abruptly into intact massive limestone or dolomite bedding with embayed and fractured contact.These 4 facies vary in thickness and spatial extent across the deposit. Repeating sequences of breccia facies frequently occur as bedding-scale collapse breccias and insoluble residue concentration sequences within an overall stratigraphic unit-scale solution brecciation. Breccia facies 1-4 decrease in volume, as (1) suprastral fracturing and (2) bedding collapse facies telescope down into tighter (3) chaotic breccias and (4) basal mud sediments. Breccia facies 3 and 4 are often absent in zones of weak brecciation, with 41only minor bedding collapse and fracturing. More competent beds, such as metre-thick dolomite lenses in Pogonip limestones appear to provide a resistant roof zone to collapse brecciation beneath and lessen further collapse along normal faults and fractures of overlying stratigraphy. This is observed in LC556C on section 12800N (Figure 3.6).Solution breccia bodies, consisting of facies 3 and 4, form horizontal pipes elongate in a SW-NE orientation along the top and bottom of Cnpd, or along boudin noses. In cross section, solution breccias laterally transition quickly into relatively unfractured and unbrecciated limestones. Breccia pipes occur as two to three parallel bodies, cross-sections of which can be observed on Figure 3.6. The breccias within limestones are spatially associated with zones of strong brittle cataclasis in underlying or overlying dolomite.The preceding description of solution breccia body and facies morphology principally describes breccias developed within the Pogonip limestones. Solution breccias within Notch Peak limestones display a drastically shortened vertical distribution of facies 1-4; often missing facies 1 and 2 completely, possibly owing to higher tensile strength and resistance to carbonate dissolution and volume loss within the overlying dolomites. Solution breccias are principally developed within and around the silty Cnplus unit as narrow pipes of chaotic breccias, and overlain by dolomite hosting significant cataclasis. The distribution of Notch Peak limestone-hosted solution breccias is not as consistent along the long axis of the deposit, as the main Pogonip limestone-hosted solution breccia pipes.It can be observed on cross-sections (e.g., Figure 3.6) that breccia zones within massive dolomite (shown to be approximately northwest dipping brittle fault zones, Figure 3.13) have fault termination sites within limestone above and below the dolomite that are occupied by mineralized solution breccias. Projecting up along a 50-65°, northwest-dipping fault plane, brecciated damage zones are observed within limestones above solution breccias.Hydrothermal solution brecciation occurred as multiple overlapping events. Earliest solution brecciation is poorly constrained, potentially occurring as early as post peak-metamorphism. Ore-stage minerals rimming breccia clasts (explained in more detail in Section 3.9.4 below) demonstrate that solution collapse breccias were well developed before the main auriferous fluid event. Mud matrix-hosted jasperoids and mineralized lamprophyre intrusive clasts indicate solution brecciation continued post 42mineralization. Cessation of solution breccia development is only constrained by the beginning of calcite cement brecciation, described below.3 .9 .1 .2 Interpretation of Ore Fluid Composition From Relative Dissolution of Calcite and DolomiteThe extent of calcite dissolution in limestone and the lack of similar dissolution in nearby dolostones suggests a specific pH range of the infiltrating ore fluid at Long Canyon. Fluid-inclusion studies indicate CO2 concentration was the major control on fluid pH in Carlin-type deposits (Cline, Hofstra et al. 2005). Fluid-inclusions from ore-stage silica and calcite at other Carlin-type deposits in the Great Basin have <4 mole percent CO2 and temperatures ranging from 250°C to 150°C (Cline, Hofstra et al. 2005). Hofstra et al (1991) using alteration and fluid-inclusion constraints from Carlin-type deposits in the Jerrit Canyon district, Nevada, calculated fluid pH of 5.07 and 5.19 at dolomite and calcite saturation, respectively, using a CO2 concentration of 4 mole percent and fluid temperature of 225°C. Cline and Hofstra et al (2005) summarized these earlier results as indicating that zones of dolomite dissolution required fluids with, under the assumed conditions, a pH of less than 5.07. With lower fluid CO2 concentrations this required pH would be higher and at lower temperatures it would be lower. Without any fluid-inclusion data from Long Canyon to indicate CO2 content or temperature of ore-stage fluids it is difficult to calculate reasonable estimates of fluid pH. However, it is a narrow pH range in which Long Canyon ore fluids must occupy where calcite dissolution is promoted, while dolomite dissolution is inhibited, as mineralogical and textural evidence indicates is the case at Long Canyon. 3 .9 .2 Quartz Grain Growth and Silicification within Ore Zones3 .9 .2 .1 Non-Ore-Stage QuartzQuartz varieties at Long Canyon are varied and while volumetrically minor are very useful for determining the geologic history of the carbonate host rocks. Authigenic quartz occurs with both the dolomite and the auriferous hydrothermal events. Paragenically early varieties present within unmineralized carbonate host rocks are briefly reviewed here for distinction with later ore-stage quartz. These include: (1) Quartz within siliciclastic laminations and calcareous siltstone beds of both the Notch Peak and Pogonip Limestone and identifiable as irregularly-shaped grains, often with undulose extinction (e.g., Figure 3.4a). These are likely detrital grains with a terrigenous origin. (2) Aggregates of intracrystal, fine-grained xenomorphic quartz (e.g., Figure 3.4c) aligned in strain shadows and foliations in metamorphosed Notch Peak and Pogonip limestones (e.g., Figure 3.4b). Chert nodules have a similar quartz texture in thin section (e.g., Figure 3.4d, Figure 3.5a). This is likely biogenic silica sourced from 43the dissolution of amorphous silica skeletal material (Maliva and Siever 1989).Metamorphism which produced the boudinage and extensional shear fracturing within the Notch Peak dolomite may have remobilized this originally biogenic silica into high angle veins along fractures (e.g., Figure 3.4e) and fault breccia cements (e.g., Figure 3.4f). The high angle fracture-filling silica veins are commonly aligned in sets and are present within all dolomite units including Cnpdl, Cnpd and Opdl. (3) Quartz occurs as clear quartz within vugs of sparry white dolomite of zebra-dolomitized Notch Peak limestone (e.g., Figure 3.5e,f). (4) Coarse quartz-carbonate veins spatially associated with lamprophyre intrusives, though not unambiguously genetically linked. These veins are often folded within host carbonate rocks.3 .9 .2 .2 Ore-Stage QuartzAuthigenic quartz grains interpreted as hydrothermal in origin have been recognized by past workers at Long Canyon (Rhys and Ross 2010; Smith 2011) as associated with the ore-stage alteration event overprinting solution breccia. In petrography they are observed as hexagonal, doubly terminated, euhedral to sub-rounded quartz grains, often inclusion-rich compared to detrital varieties and often with subtle growth zoning and possible inherited cores. They occur in all breccias with gold mineralization and occur in 3 textural forms (see examples in Figure 3.18): 1) commonly within breccia matrix; 2) rarely as druzy overgrowths on breccia clasts and wall rock; 3) overprinting earlier, unoxidized and oxidized, fine-grained (2-4 μm) arsenical pyrite in ore zones (e.g., Figure 3.22). These textures suggest that hydrothermal quartz grains were deposited from a fluid in the later stages of ore deposition within fluid-saturated breccias.Pervasive fine grained silicification occurs overgrowing earlier authigentic, doubly terminated quartz grains and sulphides within solution breccias as fabric-retentive quartz. In petrography this silicification has variably a fine-grained reticulate texture of interlocking, orthogonal quartz crystals (e.g., Figure 3.18b) to a jigsaw texture of interlocking, sub-anhedral quartz grains. Lovering (1972) describes similar quartz textures observed in jasperoid alteration around sediment replacement ore deposits. Quartz grain size is dependent upon the lithology it is replacing as very fine-grained when replacing siltstones and micritic limestone and coarser grained when breccia matrix replacement allows larger grain growth. Open space cavities lined with druzy quartz often occur within pervasively silicified material, sometimes infilled with later coarse calcite.Distribution of silicified breccias is generally localized to the mineralized cores of fluid flow paths within limestone solution breccias directly above and below Cnpd dolomite. Along the long axis of the deposit, 44occurrence of pervasively silicified breccias is localized to discrete ‘pockets’, though frequency and volumes of silicified zones increases to the southwest until silicified limestones become cliff-forming units at ‘Jasperoid Ridge’ and the extensive silicified breccias of the ‘West Zone’ (e.g., Figure 3.9). The presence of pervasive fine-grained silica overprinting coarse euhedral quartz grains within breccias, suggest several possible ore fluid temperature and silica solubility evolutions at Long Canyon. The textures are consistent with a fluid cooling due to conductive heating loss along a quartz-saturation curve, followed by rapid quartz supersaturation and precipitation. The cooling occurs within breccia owing to the high capacity for heat dissipation within the large surface area of breccias. Figure 3.19 displays several possible pathways of such a fluid. In both pathways, an infiltrating fluid nearing quartz-saturation cools onto the quartz-saturation curve at ~ 250°C. Further cooling moves the fluid along this curve, precipitating the euhedral hydrothermal quartz grains we observe in breccias (e.g., Figure 3.18). The pervasive overprinting fine grained silicification we observe in Figure 3.18b would require rapid precipitation of quartz. Two scenarios seem possible to achieve this: 1) fluid mixing of the ~225°C ore fluid and a hypothetical 100°C meteoric fluid with the resulting fluid composition becoming rapidly quartz-supersaturated (e.g., Hofstra and Cline 2000); and 2) continued cooling of the fluid past 180° to 200° C where quartz precipitation is kinetically inhibited by low temperatures, but supersaturates the fluid until rapid precipitation occurs (Fournier 1985; Hofstra and Cline 2000) Accordingly, the increased density of silicification at the southwestern end of the deposit is considered a possible outflow zone for cooling ore fluids migrating from the northeast to southwest along the long axis of the deposit. 3 .9 .3 Calcite Veining and Breccia CementationCoarse, crystalline calcite is present in many forms within the carbonate rocks of Long Canyon. Volumetrically the greatest occurrences are within late breccias as coarse calcite cementing earlier breccias; both clast supported and cement supported. These generally overprint earlier solution breccias and zones of structural damage along boudin necks and earlier faulting. Calcite veining occurs in greatest concentration in carbonate rocks above Pogonip Limestone-hosted ore bodies.3 .9 .3 .1 Coarse Calcite-Cemented BrecciaCalcite occurs as a breccia cement that cuts all earlier breccias and appears to even postdate oxidation. It is characterized by brittle fracturing of host lithologies and subsequent support of typically angular, polymictic breccia clasts in a coarse-grained, white calcite cement (Figure 3.20). When overprinting 45poorly indurated solution and cataclastic breccias, calcite cements infiltrate and re-cement earlier silty to sandy breccia matrices, several examples of this are displayed in (Figure 3.20a-f). Open space vuggy calcite growth occurs, and a single occurrence of stalagmite growth has been observed.Breccias vary from clast-supported to calcite cement-supported. They can host slightly rotated clasts from nearby wall rock source (Figure 3.20h) to transported exotic polymictic clasts (Figure 3.20g). Clast composition includes micritic limestone, silty limestone, dolomite, dolomitized limestone, unmineralized lamprophyre, mineralized lamprophyre, mineralized breccia fragments and silicified breccias. Breccia clast compositions suggest that clasts moved down-stratigraphy. Oxidized lamprophyre and silt bedding clasts hosting liesengang banding rotated relative to similar oxidation patterns in wall rock indicate calcite brecciation may have occurred after the initiation of oxidation of the deposit.Calcite-cemented breccia distribution is largely coincident with areas of previous cataclastic and solution brecciation. The highest density of calcite-cemented breccia occurs along boudin noses, where it can be definitively observed that clasts source rocks are from stratigraphically higher. The most exotic breccias occur in the West Zone (Figure 3.9) where significant extensional separation of boudins has occurred.These breccias are easy to distinguish from earlier solution breccias as they add significant volumes of new calcite, rather than dissolving carbonate. This represents a significant shift in the pH and temperature of fluids within the system from decarbonating to carbonate precipitating. The timing of this fluid event relative to the solution breccia generating/mineralizing fluid event is poorly constrained beyond simply post-dating. The calcite could be sourced from evolved ore fluids which dissolved significant quantities of carbonate during earlier solution brecciation; though the apparent timing of calcite cementation as post-initiation of oxidation does not agree with this interpretation. Similar breccias mapped by (Evans 2000) at Meikle, the breccia-hosted Carlin type deposit, were dated by Emsbo (1999) at 2 Ma by U/Pb dating methods.3 .9 .3 .2 Calcite VeiningThe most common form of calcite veining at Long Canyon are generations of thin (0.1 – 1cm wide) veins at a high angle to bedding, present predominantly in Pogonip limestone suprastratal damage zones above solution breccias (Figure 3.16a,b), with vein density significantly decreasing outwards from these fractured zones. The timing of these high angle calcite veins is unclear. They cut foliation; occupy 46fractures interpreted to have occurred during or immediately after solution brecciation/mineralization; in many places cut contiguous oxidation patterns in bedding suggesting they occurred post oxidation. In places, however, irregular oxidation patterns within wall rock on either side of veining suggests veining may predate oxidation. Locally calcite-filled fractures exhibit dilation. Many of these timing indicators are coincident with those of calcite-cemented breccias, suggesting the calcite veining, or at least one generation, and breccia cements formed as part of the same system.Subordinate calcite occurrences include coarse quartz-carbonate veins spatially associated with mafic intrusives; a coarse calcite filling vugs within silicified limestone (jasperoids); and a single occurrence of massive manganese calcite with radiating crystals occurring within a Cnpdl-hosted breccia cavity, intersected in drillhole LC556C on section 12800N.3 .9 .4 Gold MineralizationCarlin-type Gold Precipitation Mechanisms - The most common mechanism of gold precipitation in Carlin-type deposits, which results in submicron gold inclusions within pyrite lattices, is sulfidation of Fe2+ (Hofstra, Leventhal et al. 1991; Simon, Kesler et al. 1999; Hofstra and Cline 2000; Kesler, Fortuna et al. 2003). The simplified sulfidation reaction of a H2S-rich fluid interacting with reactive iron in carbonate proceeds via the reaction,2H2S + (Fe2+,Ca)CO3 = FeS2 + H2 + H2CO3+CaCO3where any reactive Fe2+ bearing mineral is dissolved, and bisulfide complexes are consumed to produce pyrite (Cline and Hofstra 2000). When an ore fluid is transporting Au complexed with bisulfide, consumption of H2S results in co-precipitation of Au. This is assumed to occur either by co-precipitation with arsenian pyrite (Simon, Kesler et al. 1999), or by adsorbtion of ionic AuI onto arsenian pyrite (Simon, Kesler et al. 1999). This ore-stage arsenian pyrite typically occurs as disseminated micron-scale grains within sediments or breccias; as overgrowths on Fe-bearing silicate minerals; or as overgrowths on pre-existing pyrite. The Fe+2 required for sulfidation to occur comes from several sources including dissolution of Fe carbonates and/or sulfidation of of silicate minerals in siliciclastic rocks and intrusives (Cline and Hofstra 2000).47Character of Long Canyon Gold Mineralization - Gold in the Long Canyon deposit is hosted predominantly within solution breccias and strongly decalcified silty limestone bedding superjacent to breccia margins. Distribution of gold in Fronteer’s 1g/t cutoff block model can be seen in Figure 3.9, hosted within breccias associated with northeast trending structures. Lamprophyres are also an important host for high grade gold mineralization at Long Canyon, becoming strongly gold-mineralized where dikes intersect zones of permeable structural damage that were utilized by later mineralizing fluids. Trace amounts of gold (0.1 – 0.5 ppm) occurs within highly comminuted faults within dolomite. Distribution of gold zones can be observed in cross section in Figure 3.21.The precise character of gold mineralization at Long Canyon is difficult to discern owing to the paucity of sulfides throughout the deposit after