Open Collections

UBC Theses and Dissertations

UBC Theses Logo

UBC Theses and Dissertations

Establishing the stable isotope and geochemical footprints associated with carbonate-hosted Zn-Pb deposits… Cook, Natalie Louise 2016

Your browser doesn't seem to have a PDF viewer, please download the PDF to view this item.

Item Metadata

Download

Media
24-ubc_2016_may_cook_natalie.pdf [ 35.97MB ]
Metadata
JSON: 24-1.0300274.json
JSON-LD: 24-1.0300274-ld.json
RDF/XML (Pretty): 24-1.0300274-rdf.xml
RDF/JSON: 24-1.0300274-rdf.json
Turtle: 24-1.0300274-turtle.txt
N-Triples: 24-1.0300274-rdf-ntriples.txt
Original Record: 24-1.0300274-source.json
Full Text
24-1.0300274-fulltext.txt
Citation
24-1.0300274.ris

Full Text

Establishing the Stable Isotope and Geochemical         Footprints Associated with Carbonate-hosted Zn-Pb Deposits of the Kootenay ArcbyNatalie Louise CookMSci., The University of Bristol, 2012A THESIS SUBMITTED IN PARTIAL FULFILLMENT OF THE REQUIREMENTS FOR THE     DEGREE OFMASTER OF SCIENCEinThe Faculty of Graduate and Postdoctoral Studies(Geological Sciences)THE UNIVERSITY OF BRITISH COLUMBIA(Vancouver)April, 2016© Natalie Louise Cook, 2016iiAbstractCarbonate-hosted Zn-Pb deposits of the Kootenay Arc in southeastern British Columbia and north-eastern Washington typically have narrow and poorly developed alteration footprints, and so generate significantly smaller visible and geochemical haloes when compared to many other mineral deposit systems. However, these carbonate-hosted deposits can have invisible alteration footprints that are much broader when detected with analyses of light stable isotopes. The intensity of this isotopic al-teration increases from peripheral regions into the centre of mineralization with greater shifts towards lighter, or more depleted, isotopic values as a result of more fluid:rock interaction. These depleted stable isotope values can therefore provide information about fluid flow during mineralization and have potential value in the search for sulphide mineralization. Studies of C and O isotopes from the host rocks of four Zn-Pb mines within the Kootenay Arc (H.B., Jersey, Remac and Pend Oreille) in addi-tion to geochemical analyses have revealed at least two separate isotopic alteration signatures. These two different isotopic alteration signatures are inferred to represent two different hydrothermal fluids; the Remac Mine exhibits a much lower temperature isotopic signature when compared to the H.B., Jersey and Pend Oreille mines. Despite the difference in isotopic signature, all four mines are thought to represent Mississippi Valley-type mineralization. The most depleted isotopic values have been in-terpreted to reflect the highest temperature and most permeable systems. At the mine-scale, isotopic depletion highlights permeable rock-types and proximity to fault zones rather than proximity to Zn-Pb mineralization. At this scale it is therefore necessary to utilize additional geological information, such as geochemistry, to delineate rock-types that could be potential hosts of Zn-Pb mineralization. These two distinct isotopic signatures, in addition to geochemistry and other geological information, could therefore be utilized to vector towards, or re-evaluate, similar Mississippi Valley-type Zn-Pb mineral-ization within this region.iiiPrefaceThis thesis is part of the Mineral Deposit Research Unit’s (MDRU’s) Carbonate Alteration Footprints project, which is sponsored by a consortium of companies. This project was di-rectly sponsored by Teck Resources Limited (“Teck”). The author and Craig Hart of the Uni-versity of British Columbia (UBC), and Lucas Marshall, Jelena Puzic, Liz Stock and Mike Takaichi all of Teck identified the research scope, project area and research objectives of the thesis. All work presented is the result of the author’s own work carried out at UBC, which included field work in southeastern British Columbia and northeastern Washington. Craig Hart acted as the author’s main supervisor and provided field assistance, advice, supervision and editorial co-authorship. Greg Dipple, Ken Hickey and Andreas Beinlich acted as committee members, contributed to comments and sugges-tions during committee meetings and discussions and provided editorial feedback on initial drafts of this thesis and published papers. All sample descriptions and data interpretations are the responsibility of the author unless noted. Samples were selected by the author for thin section production, cold-cathode cathodoluminescence, carbonate staining, ultraviolet light analysis, X-ray diffraction and mineral isotope analysis. The author, Craig Hart, Lucas Marshall, Jelena Puzic, Liz Stock and Mike Takaichi selected samples for X-ray fluores-cence and analytical geochemical analysis.Chapters 3 and 4 comprise work that formed the following paper: Cook, N.L., and Hart, C.J.R. (2015): Carbonate alteration footprints of carbonate-hosted zinc-lead deposits in southeastern British Co-lumbia (NTS 082F/03): applying carbon and oxygen isotopes; in Geoscience BC Summary of Activities 2014, Geoscience BC, Report 2015-1, p. 95–102. I am the lead author of this paper and designed the bulk of the text and figures. Co-author Craig Hart and committee member Greg Dipple provided tech-nical advice and suggestions during manuscript preparation. ivTable of ContentsAbstract................................................................................................................................................ iiPreface ................................................................................................................................................ iiiTable of Contents.............................................................................................................................. ivList of Tables ...................................................................................................................................... ixList of Figures ..................................................................................................................................... xAcknowledgements ........................................................................................................................xivChapter 1: Project Setting and Objectives ................................................................................. 1Introduction ............................................................................................................................... 1Sediment-Hosted Zn-Pb Systems ............................................................................................... 3Overview......................................................................................................................... 4Continental Rifts ............................................................................................................. 6Passive Margins .............................................................................................................. 6Destruction of Passive Margins ...................................................................................... 7Extensional Arcs .............................................................................................................. 7Continental Sag Basins .................................................................................................... 7SEDEX Deposits ............................................................................................................... 8MVT Deposits ................................................................................................................. 9Non-Mixing Model ............................................................................................. 10Mixing Model ..................................................................................................... 11Irish-type Deposits ........................................................................................................ 11Regional Geology ..................................................................................................................... 12Structure and Metamorphic History ....................................................................................... 15Mineralization .......................................................................................................................... 15Chapter 2: Global Context ............................................................................................................ 18Introduction ............................................................................................................................. 18Method .................................................................................................................................... 18Reocín, Spain ............................................................................................................................ 19Introduction .................................................................................................................. 19Regional Geology .......................................................................................................... 19Deposit Geology ........................................................................................................... 21Paragenesis ................................................................................................................... 21Carbonate Textures ....................................................................................................... 22D1 Dolomite ....................................................................................................... 23D2 Dolomite ....................................................................................................... 23D3 Dolomite ....................................................................................................... 23D4 Dolomite ....................................................................................................... 23vC and O Isotopes ........................................................................................................... 23Unaltered Limestones ........................................................................................ 24D1 Dolomite ....................................................................................................... 24D2, D3 and D4 Dolomite .................................................................................... 24Summary ...................................................................................................................... 25Pine Point, Canada ................................................................................................................... 25Introduction .................................................................................................................. 25Regional Geology .......................................................................................................... 25Deposit Geology ........................................................................................................... 28Paragenesis ................................................................................................................... 28Carbonate Textures ....................................................................................................... 29Fine Dolomite ..................................................................................................... 29Medium Dolomite .............................................................................................. 30Coarse Dolomite................................................................................................. 30C and O Isotopes ........................................................................................................... 30Summary ...................................................................................................................... 