extensive post-mineralization oxidation. Native gold grains of 2-5 μm in width have been observed by several past workers (e.g. Rhys and Ross and examples therein 2010) occurring within solution breccias encapsulated in doubly terminated quartz grains, within druzy quartz vugs, on the margin of late, coarse calcite crystals, and encapsulated within late, coarse calcite. However, free gold is not common and in most high grade samples (10-45 ppm Au), no native gold grains were observed in thin section, suggesting it is not the predominant style of hypogene mineralization. The mechanism of gold deposition at Long Canyon is suggested by several pieces of indirect evidence. Euhedral, arsenic-bearing iron oxide grains (presumably hematite after pyrite) were observed in SEM associated with silica in gold-mineralized intervals (e.g., Figure 3.22). In gold-mineralized lamprophyres, iron oxide pseudomorphs after pyrite replace Fe-silicate minerals (presumably iron-rich micas) (e.g., Figure 3.7c,d). High grade gold-mineralized limestone breccias in LC556C, hosting up to 45 ppm Au and >1% As in hand sample and shown in Figure 3.15e,f,g,h, exhibit abundant speckled iron oxides, several microns wide, along fracture margins, rimming breccia clasts and concentrated bedding within Fe-rich silty limestone bedding clasts. As no coarse gold is observed in these samples, the most likely host for gold mineralization is within what were likely disseminated arsenian pyrites; indicating that sulfidation of arsenian pyrite and subsequent deposition of submicron Au within arsenian pyrite lattices is a more likely mechanism of gold deposition, as is common in Carlin-style deposits. In many instances, gold mineralization at Long Canyon occurs in silicified breccias. There is no definitive textural evidence to conclude whether this gold precipitated with silica, or whether silicification simply overprints previously mineralized breccia.483 .9 .5 Clay Alteration and Fluid SourcesClay isotopic studies were carried out on selected samples along a cross-section through the core of the mineralized system at Long Canyon. The purpose of the study was to determine ore fluid sources, evolution and temperature. This data is used to compare Long Canyon ore fluid characteristics with those from other Carlin-type deposits across Nevada.3 .9 .5 .1 Methodology of Clay StudiesIn this study five hand samples were selected for determination of clay mineralogy and δD and δ18O composition of clay minerals. Samples were selected from drillhole LC556C as a cross-section through strongly mineralized breccia and surrounding stratigraphy that hosted significant hydrothermal fluid flow. Samples were all either silty laminated limestones or breccia containing silty laminated limestone clasts, selected for their high silicate mineral content and correspondingly high clay content following interaction with hydrothermal fluids. Sample locations are shown in Figure 3.23. Hand sample clay mineralogy was first determined by near and short wave infrared (SWIR) analysis using the TerraspecTM analytical spectral device (ASD). Multiple measurements were taken from sawn core faces and reflectance spectra interpreted with the aid of reference spectra. Clay separates of the <2 μm fraction were then prepared from samples, according to the methods outlined in Moore and Reynolds (1997) for x-ray diffraction (XRD) analysis. Following XRD analysis, clay separates were submitted to GNS Science in New Zealand for δD and δ18O stable isotope analysis. All sample preparation and analytical methods are further described in Appendix 1.3 .9 .5 .2 Results of Clay StudiesASD and XRD interpreted mineralogy results are shown in Table 3.1, with ASD reflectance spectra and XRD patterns in Appendix 1. Results from both ASD and XRD were in agreement for each sample analyzed, although XRD allowed identification of secondary minerals beyond the singular dominant mineralogy identified by ASD. Results, from the limited sample population (n=5), show dickite within the core of mineralization, flanked by illite alteration, with distal kaolinite + illite. Chlinochlore identified in distal samples is likely a product of earlier metamorphism (Camilleri and Chamberlain 1997).49A previous study of clay alteration zonation at Long Canyon involved TerraSpecTM ASD analysis of drill core across several cross sections by (Rhys and Ross 2010). They defined a zonation of clay minerals outwards from gold mineralization as follows: dickite +/- kaolinite + crystalline white mica → kaolinite + sericite → illite + kaolinite → kaolinite + smectite. This zonation is largely in agreement with clay mineral distribution relative to gold mineralization observed in this study.Measured δD and δ18O ratios of clay minerals (Table 3.1) were used to calculate the composition of waters in equilibrium with clay minerals. Equilibrium H- and O-isotope fractionation equations of Sheppard and Gilg (1996) and Marumo et al. (1980) were used for the mineral-water systems of illite-muscovite-H20 and kaolinite/dickite-H20. Fluids associated with formation of Carlin-type deposits range from 180° - 240°C. The median temperature of this range, 210°C, was used to calculate each isotopic fluid composition, with the full temperature range calculated as uncertainty bars. Results are presented in Table 3.1 and plotted in Figure 3.23, along with boxes for the range of isotopic compositions for magmatic and metamorphic fluids, as well as the meteoric water line (Taylor 1974). For sample LCWAL260 containing unknown quantities of both kaolinite and illite, potential fluid equilibrium isotopic compositions were calculated for 100% kaolinite (δDH2O=-96.0) and 100% illite (δDH2O=-74.3). The plotted value for LCWAL260 in Figure 3.23 used weighted equilibrium fractionation equations for a clay mineral containing 50% kaolinite and 50% illite with error bars representing a monomineralic composition.3 .9 .5 .3 Interpretation of Clay StudiesHydrothermal fluids responsible for argillization at Long Canyon exhibit a strongly magmatic or possibly metamorphic calculated δDH2O and δ18OH2O isotopic composition, with little evidence for initial mixing with meteoric fluids. Fluids in equilibrium with illite exhibit a strongly magmatic signature, but could also be metamorphic illite. Kaolinite present in a distal position to mineralized cores exhibits a δDH2O composition that is primarily magmatic, with a slight mixing with meteoric fluids with a lighter δDH2O composition. Dickite is the only sample which exhibits a significantly lighter δDH2O composition indicative of mixing with meteoric water. Texturally, the analyzed dickite occurs late within void space of a strongly mineralized breccia overprinted by very fine-grained, fabric-retentive quartz. This dickite appears to be late ore-stage, just prior to silicification. As discussed earlier in Section 3.9.2.2, a possible mechanism for pervasive silicification is fluid mixing between a hydrothermal fluid and cooler meteoric waters. The light δDH2O composition of dickite appears to support this fluid mixing mechanism.503 .9 .6 Oxidation of the Long Canyon DepositThe Long Canyon deposit is extensively oxidized within and immediately surrounding mineralized breccia zones. Styles of oxidation include: 1) distal yellow oxidation grading into and overprinted by; 2) orange-red oxidation, typically barren of Au-mineralization; with 3) intense red oxidation increasing into mineralized zones. Oxidation alteration occurs most prevalently within siliciclastic beds and breccia zones where Fe-bearing minerals are concentrated. Massive micritic and coarse-grained carbonate rocks (limestone and dolomite) remain largely unoxidized even when occurring as clasts within extensively oxidized breccias. Hypogene sulfides have not been observed in an unoxidized state in this study, though past workers (e.g., Rhys and Ross 2010 and references therein) have observed arsenical pyrite and As- and Hg-sulfides encapsulated within quartz. Unoxidized diagenetic pyrite is prevalent within siliciclastic beds outside of mineralized zones. Where oxidized, these diagenetic pyrites typically produce a brick-red oxidation overprinting an earlier yellow oxidation within siliciclastic beds.XRD analysis of the 5 samples from LC556C (Figure 3.23 and Appendix 2) indicates hematite as the dominant iron oxide present in samples LCWAL189 and LCWAL233, and goethite in samples LCWAL243 and LCWAL260. This distribution of hematite in the oxidized core of the system and goethite dominating distal portions of oxidation is in agreement with previous ASD studies by (Rhys and Ross 2010) who broadly defined oxide zonation as hematite > goethite in mineralized the core of the system and goethite > hematite in distal extents. Interestingly, the limited ASD data run on the five samples from LC556C showed goethite as the dominant oxide (Appendix 2). This two-oxide characterization is a gross oversimplification, as a quick visual estimation of many samples typically suggests 2-4 oxide varieties by colour. A green alteration mineral, likely scorodite, is often present in mineralized areas where geochemistry indicates elevated As levels, although its presence is not a requisite for elevated As levels. Scorodite overprints brick-red hematite oxidation in ore zones, most commonly within basal breccias of the Pogonip Group. Hematite then overprints yellow-orange goethite in the distal extents of oxidation (Figure 3.17c). In Figure 3.29 scoroditic alteration is present in equal or greater abundances than hematite through breccia zones, whereas in Figure 3.30 scorodite is barely present as an oxidation mineral. This could possibly reflect varying ore-stage As-sulfide speciation between the 50 metre spaced ore zones which, in Carlin-type deposits, include arsenical pyrite, arsenopyrite, orpiment and realgar.51Stratigraphic unit-scale oxidation patterns within Pogonip limestones are displayed in Figure 3.31, Figure 3.30 and Figure 3.29. Drillholes LC452C and LC556C are on section L12800N and LC577C on section L12750N. LC452C hosts only a few thin zones of oxidation; no significant breccia development; and geochemical zonation of ore-stage trace elements that shows very little evidence of mineralizing fluid-flow within rocks intersected in this drillhole (e.g., Figure 3.25 and Figure 3.26). The paucity of oxidation suggests the importance of earlier permeability enhancement by structural preparation, dissolution of carbonate rocks and the presence of sulfide for later oxidizing fluids to generate the oxidation patterns we see today at Long Canyon. Pyrite generated during ore-stage fluid flow would have also contributed to enhanced oxidation profiles restricted to ore-stage fluid flow paths. Phillips (2005) summarizes this observation of enhanced weathering within Carlin deposits as largely due to oxidation of introduced ore-stage pyrite by ground waters summarized by the reaction,FeS2 + 3.75 O2 + 3.5 H2O = Fe(OH)3 + 2SO4-2 + 4H+the result of which, is more acidic waters flowing through ore deposit fluid pathways than groundwaters within surrounding host rocks. 3 .10 Lithogeochemistry of the Long Canyon Deposit3 .10 .1 Reactive Iron Control on Gold MineralizationLithologies with the highest Fe content consistently display the strongest correlation with gold grade across the Long Canyon deposit. This importance of abundant reactive Fe in host rocks for gold mineralization suggests sulfidation of iron within host rocks as a principal mechanism for precipitating gold from an infiltrating solution. Relative iron contents of lithologies are described below and displayed in Figure 3.24 and Table 3.2 from drill core handsamples (n=268) off sections L12800N and L12750N. Lithologies hosting the highest gold grades can be observed in Figure 3.21 where down hole gold grades are overlain on lithological units. The three units hosting the highest gold grades are Cnplus in the Notch Peak and Oplw and Oplsm in the Pogonip. Cnplus beds appear to be the most reactive unit present beneath the Notch Peak Dolomite, as it consistently displays the strongest alteration along the margins of mineralized areas. As mentioned, it is the preferred host for high grade gold mineralization in the Notch Peak, likely owing to its high Fe content of 10,000 – 20,000 ppm.52Basal units of the Lower Pogonip Group limestone, which host the majority of gold mineralization at Long Canyon, have the following ranges of iron concentrations. In all units, iron content correlates positively with abundance of siliciclastic silty laminations, presumably owing to increased abundance in Fe-bearing silicate minerals. The Oplw Fe content is approximately 3,000 – 8,000 ppm, except where silt abundance is particularly high. Oplsm Fe content within silt beds ranges from 8,000 ppm – 18,000 ppm, and within grainstone – micritic limestone interbeds from approximately 4,000 – 8,000 ppm. Opls Fe content varies from approximately 4,000 – 12,000 ppm. The highest gold grades are coincident with breccia units Opbx and Cnpbx in Figure 3.24c. Opls, Oplw and Oplsm units are anomalously low in the same figure, considering the apparent localization of Au mineralization within these units on Figure 3.21. The apparent low Au grade within these unbrecciated units indicates the strong control of brecciation on Au mineralization. Several more conclusions can be made regarding Fe contents within these breccia units. The first is that they contain high Fe contents compared to their respective protolith lithologies, likely owing to passive concentration of Fe during pre-Au-mineralization. The second is that where mineralized, the Fe contents of Cnpbx and Opbx are even higher, possibly reflecting further dissolution of carbonate rocks and concentration of Fe-bearing silicate minerals.Geochemical comparison of the Fe-content of non-dolomitized and non-altered Cnplbu and Cnplonc with their dolomitized counterparts in adjacent holes indicates that dolomitizing brines introduced little to no new Fe into the stratigraphy. Rather, carbonate rocks maintain an average Fe range of 1,000 – 2000 ppm, whether dolomitized or non-dolomitized. This is contradictory to dolomitization observed at gold deposits along the northern Carlin-trend (Meikle, Ren, Dee-Rossi) where ferroan hydrothermal dolomites represents the most favourable host for Au mineralization (Evans 2000; Emsbo, Hofstra et al. 2003). Dolomitized limestone at Long Canyon on the other hand is only a marginal host for localizing Au mineralization (e.g., Au distribution within dolomitized limestone in Figure 3.21 rarely occurs without significant brecciation).3 .10 .2 Gold-Silver Ratio of MineralizationGold-silver ratios at Long Canyon were assessed using assays (n=1877) from drillholes (n=22) present on sections 12800N and 12750N that were analyzed for bulk isotope studies in Chapter 4. Assays present 53within all lithologies were used to calculate the ratio. The mean Au-Ag ratio was calculated as 18.2. This is on the high end of typical calculated values for Carlin-type deposits (>3-20) (Hofstra and Cline 2000). There were significantly higher calculated Au-Ag ratios within specific lithologies.3 .10 .3 Ore-Stage Trace Element ZonationPrevious studies have shown the most strongly introduced element assemblage of Carlin-type deposits to be Au, As, Sb, Tl, Hg ± Te, ± W (Hofstra and Cline 2000). At Long Canyon the same trace element assemblage is associated with the main Au mineralization event. In this study the four main ore-stage trace elements (As, Sb, Tl and Hg) were contoured from relatively contiguous 3 metre geochemical samples from all drillholes across sections L12800N and L12750N. The aims were to determine: 1) trace element distribution relative to Au (e.g., Figure 3.21); 2) their distribution relative to O and C stable isotope ratios, discussed in Chapter 4; and 3) whether their distribution elucidated any structural and/or stratigraphic controls on fluid flow paths.Geochemical thresholds used for contouring were determined for each trace element As, Sb, Tl and Hg by spatial analysis using absolute values. Statistical methods of threshold estimation such as mean plus two standard deviation and probability plots (e.g., Sinclair 1991) were attempted and discarded as ineffective. For mean plus two standard deviation, the threshold values returned were strongly anomalous and nowhere near to geochemical background. The geochemical data set has a large sample population of moderate to high concentrations of ore-stage trace elements, resulting in a high mean from which background could not be determined. Probability plots also suggested threshold values that were obviously too high, based on the limited areal extent of a target trace element’s distribution along a cross-section when using the determined threshold. The decision to discard these results required the author’s understanding of the deposit geochemistry and its relative distribution. With other less common elements, the author may not have been able to determine if a determined geochemical threshold was valid. The most effective method was also the simplest. Drill core assay values plotted in Gemcom GEMSTM were lowered until the extents of the geochemical halo stopped expanding in a consistent pattern. Below a certain threshold, assay values would extend abruptly out along the entire drillhole, indicating threshold geochemical values for that particular element. Background values for As were determined to be < 20 ppm; for Sb < 0.15 ppm; for Tl <0.25 ppm; and for Hg < 0.1 ppm.54Geochemical data plotted as drillhole bar charts was contoured by hand in Adobe IllustratorTM at the determined threshold for each trace element. An anisotropy was applied to the contouring to reflect the moderate to steeply west- to north-dipping structural fabric (e.g., Section 3.8.3). Resulting plots of raw geochemical data and contouring are presented in Figure 3.25, Figure 3.26, Figure 3.27, Figure 3.28. Thresholded and contoured trace elements and gold exhibit a zonation from narrowest to widest halo size of Au < Sb < Tl < Hg < As. Trace element distribution relative to gold is very tightly nested within the Notch Peak limestones and dolomitized limestone. Within the Pogonip limestones, trace element distribution extends the farthest upwards into stratigraphy above strongly developed solution breccias. These patterns suggest a tight structural control on ore fluid flow within the deposit, with buoyant rise and exhaust of fluids within well-developed fracture permeability above Pogonip limestone-hosted solution collapse breccias.3 .11 Interpretation and Discussion The geology of the Long Canyon deposit described in this chapter establishes the geologic framework necessary for interpreting oxygen and carbon isotope alteration patterns in the following chapter. Of particular importance is the architecture of fluid flow paths, their temporal evolution, and the characteristics of ore and supergene fluids and related alteration. Carbonate sediments hosting the Long Canyon deposit have been variably subjected low grade metamorphism and boudinage with brittle fracturing, hydrothermal fluid infiltration resulting in decarbonatization, gold mineralization, dissolution brecciation, silicification, argillization, deep oxidation and continued brecciation from late karst processes.3 .11 .1 Fluid Flow Path Development in the Long Canyon DepositA first order control on localizing fluids at the Long Canyon deposit is the boudinage of the Notch Peak Dolomite and a penetrative nearly bedding-parallel foliation within enclosing Notch Peak and Pogonip Group limestones, all attributed to contraction during southeast directed thrusting during either the Jurassic Elko Orogeny (Thorman and Peterson 2003) or the Cretaceous Sevier Orogeny (Camilleri and Chamberlain 1997). Brittle moderate – high angle fault and fracture zones within the Cnpd dolomite are aligned with solution breccias above and below the Notch Peak Dolomite indicating that brittle fault damage zones likely extended into Cnpl and Opl limestone on margins of Cnpd dolomite, in which later solution breccias formed from mildly acidic infiltrating hydrothermal fluids. 