31Lennard Shelf, Canning Basin, Australia .................................................................................. 31Introduction .................................................................................................................. 31Regional Geology .......................................................................................................... 31Deposit Geology ........................................................................................................... 33Paragenesis ................................................................................................................... 34Carbonate Textures ....................................................................................................... 34C and O Isotopes ........................................................................................................... 35Summary ...................................................................................................................... 36Nanisivik, Canada ..................................................................................................................... 36Introduction .................................................................................................................. 36Regional Geology .......................................................................................................... 38Deposit Geology ........................................................................................................... 38Paragenesis ................................................................................................................... 39Carbonate Textures ....................................................................................................... 40C and O Isotopes ........................................................................................................... 40Host Dolostone .................................................................................................. 40White Sparry Dolomite Associated with Sulphides ............................................ 40Late Carbonate ................................................................................................... 41Discussion ................................................................................................................................ 42Saddle Dolomite ........................................................................................................... 44C and O Isotopes ........................................................................................................... 44Reocín ................................................................................................................ 44Pine Point ........................................................................................................... 45Lennard Shelf ..................................................................................................... 45viNanisivik ............................................................................................................. 46Conclusion ................................................................................................................................ 46Chapter 3: Salmo-type Carbonate-Hosted Zn-Pb Deposits: Jersey and H.B. ..................48Introduction ............................................................................................................................. 48Jersey Deposit Geology ................................................................................................. 48H.B. Deposit Geology .................................................................................................... 50Sampling................................................................................................................................... 51Regional ........................................................................................................................ 51Jersey ............................................................................................................................ 51H.B. ............................................................................................................................... 51Petrography .............................................................................................................................. 53Carbonate Units ............................................................................................................ 53Mineralization ............................................................................................................... 54Cold-Cathode Cathodoluminescence ...................................................................................... 55Results .......................................................................................................................... 55Carbonate Staining and UV Light Analysis .............................................................................. 56Carbonate Staining ....................................................................................................... 56Alizarin Red S ...................................................................................................... 56Potassium Ferricyanide ...................................................................................... 57Results ................................................................................................................ 57UV Light Analysis .......................................................................................................... 58Results ................................................................................................................ 58C and O Isotope Analysis ......................................................................................................... 59Regional ........................................................................................................................ 59Jersey ............................................................................................................................ 65H.B. ............................................................................................................................... 70Distance from Mesozoic Plutons .................................................................................. 74Vein-Wall Rock Pairs ..................................................................................................... 75Discussion ................................................................................................................................ 76Conclusion ................................................................................................................................ 78Chapter 4: Salmo-type Carbonate-Hosted Zn-Pb Deposits: Remac .................................. 80Introduction ............................................................................................................................. 80Remac Deposit Geology ................................................................................................ 80Sampling................................................................................................................................... 82Petrography .............................................................................................................................. 82Carbonate Units ............................................................................................................ 82Mineralization ............................................................................................................... 83Cold-Cathode Cathodoluminescence ...................................................................................... 83viiResults .......................................................................................................................... 84Carbonate Staining and UV Light Analysis .............................................................................. 85Carbonate Staining Results ........................................................................................... 85UV Light Analysis Results .............................................................................................. 86C and O Isotope Analysis ......................................................................................................... 87Results .......................................................................................................................... 87Discussion ................................................................................................................................ 88Source of Si ........................................................................................................ 90Conclusion ................................................................................................................................ 92Chapter 5: Metaline-type Carbonate-Hosted Zn-Pb Deposits: Pend Oreille ..................95Introduction ............................................................................................................................. 95Pend Oreille Deposit Geology ....................................................................................... 96Josephine-Style Mineralization ..................................................................................... 99Yellowhead-Style Mineralization .................................................................................. 99Sampling................................................................................................................................. 100Petrography ............................................................................................................................ 100Regional ...................................................................................................................... 102Dolomite ..................................................................................................................... 102Mineralization ............................................................................................................. 102Cold-Cathode Cathodoluminescence .................................................................................... 104Results ........................................................................................................................ 104Carbonate Staining and UV Light Analysis ............................................................................ 106Carbonate Staining Results ......................................................................................... 106UV Light Analysis Results ............................................................................................ 107Geochemistry ......................................................................................................................... 108Significant Lithologies ................................................................................................. 109Results .............................................................................................................. 109Drill Hole TB731 .......................................................................................................... 111Analytical Data: Building a Chemo-Stratigraphic Framework...........................111Analytical Data: Geochemical Characterization of Lithologies .........................113Analytical Data: Hydrothermal Modification Pathways ................................... 113Analytical Data: Evaluation of Past Logging ..................................................... 113Analytical Data: Geochemical Haloes of Ore Horizons ..................................... 115Analytical Data: Trace Element Enrichments ................................................... 117Analytical Data: Element Detection Using a pXRF ............................................ 119pXRF Data ......................................................................................................... 119XRD Data .......................................................................................................... 122Drill Hole MXU573 ...................................................................................................... 124Analytical Data: Building a Chemo-Stratigraphic Framework...........................126viiiAnalytical Data: Geochemical Characterization of Lithologies .........................128Analytical Data: Geochemical Haloes of Ore Horizons ..................................... 130Analytical Data: Trace Element Enrichments ................................................... 130pXRF Data ......................................................................................................... 132C and O Isotope Analysis ....................................................................................................... 135Results ........................................................................................................................ 136Drill Hole TB731 ............................................................................................... 145Vein-Wall Rock Pairs ......................................................................................... 145Discussion .............................................................................................................................. 146Conclusion .............................................................................................................................. 151Chapter 6: Discussion and Conclusion .................................................................................... 153Discussion .............................................................................................................................. 153C and O Isotopes ......................................................................................................... 153Carbonate Textures ..................................................................................................... 155Sulphide Mineralogy and Paragenesis ........................................................................ 155Geochemistry ............................................................................................................. 156Global Context ............................................................................................................ 157Implications for Exploration ........................................................................................ 159Suggested Future Research ........................................................................................ 160Conclusion .............................................................................................................................. 161References ...................................................................................................................................... 164Appendix A: Methods .................................................................................................................. 175Appendix B: Sample Database .................................................................................................. 180Appendix C: Petrographic Descriptions and Field Photographs of each Key Lithologi-cal Unit Present within the Jersey and Remac mines, Southeastern BC........................211Appendix D: Petrographic Descriptions and Field Photographs of each Key Lithologi-cal Unit Present within the Pend Oreille Mine, Northeastern Washington..................236 Appendix E: Compiled C and O Data for Southeastern BC ................................................ 271Appendix F: Compiled Geochemistry Data for the Pend Oreille Mine, Northeastern Washington .....................................................................................................................................275Appendix G: Compiled XRD Spectra for Drill Hole TB731 .................................................. 313Appendix H: Compiled C and O Data for the Pend Oreille Mine ..................................... 324ixList of TablesTable 2.0. Comparison table of the four highlighted deposits in this study ......................................... 43Table 3.0. Carbonate phases and their associated alizarin red S and potassium ferricyanide stains. Modified from Hitzman (1999) ............................................................................................................. 56Table 3.1. Alizarin red S and potassium ferricyanide staining results for the Jersey Mine ...................57Table 3.2. UV light analysis results for the Jersey Mine ........................................................................ 58Table 4.0. Alizarin red S and potassium ferricyanide staining results for the Remac Mine ..................85Table 4.1. UV light analysis results for the Remac Mine ....................................................................... 86Table 5.0. Alizarin red S and potassium ferricyanide staining results for the Pend Oreille Mine ......  106Table 5.1. UV light analysis results for the Pend Oreille Mine ...........................................................  107Table 5.2. Summary table of proxy elements required if using pXRF data to define the chemo-stratig-raphy downhole in TB731. .................................................................................................................  122Table 5.3. Summary of results for XRD of aluminosilicate-rich horizons ...........................................  123Table 5.4. Summary of results for XRD of aluminosilicate-poor background samples ......................  123Table 5.5. Summary table of proxy elements required if using pXRF data to define the chemo-stratig-raphy downhole in MXU573 ..............................................................................................................  135xList of FiguresFigure 1.0. Location of sediment-hosted Zn-Pb deposits throughout British Columbia and Washing-ton. Modified from Fyles (1970) and Nelson et al. (2002) ..................................................................... 3Figure 1.1. Schematic diagram to show the variable tectonic settings of SEDEX and MVT deposit sys-tems. Modified from Lister et al. (1986) and Leach et al. (2010) ........................................................... 5Figure 1.2. Schematic diagram to show the formation environment of SEDEX and MVT deposit sys-tems. Modified from Leach et al. (2010) ................................................................................................ 9Figure 1.3. Regional geology of the Kootenay Arc in southeastern BC and northeastern Washington. Modified from Colpron and Price (1995), originally after Wheeler and McFeely (1991). ...................  13Figure 1.4. Schematic stratigraphic log of the regional geology in southeastern British Columbia and northeastern Washington with the Canadian and American nomenclature. Modified from Fyles (1970) ................................................................................................................................................... 14Figure 1.5. Location map of the four mineral deposits studied in this project and their corresponding geological setting. Modified from Wheeler and McFeely (1991), and Paradis et al. (2009)................  16Figure 2.0. Location of the Reocín Mine and regional geology of the Santillana syncline area in the Basque-Cantabrian basin of northern Spain. Modified from Symons et al. (2009) .............................  20Figure 2.1. Schematic stratigraphic log through the sedimentary sequence in the Basque-Cantabrian basin. Modified from Velasco et al. (2003) ..........................................................................................  20Figure 2.2. Schematic diagram to show the paragenetic sequence of events at the Reocín Mine. As described by Velasco et al. (2003) .......................................................................................................  22Figure 2.3. Graph of C and O isotopic data at the Reocín Mine. Data from Velasco et al. (2003) .......  24Figure 2.4. Location of the Pine Point region and regional geology of the Presqu’ile barrier and West-ern Canada Sedimentary Basin. Modified from Qing and Mountjoy, (1994; a.) Deposit geology of the Pine Point region. Modified from Rhodes et al. (1984; b.) ..................................................................  26Figure 2.5. Schematic stratigraphic log through the sedimentary sequence in the Presqu’ile barrier. Modified from Qing and Mountjoy (1994) ..........................................................................................  27Figure 2.6. Schematic diagram to show the paragenetic sequence of events at the Pine Point Mine. As described by Qing and Mountjoy (1990) and Gleeson and Turner (2007) ..........................................  29Figure 2.7. Graph of C and O isotopic data at the Pine Point Mine. Data from Qing (1998) ...............  30Figure 2.8. Location of the Lennard Shelf and regional geology of the Canning Basin region, north-west Australia. Modified from Arne et al. (1989), originally after Playford (1981) .............................  32Figure 2.9. Schematic stratigraphic log through the sedimentary sequence of the Lennard Shelf. Modified from Nicoll et al. (1993) and Tompkins et al. (1994) ............................................................  32Figure 2.10. Schematic diagram to show the paragenetic sequence of events at the Cadjebut deposit, Lennard Shelf, Canning Basin. As described by Tompkins et al. (1994) ...............................................  34Figure 2.11. Graph of C and O isotopic data at the Cadjebut deposit, Lennard Shelf, Canning Basin. Data from Tompkins et al. (1994) ........................................................................................................  35Figure 2.12. Location of the Nanisivik Mine and regional geology of the Borden basin and Milne Inlet region, northern Baffin Island. Modified from Arne et al. (1991) .......................................................  37xiFigure 2.13. Schematic stratigraphic log through the sedimentary sequence in the Bylot Supergroup of the Borden basin. Modified from Sherlock et al. (2004) .................................................................  37Figure 2.14. Schematic diagram to show the paragenetic sequence of events at the Nanisivik Mine. As described by Turner (2011). ............................................................................................................  39Figure 2.15. Graph of C and O isotopic data at the Nanisivik Mine. Data from Ghazban et al. (1991) .41Figure 2.16. Comparison graph of C and O isotopic data from all four key MVT systems. Data from Velasco et al. (2003), Qing (1998), Tompkins et al. (1994) and Ghazban et al. (1991) ........................  45Figure 3.0. Deposit geology of the Jersey and H.B. deposits, southeastern British Columbia. Modified from Giroux and Grunenberg (2010) ...................................................................................................  49Figure 3.1. Sampling location map on regional geology. Modified from Wheeler and McFeely (1991); Paradis et al. (2009) .............................................................................................................................  52Figure 3.2. Paragenetic diagram to highlight the sequence of events at the Jersey Mine ..................  53Figure 3.3. Photographs to display sulphide associations at the Jersey Mine ...................................... 54Figure 3.4. Cold-cathode CL images for samples JS07-08_180 and NC29 (FOV = 2 mm) ....................  55Figure 3.5. Selection of images to show results of the staining on Jersey samples .............................  