55In regards to the role of brittle damage zones as pathways for hydrothermal fluid flow, it is unclear if brittle damage zones observed within dolomite extended into overlying and underlying limestone units at the time of hydrothermal fluid infiltration. There are no definitive observations that indicate whether brittle fault damage zones developed within Pogonip limestones coincident with dolomite-hosted fault generation, and were thus potential pathways of secondary permeability for later hydrothermal fluid flow, or whether they developed as suprasratal damage during collapse of units overlying voids created by dissolution of carbonate material by infiltrating hydrothermal fluids (solution collapse brecciation). While the initial fault zones within the massive dolomite are considered to have formed during the boudinage event as incipient boudin necklines with fracturing of the dolomite, the brittle damage zones within overlying limestones could have formed or been enhanced during one, or a combination of several different events including: 1) synchronously with boudinage; 2) during reactivation of faults during the contraction which generated S2; or 3) during reactivation of these fault zones with later extension. Interestingly, the angular relationship of foliation (S1) to crenulation lineation (S2) (Figure 3.8) appears to have a similar anisotropy as the orientation of dolomite boudins to fault breccia and connected damage zones above, below and through the dolomite boudins (Figure 3.8 c,d) and Figure 3.6). It is plausible that this late shortening event overprinted existing brittle damage zones and served to further develop secondary permeability. There is also further evidence for faulting along these structural damage zones post-mineralization, as jasperoids can be observed cut by brittle faults.Breccias within Long Canyon are a first order control on localizing high grade gold mineralization. Examination of breccia body morphology, clast speciation, morphology and distribution, matrix material, cross-cutting relationships and surrounding host rock reveals a complex history of polyphase brecciation with evidence of early brittle high angle fault and breccia development, modified by later hydrothermal solution brecciation which likely occurred progressively, resulting in re-brecciation of earlier breccias and development of extensive ‘karst tectonics’ (Sass-Gustkiewicz, Dzulynski et al. 1982) within the breccias and surrounding host rocks as supradjacent strata reacted to volume loss lower in the stratigraphy from solution brecciation and collapse including gravity induced normal faulting and crackle and mosaic breccia damage zones. The role of suprastratal damage zones in providing secondary permeability for fluid access to reactive siliciclastic strata is apparent from trace element halos coincident with the margins of collapse breccias in Pogonip stratigraphy. The wide separation between gold and trace element fronts within fluid flow paths above solution collapse breccias suggests that dynamic fracture permeability 56allowed buoyant rise and exhaust of fluids depleted in gold but not trace elements. Separation of geochemical fronts relative to oxygen isotope depletion fronts within fluid flow paths is investigated in Chapter 4.3 .11 .2 Comparison of the Long Canyon Deposit to Carlin-Type Deposits in NevadaSedimentary rock-hosted gold deposits in northeastern Nevada form a distinct class of mineral deposit. Gold deposits cluster along trends (e.g., Figure 1.1) and are classified as Carlin-type deposits. Salient characteristics of Carlin-type gold deposits are summarized in Section 1.2. Long Canyon has only been classified as the descriptive “sedimentary rock-hosted gold deposit”. Similarities between the Long Canyon deposit and Carlin-type deposits include:1. Sulfidation of Fe2+ in host rocks and precipitation of submicron Au into arsenian pyrite or arsenical rims on pre-existing pyrite as a likely mineralization style.2. Silty laminated limestone units as the predominant host for Au mineralization with intrusive rocks as a secondary ore host.3. Extensive dissolution of host carbonate rocks associated with Au mineralization.4. Argillization of silicate minerals in silty carbonate and intrusive host rocks, with similar temperature clay mineral assemblages.5. Variable amounts of hypogene quartz and silicification as jasperoid near orezones.6. Dissolution breccia and high angle faults as controlling features on mineralization.More work is required to determine the ore-fluid, Au and trace element source at Long Canyon in order to classify it as Carlin-type. However, even between Carlin-type deposits, the source of fluids and metals is not definitively characterized. Proposed models vary between magmatic, meteoric and metamorphic sources of fluids, and metals sourced from either scavenging of sediments or deep magmatic sources (e.g., Hofstra and Cline 2000). At Long Canyon, there is no evidence suggesting significantly different fluid chemistry, as alteration in carbonates shows many similarities to that in Carlin-type deposits. Accordingly, Carlin-type deposits are used as a basis for comparison of Long Canyon in this study, in particular past studies of oxygen and carbon isotope alteration in host carbonate host rocks, explored in Chapter 4.57BrecciaRelative Sea Level DeepeningCFigure 3 .1: Profile of a distally steepened carbonate ramp characteristic of the depositional environment present in the Great Basin during Upper Cambrian-Lower Ordovician.  Long Canyon stratigraphy repre-sent near-Outer to Inner Platform carbonates and Continental Clastics. GS=Grainstone, PS=Packstone, WS=Wackestone, MS=Mudstone.  Modified from Cook and Corboy (2004).BrecciaRelative Sea Level DeepeningCFigure 3 .2: Long Canyon stratigraphic sequence modified from Smith (2010). a) Upper Cambrian-Lower Orodovician stratigraphic sequence, with larger 3rd order cycles and smaller 4th and 5th order shallowing cycles interpreted from sequence stratigraphy. Deposit scale stratigraphy in B and C, determined by Fron-teer Gold geologists through drillcore; b) Lower Ordovician Pogonip Group; c) Upper Cambrian Notch Peak Formation. a bc58Figure 3 .3: Long Canyon property scale geology consists of Upper Cambrian to Permian carbonate and clastic sediments, as well as Cenozoic volcanics. Mapping by Smith and Thompson (2009). Ore deposit outline in yellow as defined in 2010 Q1 by Fronteer Gold.59adetrital qtzb500 μmpyd1000 μmdolribbon chertcdolqtzfdolqtz matrixeqtzdoldolmm qtz200 μm400 μm1000 μmlstsltsltdolFigure 3 .4: Quartz varieties observable in host rocks at Long Canyon  a) Photomicrograph LCWAL026: Detrital quartz grains, subrounded to subhedral, irregular and broken within calcareous siliciclastic siltstone interbed of the Pogonip Limestone (Opls).  b) Photomicrograph LCWAL411: Metamorphic quartz aggre-gates aligned in layers within strongly foliated calcareous siltstones (Opls) and as pressure shadows on a euhedral pyrite grain.  c) Photomicrograph LCWAL037: Siliceous dolomite (Cnpd) with fine grained intrac-rystalline quartz aggregates between dolomite crystals, similar morphology to quartz observed in chert.  d) Photomicrograph LCWAL287: Ribbon chert margin in diagenetic dolomite (Cnpd).  e) Photomicrograph LCWAL294: Mottled Notch Peak Dolomite (Cnpd) of varied grain size with commonly observed high angle quartz veins crosscutting.  f) 6.5 cm diameter (HQ) drill core photo LC083C 205’-213’: Fault zone within Notch Peak Dolomite (Cnpd) with pervasively silicified matrix/cement of white quartz.  60a1 cmchert b400 μmdold2000 μmdol’d sltsparry dol fe oxfe oxg1 cmfe ox dolh500 μmdolfe oxc1 cmddol’d sltdolfe oxf500 μmdolqtze1 cmfqtzdolqtzdoldoldolFigure 3 .561Figure 3 .5: Dolomitization of Notch Peak Limestone a) Slabbed core LCWAL117: Well developed zebra dolomite replacing burrowed, cherty Notch Peak Limestone (Cnplbu). Note open space crystal growth of sparry white dolomite in vug. b) Photomicrograph LCWAL280: Saddle dolomite crystal growth of zebra dolomite. Note sweeping extinction and curved crystal faces and cleavages. c) Slabbed core LCWAL070: Dolomitized Upper Silty unit of Notch Peak Limestone (Cnplus) altered to typical tan colour with sparry white dolomite veins and oxidized sulphide pseudomorphs. d) Photomicrograph LCWAL070: Grain-size difference between dolomitized siltstone (lower) and dolomite vein (upper). Iron oxide within vein indicates contemporaneous sulphide growth. e) Slabbed core LCWAL278: Zebra dolomite alteration of oncolitic Notch Peak Limestone (Cnplonc) with late quartz growth in open space vugs. f) Photomicrograph LCWAL278: Paragenetically late quartz growth in dolomite vug with slightly curved saddle dolomite crys-tal faces with compositional zoning. g) Slabbed core LCWAL119: Sulphide growth (oxidized) syn or post dolomitization as coarse clots within dolomitized Cnplonc. Sulphide growth generally occurs as isolated interstitial grains or on dolomite growth lamellae in crystals. This coarse variety is rarely observed. No significant elevation in base metal content over the 5 ft interval to indicate it was anything other than pyrite. h) Photomicrograph LCWAL278: Sulphide (oxidized) along growth lamellae on dolomite rhomb PREVI-OUS PAGE.Figure 3 .6: Breccia distribution mapped from drill core photos from all diamond drill holes across sections L12800N (top) and L12750N (bottom). Contoured breccia outlines include multiple breccia types including cataclastic, carbonate dissolution collapse breccias and late calcite-cemented breccias. Some breccia out-lines are drawn with a prefered orientation of a moderately to steeply northwest dipping fabric intersecting the plane of the W40N - E40S oriented cross-section. NEXT PAGE. See text for details. 62LC480LC733CLC537CLC541CLC462CLC452CLC533CLC737CLC446LC548CLC445CLC524CLC673CLC564CLC556CLC666CLC555CLC442LC396LC516LC448Alteration Opdl Replacement Dolomite of Pogonip LimestoneCnpdl Hydrothermal Replacement Dolomite of Notch Peak LimestoneStructure Breccia Brecciation including Fault Breccia, Solution Breccia and Post-ore Calcite Cement BrecciaQalOplsCnpdOplwCambrian Notch Peak Group LithologiesOrdovician Pogonip Group LithologiesOplsmCnplbuCnploncCnplwCnplusAlluvium/ColluviumQuaternary LithologyLaminated to thin-bedded silty limestone and limy siltstoneMassive dolomiteMassive limestone with silt wispsSilty limestone alternating with thick-bedded massive limestone and at pebble conglomeratesMassive, burrowed, bioturbated, mottled limestone, often dolomitizedMassive oncolitic limestone, often dolomitizedMassive limestone with silt wispsUpper siltstone, 0.5-1 m beds of silty limestoneLong Canyon Section L12800NLooking N40E12800N12750N500 metresSection Location 3 g/t Au Cuto100 metres 1000 mX800 mX1750 mRLLC557CLC582CLC571CLC692CLC466CLC471CLC577CLC680CLC518LC539CLC412LC525LC424LC613CLC420LC405Structure Breccia Brecciation including Fault Breccia, Solution Breccia and Post-ore Calcite Cement Breccia12800N12750N500 metresSection Location3 g/t Au CutoAlteration Opdl Replacement Dolomite of Pogonip LimestoneCnpdl Hydrothermal Replacement Dolomite of Notch Peak LimestoneQalOplsCnpdOplwCambrian Notch Peak Group LithologiesOrdovician Pogonip Group LithologiesOplsmCnplbuCnploncCnplwCnplusAlluvium/ColluviumQuaternary LithologyLaminated to thin-bedded silty limestone and limy siltstoneMassive dolomiteMassive limestone with silt wispsSilty limestone alternating with thick-bedded massive limestone and at pebble conglomeratesMassive, burrowed, bioturbated, mottled limestone, often dolomitizedMassive oncolitic limestone, often dolomitizedMassive limestone with silt wispsUpper siltstone, 0.5-1 m beds of silty limestoneLong Canyon Section L12750NLooking N40E100 metres 1000 mX800 mX1750 mRLFigure 3 .663a b400 μmdc200 μm 200 μm200 μmdolomitedolomitecchemfracture contactcontactbt lathchertbt lathbt lathhemlamprophyree fAz Dip288 53311 19341 26314 322 2305 3314 23307 20295 46Figure 3 .7: a,b) Photomicrographs of LCWAL046 and LCWAL047 showing chlorite altered biotite laths and aggregates in unmineralized lamprophyre intrusive. c) Photomicrograph LCWAL073, XPL, oxidized, 4 ppm gold-mineralized lamprophyre intrusive with opaques overprinting biotite laths. d) Photomicrograph, reflected light, same image as figure c, hematite psuedomorphs overgrown on biotite laths, likely after pyrite, evidence for sulfidation. e) Mafic intrusive in drill hole LC555C, weakly gold mineralized, cutting dolomite at a high angle to bedding. Adjacent fractures cutting chert nodule of similar orienation indicate orientation of structural damage exploited by intruding mafic dyke. f) Stereographic projection of dike and sill orientations down drillhole LCG02.64aac750 μmbd200 μmearly foliationfoliationboudinageisoclinal rootless fold of beddinglate crenulation1 cm1 cmFigure 3 .8: a) Thinsection Scan LCWAL315: Well foliated silty limestone with elongated limestone bou-dins and coarse calcite and pyrite clots. Later cleavage crenulating at high angle to foliation.  b) Coreslab LCWAL274 Folded and foliated silty limestone with stretched limestone boudins.  c) Photomicrograph LCWAL332 Foliated siliciclastic silty limestone layers with elongate boudins of coarse calcite, all crenulat-ed by a later, high angle cleavage.   