57Figure 3.6. Selection of images to show results of the UV light analysis on Jersey samples ................59Figure 3.7. Plot to display the δ13C and δ18O values of all regional samples ....................................  60Figure 3.8. Plots to display the δ13C and δ18O values of all regional BC and Washington samples ac-cording to their rock-type and Formation ............................................................................................ 61Figure 3.9. Plot to display the δ18Ofluid and corresponding temperature for regional BC calcite samples ...............................................................................................................................................  64Figure 3.10. Plots to display the δ18Ofluid and corresponding temperature for the remaining regional BC and Washington samples ...............................................................................................................  66Figure 3.11. Plots to display the δ13C and δ18O values of all BC regional, Jersey, H.B. and Remac samples ...............................................................................................................................................  68Figure 3.12. Plots to display the δ13C and δ18O values of all Jersey and H.B. samples according to their rock-type .....................................................................................................................................  69Figure 3.13. Plots to display the δ18Ofluid and corresponding temperature for the Jersey and H.B. samples ...............................................................................................................................................  71Figure 3.14. Plots to display the δ18O and δ13C values of regional, Jersey, H.B. and Remac samples with distance to proximal Mesozoic intrusions ...................................................................................  75Figure 3.15. Plots to display the δ18O and δ13C values of vein-wall rock pairs. .................................  76Figure 4.0. Deposit geology of the Remac deposit, southeastern British Columbia. Modified from Kushner (2009) ....................................................................................................................................  81Figure 4.1. Paragenetic diagram to highlight the sequence of events at the Remac Mine .................  82Figure 4.2. Photographs to display sulphide associations at the Remac Mine ....................................  83Figure 4.3. Cold-cathode CL images for samples NC14-1 and NC14-3 (FOV = 2 mm) .........................  84Figure 4.4. Selection of images to show results of the staining on Remac samples. ...........................  85xiiFigure 4.5. Selection of images to show results of the UV light analysis on Remac samples ...............86Figure 4.6. Plot to display the δ13C and δ18O values of all Remac samples according to their rock-type .....................................................................................................................................................  88Figure 4.7. Plot to display the δ18Ofluid and corresponding temperature for the Remac samples ...  89Figure 4.8. Plot to display the δ18Ofluid and corresponding temperature for the Remac samples and chert ....................................................................................................................................................  93Figure 5.0. Deposit geology of the Pend Oreille Mine in the Metaline District, northeastern Washing-ton. Modified from St. Marie and Kesler (2000) .................................................................................  95Figure 5.1. Schematic stratigraphic log of the deposit geology at the Pend Oreille Mine, northeastern Washington. Modified from Otto (2001). ............................................................................................  97Figure 5.2. Schematic lithofacies model of the deposit geology at the Pend Oreille Mine, northeast-ern Washington. Modified from Bush et al. (1992) .............................................................................  98Figure 5.3. Sampling location map on regional geology. Modified from St. Marie and Kesler (2000) ............................................................................................................................................................ 101Figure 5.4. Paragenetic diagram to highlight the sequence of events at the Pend Oreille Mine .......101Figure 5.5. Photographs to display petrographic relationships at the Pend Oreille Mine ..................103Figure 5.6. Cold-cathode CL images for samples POM Limestone, POM Fine Grained Dolomite, POM 4707-1792; Josephine mineralization, POM 4862-1246; Yellowhead 1 mineralization and POM MX700S; Yellowhead 2 mineralization (FOV = 2 mm) ........................................................................ 105Figure 5.7. Selection of images to show results of the staining on Pend Oreille samples ..................106Figure 5.8. Selection of images to show results of the UV light analysis on Pend Oreille samples ....108Figure 5.9. Ternary plots to geochemically characterize the 12 significant lithologies observed at the Pend Oreille Mine. .............................................................................................................................. 110Figure 5.10. Downhole plots from drill hole TB731 to display element abundances with depth ......112Figure 5.11. Ca:Mg and Ca:Mn plots to display different stratigraphic units defined through geochem-istry for TB731 .................................................................................................................................... 114Figure 5.12. Plots to highlight the ore horizons in drill hole TB731 .................................................... 116Figure 5.13. Downhole plots to display trace element enrichments with depth in hole TB731 ........118Figure 5.14. Pb:Sb, Pb:Ag and Pb:Mo plots to display the positive correlations between these ele-ments for drill hole TB731 .................................................................................................................. 119Figure 5.15. Plots to display analytical and pXRF data comparisons for several elements for drill hole TB731 ................................................................................................................................................. 121Figure 5.16. Plots to determine the geochemical signature of each individual clay layer in TB731 ...125Figure 5.17. Downhole plots from drill hole MXU573 to display element abundances with depth ..127Figure 5.18. Plots to determine the geochemical signature of each individual clay layer in MXU573 ............................................................................................................................................................ 129Figure 5.19. Ca:Mg and Ca:Mn plots to display different stratigraphic units defined through geochem-istry for MXU573 ................................................................................................................................ 130Figure 5.20. Plots to highlight the ore horizons in drill hole MXU573 ................................................ 131xiiiFigure 5.21. Downhole plots to display trace element enrichments with depth in hole MXU573 .....133Figure 5.22. Plots to display analytical and pXRF data comparisons for several elements for drill hole MXU573.............................................................................................................................................. 134Figure 5.23. Plots to display the δ13C and δ18O values of all Pend Oreille samples according to their location ............................................................................................................................................... 136Figure 5.24. Plots to display the δ13C and δ18O values of all non-mineralized Pend Oreille samples according to their rock-type ............................................................................................................... 138Figure 5.25. Plots to display the δ13C and δ18O values of all Pend Oreille mineralized samples ac-cording to their rock-type ................................................................................................................... 139Figure 5.26. Plots to display the δ18Ofluid and corresponding temperature for the Pend Oreille samples .............................................................................................................................................. 141Figure 5.27. Average δ13C and δ18O values from all drill holes sampled at Pend Oreille plotted onto the deposit geology. Modified from St. Marie and Kesler (2000) ...................................................... 144Figure 5.28. δ18O and δ13C values with depth for drill hole TB731 .................................................. 145Figure 5.29. Plots to display the δ18O and δ13C values of vein-wall rock pairs .................................146Figure 5.30. Al:Cu and Al:In plots for drill holes TB731 and MXU573 to display the association of these elements with mineralization and clay-rich horizons ............................................................... 149Figure 6.0. Plot to display the δ13C and δ18O values from all 8 deposits studied .............................158xivAcknowledgementsI would like to thank Craig Hart for providing the opportunity to study with the Mineral Deposit Re-search Unit and for all of his help in designing and completing this particular project. I am also thankful for the support received from my supervisory committee, Greg Dipple, Ken Hickey and, in particular, Andreas Beinlich. Equally important was the support and advice offered by Lucas Marshall, Jelena Puzic, Liz Stock, Mike Takaichi and John Lamsma of Teck Resources Limited (“Teck”), Ed Lawrence of Sultan Minerals Inc. and Jack Denny of the Chamber of Mines of Eastern British Columbia, thank you to all of you.Geoscience BC, the Society of Economic Geologists and the Canadian Institute of Mining, Metallurgy and Petroleum are acknowledged and thanked for the scholarships provided to aid my efforts in work-ing on this project; this research would not have been possible without this support. Finally I would like to thank all of my fellow EOAS Department and MDRU graduate students, in par-ticular those who attended ladies’ wine nights. We will all get there eventually.1Chapter 1: Project Setting and ObjectivesIntroductionVisible and geochemical alteration of host rocks is common during the formation of many hydrother-mal mineral deposit systems, such as: porphyry, volcanogenic massive sulphide (VMS), epithermal and orogenic gold. These hydrothermal deposits typically form from high temperature fluids with the host rocks undergoing extensive fluid:rock interaction, which enables the formation of large, concen-trically variable haloes of visibly and geochemically altered rock. These haloes can subsequently be used to define the alteration footprint of the specific ore deposit type (for example, lateral and vertical alteration-mineralization haloes in porphyry deposits; Lowell and Guilbert, 1970). By contrast lower temperature hydrothermal deposits, such as base metal deposits that form in carbonate host rocks, typically have narrow and poorly developed visible and geochemical haloes. As such, these hydrother-mal systems generate significantly smaller alteration footprints, making exploration for these deposit types more challenging.Despite this, carbonate-hosted deposits can display invisible stable isotopic alteration haloes that are typically larger than the limits of the orebody, visible alteration or even geochemical anomalies (e.g. Barker et al., 2013). Stable isotopic alteration haloes can therefore be used to define the alteration footprint of these specific deposit types. Light stable isotopes of common elements in ore systems, such as C, O, H and S, have been utilized to understand fluid:rock interactions in and around ore de-posits for more than 40 years (Nesbitt, 1996). The intensity of stable isotopic alteration in ore systems increases from peripheral regions into the centre of mineralization with greater shifts towards lighter, or more depleted, isotopic values resulting from more fluid:rock interaction. Studies of light stable isotopes (C, O, H and S) have been used in an attempt to characterize the alteration footprints of a variety of base metal carbonate-hosted ore systems (e.g.: Nanisivik, Canada; Ghazban et al., 1991; Cadjebut, Australia; Tompkins et al., 1994; Pine Point, Canada; Qing, 1998; Pend Oreille, US; St. Marie and Kesler, 2000; Reocín, Spain; Velasco et al., 2003) with the use of stable isotopes in these systems being enabled by the fact that their host rocks undergo considerable fluid:rock interaction. The stable isotope composition of host rocks that have interacted with hydrothermal fluid will depend on a vari-ety of factors, including: •	 the isotopic composition of unaltered host rock,•	 the isotopic composition of the hydrothermal fluid, and•	 the temperature of dissolution and precipitationIt has been shown that light stable isotope studies, when used in conjunction with geological data, fluid inclusion studies and geochemical data, can not only identify fluid components, but also place im-portant constraints on their evolution in the system, e.g., origin of the ore fluid (Rye, 1993). Stable iso-tope alteration haloes can therefore provide information about fluid flow during mineralization, which in turn can enable mapping the extent of fluid interactions, discern fracture-controlled versus perva-2sive permeability, determine alteration temperatures, assess alteration intensity, and contribute to the development of ore deposit and exploration models (Barker et al., 2013). In general, host rocks that have seen higher degrees of fluid:rock interaction, or where fluid:rock interaction occurred at higher temperatures, will have a greater shift towards lighter, or more depleted, isotopic ratios. Therefore, host rocks in the vicinity of mineralization should have the most depleted isotopic values compared to rocks that are more distal. Furthermore, the size and distribution of stable isotopic alteration is likely controlled by the total flux of hydrothermal fluid and as such, is itself controlled by permeability, min-eralogy, grain size, temperature and fluid:rock ratios in the surrounding host rocks (Barker et al., 2013).Stable isotope analyses are traditionally measured using gas-source isotope ratio mass spectrometry (IRMS). However, C and O isotope ratios can now be analyzed using infrared absorption to measure isotopic signatures in different gas species. One such infrared absorption technique, that has been developed and applied at the Mineral Deposit Research Unit (MDRU) at the University of British Co-lumbia (UBC), is off-axis-integrated cavity output spectroscopy (OA-ICOS). The OA-ICOS uses a laser source, which produces light at an infrared wavelength suitable for interacting with the gas species of interest (Barker et al., 2013). This Masters thesis focuses on characterizing the alteration footprints of the carbonate-hosted Zn-Pb deposits in southeastern British Columbia and northeastern Washington; utilizing C and O isotopic alteration. These deposits are poorly classified and have been variably attributed to a range of mineral deposit models, such as: Mississippi Valley-type, Sedimentary Exhalative and Irish-type (Fyles, 1970; Sangster, 1990; Goodfellow and Lydon, 2007; Paradis, 2008; Simandl and Paradis, 2009), and further complications on classification arise because many deposits also contain non-sulphide ores and/or skarn mineralization (Simandl and Paradis, 2009; Giroux and Grunenberg, 2010). The four deposits studied in this thesis are H.B., Jersey and Remac in British Columbia and Pend Oreille in Washington (Figure 1.5). The project objectives are as follows:•	 determine the size and intensity of alteration surrounding different types of carbonate-hosted Zn-Pb deposits within southeastern British Columbia and northeastern Washington,•	 characterize and map fluid flow pathways and assess the intensity of fluid:rock interactions as vectoring tools,•	 assess stable isotope alteration from proximal intrusion-related through to distal carbonate-hosted ore systems, and•	 determine the optimal sampling protocols and strategies to utilize stable isotopes as an explo-ration toolIt is anticipated that a study of the C and O isotopic alteration halo of these deposits, when used in conjunction with other geological data, will be able to place better constraints on the fluid components and fluid source and so aid in clarifying the current deposit models for these systems. In addition, the definition of a characteristic alteration footprint for these systems may aid in vectoring towards similar 3deposits in this region. Methods utilized in this Masters thesis are outlined in Appendix A.In order to better comprehend and interpret the C and O isotope data collected during this study, it is essential to have a solid understanding of the several deposit models for sediment-hosted Zn-Pb ore systems. The following presents a summary of the formation of Sedimentary Exhalative, Mississippi Valley-type and Irish-type Zn-Pb deposits. Sediment-Hosted Zn-Pb SystemsSediment-hosted stratiform and stratabound Zn-Pb deposits occur in a series of belts in the Canadian Cordillera (Nelson et al., 2002). In northeastern British Columbia, Sedimentary Exhalative (SEDEX) Zn-Pb deposits are hosted in Devonian to Mississippian shales of the Kechika Trough and Selwyn Basin (Nelson et al., 2002). Further east, Mississippi Valley-type (MVT) Zn-Pb deposits are hosted in Silurian to Devonian dolomites (Nelson et al., 2002). In southeastern British Columbia and northeastern Wash-ington additional carbonate-hosted Zn-Pb deposits form a belt throughout the Cambrian to Ordovi-cian Kootenay Arc (Fyles, 1970). Figure 1.0 displays the location of these deposits throughout British Columbia and Washington. INSULAR BELTCOAST BELTINTERMONTANE BELTOMINECA BELTFORELAND BELTBRITISHCOLUMBIAWASHINGTONALBERTA117°00’ W49°00’ NIDAHO0 200kmNSEDEX DepositsMVT DepositsKootenay Arc DepositsVancouverSeattle SpokaneNelsonKelownaPrince GeorgeDevonian Carbonate FrontAncestral North America Basinal FaciesPericratonic Terrane with Devonian-Mississippian Igenous ComponentCassiar PlatformSullivanMineFigure 1.0. Location of sediment-hosted Zn-Pb deposits throughout British Columbia and Washington. Modified from Fyles (1970) and Nelson et al. (2002).4In addition to Zn-Pb deposits in the northeast and southeast, British Columbia is also host to the Sulli-van Fe-Pb-Zb-Ag deposit (Figure 1.0); one of the world’s largest SEDEX deposits. The Sullivan deposit is located near Kimberley in southeastern British Columbia and is hosted in Mesoproterozoic metasedi-mentary rocks of the Aldridge Formation, Purcell Supergroup (Hamilton et al., 1983). The ore is found at the contact between the lower and middle Aldridge Formation, which is formed of argillaceous and quartzitic turbidites, and comprises pyrrhotite, sphalerite, galena and lesser pyrite (Hamilton et al., 1983). The Sullivan Mine closed in 2001 after having produced more than 160 Mt of ore grading 6.5 % Pb, 5.6 % Zn, 67 g/t Ag and 25.9 % Fe (Höy et al., 2003).Oxygen isotope data collected by Nesbitt and Longstaffe (1982) from regional and ore-proximal rocks of the Aldridge Formation indicate that there is no shift in δ18O between mineralized and non-miner-alized host rocks in the vicinity of the Sullivan deposit. Furthermore, the δ18O isotope data (9.90 ‰ to 13.60 ‰; Nesbitt and Longstaffe, 1982) are considerably more depleted compared to δ18O data for the Zn-Pb systems presented in this thesis. As there is an abundance of these deposit-types throughout the Canadian Cordillera, it is important to understand the separate deposit models for each sediment-hosted Zn-Pb system.OverviewSediment-hosted Zn-Pb ore deposits are significant contributors to global Zn and Pb resources. Ac-cording to Leach et al. (2005; 2010) they may be split into two major subtypes: Sedimentary Exhalative (SEDEX) and Mississippi Valley-type (MVT) deposits. The most significant difference between these two deposit types is their depositional environment (Figure 1.1). SEDEX deposits form in passive margins, back-arc basins and continental rifts, whereas MVT deposits form in platform carbonate sequences, typically in passive margin tectonic settings (Leach et al., 2010). This fundamental difference in deposi-tional environment in turn controls the host rock lithology, deposit morphology and ore textures that further define each deposit type (Leach et al., 2005).The uneven global distribution of sediment-hosted Zn-Pb deposits in space and time is a direct reflec-tion of their genesis, which is linked to the Earth’s evolving tectonic and geochemical systems includ-ing the recycling of the sedimentary rock record. As such, most sediment-hosted Zn-Pb deposits have most likely been destroyed by subduction and erosion, or modified by metamorphism and tectonism, so that they are no longer recognizable (Leach et al., 2010). Despite this challenge, several papers (Sangster, 1990; Goodfellow et al., 1993; Leach et al., 2001; Leach et al., 2005; Lydon, 2004; Kesler and Reich, 2006; Lyons et al., 2006; Goodfellow and Lydon, 2007) have focused on these deposit systems and have attributed their distribution to the following: 1) long-term changes in the tectonic processes of the Earth that are related to the cooling planet, 2) the change in oxidation state of the atmosphere and hydrosphere, 3) the evolution of life, and 4) the tectonic recycling and/or destruction of their car-bonate host rocks.Sediment-hosted Zn-Pb deposits may be hosted in a wide range of siliciclastic and carbonate rocks, 5Figure 1.1. Schematic diagram to show the variable tectonic settings of SEDEX and MVT deposit systems. A single con-tinent is undergoing extension (A). Rifting occurs (B) and is accompanied by the opening of a new ocean, which is bordered on either side by a passive margin. The ocean widens (C). After some time, an arc approaches from the west (D). Collision occurs between the approaching arc and the subducting passive margin (E). Final stage of the arc-passive margin collision (F). Modified from Lister et al. (1986) and Leach et al. (2010).ARifting200 kmUpper continental crustMiddle continental crustLower continental crustLithosphereBPassive margins facing a new, wide ocean basinCPassive margins facing a wide ocean basinOceanic crustDArc approaches a passive marginArcEInitial arc-passive margin collisionFCollisionSEDEX host MVT host MVT mineralizationShale and deep-watercarbonatesCarbonate andminor clastic rocks Rift facies6which are generally not associated directly with any igneous activity (Leach et al., 2010). The ore min-erals consist mainly of sphalerite and galena with lesser amounts of pyrite, marcasite, pyrrhotite and silver. Gangue minerals commonly comprise dolomite, siderite, ankerite and calcite with lesser barite and silicification of the host rocks (Leach et al., 2010). The ore deposits themselves have a variety of relationships with their host rocks that include stratiform, stratabound and discordant ores (Leach et al., 2010).The tectonic environments (Figure 1.1) of these sediment-hosted deposits are important in the forma-tion and preservation of the Zn-Pb ore as the tectonic setting controls a variety of key factors: host rock type, ore controls, temperature, pressure and preservation potential (Leach et al., 2010). The following few paragraphs present a summary of the significant tectonic environments, mainly rift or passive margin settings, in which sediment-hosted Zn-Pb deposits are formed as described by Leach et al. (2010).Continental RiftsThe key characteristics of a continental rift, such as the East African rift system, are an extended con-tinental basement, a topographically low