d) Thinsection Scan LCWAL406 Foliated silty limestone layering with thin micritic limestone interbeds boudinaged, rotated and overprinted by high angle cleavage.65Jasperoid ZoneRC Crack ZoneD iscover y ZoneCrevasse ZoneShadow ZoneL12000NL11900NWestern Long S ec tionL12750NL12800NWest ZoneCnplwQalCnpdOplsmOplwOplsCnpdCnpdCnpdCnpdCnplCnplCnplOplOplOplQalQalStretching LineationCrenulationLineation12800N12750N12000NSection Line11900NWestern Long Section54321Block Model Au Grade Cuto g/t1 g/t cuto500 metresLooking DownQal - Quaternary Alluvium/ColluviumOpls - Laminated to thin-bedded silty limestone and limy siltstoneCnpd - Massive DolomiteOplw - Massive limestone with silt wispsUpper Cambrian Notch Peak GroupLower Ordovician Pogonip Group Oplsm - Silty limestone alternating with thick-bedded massive limestone and at pebble conglomeratesCnplbu - Massive, burrowed, bioturbated, mottled limestone, often dolomitizedCnplonc - Massive oncolitic limestone, often dolomitizedCnplw - Massive limestone with silt wispsCnplus - Upper siltstone, 0.5-1.0 m beds of silty limestoneQuaternaryN50EFigure 3 .966Figure 3 .9: Top) Block model image of Long Canyon lithologies over the resource area. Middle Left) Gold distribution trends in 1 g/t cutoff block model with cross section locations of L12750N, L12800N, L12000N, L11900N and Western Long Section shown. Block model data current as of December 15, 2010 from Fronteer Gold. Middle Right) Local simplified geology in the vicinity of the deposit, modified from M.T. Smith presentation. Anticlines occur in Cnp limestones and synclines in Op limestones in between dolomite blocks due to limestone beds flowing into bouding necklines during extension. Stretching linea-tion is roughly perpendicular to boudin block orientation consistent with extension in a NW-SE direction. Crenulation lineation overprinting foliation indicates a second NW-SE shortening event. Normal fault on western margin of map area places Upper Ordovician over Lower Ordovician stratigraphy. Lower) Ordo-vician Pogonip lithologies are removed to show coincident gold block model trend with dolomite boudin necks and fractures. Block models created primarily in Gemcom GEMS by Fronteer Gold geologists. Main mineralized zones are displayed for reference, as well as cross section lines 12000N, 127500N and 12800N. PREVIOUS PAGE.Figure 3 .10: a) Long Canyon interpreted cross section 12000N from drill core, modified from Rhys (2010), displaying boudin block distribution and limestone bedding/S1. b) Oplsm silt beds display alpha angles that with the drill hole orientation and underlying contact with Cnpd indicate sub-horizontal bedding. c) Cnplus and dolomitized oncolite bedding displaying sub-horizontal bedding at bas of Cnpd boudin block. d) Oplsm bedding with near 0° degree alpha angles show bedding/S1 plunging over nose of dolomite boudin. e) Fault zone where Cnpd and most of Cnplbu is missing between Oplw and Cnplonc. f) Opdl clasts within solution breccia, below the level of Cnpd boudin block on SE side of intervening fault zone, indicating clockwise rotation of Cnpd blocks. g) Cnpdl bedding with zebra-texture and high alpha angles displaying near hori-zontal bedding adjacent to fault to NW which accomodated rotation of dolomite blocks. NEXT PAGE.67Lim est one,  silt y, massive(Oplsm) Limestone, silty, laminated(Opls) Limestone, silty, wispy(Oplw)Dolom it e (C np d)Dolomitized Limestone (Cnpdl)CnploncBrecci at ed Dolom it eLim est one, undifferentiated (C np l)H em at it ic  b recc ias Bedding formlinesDrill Hole (DDH +RC)Ordovician Pogonip Group Cambrian Notch Peak FormationOverburdenModifed from Rhys 2010 and Mandal 2000 & Druguet 2009Tensile Fracture Boudinage Extensional Shear Fracture BoudinageNW SETensile Fracture Boudinage Extensional Shear Fracture Boudinage1 cmUPbbacdefgcdgfeModified from Rhys (2010)Figure 3 .1068Lim est one,  silt y, massive(Oplsm) Limestone, silty, laminated(Opls) Limestone, silty, wispy(Oplw)Dolom it e (C np d)Dolomitized Limestone (Cnpdl)CnploncBrecci at ed Dolom it eLim est one, undifferentiated (C np l)H em at it ic  b recc ias Bedding formlinesDrill Hole (DDH +RC)Ordovician Pogonip Group Cambrian Notch Peak FormationOverburdenModifed from Rhys 2010 and Mandal 2000 & Druguet 2009Tensile Fracture Boudinage Extensional Shear Fracture BoudinageNW SETensile Fracture Boudinage Extensional Shear Fracture Boudinage1 cmUPFigure 3 .11: Above) Long Canyon interpreted cross sec-tion 12000N from drill core, modified from Rhys (2010), displaying different boudin failure modes recognized by Mandal (2000).    Middle) Outcrop scale analogues in Cap de Creus, NE Spain, displaying similar boudin failure modes as above, observed by Druguet (2009) in a marble-metapsammite multilayer. Normal faulting accompanies boudin failure and boudin neck infill. Similar normal fault orientations hosting mineralization are observed at Long Canyon, although their timing is generally attrib-uted to extension occuring post boudinage deformation. These boudin/fault timing examples show initial damage zones may have been established during earlier deforma-tion. Close examination of extensional boudin neck (left) reveals predominantly inflow of overlying stratigraphy.    Lower Left) Unoriented rill core sample LCWAL274 from LC533C. Foliation development has folded lime-stone beds as tight isoclinal folds in middle of core. At base of drillcore a limestone bed has been stretched into boudins with an apparent antithetic rotation on boudins. Close examination of boudin necks reveals predominantly inflow of underlying silty limestone.69Figure 3 .12: Notch Peak Dolomite and dolomitized Limestone fault breccia and overprinting alteration.  a) Slabbed core LCWAL196: Subangular matrix supported breccia within dolomitized burrowed and cherty limestone (Cnplbu). Matrix is weakly siliceous, oxidized and comminuted. Handsample contains 2.6 ppm Au and interestingly only 24 ppm As.  b) Photomicrograph LCWAL196: Comminuted dolomite breccia matrix with subhedral quartz grain growth, finer grained quartz as cement and iron oxide after sulphide pseudomorphs.  c) Slabbed core LCWAL380: High angle fault damage zone in Notch Peak Dolomite (Cnpd). Comminuted matrix strongly oxidized and weakly silicified. Handsample contains 0.13 ppm Au.  d) Photomicrograph LCWAL380: Dolomite fault breccia matrix with sub – euhedral quartz grain growth and finer-grained quartz cementing matrix.  e) Thin Section Scan LCWAL297: Notch Peak Dolomite faulted and milled along high angle fracture zones. Late high angle stylolites after brecciation.  f) Photomi-crograph LCWAL297: Dolomite fault breccia of milled subangular dolomite clasts aligned along fractured damage zones supported in a comminuted matrix. High angle stylolite oxidized and cross cutting breccia.  g) Slabbed core LCWAL298: Notch Peak Dolomite fault breccia of angular – subangular clasts aligned in a high angle fabric. Handsample contains 0.13 ppm Au.  h) Photomicrograph LCWAL286: Notch Peak Do-lomite fault fabric of comminuted dolomite matrix supporting subangular siliceous dolomite breccia clasts. NEXT PAGE.70achertb400 μmqtzddolfe oxqtzg dol hdolqtzcbdoldolfdolbxe1 cmddol1 cm400 μmdol1 cm1000 μm1000 μm1 cmdoldoldolbxdoldoldolbxstylolitedolbxUp HoledolbxstylolitefdolbxFigure 3 .1271a 12800N b 12750NDrillhole Azimuth Plunge Median α <LC556C 310° -80° 32.5°LC555C 310° -60° 15°LC564C 310° -72° 25°LC541C 130° -77° 50°Drillhole Azimuth Plunge Median α <LC577C 310° -68° 15°LC557C 130° -60° 40°LC571C 310° -90° 25°LC692C 130° -60° 40°Figure 3 .13: Equal Area stereonet plots utilizing GEOrient (Holcombe 2011), depicting selected drillholes (open dots) from sections 12800N and 12750N with median alpha angle of dominant fault/fracture fabric within Cnpd plotted as small circles around drillholes. Alpha angles were measured from core photos, with an estimated +/- 10 degree error.  Intersections of small circles represent range of poles to likely fault plane orientations. a) 12800N, possible fault orientations dip 50 degrees towards 273 to 350 degrees. b) 12750N, possible fault orientations dip 65 degrees towards 258 to 002 degrees.72Figure 3 .14: a) Modelled dolomite with drillhole inter-cepts of Opdl and Cnpdl above and below Cnpd. Where boudin necks occur with wide extensional separation, there is a notable absence of Opdl drill intercepts. Conversely where Opdl intercepts are more frequent, boudinage mode is by shear failure with fracturing of the dolomite without extension. b) Variation of boudinage failure mode with respect to layer-thickness ratio and strength ratio between brittle and ductile layers. c) Variation of boudin aspect ratio with layer-thickness ratio. From Mandal et al, 2000.Tensile Fracture Boudinage DomainCnpdl - Dolomitized Notch Peak LimestoneOpdl - Dolomitized Pogonip LimestoneDrillhole DistributionShear Fracture Boudinage Domainbca73Figure 3 .15: Ore stage solution breccia examples and alteration from Pogonip and Notch Peak Limestone breccias  a) Coreslab LCWAL021: Matrix supported solution breccia in Oplw limestone with concentra-tion of undissolved detrital, metamorphic and hydrothermal quartz grains shown in inset. Matrix is weakly calcite cemented. Sample interval contains 10.25 ppm Au.  b) Photomicrograph LCWAL071: Silicified solution breccia within Cnplw and stratigraphically below dolomitized limestone. Large euhedral and doubly terminated quartz grains in breccia. Breccia margin shows contact with later calcite cement breccia. Sample interval contains 0.58 ppm Au. Inset shows thin section source coreslab with early silicified breccia on left and later comminuted, oxidized, calcite cemented breccia on right.  c) Coreslab LCWAL150: Cnplw limestone solution breccia, with oxidized matrix supporting grey apparently unaltered limestone clasts, all cut by later coarse calcite cement breccia. Sample interval contains 1.96 ppm Au.  d) Photomicrograph LCWAL150: Limestone solution breccia exhibiting ‘overpacked’ texture with interpenetration of limestone clasts along stylolite contacts. Iron oxide and clay matrix host both detrital and hydrothermal quartz grains, with calcite cement. Hydrothermal quartz growth occurs out of frame as insitu growth on wall rock and limestone clast margins indicating hydrothermal quartz grains authigenic.  e) Coreslab LCWAL220: Oplsm limestone solution breccia from within high grade mineralized core of drillhole LC556C on 12800N. Cal-careous silt bedding fragments in matrix supported breccia with sulphide (now iron oxide pseudomorphs) deposition predominantly along clast margins and clast fractures indicating sulfidation occurred late in solution brecciation event. Thin section shows similar pattern of hydrothermal quartz grain growth, with late pervasive silicification of bedding clasts and matrix. Handsample contains 3.97 ppm Au. f) Coreslab LCWAL221: Oplsm limestone solution breccia above figure e). Stratified insoluble solution breccia resid-uae and silt bedding clasts. Dark maroon areas are very fine-grained iron oxide pseudomorphs after sulfide clots. Bright white clay is dickite. Thin section study shows abundant hydrothermal quartz, locally appear-ing as overgrowths on earlier sulfide. Handsample contains 6.4 ppm Au. g) Coreslab LCWAL224: Oplsm limestone bedding collapse breccia. Iron oxide pseudomorphs after sulfide distribution pattern indicates sulfidizing fluids preferentially exploited fractures and bedding. Sample is pervasively silicified. Handsam-ple contains 10.95 ppm Au.  h) Coreslab LCWAL226: Oplsm limestone pervasively replaced by iron oxide pseudomorphs after sulfide. The extent of pre-sulfidation brecciation is difficult to determine, although surrounding strata relatively intact and hosting similar apparent sulfidation along bedding. Handsample contains 45.10 ppm Au and greater than >1% As, providing a strong geochemical and textural argument for gold deposition through sulfidation with arsenical pyrite. NEXT PAGE.74a1 cmb400 μmd1000 μmclastqtzstylolitegcdcclstf1 cme1 cm1 cm1 cmdkthydroqtzfe oxsilicified brecciabrecciamarginh1 cmUp Holefe ox fe oxlstclastclastUp HoleUp Holesolution breccia residuumsilicified brecciasilicified brecciasilicified brecciaFigure 3 .15751) LC591C 62 - 67 ft, Au below detection limit. Contiguous undisturbed bedding overlying zone of dissolu-tion. Thin high-angle calcite veins exploiting fractures characteristic of strata suprajacent to solution brecciation.2) LC591C 87 - 92 ft, 4 ppm Au. Relatively contiguous bedding, overprinted by crackle and mosaic breccia-tion and small normal faults. Volume loss in silty beds leads to bedding collapse clay rich seams.3) LC591C 116 .5 - 122 ft, 11.65 ppm Au. Chaotic matrix to clast-supported breccia. Rotated and displaced bedding and limestone clasts, mud matrix, limestone clasts rounded and embayed from acid attack.4) LC591C 178 - 182 .5 ft, 9.16 ppm Au. Mud matrix-supported basal solution breccia. Matrix consists domi-nantly of insoluble silicate minerals and clays. Cave floor bedding contiguous with matrix-filled seams and fractures.abcdFigure 3 .16: Drillhole LC591C. Commonly observed solution collapse breccia facies.76a1 cmbc1 cm1 cmFigure 3 .17: Extensive fracturing and  small-scale normal faulting occurs in Pogonip limestone overlying areas of volume loss through solution brecciation. Occuring ubiquitously above solution breccias, this suprastratal deformation is notably absent where underlying limestones are unbrecciat-ed. Examples from drillhole LC533C on L12800N a) LCWAL349, 224 ft: small scale fracturing and normal faulting of Opls limestone. Silty lamina-tions exhibit carbonated dissolution and flow of siliciclastic material into fracture. Inset shows trashzone of limestone bedding chips, largely cal-cite with little quartz, amidst siliciclastic material.   b)LCWAL330, 461 ft: fracturing of Opls lime-stone with normal slip. Dilatant zones filled with siliciclastic residuum, indicating likely carbonate dissolution along fractures, which are filled with calcite veinlets. c) LCWAL339, 330 ft: fracturing within Opls limestone, exhibiting 2-3 distinct fluid flow events, consisting of two oxidation events and calcite deposition.77Figure 3 .18: Hydrothermal quartz grains and silicification observable at Long Canyon  a) Thin Section Scan LCWAL021: Ore stage solution breccia at base of Oplw consisting of dissolved limestone and cal-careous siltstone residuum with detrital and hydrothermal quartz grains within iron oxide and clay matrix. Sample interval contains 10.25 ppm Au. Inset photomicrograph of solution breccia residuum with doubly terminated and sub-rounded to sub-euhedral hydrothermal quartz grains amongst irregularly shaped detrital quartz hosted in iron oxide and clay matrix with minor calcite cement. b) Photomicrograph LCWAL220: Hydrothermal quartz, doubly terminated subhedral to euhedral grains grown within Oplw limestone dis-solution breccia, overprinted by fine grained-microcrystalline quartz. Handsample contains 4 ppm Au.  c) Thin Section Scan LCWAL059: Bedding collapse solution breccia of Opls calcareous siltstone bedding with oxidation and rotation of silt bedding fragments and concentration of quartz grains.  d) Photomicro-graph LCWAL059: Quartz species in iron oxide and clay matrix amidst bedding fragments. Quartz varieties include fine grained aggregates of metamorphic quartz, detrital quartz and doubly terminated hydrothermal quartz grains. Sample interval contains 0.13 ppm Au.  e) Coreslab LCWAL037: Several breccia generations within dolomitized Cnplbu, with an earlier light brown coloured, more competent breccia cut by a later ma-roon, iron oxide rich breccia, all cut by late calcite veins.  f) Photomicrograph LCWAL 037: Dolomitized limestone clast in iron oxide matrix breccia with doubly terminated, euhedral hydrothermal quartz grain. Sample interval contains 0.25 ppm Au.  g) Thin Section Scan LCWAL034: Several generations of dolo-mitized limestone (Cnplbu) breccia. Earlier breccia clast appears to host tectonic fabric of aligned dolomite fragments within comminuted matrix.  h) Photomicrograph LCWAL034: Early tectonic dolomite fault brec-cia clast within a later breccia hosting hydrothermal quartz grains and possibly a weakly silicified matrix. Sample interval contains 0.40 ppm Au. NEXT PAGE.78500 μmqtzqtza1 cmbd400 μmhydro qtzg1 cmh1000 μmearly dol bxc1 cmdf1000 μme5 mmmm qtzfe ox qtzhydro qtzhfearlydol bxlatedol bxlate dol bxhydro qtzearlybxlatebxlatebxslt bed fragmentFigure 3 .1879?MixingCooling w/o quartz precipitation12344Figure 3 .19: Quartz and amorphous silica solubility as a function of temperature in a 5 wt percent NaCl solution at 500 bars, modified from Hofstra (2000). Displayed are two possible cooling paths of Long Canyon ore fluids as indicated by quartz textures: 1) a silica-rich fluid cools until reaching the quartz-satu-ration curve, after which it cools along the curve while precipitating euhedral hydrothermal quartz grains; 2) mixing (red path) with a cooler meteoric fluid occurs and the resulting fluid becomes rapidly quartz-supersaturated; 3) no fluid mixing (blue path) occurs, fluid continues cooling until 180-200 °C, after which quartz precipitation stops (Fournier, 1985; Hofstra, 2000) and fluid continues cooling in a quartz supersatu-rated state; 4) quartz supersaturation results in precipitation of large volumes of fine-grained, non-texturally destructive silicification within breccias.801 cma1000 μmbd400 μmdolsltfe oxbx matrixg1 cmsilicifiedbxcch2 cmccc1 cmdbxf1000 μmmm qtze1 cmsltclastccUp Holefsltclastsltclastbxclasthydro qtzccbsltclastdolbx matrixbx matrixcchydro qtzccsltdoldolccccdolFigure 3 .2081Figure 3 .20: Post Ore Calcite Cement Breccias within Pogonip and Notch Peak lithologies. a) Thin Section Scan LCWAL033: Dolomitized wispy limestone (Cnplw) clasts with early oxidized and silicified breccia matrix on margins, all brecciated by a later coarse calcite cement breccia. Inset of the thin section source coreslab. Sample interval contains 2.5 ppm Au. b) Photomicrograph LCWAL033: Calcite cement breccia with clasts of an earlier breccia of dolomitized limestone with a comminuted matrix hosting euhedral, dou-bly terminated quartz grains and iron oxides. c) Thin Section Scan LCWAL035: Calcite cement polymictic breccia within dolomitized Cnplbu. Subangular dolomite clasts within an earlier iron oxide rich breccia matrix occur together as large composite clasts within coarse calcite cement. Isolated iron oxide breccia matrix as clasts indicate coarse calcite was not just a cement to the earlier breccia. Sample interval contains 6.79 ppm Au. d) Photomicrograph LCWAL035: Early iron oxide breccia matrix hosting euhedral, doubly terminating quartz grains. Clasts of dolomitized siltstone (Cnplus) identified based on dolomite grainsize indicates early iron oxide rich polymictic breccia was the likely the result of reverse faulting as Cnplus is stratigraphically below Cnplbu. Hydrothermal quartz likely associated with 6.79 ppm Au mineralization in sample interval. All brecciated by later coarse calcite.  