depression and elevated flanks. These elevated flanks are commonly bordered by normal faults, which may rise to several kilometres elevation (Sengor and Natal’in, 2001). Coarse, clastic sediment sheds off these higher elevations to form alluvial fans that may either be marine or more commonly non-marine depending on global sea level at the time of rift formation (Sengor and Natal’in, 2001). Subsidence is commonly rapid enough for the formation of deep-water, off-shelf environments in which SEDEX deposits may be generated (Leach et al., 2010). Geophysical modelling of symmetrical extension has not been able to explain why opposing margins do not typically show identical structures (Lister et al., 1986). An asymmetrical model of extension has been suggested (Lister et al., 1986; Lister et al., 1991) suggesting that detachment faults (low-angle normal faults) play an important role in continental rifting (Wernicke 1981; Wernicke 1985; Lister et al., 1986; Lister et al., 1991). Asymmetric detachment models during continental rifting produce highly asymmetric structures as the middle and lower crust are pulled out from underneath a fracturing, brittle upper crust (Figure 1.1; Lister et al., 1986; Lister et al., 1991). During detachment faulting, the lower plate will warp upwards in response to the removal of the upper plate (Lister et al., 1986; Lister et al., 1991). In the simplest model, continental rifting will occur within, or close to the peak distortion of the lower plate as this is where the crust is thinnest (Lister et al., 1986; Lister et al., 1991). Passive MarginsContinental rifting has the potential to evolve into sea-floor spreading, the formation of oceanic crust and a mid-oceanic ridge. Over time, the continents on either side of the mid-oceanic ridge move further apart and allow for the formation of passive margins on the rifted edges of both diverging continents. Due to the asymmetrical model of extension (Lister et al., 1986; Lister et al., 1991), these resultant passive margins will also display a marked asymmetry (Figure 1.1; Lister et al., 1986; Lister 7et al., 1991). The asymmetry of the margin will depend on whether the underlying detachment fault dipped towards or away from the ocean (Lister et al., 1986) and so will result in two classes of passive margin as described by Lister et al. (1986) and Lister et al. (1991): 1) upper-plate margins composed of rocks above the detachment fault, and 2) lower-plate margins composed of deeper crystalline rocks overlain by highly faulted remnants of the upper crust. The upper-plate margin will have relatively little structure, contrasting with the lower-plate margin which will be highly structured (Lister et al., 1986; Lister et al., 1991).These passive margins undergo thermal subsidence through cooling and allow for the burial of the earlier rift-related sediments by passive margin strata (“the rift-to-drift transition”; Leach et al., 2010). Over time, the passive margin sediments build upward and outward to form a seaward thickening se-quence referred to as the shelf, slope and rise. Sea level, climate, distance from shore and water depth all govern the character of the sediments deposited on the shelf, slope or rise (Leach et al., 2010). High latitude shelf environments will be dominated by mature sandstones and siltstones whereas low lati-tude environments will be dominated by platform carbonates, which may eventually become the host of most MVT deposits (Leach et al., 2010). Further offshore, on the continental slope and rise, sedi-ments are commonly composed of finer grained clastic rocks (siltstone and shale), which may become host to syngenetic or diagenetic SEDEX deposits (Leach et al., 2010).Destruction of Passive Margins Most passive margins will be destroyed through a collision with an arc (Bradley, 2008). The sediments that form the slope, shelf and rise along with the underlying continental basement are pushed down-wards into the subduction zone. As subduction and plate convergence proceeds, the foreland ba-sin fills with sediment. Therefore beneath the foreland basin deposits, in buried platform carbonates of previous passive margins, is a highly favourable environment in which to find MVT mineralization (Bradley and Leach, 2003).Extensional ArcsSome SEDEX deposits have been found in back-arc basin tectonic settings (Leach et al., 2010). Back-arc rifting can result in a back-arc basin formed from oceanic crust (Leach et al., 2010). One side of the resulting passive margin is likely affected by subsequent arc magmatism, so passive margin sedimen-tation is commonly overprinted by igneous activity (Leach et al., 2010). The other margin, however, will behave similarly to a classic passive margin and could be the potential host of SEDEX and/or MVT mineralization (Leach et al., 2010).Continental Sag BasinsThermal subsidence dominates continental rifts in which rifting is halted before sea-floor spreading and the formation of a mid-oceanic ridge. The rift-related sediments are then buried in a similar pro-8cess to that of passive margins (Leach et al., 2010).SEDEX DepositsSedimentary Exhalative (SEDEX) massive sulphide deposits are an economically important source of Pb, Zn, and Ag worldwide. These deposits are typically found in Middle Proterozoic to Middle Paleo-zoic fault-controlled passive margins, back-arc basins or continental rifts (MacIntyre, 1991). They are characterized by thin to massive beds of relatively simple sulphide mineralogy comprising sphalerite, galena, pyrite and pyrrhotite, although barite may be a major component in deposits of Devonian to Mississippian age (MacIntyre, 1991). They are commonly hosted in carbonaceous cherts, fine grained turbidites or shelf carbonates (MacIntyre, 1991). The abundance of Cu is commonly low; however, some deposits do show significant chalcopyrite mineralization. The presence of Cu is attributed to higher temperature ore formation and often exhibits the most abundance in deposits associated with volcanic rocks (MacIntyre, 1991).The formation of SEDEX deposits is commonly associated with syn-sedimentary faulting, as indicated by the presence of slump breccias, conglomerates and coarse-clastic turbidites (MacIntyre, 1991). These faults are thought to be pathways that permitted heated basinal brines to move up into the overlying sedimentary sequence. As the brines are dense, upon exhalation they settle into submarine depressions (Sato, 1972). Sulphide minerals and/or barite may then precipitate from these brines due to changing physiochemical conditions (MacIntyre, 1991). The source of Zn, Pb and Ba in these SEDEX deposits is thought to be the underlying sedimentary se-quence or basement rocks (Badham, 1981) through the reaction of heated basinal fluids with cratonic derived, clastic material. This is due, for the most part, to the absence of igneous activity and the relatively radiogenic nature of the Pb and Sr isotopes from these SEDEX systems, suggesting an upper crustal origin (Godwin and Sinclair, 1982; Goodwin et al., 1982; Modene and Ryan, 1986; Turner, 1990).The stockwork zones beneath many SEDEX systems suggest the upflow of the mineralizing fluid is di-rectly below the deposit itself (MacIntyre, 1991). Mineralizing fluids are most likely derived from the dewatering of a basinal sedimentary sequence due to rifting, subsidence and an increased heat flow (Figure 1.2). These fluids are then able to reach the seafloor through permeable sedimentary horizons and syn-sedimentary faults. Fluid inclusion studies in SEDEX systems indicate that the mineralizing fluids are moderately to highly saline with an average salinity of 9.4 wt. % NaCl and range from 150°C to 275°C in temperature (Mako and Shanks, 1984; Gardener and Hutcheon, 1985; Shaw and Hodgson, 1986).  9Figure 1.2. Schematic diagram to show the formation environment of SEDEX and MVT deposit systems. Modified from Leach et al. (2010).The source of S in SEDEX deposits has been studied using δ34S isotopes. SEDEX deposits usually have δ34S values that are the same or heavier than those of coeval seawater (Claypool et al., 1980; Cecile et al., 1983; Goodfellow and Jonasson, 1984). Heavier δ34S values in sulphide and sulphate phases are thought to be caused by the progressive reduction of sulphate to H2S within an anoxic water column (MacIntyre, 1991). Heavier values may also be produced by preferential removal of 32S through the for-mation of diagenetic pyrite and related phases. These processes coincide with the host lithologies of many SEDEX deposits, which also indicate a strongly reduced basinal environment (MacIntyre, 1991).MVT DepositsMississippi Valley-type (MVT) deposits are also an economically important source of Pb and Zn world-wide. The defining characteristics of MVT deposits as described by Leach and Sangster (1993) and Leach et al. (2005) are as follows: 1) epigenetic, 2) not associated with igneous activity, 3) hosted in dolostone or limestone and rarely in sandstone, 4) the main ore minerals are sphalerite, galena, pyrite, marcasite and gangue minerals are dolomite and calcite, 5) occur in platform carbonate sequences commonly at the margins of basins or foreland thrust belts, 6) typically stratabound, but may be locally stratiform, 7) typically form in large districts, 8) the ore fluids are basinal brines (evaporated seawater) with roughly 10 to 30 wt. % salt, 9) a crustal source of metal and S, 10) temperature of ore deposition is typically between 75°C and 200°C, 11) the most important ore controls are faults and fractures, dis-solution collapse breccias and lithological transitions, 12) sulphides are coarse to fine grained, massive to disseminated, 13) the sulphides occur mainly as replacement of carbonate rocks and to a lesser extent, open-space fill, and 14) alteration consists mainly of dolomitization, host-rock dissolution and SEDEX HostMVT Hostsea levelMac and ultramac rocksShale and deep-water carbonatesCarbonate and minor clastic rocks (passive-margin platform)Rift faciesBasement10brecciation.Most MVT systems are hosted in Phanerozoic rocks with the ore type significantly less common in Proterozoic rocks and only one known deposit, Pering-Bushy Park in South Africa, hosted in Neoar-chean rocks (Leach and Sangster, 1993; Leach et al., 2005). MVT deposits form in platform carbonate sequences in passive margin environments and are commonly associated with pre-existing extensional structures inboard of Phanerozoic contractional tectonic belts (Figure 1.2; Leach et al., 2001; Leach et al., 2005). The deposits commonly occur in extensive districts and the average size and grade for in-dividual MVT deposits is 7.0 Mt ore grading at 6.0 % Zn and 1.9 % Pb (Leach et al., 2010). The general consensus is that most Phanerozoic MVT deposits formed during major contractional events (Leach et al., 2001; Leach et al., 2005). This relationship applies to most MVT districts with the exception of the Canning Basin district in Australia (Leach et al., 2010), which formed in a large scale extensional environment (Leach et al., 2010).The source of Zn and Pb in MVT deposits is thought to be the same as for SEDEX deposits, i.e., scav-enged from the underlying sedimentary sequence or basement rocks (Badham, 1981). Again, this is due to the absence of igneous activity and the relatively radiogenic nature of the Pb and Sr isotopes from these MVT systems suggesting an upper crustal origin (Godwin and Sinclair; Goodwin et al., 1982; Modene and Ryan, 1986; Turner, 1990).Mineralizing fluid flow can be attributed to three different processes (none of which successfully ex-plain all features of an MVT system); 1) topographic or gravity-driven flow whereby meteoric water flushes mineralized brines through fault networks or permeable lithologies from an elevated recharge area, 2) sediment compaction or tectonically driven expulsion of the mineralized fluid during compres-sion, or 3) hydrothermal convection. Fluid inclusion studies in MVT systems indicate that the mineraliz-ing fluids are moderately to highly saline with a salinity range of 10 to 30 wt. % salt and a temperature of 90°C to 150°C (Basuki and Spooner, 2004; Leach et al., 2005). As with the source of S in SEDEX systems, the δ34S isotopic values in MVT deposits are generally heavy and indicate a crustal, ultimately seawater origin that has been produced through the progressive re-duction of sulphate to H2S within an anoxic water column (Anderson and Macqueen, 1982).The deposition of sulphides in an MVT deposit is primarily controlled by two end member scenarios; the non-mixing and the mixing models (Anderson and Macqueen, 1982). The non-mixing model re-quires that the metals and reduced S travel together in the ore-forming fluid whereas in the mixing model, S is derived at the site of ore deposition through the reduction of sulphate carried in the ore-forming fluid and subsequent reaction with the metal-rich host rock (Anderson and Macqueen, 1982).Non-Mixing ModelIn the non-mixing model the metals and reduced S are required to travel together in the same ore-forming fluid; however, this instigates the problem of metal solubility. In fluids that contain appre-ciable amounts of reduced S metal solubility will be low, hence there will be a low potential of ore 11deposition (Anderson and Macqueen, 1982). In order for this non-mixing model to operate under the normal pH conditions of MVT deposits, chloride complexing of metals cannot be the only form of met-al transport. An alternative transport method is thought to be through organic complexing (Anderson and Macqueen, 1982). In this model, sulphide precipitation would then be driven by cooling, a change in pH or dilution by groundwater (Anderson, 1975). Geologically, the non-mixing model is evidenced through extensive sphalerite banding that could not otherwise be produced so consistently with an individual, local source of S (southwestern Wisconsin, USA; McLimans et al., 1980) and the extensive dissolution and replacement of previously deposited sulphides in some MVT districts (southeastern Missouri, USA; Anderson and Macqueen, 1982). This replacement is difficult to reconcile with the in-soluble nature of sulphide minerals required by the mixing model (Anderson and Macqueen, 1982). Furthermore, coarse and well-formed sulphide crystals are thought to indicate the absence of high degrees of supersaturation or large concentration gradients necessary to the mixing model (Anderson and Macqueen, 1982).Mixing ModelAs a low pH is required to transport metals and sulphide together, an alternative model in which the sulphide and metals are transported separately has been proposed. In particular, mixing models in which H2S is added to metal-bearing brines at the site of ore deposition are more hydrodynami-cally favourable than those mixing models that require the sulphide and metals to be transported in two separate fluids that subsequently react with each other (Anderson and Macqueen, 1982). A key source of reduced S is from the reduction of locally available sulphate minerals or seawater sulphate. Organic matter that is common in MVT deposits is the most likely reducing agent in this process at the typical temperatures of these systems (Anderson and Macqueen, 1982). The mixing of metals and H2S at the site of ore deposition would give rise to the large concentration gradients, high degrees of supersaturation and rapid precipitation of sulphides required for the fine grained sulphides observed in some MVT deposits (e.g., Pine Point, Canada and Pend Oreille, USA). In addition, the mixing process is also able to take place over a longer timescale by the diffusion of H2S through host rocks into the hydrothermal system. This longer timescale may result in the typical, coarse grained and well-formed sulphide minerals that were previously thought to be indicative of the non-mixing model (Anderson and Macqeen, 1982).Irish-type DepositsAn additional sediment-hosted Zn-Pb deposit style is the Irish-type deposit, so named for its charac-terization in the Irish Midlands. There has been substantial discussion regarding the origin of these de-posits with several authors supporting an MVT-style model (Hitzman, 1999; Leach et al., 2001, Leach et al., 2010) and several authors supporting a SEDEX-style model (Andrew and Ashton, 1985; Wilkinson, 2003; Wilkinson, 2010; Wilkinson, 2014). As Irish-type deposits display characteristics of both MVT- and SEDEX-style mineralization, it has been suggested that these Irish-type deposits are transitional (Wilkinson, 2014).12The Irish-type deposits display very similar mineralization styles to SEDEX and MVT systems with ore minerals dominated by sphalerite, galena, pyrite and minor chalcopyrite (Wilkinson, 2010; 2014). Sul-phide mineralization is commonly associated with Carboniferous dolomites and limestones and oc-curs as stratiform to stratabound lenses at permeable or reactive host rock horizons (Wilkinson, 2010; 2014). The source of Zn and Pb in Irish-type deposits is thought to be the same as for SEDEX and MVT deposits, i.e., scavenged from the underlying sedimentary sequence or basement rocks. The ore fluids have a seawater origin with roughly 8 to 19 wt. % salt and an ore forming temperature of 70°C to 200°C (Wilkinson, 2003; 2010; 2014). It has been suggested that the formation of the major Irish ore bodies was due to hydrothermal fluid flow linked to limited extension in the Lower Carboniferous (Wilkinson, 2014).Regional GeologySoutheastern British Columbia (BC) and northeastern Washington are underlain by the central part of the Kootenay Arc (Figure 1.3), a curving belt of lower to mid-Paleozoic complexly deformed sed-imentary, volcanic and metamorphic rocks extending to the southeast, south and southwest from Revelstoke across into northeastern Washington (Fyles, 1967). This belt of rocks lies along the suture between ancestral North America and the Intermontane superterrane (Colpron and Price, 1995). The Kootenay Arc spans a stratigraphic and structural transition between rocks that consist, from west to east, of: 1) an Early Mesozoic island arc assemblage, 2) a complex succession of upper Paleozoic to lower Mesozoic marginal basin assemblages, 3) the lower Paleozoic North American outer continental margin that has experienced several stages of deformation, and 4) the upper Proterozoic to middle Paleozoic North American miogeocline (Einarson, 1994). Structurally, the arc marks a distinct change in structural style from the upright folds of the Purcell Anticlinorium in the east to coaxially refolded, west-verging recumbent isoclines farther west (Fyles, 1964; Reesor, 1973; Warren and Price, 1993). In general, polyphase deformation along the Kootenay Arc has obscured primary stratigraphic relation-ships (Colpron and Price, 1995). However, on a local scale there is well-preserved gradational inter-bedding, consistent facing directions and lateral continuity that have enabled the construction of a reasonably reliable regional stratrgraphic sequence (Colpron and Price, 1995).13CanadaUnited States52°00’ N116°00’W120°00’ W48°00’ NMesozoic plutonsDevonian plutonsKootenay TerraneEagle Bay AssemblageActive, Nelway and Laib formationsNorth AmericaMiogeoclineBelt - Purcell Supergroup0 100kmNMonashee ComplexValhalla ComplexIntermontaneSuperterraneINSULAR BELTCOAST BELTINTERMONTANE BELTOMINECA BELTFORELAND BELTBRITISHCOLUMBIAWASHINGTONALBERTAIDAHOVancouverSeattle SpokaneNelsonKelownaPrince GeorgeNelsonKelownaFigure 1.3. Regional geology of the Kootenay Arc in southeastern BC and northeastern Washington. Modified from Col-pron and Price (1995), originally after Wheeler and McFeely (1991).14The rocks of the central Kootenay Arc comprise a thick clastic assemblage of Neoproterozoic Wind-ermere Supergroup, upon which Lower Cambrian argillaceous quartzites of the Reno Formation are overlain by the Cambrian Laib Formation, containing the Truman, Reeves and Emerald members (Fig-ure 1.4). The Truman member of the lower Laib Formation consists of a thin sequence of interbedded phyllite and limestone. The Reeves member mainly consists of fine to medium grained limestone, which has been locally altered to dolostone. This limestone characteristically displays grey, black and white banding typically a few centimetres in width. The dolostone commonly weathers buff, is poorly banded or massive, and is normally finer grained than the limestone. The Emerald member overlies the Reeves member limestone and is characterized by a black to grey, foliated, carbonaceous, often crenulated, phyllite unit. At the top of the Laib Formation, the upper Laib Formation is an undivided series of phyllite units with lesser intercalated beds of micaceous quartzite and limestone (Fyles and Hewitt, 1959).The Laib Formation is overlain by the Middle Cambrian Nelway Formation (Figure 1.4), a second unit of limestone and dolomite. Above the Nelway Formation is the Ordovician Active Formation (Figure 1.4); a unit of black argillite slate with minor calcareous lenses (Fyles and Hewitt, 1959). Throughout southeastern BC and northeastern Washington there are numerous intrusions of Mesozoic granite, leucocratic granite and syenite. These have traditionally been considered part of the Nelson batholith suite, which has been dated as Middle Jurassic (Little, 1960; Höy and Dunne, 1997).LowerCambrianMiddleCambrianOrdovicianQuartzite Range FormationReno FormationTrumanmemberReevesmemberEmeraldmemberNelwayFormationActiveFormationupperLaibFormationLaib FormationPeriodSoutheastern British ColumbiaNortheastern WashingtonLedbetterSlateMetalineFormationMaitlenPhylliteGypsyQuartziteReevesmemberFigure 1.4. Schematic stratigraphic log of the regional geology in southeastern British Columbia and northeastern Wash-ington with the Canadian and American nomenclature. Modified from Fyles (1970).15Although the rock units in southeastern BC and northeastern Washington are generally the same, Canadian and American nomenclature of these units differ. Figure 1.4 shows a schematic stratigraphic log of the regional geology for these areas with the Canadian and American equivalents. For simplicity, the American nomenclature will only be used in Chapter 5; all other chapters will use the Canadian nomenclature.Structure and Metamorphic HistoryThe rocks of the Kootenay Arc have a complex structural history involving at least three phases of fold-ing, major regional low angle thrust faults and multiple smaller faults spanning an age interval from the mid-Mesozoic to the Eocene (Fyles and Hewitt, 1959; Höy, 1977). Notably, the regional structural trend is to the northeast, which is contrary to the typical northwesterly trend of the Cordillera. In gen-eral, there is an increase in the structural complexity and intensity of deformation from the southwest to the east and north (Webster and Pattison, 2013). This increase in structural complexity and defor-mation corresponds to a deepening in structural level (Webster and Pattison, 2013).Regional metamorphism reaches to lower greenschist facies and is thought to have been synchronous with the earliest phase of deformation. Contact metamorphism is locally associated with the intrusion of the Middle Jurassic igneous rocks and postdates all phases of folding (Fyles and Hewitt, 1959; Höy, 1977).MineralizationCarbonate-hosted Zn-Pb deposits occur along the entire length of the Kootenay Arc with the largest deposits occurring in the vicinity of Salmo, BC and Metaline Falls, Washington, on either side of the Canada-United States border. From north to south, the major carbonate-hosted Zn-Pb deposits of the Kootenay Arc include the Duncan, H.B., Jersey, Remac and Pend Oreille deposits (Figure 1.5). These deposits can be split into two endmembers (Fyles, 1970):•	 Salmo-type deposits consist of stratiform lenses of pyrite, sphalerite and galena in zones of dolomite in the highly deformed Lower Cambrian Reeves member limestone. These types of deposits include the Duncan (near the north end of Kootenay Lake), the H.B., Jersey and Remac orebodies.