e) Thin Section Scan LCWAL060: Early calcareous siltstone bedding collapse breccia within Oplsm with comminuted iron oxide and quartz grain-rich matrix, all brecciated by a later coarse calcite cement. Sample interval contains 0.75 ppm Au. f) Photomicrograph LCWAL060: Coarse, late calcite cement brecciating an earlier quartz, clay and iron oxide-rich matrix supporting calcareous siltstone clasts. Isolated clasts of early breccia matrix indicates coarse calcite was a later event and not a progressive cementation of quartz-bearing breccia. g) 6.5 cm (HQ) diameter drill core LC173C 153.1’: Coarse, late calcite breccia cement within Cnplw with clasts of silicified early calcareous silt bedding collapse breccia. h) 8.5 cm (PQ) diameter drill core LCBS2A 68’: Coarse, late calcite cement within Oplw brecciating oxidized calcareous siltstone and grainstone interbeds. Liesengang banding which occurs elsewhere nucleating from oxidizing fluid flow along fractures is present here, indicating extension and calcite matrix infill occurred along preexisting fractures. PREVIOUS PAGE.Figure 3 .21: Contoured Au grades >0.01ppm on sections L12800N (top) and L12750N (bottom) overlain on lithology and contoured breccia distribution. Widths of down hole bar chart displaying Au grade is con-strained to 5 ppm Au or less to maintain clarity on the figure. Au contours are often drawn with a prefered orientation of a moderately to steeply northwest-dipping fabric intersecting the plane of the W40N - E40S oriented cross-section. This orientation follows that of the breccias, which are the prefered host and first order control on the distribution of Au mineralization. NEXT PAGE.82LC480LC733CLC537CLC541CLC462CLC452CLC533CLC737CLC446LC548CLC445CLC524CLC673CLC564CLC556CLC666CLC555CLC442LC396LC516LC44810.0 -    63.05.0 -    10.00.3 -    1.00.1 -    0.31.0 -    3.03.0 -    5.0Drillhole Gold Assay in ppmQalOplsCnpdOplwCambrian Notch Peak Group LithologiesOrdovician Pogonip Group LithologiesOplsmCnplbuCnploncCnplwCnplusAlluvium/ColluviumQuaternary LithologyLaminated to thin-bedded silty limestone and limy siltstoneMassive dolomiteMassive limestone with silt wispsSilty limestone alternating with thick-bedded massive limestone and at pebble conglomeratesMassive, burrowed, bioturbated, mottled limestone, often dolomitizedMassive oncolitic limestone, often dolomitizedMassive limestone with silt wispsUpper siltstone, 0.5-1 m beds of silty limestoneAlteration OpdlAuReplacement Dolomite of Pogonip LimestoneCnpdl Hydrothermal Replacement Dolomite of Notch Peak LimestoneGold mineralization outline greater than 0.1 ppmStructure Breccia Brecciation including Fault Breccia, Solution Breccia and Post-ore Calcite Cement BrecciaLong Canyon Section L12800NLooking N40E12800N12750N500 metresSection Location 3 g/t Au Cuto100 metres 1000 mX800 mX1750 mRLLC557CLC582CLC571CLC692CLC466CLC471CLC577CLC680CLC518LC539CLC412LC525LC424LC613CLC420LC405QalOplsCnpdOplwCambrian Notch Peak Group LithologiesOrdovician Pogonip Group LithologiesOplsmCnplbuCnploncCnplwCnplusAlluvium/ColluviumQuaternary LithologyLaminated to thin-bedded silty limestone and limy siltstoneMassive dolomiteMassive limestone with silt wispsSilty limestone alternating with thick-bedded massive limestone and at pebble conglomeratesMassive, burrowed, bioturbated, mottled limestone, often dolomitizedMassive oncolitic limestone, often dolomitizedMassive limestone with silt wispsUpper siltstone, 0.5-1 m beds of silty limestoneLong Canyon Section L12750NLooking N40EAlteration OpdlAuReplacement Dolomite of Pogonip LimestoneCnpdl Hydrothermal Replacement Dolomite of Notch Peak LimestoneGold mineralization outline greater than 0.1 ppmStructure Breccia Brecciation including Fault Breccia, Solution Breccia and Post-ore Calcite Cement Breccia12800N12750N500 metresSection Location3 g/t Au Cuto10.0 -    63.05.0 -    10.00.3 -    1.00.1 -    0.31.0 -    3.03.0 -    5.0Drillhole Gold Assay in ppm100 metres 1000 mX800 mX1750 mRLFigure 3 .2183Figure 3 .22: SEM backscatter images of iron oxide pseudomorphs with detectable As. LCWAL040 (top) 35 μm Fe-oxide psuedomorph, likely after pyrite, within quartz-replaced Cnpdl (0.21 ppm Au/ft). Image displays an alteration rim, spot analysis indicates arsenical pyrite with As as an identifiable ore-stage trace element. LCWAL061 (bottom) 5 x 10 μm zoned Fe-oxide psuedomorph after pyrite grain encapsulated within silica overprinting Oplw solution breccia matrix (15.75 ppm Au/5ft). Spot analysis indicates Fe, O and As suggesting oxidation of arsenian pyrite.84cdfgeLC480LC733CLC537CLC541CLC462CLC452CLC533CLC737CLC446LC548CLC445CLC524CLC673CLC564CLC556CLC666CLC555CLC442LC396LC516LC4483.74 -    8.008.00 -    12.0017.50 -    19.0019.00 -    25.1616.00 -    17.5012.00 -    16.00MDRU-MIA Oxygen Isotope 18O/16O Drillhole Values QalOplsCnpdOplwCambrian Notch Peak Group LithologiesOrdovician Pogonip Group LithologiesOplsmCnplbuCnploncCnplwCnplusAlluvium/ColluviumQuaternary LithologyLaminated to thin-bedded silty limestone and limy siltstoneMassive dolomiteMassive limestone with silt wispsSilty limestone alternating with thick-bedded massive limestone and at pebble conglomeratesMassive, burrowed, bioturbated, mottled limestone, often dolomitizedMassive oncolitic limestone, often dolomitizedMassive limestone with silt wispsUpper siltstone, 0.5-1 m beds of silty limestoneAlteration OpdlAuReplacement Dolomite of Pogonip LimestoneCnpdl Hydrothermal Replacement Dolomite of Notch Peak LimestoneGold mineralization outline greater than 0.1 ppmStructure Breccia Brecciation including Fault Breccia, Solution Breccia and Post-ore Calcite Cement BrecciaLong Canyon Section L12800NLooking N40E12800N12750N500 metresSection Location 3 g/t Au Cuto100 metres 1000 mX800 mX1750 mRLcdegf- 1 6 0- 1 4 0- 1 2 0- 1 0 0- 8 0- 6 0- 4 0- 2 00- 2 0 - 1 5 - 1 0 - 5 0 5 1 0 1 5 2 0 2 5 3 0LCWAL2 2 1LCWAL2 3 3LCWAL1 8 9LCWAL2 4 3LCWAL2 6 0MAGMATICMETAMORPHICMETEORICWATER LINECnozoic      Cretaceous      JurassicδDH2O(per mil) VSMOWδ18OH2O(per mil) VSMOWLCWAL189Section L12800N  Drillhole LC556CLCWAL221LCWAL233LCWAL243LCWAL260Figure 3 .23: a) plotted values of calculated fluid δD and δ18O values for H20 in equilibrium with silicate minerals. Error (uncertainty) bars represent possible fluid istopic compositions at temperatures from 180 - 240 °C. XRD analysis of clay separate mineralogy in Table 3.1. Additional y-axis error bars on sample LCWAL260 represent endmember isotopic compositions for 100% illite and 100% kaolinite in sample, as XRD indicates presence of both minerals but not relative abundances. Ceno-zoic, Cretaceous and Jurassic meteoric water lines (acounting for Paleo-latitudes) taken from Hofstra (1999). Also shown are magmatic and metamor-phic fluid isotopic compositions, and the meteoric water line from Taylor (1974). Calculated isotopic composition of fluids in equilibrium with clay minerals fall largely in the magmatic fluids field, while the dickite sample LCWAL221 appears to have a mixed with a component of meteoric fluid. b-g) - handsamples and their location on drillhole LC556C, from which clay separates were prepared and analyzed for δD and δ18O values. ab85Table 3 .1: Calculated fluid δD and δ18O values for H20 in equilibrium with silicate minerals in clay sepa-rates collected in this study from 5 hand samples down drill hole LC556C. Whole rock mineral identi-fication by TerraspecTM analytical spectral device (ASD). Clay separate mineral identification by x-ray diffraction (XRD).ASD XR D Fluid T δD δ18 O δDH 2O* δ18 OH 2O*Sample Minerals Minerals (° C) (‰) (‰) (‰) (‰)LCW AL18 9 illite phengitic illite 210 -8 3.6 13.3 -63.6 6.8LCW AL221 dickite dickite 210 -117.4 7.4 -119.1 2.3LCW AL233 illite phengitic illite 210 -8 4.6 14.3 -64.6 7.8LCW AL243 - phengitic illite + /-chlinochlore 210 -95.1 13.7 -75.1 7.2LCW AL260 kaolinite kaolinite +  phengitic illite + /- chlinochlore 210 -94.3 13.7 -8 5.2 7.9* Isotope fractionation equations after Gilg and Sheppard (1996)Notes: Clay separates mineralogy determined by XRD, spectra in Appendix.86CnpbxCnpd CnpdlCnpdzCnploncCnplwCnplta Opbx Opdl OplsOplwOplsmLampCnpbxCnpd CnpdlCnpdzCnploncCnplwCnplusCnplta Opbx Opdl OplsOplwOplsmLampCnpbxCnpd CnpdlCnpdzCnploncCnplwCnplusCnplta Opbx Opdl OplsOplwOplsmLamplog Fe ppmlog Au ppmlog Fe ppma bc Figure 3 .24: Box and whisker plots of Fe and Au contents of lithologies sampled from drill-holes LC533C, LC555C and LC556C on section L12800N.   Boxes represent middle 50% of data. Whiskers are the 5% and 95% values. Black circles are the mean. Black lines are the median. Open circles are top or bottom 5% of data (outli-ers). Specific values displayed in Table 3.1. a) Fe content range of samples where Au is below detec-tion limit (0.005 ppm Au). b) Fe content range of all samples, Au-mineralized and unmineralized. c) Au content range of all samples. All lithology abbreviations as described in text, and in Table 3.1 footnotes. Of particular note are highest Au con-tents contained in units Opbx and Cnpbx - breccias within Pogonip Group and Notch Peak Formation rocks, respectively. Lamprophyres sampled on these sections significantly under-represent typical Au grade observed in mineralized lamprophyres elsewhere in the deposit.87Table 3 .2: Fe values of handsamples from drillholes LC533C, LC555C and LC556C on section L12800N. Two sample populations are presented: samples with Au assays below detection limit (0.005 ppm Au) (n=171); and samples with both no Au and detectable Au assays (n=268).Lithology Min Median Max n Min Median Max n Min Median Max nCnpbx 1400 1400 1400 1 900 2200 35800 30 0.005 0.138 6.42 30Cnpd 400 633 1200 12 300 900 48100 24 0.005 0.007 3.41 24Cnpdl 400 1300 2800 7 400 1550 2800 10 0.005 0.005 0.212 10Cnpdz 1200 1500 1600 3 1200 1550 3400 6 0.005 0.013 3400 6Cnplonc 3500 3500 3500 1 3500 3500 3500 1 0.005 0.005 0.005 1Cnplta 12200 12200 12200 1 12200 13550 14900 2 0.005 0.02 0.034 2Cnplus - - - - 20000 20000 20000 1 3.53 3.53 3.53 1Cnplw 1400 3600 16500 13 1100 3000 16500 23 0.005 0.005 4.29 23Opbx 9400 9400 9400 1 4800 15700 51200 8 0.005 5.185 10.95 8Opdl 1600 4150 5100 4 1400 3550 6800 12 0.005 0.012 2.41 12Opls 1200 9900 23500 95 1200 9950 23500 96 0.005 0.005 0.013 96Oplw 3200 5600 10000 7 3200 6100 10000 9 0.005 0.005 0.902 9Oplsm 1600 11400 24500 24 1600 12300 85300 41 0.005 0.005 45.1 41Lamp 51200 51650 52100 2 51200 53300 59000 5 0.005 0.044 1.375 5Au (ppm)[Fe] (ppm), all samples[Fe] (ppm), sample no AuLithology abbreviations as presented previously. Abbreviations not previously described include Cnpbx - Notch Peak Breccia; Cnpdz - Notch Peak Zebra Dolomite (Cnpdl equivalent); Cnplta - Notch Peak Tan Altered Lime-stone (altered Cnplus equivalent); Opbx - Pogonip Breccia; Lamp - Lamprophyre.88Figure 3 .25: Cross-sections L127800N (top) and L12750N (bottom) displaying hand contoured As from drillhole sampling at a cutoff of >20 ppm As. Overlain on contoured Au>0.1ppm (red outline), lithologies and mapped breccia. Figure 3 .26: Cross-sections L127800N (top) and L12750N (bottom) displaying hand contoured Sb from drillhole sampling at a cutoff of >0.15 ppm Sb. Overlain on contoured Au>0.1ppm (red outline), lithologies and mapped breccia. Figure 3 .27: Cross-sections L127800N (top) and L12750N (bottom) displaying hand contoured Tl from drillhole sampling at a cutoff of >0.1 ppm Tl. Overlain on contoured Au>0.1ppm (red outline), lithologies and mapped breccia. Figure 3 .28: Cross-sections L127800N (top) and L12750N (bottom) displaying hand contoured Hg from drillhole sampling at a cutoff of >0.25 ppm Hg. Overlain on contoured Au>0.1ppm (red outline), lithologies and mapped breccia. 89123456A B C D E FLC480LC733CLC537CLC541CLC462CLC452CLC533CLC737CLC446LC548CLC445CLC524CLC673CLC564CLC556CLC666CLC555CLC442LC396LC516LC4481,000 -   15,000500 -    1,00075 -    10050 -    7520 -    50100 -    200200 -    500Drillhole Arsenic in ppm Alteration OpdlAuReplacement Dolomite of Pogonip LimestoneCnpdl Hydrothermal Replacement Dolomite of Notch Peak LimestoneGold mineralization outline greater than 0.1 ppmAs Arsenic outline greater than 20 ppmQalOplsCnpdOplwCambrian Notch Peak Group LithologiesOrdovician Pogonip Group LithologiesOplsmCnplbuCnploncCnplwCnplusAlluvium/ColluviumQuaternary LithologyLaminated to thin-bedded silty limestone and limy siltstoneMassive dolomiteMassive limestone with silt wispsSilty limestone alternating with thick-bedded massive limestone and at pebble conglomeratesMassive, burrowed, bioturbated, mottled limestone, often dolomitizedMassive oncolitic limestone, often dolomitizedMassive limestone with silt wispsUpper siltstone, 0.5-1 m beds of silty limestoneBreccia Brecciation including Fault Breccia, Solution Breccia and Post-ore Calcite Cement BrecciaStructure Long Canyon Section L12800NLooking N40E12800N12750N500 metresSection Location 3 g/t Au Cuto100 metres 1000 mX800 mX1750 mRL123456A B C D E FLC557CLC582CLC571CLC692CLC466CLC471CLC577CLC680CLC518LC539CLC412LC525LC424LC613CLC420LC405QalOplsCnpdOplwCambrian Notch Peak Group LithologiesOrdovician Pogonip Group LithologiesOplsmCnplbuCnploncCnplwCnplusAlluvium/ColluviumQuaternary LithologyLaminated to thin-bedded silty limestone and limy siltstoneMassive dolomiteMassive limestone with silt wispsSilty limestone alternating with thick-bedded massive limestone and at pebble conglomeratesMassive, burrowed, bioturbated, mottled limestone, often dolomitizedMassive oncolitic limestone, often dolomitizedMassive limestone with silt wispsUpper siltstone, 0.5-1 m beds of silty limestoneLong Canyon Section L12750NLooking N40EAlteration Opdl Replacement Dolomite of Pogonip LimestoneCnpdl Hydrothermal Replacement Dolomite of Notch Peak LimestoneAs Arsenic outline greater than 20 ppmAu Gold mineralization outline greater than 0.1 ppmBreccia Brecciation including Fault Breccia, Solution Breccia and Post-ore Calcite Cement BrecciaStructure 12800N12750N500 metresSection Location3 g/t Au Cuto1,000 -   15,000500 -    1,00075 -    10050 -    7520 -    50100 -    200200 -    500Drillhole Arsenic in ppm100 metres 1000 mX800 mX1750 mRLAsAsFigure 3 .2590123456A B C D E FLC480LC733CLC537CLC541CLC462CLC452CLC533CLC737CLC446LC548CLC445CLC524CLC673CLC564CLC556CLC666CLC555CLC442LC396LC516LC412LC448LC537CLC462CLC533CLC548CLC524CLC44210 -   505 -   101.0 -    1.251.25 -    1.51.5 -    50.75 -   1.00.5 -   0.750.15 -    0.20.2 -    0.30.3 -    0.5Drillhole Antimony in ppm Alteration OpdlAuReplacement Dolomite of Pogonip LimestoneCnpdl Hydrothermal Replacement Dolomite of Notch Peak LimestoneGold mineralization outline greater than 0.1 ppmSb Antimony outline greater than 0.15 ppmQalOplsCnpdOplwCambrian Notch Peak Group LithologiesOrdovician Pogonip Group LithologiesOplsmCnplbuCnploncCnplwCnplusAlluvium/ColluviumQuaternary LithologyLaminated to thin-bedded silty limestone and limy siltstoneMassive dolomiteMassive limestone with silt wispsSilty limestone alternating with thick-bedded massive limestone and at pebble conglomeratesMassive, burrowed, bioturbated, mottled limestone, often dolomitizedMassive oncolitic limestone, often dolomitizedMassive limestone with silt wispsUpper siltstone, 0.5-1 m beds of silty limestoneBreccia Brecciation including Fault Breccia, Solution Breccia and Post-ore Calcite Cement BrecciaStructure Long Canyon Section L12800NLooking N40E12800N12750N500 metresSection Location 3 g/t Au Cuto100 metres 1000 mX800 mX1750 mRL123456A B C D E FLC448LC557CLC582CLC571CLC692CLC466CLC471CLC577CLC680CLC518LC539CLC412LC525LC424LC613CLC420LC405LC571CQalOplsCnpdOplwCambrian Notch Peak Group LithologiesOrdovician Pogonip Group LithologiesOplsmCnplbuCnploncCnplwCnplusAlluvium/ColluviumQuaternary LithologyLaminated to thin-bedded silty limestone and limy siltstoneMassive dolomiteMassive limestone with silt wispsSilty limestone alternating with thick-bedded massive limestone and at pebble conglomeratesMassive, burrowed, bioturbated, mottled limestone, often dolomitizedMassive oncolitic limestone, often dolomitizedMassive limestone with silt wispsUpper siltstone, 0.5-1 m beds of silty limestoneLong Canyon Section L12750NLooking N40EAlteration Opdl Replacement Dolomite of Pogonip LimestoneCnpdl Hydrothermal Replacement Dolomite of Notch Peak LimestoneSb Antimony outline greater than 0.15 ppmAu Gold mineralization outline greater than 0.1 ppmBreccia Brecciation including Fault Breccia, Solution Breccia and Post-ore Calcite Cement BrecciaStructure 12800N12750N500 metresSection Location3 g/t Au Cuto10 -   505 -   101.0 -    1.251.25 -    1.51.5 -    50.75 -   1.00.5 -   0.750.15 -    0.20.2 -    0.30.3 -    0.5Drillhole Antimony in ppm100 metres 1000 mX800 mX1750 mRLSbSbFigure 3 .2691123456A B C D E FLC480LC733CLC537CLC541CLC462CLC452CLC533CLC737CLC446LC548CLC445CLC524CLC673CLC564CLC556CLC666CLC555CLC442LC396LC516LC448LC537CLC462CLC533CLC737CLC548CLC445CLC524CLC673CLC44210 -   206 -   100.6 -    10.4 -    0.61 -    20.2 -    0.30.1 -    0.20.3 -    0.42 -    6Drillhole Thallium in ppm Alteration OpdlAuReplacement Dolomite of Pogonip LimestoneCnpdl Hydrothermal Replacement Dolomite of Notch Peak LimestoneGold mineralization outline greater than 0.1 ppmTl Thallium outline greater than 0.1 ppmQalOplsCnpdOplwCambrian Notch Peak Group LithologiesOrdovician Pogonip Group LithologiesOplsmCnplbuCnploncCnplwCnplusAlluvium/ColluviumQuaternary LithologyLaminated to thin-bedded silty limestone and limy siltstoneMassive dolomiteMassive limestone with silt wispsSilty limestone alternating with thick-bedded massive limestone and at pebble conglomeratesMassive, burrowed, bioturbated, mottled limestone, often dolomitizedMassive oncolitic limestone, often dolomitizedMassive limestone with silt wispsUpper siltstone, 0.5-1 m beds of silty limestoneBreccia Brecciation including Fault Breccia, Solution Breccia and Post-ore Calcite Cement BrecciaStructure Long Canyon Section L12800NLooking N40E12800N12750N500 metresSection Location 3 g/t Au Cuto100 metres 1000 mX800 mX1750 mRL123456A B C D E FLC557CLC582CLC571CLC692CLC466CLC471CLC577CLC680CLC518LC539CLC412LC525LC424LC613CLC420LC405LC571CLC692CLC466CLC420LC405QalOplsCnpdOplwCambrian Notch Peak Group LithologiesOrdovician Pogonip Group LithologiesOplsmCnplbuCnploncCnplwCnplusAlluvium/ColluviumQuaternary LithologyLaminated to thin-bedded silty limestone and limy siltstoneMassive dolomiteMassive limestone with silt wispsSilty limestone alternating with thick-bedded massive limestone and at pebble conglomeratesMassive, burrowed, bioturbated, mottled limestone, often dolomitizedMassive oncolitic limestone, often dolomitizedMassive limestone with silt wispsUpper siltstone, 0.5-1 m beds of silty limestoneLong Canyon Section L12750NLooking N40EAlteration Opdl Replacement Dolomite of Pogonip LimestoneCnpdl Hydrothermal Replacement Dolomite of Notch Peak LimestoneTl Thallium outline greater than 0.1 ppmAu Gold mineralization outline greater than 0.1 ppmBreccia Brecciation including Fault Breccia, Solution Breccia and Post-ore Calcite Cement BrecciaStructure 12800N12750N500 metresSection Location3 g/t Au Cuto10 -   206 -   100.6 -    10.4 -    0.61 -    20.2 -    0.30.1 -    0.20.3 -    0.42 -    6Drillhole Thallium in ppm100 metres 1000 mX800 mX1750 mRLTlTlFigure 3 .2792123456A B C D E FLC480LC733CLC537CLC541CLC462CLC452CLC533CLC737CLC446LC548CLC445CLC524CLC673CLC564CLC556CLC666CLC555CLC442LC396LC516LC448LC537CLC462CLC533CLC737CLC548CLC524CLC4428 -   204 -   80.8 -    10.6 -    0.80.5 -    0.60.4 -    0.50.3 -    0.41 -    22 -    4Drillhole Mercury in ppm Alteration OpdlAuReplacement Dolomite of Pogonip LimestoneCnpdl Hydrothermal Replacement Dolomite of Notch Peak LimestoneGold mineralization outline greater than 0.1 ppmHg Mercury outline greater than 0.25 ppmQalOplsCnpdOplwCambrian Notch Peak Group LithologiesOrdovician Pogonip Group LithologiesOplsmCnplbuCnploncCnplwCnplusAlluvium/ColluviumQuaternary LithologyLaminated to thin-bedded silty limestone and limy siltstoneMassive dolomiteMassive limestone with silt wispsSilty limestone alternating with thick-bedded massive limestone and at pebble conglomeratesMassive, burrowed, bioturbated, mottled limestone, often dolomitizedMassive oncolitic limestone, often dolomitizedMassive limestone with silt wispsUpper siltstone, 0.5-1 m beds of silty limestoneBreccia Brecciation including Fault Breccia, Solution Breccia and Post-ore Calcite Cement BrecciaStructure Long Canyon Section L12800NLooking N40E12800N12750N500 metresSection Location 3 g/t Au Cuto100 metres 1000 mX800 mX1750 mRL123456A B C D E FLC448LC557CLC582CLC571CLC692CLC466CLC471CLC577CLC680CLC518LC539CLC412LC525LC424LC613CLC420LC405L5LLC405QalOplsCnpdOplwCambrian Notch Peak Group LithologiesOrdovician Pogonip Group LithologiesOplsmCnplbuCnploncCnplwCnplusAlluvium/ColluviumQuaternary LithologyLaminated to thin-bedded silty limestone and limy siltstoneMassive dolomiteMassive limestone with silt wispsSilty limestone alternating with thick-bedded massive limestone and at pebble conglomeratesMassive, burrowed, bioturbated, mottled limestone, often dolomitizedMassive oncolitic limestone, often dolomitizedMassive limestone with silt wispsUpper siltstone, 0.5-1 m beds of silty limestoneLong Canyon Section L12750NLooking N40EAlteration Opdl Replacement Dolomite of Pogonip LimestoneCnpdl Hydrothermal Replacement Dolomite of Notch Peak LimestoneHg Mercury outline greater than 0.25 ppmAu Gold mineralization outline greater than 0.1 ppmBreccia Brecciation including Fault Breccia, Solution Breccia and Post-ore Calcite Cement BrecciaStructure 12800N12750N500 metresSection Location3 g/t Au Cuto8 -   204 -   80.8 -    10.6 -    0.80.5 -    0.60.4 -    0.50.3 -    0.41 -    22 -    4Drillhole Mercury in ppm100 metres 1000 mX800 mX1750 mRLHgHgFigure 3 .2893OplsmOplsOpdlOpdlOpdlOpls 31’ 122’ 236’ 314’ 406’Figure 3 .29: Brecciated, mineralized and oxidized Ordovician Pogonip Limestone lithologies from 28’ - 499’ from hole LC577C on cross section 12750N. Demonstrated is the extensive brecciation and altera-tion above solution breccias, scorodite overprint of hematite and limonitic upper. Arranged as downhole columns from left to right. Lithologies going down stratigraphy include Opls - silty laminated limestone, Oplsm - silty massive limestone, Oplw - wispy limestone and overprinting Opdl - dolomitized grainstone and minor silt interbeds. 94OplsmOplsOpdlOpdlOpdl106’ 200’ 294’ 387’ 475’Figure 3 .30: Brecciated, mineralized and oxidized Ordovician Pogonip Limestone lithologies from 104’ - 569’ from hole LC556C on cross section 12800N. Demonstrated is the strong solution brecciation adjacent to dolomitized limestone which are not decarbonated. Arranged as downhole columns from left to right. Lithologies going down stratigraphy include Opls - silty laminated limestone, Oplsm - silty massive lime-stone, Oplw - wispy limestone and overprinting Opdl - dolomitized grainstone and minor silt interbeds.95OplsmOplsOplwOplsmOpdlOpdlOpdl120’ 209’ 307’ 398’ 488’Figure 3 .31: Relatively unaltered Ordovician Pogonip Limestone lithologies from 119’ - 584’ from hole LC452C on cross section 12800N. Demonstrated is the lack of structure within limestones and a corre-sponding lack of decalcification and alteration. Arranged as downhole columns from left to right. Litholo-gies going down stratigraphy include Opls - silty laminated limestone, Oplsm - silty massive limestone, Oplw - wispy limestone and overprinting Opdl - dolomitized grainstone and minor silt interbeds.96hydrothermal dolomitizing brineDolomitizationShor tening, B oudinage and FoldingEarly BoudinageLateCrenulationCleavageLower O rdovician Po gonip Silt y LimestonePogonip Diagenetic Dolomite lensesUpp er Cambrian Notch Peak DolomiteNotch Peak Hydrothermal DolomiteUpp er Cambrian Notch Peak Silt y LimestoneEx tension and ExhumationLamprophyreDikes and SillsABCD? ? ?? AuriferousHydrothermal Fluid FlowA: Circulating basinal brines associated with Devonian - Mississipian age MVT and Sedex style mineralization across the Great Basin have been attributed to hydrothermal dolomitization (Diehl 2010). Limestone beneath Notch Peak Dolomite is dolomitized producing bedding parallel hydrothermal zebra dolomite with anomalous base metal geochemistry (W, Zn). This will be overprinted by later deformation fabrics (B + C).C: Extension occurs along moderate - low angle normal faults likely owing to decompressional detachment faulting to the West (Camilleri et al 1997). Lamprophyre dykes and sills which overprint earlier deformation fabrics likely intrude during this period of extension and are observed at the deposit scale to concentrate in brittle damage zones and boudin necks.D: Mineralization is not specically dated, yet is considered to be Eocene age owing to a favourable extensional environment, active magmatic centers and similar mineralization type to Eocene Carlin-type deposits to the West. Mineralizing uids travel along core uid channelways in fault and fracture zones. Increase in volume of silicication to the southwest indicates cooling ore uids and a likely uid source from the northeast.B: Thrust faulting and tectonic burial occurs during crustal shortening (Camilleri et al 1997). Boudinage of brittle dolomite units occurs with NE trending necklines. Enclosing limestones are foliated and ow into boudin necklines. Open folding and a crenulation cleavage overprint foliated limestones with coincident steep NW dipping axial planes.Cooling OreFluidsAuriferousFluidsFigure 3 .32: Block model schematic evolution of the Long Canyon deposit.97Chapter 4: Mapping Hydrothermal Fluid Flow With High Density Sampling of Oxygen and Carbon Isotope Ratios in Carbonate Rocks of the Long Canyon Deposit4 .1 IntroductionStable isotopes have been used in the study of hydrothermal mineral deposits for over 50 years (Engel, Clayton et al. 1958; Taylor 1974). Such studies have most commonly used the light stable isotope systems (e.g. 18O/16O, H/D, 13C/12C and 34S/32S) to identify fluid source and alteration processes – ultimately characterizing mineral deposits after they have been discovered. Using stable isotopes as tools aiding discovery of new ore bodies on the other hand, have until recently, not been considered useful in mining companies’ exploration programs owing to: 1) long turn-around times on analyses relative to the short decision making timeline required by the exploration geologist; 2) high analytical costs; and 3) poorly defined exploration criteria for how to effectively interpret the isotopic ratios in relation to the fossil hydrothermal system and related mineralization. Important questions include how does the intensity of isotopic alteration relate to distance to ore-body? Is there a correlation between size of isotopic alteration halo to size of target orebody? In hydrothermally-altered carbonate rocks hosting Carlin-type deposits, fluid-mineral reactions typically result in a lowered 18O/16O ratio within the altered carbonate minerals (Radtke, Rye et al. 1980; Nesbitt 1996; Stenger, Kesler et al. 1998; Arehart and Donelick 2006). However, controls on spatial variation of isotopic ratios at the individual ore-shoot-scale within the hydrothermal system have been poorly defined owing to the sparse sample suites that have been collected over deposits. Resulting fluid flow models based on stable isotope data are often idealized and the product of extrapolation between widely spaced data points, which ignore heterogeneities of structural and stratigraphic controls on fluid flow.In this study we test the extents of hydrothermal fluid infiltration into the multilayered carbonate sequence hosting the Long Canyon deposit and define the structural and stratigraphic controls on fluid flow. This is achieved principally by mapping out oxygen isotope depletion patterns as an indicator of ore fluid-rock interaction through high density sample collection across inferred fluid flow paths. The project benefitted from a understanding of the deposit scale geology and closely-spaced drilling, while being hindered in direct observation of the effects of Au-mineralizing fluid flow by the extensive modification by pervasive post-ore oxidation processes. MDRU’s desktop mineral isotope analyzer (MIA) was used for rapid, inexpensive oxygen and carbon isotope analyses and allowed the largest stable isotope study of 98a hydrothermal ore deposit its kind to date. Results from the study support the utility of oxygen isotope analysis in an active exploration program paired with traditional pathfinder geochemistry to improve understanding of fluid flow paths of the hydrothermal system and assist in targeting ore bodies.4 .2 Previous StudiesRadke et al. (1980) conducted the first carbon and oxygen isotope study of Carlin-type deposits at the partially oxidized Carlin deposit. Their work focused on isotopic variation between lithologies and alteration types taken from surface and drill core samples up to 2 km outwards from the ore body. The sample suite consisted of approximately 24 whole-rock carbonate samples for 18O/16O and 13C/12C determination as well as silicate minerals for 18O/16O and H/D. They concluded: 1) the degree of recrystallization is related to the degree of isotopic depletion from higher background δ18O values in unrecrystallized rocks; 2) where calcite is strongly depleted in δ18O values and recrystallized, dolomite in the vicinity is only very weakly recrystallized and shows little 18O depletion; and 3) isotopic exchange (18O depletion) in carbonate rocks occurred during main-stage mineralization prior to supergene oxidation, and that post-hydrothermal supergene oxidation and weathering had no detectable effect on δ18O values of altered or unaltered calcite or dolomite of the host rocks.Stenger et al. (1998) studied carbon and oxygen isotope variation in 108 whole-rock samples of calcareous shale and limestone host rocks across several drill cross-sections at the Twin Creeks Carlin-type gold deposits. They determined a decrease in δ18O towards ore zones within the shales, but within limestones they did not observe a consistent spatial variation on δ18O values, despite an absolute variation from 0 to +24‰ δ18O (VSMOW). They attributed this inconsistent spatial variation in the limestones to post-ore modification. Arehart et al. 2006 studied carbon and oxygen isotope variation in 132 whole-rock carbonate samples from surface and drill core from two cross-sections across the Pipeline Carlin-type gold deposit. They recognized broad depletion in δ18O values from +21.9 to -1.7‰ (VSMOW) towards the ore body across several kilometres. They concluded the need for more detailed studies to better understand fluid flow paths.Vaughan et al. (2010) studied carbon and oxygen isotopes in carbonate host rocks around the Banshee, Meikle, Screamer and Betze-Post Carlin-type deposits. At the time it was the largest carbon and oxygen 99isotope study of carbonate-hosted ore deposits. The initial Banshee study consisted of micro-drilled samples from 12 drillholes across several cross-sections at 5-20 metre intervals for a total of 349 samples. They observed depletion in δ18O values towards the ore zones ranging from +25 to +4‰ (VSMOW), with isotopic depletion from background values extending ~30 metres outside of visual alteration features. As with past studies, they identified a broad oxygen depletion halo and a correlation between lighter δ18O values and Au and trace elements, however, the density of sampling did not elucidate small-scale controls on fluid flow paths. Their second study across the Goldstrike deposit aimed to remedy this sample density issue by analyzing 1,658 drill assay pulps at 5-20 ft intervals from 29 drillholes along 2 approximately 4 km long cross-sections and comparing this dataset to aqua regia geochemistry. While sample density was closely spaced down drillholes, drillhole spacing was often 200-400 metres. This was sufficient to define broad patterns of fluid flow related Au mineralization, with the majority of δ18O values ranging from +10 to +27‰, but defining small scale variations (e.g. channelized fluid flow paths) was not possible at the scale of sampling. Of particular note, the initial Banshee study samples were analyzed by isotope ratio mass spectrometry (IRMS), while the second larger study was analyzed by laser spectroscopy as the first large isotope study carried out on the prototype MDRU-MIA.Carbonate-hosted Pb-Zn deposits have also seen several stable isotope studies which impacted exploration models for orebodies in their respective mining districts. Naito et al. (1995) in brownfields prospecting for carbonate- and silicate-hosted Pb-Zn skarn deposits in Japan, took 33 carbonate samples across 9 km2. The resulting gridded δ18O profile showed oxygen depletion around a fault hosting known mineralization, and oxygen depletion surrounding a second, apparently unmineralized fault. Drill testing of this fault intersected a significant Pb-Zn orebody. Vazquez et al. (1998) in a study at El Mochito in Honduras, a limestone-replacement lead-zinc-silver skarn deposit, investigated isotopic depletion in host carbonate rocks as far as 7.5 km away from the deposit from 82 samples. They observed a progressive decrease of up to 18‰ in δ18O value within carbonate rocks as far as 4 km away from distal background values to within the ore deposit and immediate wall rocks. Comparisons between micro-drilled analyses and bulk samples yielded similar results between matrix and bulk samples, but significant variance between sampled fossils, bioclasts and bulk samples.1004 .3 Project DesignStable isotope alteration intensity and distribution in host carbonate rocks across the Long Canyon Deposit was determined by sampling at four different scales in this study (Figure 4.1):1. 1.5 m drill assay pulps, sampled either contiguously or every second assay pulp down diamond drill and reverse circulation (RC) drillholes.2. 15 cm long x 7.6 cm diameter drill core hand samples were sampled for geochemistry, with returned whole-rock pulps analyzed for carbon and oxygen isotope ratios; sampled approximately every 3 metres (or approximately 1 for every two drill assay pulps).3. Micro-drilled carbonate material from drill core hand samples.4. Surface hand samples, micro-drilled for carbonate material for analysis.Drillholes selected for sampling were typically drilled either straight through or on the margins of breccia bodies. Where possible, entire drill fans intersecting breccia bodies were selected for sampling. This allowed for tightly constrained studies of fluid flow along the sides, as well into damage zones of breccia body tops and bottoms.Samples were collected in two phases over two field visits. Phase I samples were selected to 1) establish isotopic ranges of unaltered and altered lithologies from hydrothermal alteration, post-mineralization oxidation, proximity to lamprophyre intusive, breccia varieties and calcite vein and breccia cement species; and 2) determine the effectiveness of modeling fluid flow across the deposit from analyzing 5 to 25 metre spaced hand samples down drillholes at ~150 metre intervals along cross-sections, i.e., was there a broad halo of alteration from fluid-flow in host-rocks that could be discerned at this scale, or was it structurally controlled, requiring denser sample collection. The majority of the samples in this study were collected in Phase II. Sampling in this phase was focused across 2 closely-spaced drill cross-sections. Dense drilling patterns on these sections delineate multiple strongly mineralized and altered zones, allowing study of the extent of, and controls on, fluid infiltration into wall rock on the margins of hydrothermal flow paths. Contiguous sampling of 1.5 m drill assay pulps down drillholes allows for identification of heterogeneities in permeability and attendant isotopic observation not observable in previous studies which composited wide intervals or extrapolated between widely spaced samples. High density hand sampling down select drillholes was undertaken to avoid contamination of wall-rock isotopic ratios by calcite veins and provide geological control on 101homogenized drill assay pulps at corresponding intervals. A further check on contributions to the isotopic ratios observed in homogenized whole-rock hand sample pulps and drill assay pulps comes from micro-drilling of hand samples within specific carbonate beds, veins, breccias clasts and matrix, with a principal focus on discerning variability in oxygen isotope depletion between laminated siliciclastic carbonaceous units and massive micritic limestones interbeds. Finally, isotope alteration patterns were compared to traditional indicators of hydrothermal alteration around mineral deposits (i.e. trace element geochemical halos, mineralogical zonation and oxidation profile). A parallel study was carried out testing isotopic alteration in carbonate rocks outcropping above ore zones or prospective structural environments.4 .3 .1 Analytical MethodsTwo analytical methods were used for measuring carbon and oxygen isotope ratios in this study 1) traditional isotope ratio mass spectrometry and 2) laser spectroscopy.4 .3 .2 Light Stable Isotope Ratio Mass SpectrometryCarbonate samples were analyzed during Phase I by light stable isotope ratio mass spectrometry (LS-IRMS) at the Pacific Centre for Isotopic and Geochemical Research (PCIGR) at the University of British Columbia. Analytical methods are described in Appendix 1.4 .3 .3 Off-Axis Integrated Cavity Output SpectroscopyLaser spectroscopy of CO2 gas liberated from carbonate rocks was a prototype analytical method developed by the Mineral Deposit Research Unit and Los Gatos Research that became available for analyses during this study (early 2010). The MDRU Mineral Isotope Analyzer (MIA) is based on off-axis integrated cavity output spectroscopy (OA-ICOS), a cavity enhanced laser absorption technique (Barker, Dipple et al. 2011). Analytical methods and data processing are described in Appendix 1.  4 .4 Sample Suites and Sampling Methodology4 .4 .1 Phase I Micro-DrillingDrill core hand samples were collected 5 to 25 metres apart down 13 drillholes, averaging 10 to 20 samples per drillhole, across 7 holes on a SW-NE oriented long section (Western Long Section), 3 holes on a NW-SE oriented cross-section (L11900N), as well as several more drillholes to the northeast around section 12000N, with section locations shown in Figure 4.2 and Figure 3.9. Carbonate material was collected by micro-drilling 1-3 specific sites from each hand sample with a Dremel® drill. For 168 hand 102samples collected, 198 carbonate analyses were drilled and analyzed by Isotope Ratio Mass Spectrometry (IRMS) at PCIGR, UBC for carbon and oxygen isotopes. Hand sample descriptions are presented in Table B.1 located in Appendix 2. 4 .4 .2 Phase II Drill Assay Rock PulpsOne hundred and sixty seven drill assay pulp samples from drillholes LC372 and LC402 on section 13100N were selected as a pilot study for analyzing whole drill assay pulps on the MIA. Following positive results, drillholes on cross-sections L12750N (10 drillholes) and L12800N (11 drillholes) were then selected for the main study area for Phase II (Figure 4.2, Figure 4.9 and Figure 4.15). From these drillholes, drill assay rock pulps originally assayed for Au and trace element geochemistry by Fronteer were now analyzed for carbon and oxygen isotopes. Drillcore rock pulps returned from the assay lab were split at the core warehouse in Wells, Nevada and a portion brought back to UBC for carbon and oxygen isotope analysis on MDRU’s MIA. On section L12800N, pulps were sampled contiguously for a total of 1,091 analyses. On section L12750N, every second pulp was sampled for a total of 529 analyses. Original sampling of drillholes by Fronteer geologists was not always contiguous down drillholes as they often avoided sampling visually unaltered zones. Consequently there are sampling gaps, particularly in the most unaltered zones, which is unfortunate as the most distal extent of alteration was a particular area of interest in this study.4 .4 .3 Phase II Drill Core Hand SamplesThree drillholes were selected on section L12800N for detailed sampling of 15 cm core intervals approximately every 3 metres (Figure 4.15). As drill core was already split in half, quarter splits were taken and split again into eighths, with one of the eighths kept as a reference and the other submitted for geochemistry. Samples were submitted to ALS Chemex for fire assay Au and aqua regia ICPMS analysis for 41 elements. Rock pulps were returned to UBC for carbon and oxygen analysis on MDRU’s MIA, which yielded 259 carbon and oxygen isotope analyses. Drillhole LC556C intersected strongly Au-mineralized breccias above and below the Cnpd (Figure 3.21), with significant alteration and superjacent brecciation outwards from the main mineralized breccias. 103 samples were taken from this 310 m drillhole. Drillhole LC555C intersected the margin of the main breccia zone intersected in LC556C within Lower Pogonip rocks, and ended in strongly faulted Cnpd. 69 samples were taken from this 219 m drillhole. Drillhole LC533C intersected the least visibly altered rocks 103out of the three drillholes sampled on the cross-section. It contains a narrow breccia zone above the Cnpd within the Lower Pogonip, but is without the extensive overlying fracturing and faulting observed above breccia zones in LC556C and LC555C. 101 samples were taken from this 316 m drillhole.4 .4 .4 Phase II Micro-DrillingHand samples from drillholes LC556C, LC555C and LC533C were micro-drilled for carbonate for carbon and oxygen isotope analysis on MDRU’s MIA. Sampling was designed primarily to determine variance in δ18O values between silty laminae and micritic limestone beds as an indicator of relative fluid:rock ratios. A second aim of the micro-drilling was to identify specific contributions of individual rock types to the bulk δ18O values determined from rock pulp analysis. 70 analyses were drilled in hand samples from LC556C, 9 analyses from LC555C and 31 analyses drilled in hand samples from LC533C. Descriptions and samples are recorded in Table B.3 in Appendix 2.4 .4 .5 Surface Hand SamplesSurface sampling for carbon and oxygen isotopes in carbonate rocks was completed over two 800 – 1000 m traverses (Figure 4.2). Along both traverses, historical rock geochemical sampling existed from ridge and spur sampling programs and prospecting. This allowed comparison between oxygen isotope and trace element geochemical profiles. Rock hand samples were micro-drilled for carbonate material, later analyzed on the MDRU-MIA, and the rock type recorded as silty calcareous laminae, micritic limestone, dolomitized limestone, limestone breccia or dolomite breccia.Section LCS was carried out in a down-faulted block south of the main Long Canyon deposit through the basal units of Lower Pogonip carbonate rocks above the Notch Peak Dolomite (Figure 4.2). This is a similar structural environment to the Long Canyon deposit host rocks with boudinage of the Cnpd and breccia developed in Pogonip limestones, however, no Au mineralization occurs in the breccias. 50 samples were collected at 15-20 metre intervals over ~1000 m.Section LC12800 was carried out through Lower Pogonip carbonate rocks approximately aligned with, and to the west of, the fan of drillholes sampled on drill cross-section L12800N (Figure 4.2). The sampling traverse cut across a strongly Au-mineralized dolomite-hosted breccia zone. 59 samples were collected at 10-15 metre intervals over ~800 m. 1044 .5 Results4 .5 .1 δ18O and δ13C Values of Carbonate Rocks at the Long Canyon DepositTwo thousand four hundred and forty five carbon and oxygen isotope analyses were obtained during this study, excluding quality control (QC) samples. Oxygen isotope values ranged from +3.7 to +25.2‰ δ18O relative to Vienna Standard Mean Ocean Water (VSMOW). Carbon isotope values ranged from -7.0 to +4.4 ‰ δ13C relative to Vienna Peedee Belemnite (VPDB)Background δ18O and δ13C ranges of Upper Cambrian to Lower Ordovician limestone globally are +20.5 to +25.5‰ (VSMOW) and 0.0 to +2.0‰ (VPDB) respectively, with mean dolomite δ18O typically several per mil higher. These values are tabulated from globally distributed limestone and dolomite sources in the Precambrian marine carbonate isotope database of Shields and Veizer (2002). Background δ18O ranges for carbonate rocks at Long Canyon were obtained by identifying unaltered samples through petrographic examination and geochemistry. Two confidence intervals of background were defined as distinctly background (+19 to +23‰) and uncertain background (+17.5 to +19‰). The uncertain background interval was defined to account for samples with slightly lower than +19‰ δ18O values yet showed no evidence of hydrothermal fluid interaction. A cumulative frequency distribution plot of δ18O values from micro-drilled silty laminae was used to help distinguish background from depleted populations in the samples (Figure 4.3). For each data point the cumulative frequency distribution plot displays a score on the x-axis of the probability of another random data point in the population being of lesser value, with the highest value having a score of 1. Shallow slopes separated by steepened inflections represent coherent populations of data of similar valueBackground carbon isotope ratios of carbonate rocks at Long Canyon vary by lithology, likely related to secular variations of 13C/12C in seawater. These secular variations likely reflect changes in the biogeochemical redox cycling and organic carbon burial and storage rates (Shields and Veizer 2002). Figure 4.4 plots δ13C values against δ18O in hand sample pulps for units Cnplw, Cnplus, Oplw, Oplsm and Opls. Dolomite units were excluded owing to replacement of original limestone by a dolomitizing brine. Breccia units were also excluded as there was no control on the contribution of isotopically heavy or light δ13C values from paragenetically late calcite. The median δ13C value for each of these units (the cross coloured by lithology, highlighted by shaded bars) decreases with stratigraphic younging. In Figure 4.4b 105plots of isotopically depleted samples (δ18O < +17.5‰) do not shift median δ13C values by a statistically significant amount from the plot in (a), which only plots data from samples with δ18O > +17.5‰.The reproducibility of carbon and oxygen analyses from rock pulps was assessed by analyzing samples twice from drillhole LC424 on section L12750N, with analyses separated by 6 months. Data is presented in Figure 4.5 and calculations in Table 4.1. Overall, the reproducibility of δ18O and δ13C data is quite good, apart from divergence in the last six samples of carbon analyses, which is likely the result of a change in operating conditions of the MIA, perhaps water vapor not sequestered in the cold trap. Carbon isotope analyses have a precision (1 SE) of 0.29‰ (0.19‰ if the last six samples are excluded). Oxygen isotope analyses have a precision (1 SE) of 0.47‰, similar to that estimated from replicate reference materials by (Barker, Dipple et al. 2011).4 .5 .2 Drill Assay Rock Pulp δ18O Values and Distribution at the Long Canyon DepositOne thousand seven hundred and sixty seven samples from drill assay rock pulps were analyzed for δ18O and δ13C with ranges of +3.7 to +25.2‰ (VSMOW) and -4.2 to +4.4‰ (VPDB) respectively. Results are presented as δ18O versus δ13C cross-plots grouped by lithology in Figure 4.6a-e and in Table_DrillAssayPulp in Digital Appendix 3. Dolomite units and Notch Peak limestones display much less oxygen depletion than Pogonip limestones while Notch Peak breccias display only slightly less oxygen depletion than Pogonip breccias.Statistics of the sample population are presented in Table 4.2 and as box and whisker plots in Figure 4.8. Ordovician Pogonip limestones, Opbx, Oplsm, Opls and Oplw host more oxygen isotope depletion than any other lithology, and dolomite units, Opdl, Cnpd and Cnpdl are rarely significantly depleted in oxygen from their background values. Carbon isotope means (e.g., Figure 4.8c) display a correlation of younging lithologies with decreasing mean δ13C values, as observed with select lithologies in Figure 4.4. There is some variation in mean δ13C ordering by lithology as subunits of Pogonip Group and Notch Peak Formation limestones show internal variation of decreasing δ13C value with stratigraphic younging.Drill assay rock pulp oxygen isotope analyses are plotted down drillholes across sections L12800N and L12750N in Figure 4.8 and Figure 4.9 respectively. Also displayed are lithological units, breccias, dolomite alteration and Au outline >0.1 ppm. Sample gaps on drillhole LC556C on section 12800N, quadrant D4, and drillhole LC577C on section 12750N, quadrant E4, represent areas of extensive silica replacement of carbonate and a lack of matrix to sample. Other gaps represent areas where no assay pulps 106were available for sampling. Observations on the spatial distribution of oxygen isotope depletion relative to host lithology, structure and Au mineralization include:1. Au mineralization is coincident with the areas of greatest 18O depletion.2. The extent of 18O depletion beyond the edge of Au mineralization is variable. A significant Au zone in quadrant C4 on section 12800N has minimal associated 18O depletion surrounding it, while the adjacent Au zone in D4 and E4 displays widespread depletion.3. Areas of greatest 18O depletion are commonly associated with zones of brecciation, although, because breccia zones cannot be mapped from RC drillholes as they can in core holes (denoted by the suffic ‘C’), this association is not always evident. 4. Dolomite units, Cnpdl, Cnpd and Opdl are consistently less depleted from background δ18O values than adjacent depleted limestones.5. Where brecciated, dolomite can locally exhibit depletion from background δ18O values.6. The silty Cnplus unit hosts the greatest 18O depletion in the Notch Peak limestone.Oxygen isotope depletion down drillholes is compared to previously contoured trace element geochemical halos in Figure 4.10, Figure 4.11, Figure 4.12 and Figure 4.13. Spatially-determined geochemical cutoffs used are As>20ppm, Hg>0.25ppm, Tl>0.1ppm and Sb>0.15 ppm. Observations from these plots include:1. δ18O depletion occurs in limestone beyond the furthest extents of anomalous pathfinder element distribution.2. The shape of pathfinder element distribution patterns at these defined concentration cutoffs largely mirror δ18O depletion patterns though they are slightly contracted in absolute extents. 3. Pathfinder elements display anomalous values in dolomite despite a lack of δ18O depletion.The relationship between As, Hg, Tl and Sb pathfinder element concentrations to their corresponding δ18O values in drill assay rock pulps is presented in cross-plots in Figure 4.14. Data points are coloured by both Au grade (a to d) and lithology (e to h). An apparent trend in this data for each element is a sub-horizontal trend through high δ18O and low element concentrations, before a distinct inflection as the data trends downwards to low δ18O values and high element concentrations. This is best developed in Hg plots (Figure 4.14 b and f). This relationship was assessed by assigning 18.0‰ as the lower range for average background values and drawing this line across to intersect the trendline for all the data. For all four element plots, the geochemical value of this intersection point is approximately where geochemical 107background was defined through spatial assessment (As>20ppm, Hg>0.25ppm, Tl>0.1ppm and Sb>0.15 ppm, see Chapter 3 and Figure 4.10, Figure 4.11, Figure 4.12 and Figure 4.13).When a line is drawn down to this point on the x-axis, the lower left quadrants contain data points representing samples which are depleted in 18O yet are not anomalously enriched in the pathfinder element. Spatially, these samples plot down the drillholes on the margins of the most 18O-depleted and Au-rich zones. Many higher-grade (>1ppm Au) samples occur without strong 18O depletion, located in the upper right quadrants. Corresponding plots of lithological units show that many of these high grade Au occurrences are within dolomitized limestone units, Cnpdl and Opdl and breccia units, Opbx and Cnpbx. Both brecciated limestones and dolomites of Pogonip Group and Notch Peak Formation units are logged as Opbx and Cnpbx without distinction. 4 .5 .3 Hand Sample Rock Pulp δ18O and Geochemical Values and DistributionsTwo hundred and fifty nine samples from hand sample rock pulps were analyzed for δ18O and δ13C with ranges of +4.4 to +22.4‰ (VSMOW) and -5.8 to +3.6‰ (VPDB) respectively. Results are presented as δ18O versus δ13C cross-plots grouped by lithology in Figure 4.6f-j and in Table B.1 in Appendix 2. Sample spread is similar to drill assay pulp data plotted in Figure 4.6a-e. The most notable difference is in hand sample carbon isotope data, with lower absolute δ13C values in Pogonip limestones (Opl prefix) and higher median δ13C values in Notch Peak limestones (Cnpl prefix).Hand sample pulp δ18O values are plotted down drillholes LC556C, LC555C and LC533C against δ18O values for drill assay pulps in Figure 4.15. Hand sample δ18O values generally present a similar degree of depletion to the nearest corresponding drill assay pulp interval. Significant zones of δ18O depletion or background values in drill assay pulps are generally mirrored in hand samples. However, there are a few zones where significant depletion in drill assay pulps displays at or near background values in hand sample (e.g. quadrants A4 and D4). Conversely, there are several zones that display significant depletion in hand samples, but background values in drill assay pulps (e.g. A2 and A3).Oxygen isotope hand sample pulp values are plotted against δ18O values of the drill assay pulp interval from which the hand sample was taken in Figure 4.16a,b,c. Of 259 hand samples analyzed, 185 were sampled within a 1.5 m interval previously sampled for geochemistry by Fronteer geologists, and are included in these plots. Also included in the plots are the correlation coefficients (r) and coefficients of 108determination (r2) with LC555C displaying the strongest r2 value followed by LC556C and LC533C. Reproducibility of sampling at the two different scales is assessed by measuring precision between the two values by taking the absolute value of the difference between a sample and the mean.Calculated results are presented in Table B.2 in Appendix 2, with a summary table of average precisions for each drill hole and the entire sample population in Table 4.3. Average precision of all samples is 1.26‰ for δ18O and 0.6‰ for δ13C.Pathfinder element As, Hg, Sb and Tl as well as Au concentrations of the 185 hand samples with corresponding drill assay pulp geochemistry are compared as log-log cross-plots in Figure 4.17. Correlation coefficients (r) and coefficients of determination (r2) show a minor spread, with increasing goodness of fit as Tl<Hg<Sb<Au<As. Reproducibility of sampling at the two different scales is again assessed as with carbon and oxygen isotopes above by measuring precision between the two values by taking the absolute value of the difference between a sample and the mean. Calculated results are presented with isotope sampling precision in Table B.2 in Appendix 2, with a summary table of average precisions by element in ppm for each drill hole and the entire sample population in Table 4.3. Average precision in ppm for Au is 0.06, As is 25.95, Hg is 0.50, Sb is 0.80 and Tl is 0.36.Trace Element Distribution Relative To 18O DepletionCarbonate 18O depletion in drill cross-section occurs beyond anomalous pathfinder element halos. This distal transition zone where pathfinder element precipitation ceases and 18O depletion in carbonate continues is best observed in hand samples. Arsenic is used as a proxy for all four pathfinder elements examined in this study as their distribution halos are largely coincident. Samples displaying As values from 6-12 ppm have bulk and micro-drilled δ18O values from +13.0 to 14.5‰ in Figure 4.18b and Figure 4.19d and e. These are significantly 18O-depleted rocks with background As values. Similarly, significant depletion in δ18O values of can also be observed in the hand samples from the upper part of drillhole LC533C, shown in quadrant A2 of Figure 4.15. This is examined further in Figure 4.20 showing hand samples from 12.0-44.0 m down the drillhole. Silty limestone units display 18O depletion ranges from +4.41 to +14.29‰ with a corresponding As range of 7-9 ppm. Hand samples do not display any significant permeable fracture networks. This zone shows evidence of significant fluid flow outside of 109anomalous geochemistry. Fluids appear to be travelling along stratigraphy an unknown distance from a fluid conduit, and quite distal to the main defined fluid flow paths of the deposit in stratigraphy below. Variance in 18O Depletion Between Hand Sample Pulps and Drill Assay PulpsThe extreme range of δ18O values observed in Figure 4.20 between adjacent hand samples highlights a likely reason for the variability between hand samples and their corresponding drill assay pulp interval. Drill assay intervals homogenize both undepleted micritic limestone beds along with depleted silty limestones for a bulk slightly depleted to undepleted δ18O signature, while a single hand sample, selecting the most interesting alteration returns an extremely depleted δ18O signature. This sample bias works to the opposite effect in strongly depleted zones as indicated by drill assay pulps. In such intervals of brecciated and friable rock, the tendency is to select a competent sample from which petrographic samples can be cut, and a reference sample that will not be pulverized during transport. As a result, hand sample δ18O often underestimate the degree of depletion relative to drill assay pulp δ18O values. This trend is observed in plots between hand sample and corresponding drill assay pulps in Figure 4.16.4 .5 .4 Micro-Drilled δ18O and δ13C Values of Long Canyon Carbonate RocksTwo hundred and ninety one micro-drilled samples were analyzed for δ18O and δ13C with ranges of +4.8 to +24.0‰ (VSMOW) and -7.0 to +3.8‰ (VPDB) respectively. Samples were analyzed by both IRMS and OA-ICOS and results are reported here together and in Table B.3 in Appendix 2. Micro-drilled δ18O analyses are plotted against their hand sample pulp δ18O value in Figure 4.21 with samples grouped by drilled rock type. Variance from the y=x line of data points displays the difference between bulk δ18O value of the sample and point analysis. Observations from these plots of micro-drilled and hand sample δ18O values include:1. Plot (a) shows micro-drilled dolomites are generally less depleted than their hand sample δ18O values, although variance is generally within background range.2. Plot (b) displays considerable variability, predominantly with micro-drilled δ18O values higher than their hand sample. This likely reflects contribution to the bulk hand sample δ18O value from more depleted silty limestone beds that were not micro-drilled.3. Plot (c) displays two populations. In both populations, silty limestone micro-drilled samples range from slightly to strongly depleted, but hand sample bulk δ18O values are either more depleted or less depleted. The less depleted hand sample population represents samples which have only been 110depleted in δ18O along silty laminae, while micritic limestone interbeds are relatively undepleted. The more depleted hand sample population represents samples where both micritic limestone interbeds and silty limestone lamine have been depleted. 4. Plot (f) shows variability between hand sample and micro-drill δ18O values which likely depends on whether matrix or breccia clast was micro-drilled within limestone breccias.Results are presented in Figure 4.22 in δ18O and δ13C cross-plots of the particular rock type from which the sample was drilled, and in Table B.3 in Appendix 2. Rock type classifications are distinguished for their inherent permeability characteristics and carbonate mineral type (dolomite versus calcite). They are listed in decreasing apparent permeability as: limestone breccia (lstbxa), dolomite breccia (dolbxa), silty limestone (slt), massive/micritic limestone (lst), and dolomite (dol). Samples are coloured in plots by lithology in Figure 4.22. Presented here are several collections of samples from significant drillhole sections or geological units.LC287C Micro-Drilling: Micro-drilled samples from drill hole LC287C are displayed in Figure 4.23. LC287C was originally selected to try to establish background δ18O values owing to its portions of visually and geochemically unaltered appearance. Observations from these samples include:1. Photos (a) through (g) show samples through Pogonip Group limestones with predominantly background, undepleted δ18O values of +17.96 to +21.16‰ in both silty and micritic limestone units.2.  Photos (b, f and g) show samples with both background and more depleted δ18O values within thin silty limestone beds ranging from +13.28 to +17.32‰3. Photo (e) shows a strongly Fe-oxidized silty limestone with a background δ18O value of +17.96‰ indicating supergene oxidation does not result in depleted δ18O values.4. Photo (f) shows a bleached micritic limestone unit sampled directly above an unmineralized lamprophyre intrusive with a δ18O value of +20.33‰. An adjacent silty limestone bed has a δ18O 111value of +16.02‰, indicating intrusion of lamprophyres did not have a significant impact on host carbonate δ18O values.5. Photo (g) shows a rare contact between Oplw and Cnpd, where the fractured limestone is significantly depleted (+14.84‰) adjacent to undepleted dolomite. 6. Photos (h and i) show two Cnplus samples 5.5 m apart displaying significant depletion in the upper unit and background δ18O values in the lower silty unit. This suggests tightly constrained fluid flow paths at Long Canyon.7. Photo (j) shows an unaltered Cnplw sample with background δ18O values of +19.2‰.LC556C Pogonip Limestone Micro-Drilling: Bulk hand sample δ18O, As and Au values along with micro-drilled δ18O analyses for selected samples from drill hole LC556C are shown in Figure 4.18. These samples represent an intersection through Pogonip stratigraphy from near surface to directly above a well-developed and Au-mineralized solution collapse breccia. Bulk δ18O values generally reflect a composite value of individual drilled isotope values. This relationship is examined further in Figure 4.21. Observations from these samples include:1. Photos (e, f and g) display background δ18O values in clean limestone interbeds (+20.33 to +21.03‰) proximal to strongly depleted silty limestone laminae (+11.90 to +13.89‰).2. Photo (d) displays 18O depletion in both silty and micritic limestone interbeds, presumably reflecting a higher bulk fluid:rock ratio than is assumed from variable depletion patterns in photos (e,f and g).3. Photos (a, b, and g) display thin calcite veins filling fractures with δ18O values in adjacent silty and micritic limestone units becoming progressively more depleted in δ18O values closer to the fracture. The opposite progression occurs in photo (c) where micro-drilled rock δ18O values increase towards the vein.LC533C Pogonip Limestone Micro-Drilling: Bulk hand sample δ18O, As and Au values along with micro-drilled δ18O analyses for selected samples from drill hole LC533C are shown in Figure 4.18. These represent an intersection through Pogonip stratigraphy from near surface to directly above the Notch Peak dolomite, with minor brecciation along the dolomite - limestone contact. Observations from these samples include:1. Photos (b, e and h) display surprisingly low bulk δ18O values (+11.54 to +16.30‰) despite micro-drilling suggesting background values (+17.52 to +20.67‰). These hand samples were re-112analyzed as duplicates as analytical error was suspected, but duplicate data was similarly depleted in δ18O. The micro-drilled samples were taken from within micritic limestone lenses rather than within the silty laminae. It is likely that 18O depletion occurs predominantly within this silty carbonate material, and that they together drive the bulk δ18O value down.2. Photo (b) does contain anomalous As (29 ppm) which suggests that despite the high micro-drilled δ18O value, this rock hosted significant hydrothermal fluid flow, as the bulk δ18O value suggests.3. Photo (e) displays a micro-drilled analysis of +19.9‰ taken within the oxidized margin of a fracture. Two conclusions can be derived from this: (1) oxidation does not cause depletion in δ18O values; and (2) oxidizing fluids accessed stratigraphy along fracture networks.4. Photo (f) displays a micro-drilled δ18O value of +19.2‰ from a dolomitized Pogonip limestone (Opdl) and yellow carbonate veining with a δ18O value of +10.9‰. The bulk hand sample δ18O value of +18.9‰ is consistent with the slight 18O-depletion observed in dolomite samples with brecciation and carbonate veining.Dolomite-Hosted Fault Breccia Micro-Drilling: Bulk hand sample δ18O, As and Au values along with micro-drilled δ18O analyses for selected dolomite breccia samples from drillholes LC533C and LC556C are shown in Figure 4.24. Samples were selected to highlight variation in δ18O values within dolomite faults and breccia. Observations from these results include:1. (a) displays micro-drilled analyses within zebra dolomite showing no significant difference in δ18O values between dark and light dolomite layers.2. Au-mineralized fault breccias (c, d, and f) display bulk δ18O values from +14.6 to +16.8‰. These highly comminuted breccias host a calcareous cement in their matrix. Comparatively, (e) is a Au-mineralized breccia without a calcareous cement and a bulk δ18O value of +18.9‰. This comparison suggests that significantly depleted δ18O values in dolomite fault breccias may be the result of isotopically light, later calcite, rather than oxygen exchange between dolomite and hydrothermal aqueous fluid.Calcite Spar Micro-Drilling: Calcite spar occurring as veins and breccia cement exhibits a wide range of carbon and oxygen isotope values at Long Canyon, with δ18O values from +4.81 to +19.29‰ and δ13C from -5.84 to +2.54‰. Twenty veins and six breccia cements were analyzed by micro-drilling carbonate material from hand samples for analysis by both IRMS and OA-ICOS methods. Table B.3 in Appendix 2 contains descriptions and results, identifiable from other micro-drilled carbonate analyses under 113‘Rocktype’ heading as ‘vn’. Results are presented in Figure 4.25 as two plots; vein δ18O versus δ13C and hand sample bulk δ18O versus micro-drilled values of veins; as well as selected photos of vein and host hand samples. Calcite spar morphologies are distinguished by data point shape in plot (a) and include: 1) thin white calcite veins typically at a high angle to bedding; 2) veins at a low angle to bedding, often banded; 3) veins of coarse yellow and white carbonate; 4) quartz carbonate veins; and 5) coarse breccia cement. Observations from these results include:1. In plot (a), vein types do not have a discernable trend or grouping relative to δ18O values, though there is a weak correlation between lower δ18O and higher δ13C values.2. Thin, high angle calcite veins are dominantly in the δ18O range of +8.0 to +14.0‰3. Veins are predominantly more depleted in δ18O than their hand sample host carbonate4. (j) white carbonate material intergrown with coarse yellow carbonate vein material displays significantly different δ18O values of +19.13 and +11.07‰ respectively.5. (k and l) δ18O values become progressively more depleted in limestone wall rock approaching veins. Although (i) does not display any change in δ18O values with progressive distance from fracture and vein selvage.4 .5 .5 Surface Traverse δ18O DistributionTwo surface sampling traverses carried out at the Long Canyon deposit yielded 58 (Section LC12800) and 50 (Section LCS) carbon and oxygen isotope analyses with δ18O and δ13C ranges of +10.0 to +24.2‰ (VSMOW) and -6.1 to +0.9‰ (VPDB), respectively. Results are presented in Table B.4 in Appendix 2, and displayed in Figure 4.26 and Figure 4.27 along with geochemical analyses from previous sampling by Fronteer geologists. Similar to drill core hand samples, rock type from which carbonate material was micro-drilled is distinguished by their permeability characteristics, listed in decreasing apparent permeability, as limestone breccia (lstbxa), dolomite breccia (dolbxa), silty limestone (slt), massive/micritic limestone (lst), and dolomite (dol).Section LCS: A broad zone of δ18O depletion occurs from 705,600 to 706,000 mE through Pogonip limestone directly above the contact with Cnpd. Again, depletion mainly occurs in silty limestone units, from +17 to +11‰ δ18O, with several micritic limestones exhibiting δ18O depletion from +17 to +14.5‰. The isotopic samples were collected along the same line as previous ridgeline rock sampling that returned no significant Au values and only two samples with As greater than 20 ppm and Sb greater than 2 ppm. 114Oxygen isotope depletion in silty layering indicates fluid flow occurred within these rocks in a similar stratigraphic and structural environment as the Long Canyon deposit to the NE. Section LC12800: Moderate depletion in δ18O is observed throughout the studied section, occurring principally within silty laminae. There are two zones of stronger depletion. The easternmost zone of depletion does not have any associated geochemistry for comparison, but is sampled in unit Op2, which is just above the base of the Pogonip Group limestone, Op1, in which the majority of depletion and Au mineralization has been defined in this study. This zone of δ18O depletion may be the result of exhausting hydrothermal fluids. The westernmost zone is associated with a known Au-mineralized, dolomite-hosted breccia zone. As and Hg roughly match the extents of significant depletion from 707,850 to 708,050 mE. Sb and Tl appear to have detection limits set too high in the historic geochemical database, reducing their usefulness in discerning the extent of fluid flow from low-level geochemical anomalies. Aside from two samples, δ18O values from micritic limestones do not record evidence for hydrothermal fluid flow.Interpretation of Stable Isotope Results at the Long Canyon DepositResults from this stable isotope study of carbonate host rocks of the Long Canyon deposit reveal several important lithologic and sampling-scale controls on isotopic depletion patterns. These controls highlight the importance of understanding the geology of the study area under isotopic investigation, before designing a sampling program and interpreting hydrothermal fluid flow paths from isotope depletion patterns. 4 .5 .6 Lithologic Control on Fluid Flow and Oxygen Isotope Depletion PatternsLithology type influenced both the degree of oxygen isotope exchange that occurred between fluid and rock and also the extent of hydrothermal fluid infiltration into carbonate host rocks at Long Canyon. The varying degrees of isotopic exchange observed between adjacent calcite and dolomite within the same fluid flow paths indicate differing chemical reactivity of the two minerals with the infiltrating hydrothermal fluid. Within m