•	 Metaline-type deposits consist of stratiform lenses in the relatively undeformed, stratigraphi-cally younger, Middle Cambrian Nelway Formation carbonate rocks. These types of deposits include those situated near Metaline Falls, Washington, and in particular, the Pend Oreille de-posit.Paradis (2008) and Paradis et al. (2014) recently concluded that Salmo-type deposits represent MVT-style mineralization, based on field-observations and Re-Os dating of the sulphide mineralization, which produced a Devonian age. In addition, previous work by Paradis et al. (2015; 2016) has con-16Figure 1.5. Location map of the four mineral deposits studied in this project and their corresponding geological setting. Modified from Wheeler and McFeely (1991), and Paradis et al. (2009).SalmoStockWallack CreekStockLost Creek            PlutonSpirit BatholithPend Oreille RiverSlate Creek ThrustFlume Creek FaultSheep Creek AnticlineLeadpoint FaultBlack Canyon ThrustColumbia ThrustSalmo Valley FaultBlack Blu ThrustLeadpoint FaultSheep Creek AnticlineSalmo River AnticlineSalmo River AnticlineBritish ColumbiaWashingtonIdaho49°00’ N117°30’ W 117°00’ W49°00’ N0 5kmNCretaceous IntrusionsLower to Middle Jurassic Elise FormationDevonian to Lower Triassic Charbonneau Creek AssemblageOrdovician Active FormationMiddle Cambrian Nelway FormationLower Cambrian to Ordovician Upper Laib FormationLower Cambrian Laib FormationLower Cambrian Reno FormationLower Cambrian Quartzite Range FormationLower Cambrian and Proterozoic Three Sisters FormationLower Cambrian and Proterozoic Monk FormationPend Oreille MineRemacMineJerseyMineH.B.MineINSULAR BELTCOAST BELTINTERMONTANE BELTOMINECA BELTFORELAND BELTBRITISHCOLUMBIAWASHINGTONALBERTAIDAHOVancouverSeattle SpokaneNelsonKelownaPrince GeorgeRoadFaultMetalineFallsSalmo17centrated on the C and O isotopic signature of non-sulphide mineralization present at the Jersey, H.B. and Remac mines. These carbonate-hosted non-sulphide deposits form during supergene enrichment of the primary base metal MVT deposits (Paradis et al., 2015; 2016). In the case of these Salmo-type deposits, non-sulphide minerals comprise of hemimorphite (Zn4Si2O7(OH)2.H2O), iron oxides, willem-ite (Zn2SiO4), hydrozincite (Zn5(CO3)2(OH)6) and cerussite (PbCO3; Paradis et al., 2015; 2016). C and O isotopic data collected by these authors from dolomites hosting the primary MVT-style mineralization display increased depletion from non-mineralized to mineralized host rocks due to increased fluid:rock interaction. The non-sulphide mineralization displays a much greater depletion in δ13C compared to the primary MVT-style mineralization (Paradis et al., 2015; 2016). This depletion in δ13C is attributed to the presence of a depleted C-source resulting from the incorporation and subsequent oxidation of organic C (Paradis et al., 2015; 2016).In order to attain a better understanding of the poorly defined Kootenay Arc carbonate-hosted Zn-Pb deposits, it is essential to consider the global context of these systems. Four key deposits with a range in geographic location, age and minimal deformation have been studied, as a basis for comparison with the deposits investigated in the current study.18Chapter 2: Global ContextIntroductionTextural and isotopic signatures have been extensively used in the past to infer specific controls on the formation environments of sediment-hosted Zn-Pb deposits; both Mississippi Valley-type (MVT) and Sedimentary Exhalative (SEDEX). This is because both textural and isotopic signatures can point towards certain characteristic environmental factors such as temperature, pressure, fluid flow, etc. Deposit models for both deposit types indicate that these systems form as a result of very similar processes. For example, most MVT deposits tend to form in platform carbonate sequences in pas-sive margin environments and are often associated with pre-existing extensional structures inboard of contractional tectonic belts (Leach et al., 2001; 2005). As such, it is anticipated that the textural and isotopic characteristics of these deposits should also be relatively consistent in order to reflect their similar genesis. The texture of dolomite host rocks in both MVT and SEDEX deposits has been extensively described and linked with specific environmental conditions such as geological environment, fluid temperature and pressure, fluid flow and Zn-Pb mineralization. For example, previous studies have shown that sparry dolomite in carbonate rocks may be produced by the generation and movement of hot basinal brines in response to arid paleoclimates and tectonism (Diehl et al., 2010). Furthermore, it has been shown that these brines served as the transport medium for metals deposited in MVT and SEDEX sys-tems (Diehl et al., 2010). In addition to textural characteristics, stable isotopic characteristics of altera-tion may also be used to infer specific environmental conditions, and furthermore can potentially be utilized as an exploration tool in discovering new mineral deposits.In order to provide a global context of carbonate-hosted Zn-Pb systems, four key MVT systems; Reocín, Spain (Early Cretaceous), Pine Point, Canada (Middle Devonian), Lennard Shelf, Canning Basin, Austra-lia (Devonian) and Nanisivik, Canada (Mesoproterozoic) have been studied. These four deposits were chosen due to their range in geographic location and age, but more importantly because they show minimal deformation. This lack of deformation, which is usually the main factor in the destruction of such deposits, has allowed for the characterization of a detailed paragenetic sequence for each of these systems. It is anticipated that the variables associated with the specific textural and isotopic sig-natures observed in each of these deposits could be used as vectoring tools and so aid in the explora-tion for similar MVT-style deposits.MethodThe paragenetic sequence, textural characteristics and isotopic signature of each of the four deposits listed above will be summarized and then finally compared to one another. It is anticipated that there will be some overlap in the controlling factors of both textural and isotopic signatures that may be used as a vectoring tool and so aid in the exploration of similar carbonate-hosted Zn-Pb deposits. 19Reocín, SpainIntroductionThe Early Cretaceous Reocín Zn-Pb deposit is hosted in the Basque-Cantabrian basin of northern Spain (Figure 2.0). This deposit is the largest stratabound carbonate-hosted Zn-Pb deposit in Spain (Velasco et al., 2003) and is one of the largest MVT deposits globally (Leach et al., 2005). Before its closure in 2003, the deposit produced roughly 62 Mt of ore grading at 8.7 % Zn and 1.0 % Pb following over 100 years of mining (Symons et al., 2009).Regional GeologyThe Iberian Peninsula exhibits regionally extensive Zn-Pb mineralization that occurs in Cretaceous car-bonate rocks in two major Mesozoic basins (Grandia et al., 2000). The major economic Zn-Pb districts are in the Basque-Cantabrian basin with more minor deposits located in the Maestrat basin in eastern Spain (Grandia et al., 2000). These basins were developed in several stages of rifting from the Triassic to the Cretaceous during the break-up of Pangea (Garci’a-Monde’jar, 1990; Salas et al., 2000) and are filled with shallow-marine carbonates (Grandia et al., 2000). In the Late Jurassic to Early Cretaceous, these two basins were further subdivided into smaller basins with the MVT mineralization in this area restricted to Aptian to Albian age (125 Ma – 100.5 Ma) rocks (Grandia et al., 2000).The sedimentary sequence (Figure 2.1) in the Basque-Cantabrian basin is up to 4000 m thick (Symons et al., 2009). The sequence sits on a deformed Paleozoic basement and contains the following: 1) Trias-sic red sandstones (Buntsandstein facies), 2) dolostones and shales (Muschelkalk facies), 3) continen-tal clays and evaporites (Keuper facies), 4) Jurassic limestone and shales, and 5) Cretaceous to Tertiary siliciclastics and carbonates (Symons et al., 2009).The Cretaceous rock sequence (Figure 2.1) begins with syn-rift fluvial sandstones and lacustrine clays of the Wealden facies (Symons et al., 2009). The Wealden facies is then overlain by a Lower Creta-ceous Urgonian platform sequence, which consists of 8 units (U1 – U8; Symons et al., 2009). The lower units from U1 – U4 are roughly 240 m thick and consist of marls, limestones and mudstones (Symons et al., 2009). These units are overlain by 100 m of limestone from unit U5 that has been pervasively dolomitized and is the main host of Zn-Pb mineralization in the basin (Symons et al., 2009). The up-per most 100 m consists of units U6 – U8 that contain sandstones, limestones and marls (Symons et al., 2009). Rifting of the North Atlantic (Bay of Biscay; Velasco et al., 2003) caused block fault tilting in the Urgonian carbonate platform, which resulted in several subaerial exposures of the platform and subsequent karstification of the carbonate rocks (Symons et al., 2009). Above the Urgonian platform are fluvial-deltaic rocks of the Supra-Urgonian facies that are up to 350 m thick and are themselves overlain by Upper Cretaceous to Tertiary carbonates, marls and sandstones (Symons et al., 2009).20Figure 2.0. Location of the Reocín Mine and regional geology of the Santillana syncline area in the Basque-Cantabrian basin of northern Spain. Modified from Symons et al. (2009).Figure 2.1. Schematic stratigraphic log through the sedimentary sequence in the Basque-Cantabrian basin. Modified from Velasco et al. (2003).SpainPortugalFranceMadridReocinBasque-Cantabrian Region Reocin mineCantabrian SeaSantander Bay43°30’N3°45’W0 10km NReocin mineComillasSantanderLa CavadaUdiasTertiaryUpper CretaceousLower CretaceousWealdenJurassicPermian-TriassicKeuperOlder PhanerozoicLower CretaceousSupra-UrgonianLower CretaceousUrgonianOre Deposits and Prospects~ 4000Thickness in mBuntsandstein faciesMuschelkalk faciesKeuperfaciesJurassicWealdenfaciesUrgonian platform sequenceU1U2U3U4U5U6U7U8Supra-UrgonianTriassic red sandstonesDolostones and shalesContinental clays and evaporitesJurassic limestones and shalesCretaceous to Tertiary siliciclastics and carbonatesMarls and limestonesLimestones (wackstones, packstones and marly limestones)Marls and marly limestonesMarly limestones and marlsLimestones with pervasive dolomitizationMain host of ore mineralizationSandstones and marlsLimestonesMarls and sandstonesSandstones21Deposit GeologyThe Basque-Cantabrian basin of northern Spain is host to the Santillana syncline with the Reocín de-posit located on the southeastern flank of this syncline (Symons et al., 2009). Cross-sectional mapping of this area indicates a stratigraphic and structural control on the deposit (Velasco et al., 2003). In the western section of the ore body, where the majority of economic Zn-Pb mineralization is located (Capa Sur), the mineralization is stratiform. In the eastern part (Barrendera), however, it is highly discordant but still stratabound (Velasco et al., 2003). Due to this split, the Reocín orebody has been subdivided laterally into three areas by Velasco et al. (2003): 1) Capa Sur in the west is mainly stratiform and ex-tends more than 2000 m with a thickness of 1 m to 5 m. It is enclosed between a set of faults at 45° and 60°. This western extent of mineralization can also be split vertically into Capa Sur at the base, Capa Norte approximately 12 m above and Tercera Capa approximately 10 m above again, 2) the central mineralized zone is both stratabound and discordant and displays characteristic geometries associated with infilling of open karstic cavities. Mineralization occurs as small lenses and pods of ore that are laterally discordant, and 3) Barrendera in the east consists of numerous small lenses of disseminated ore that formed through replacement and infilling of pre-existing vugs and pore spaces in breccias.Independent of morphology, disseminated sphalerite and/or galena that has replaced the host do-lostone and open-space filling of veins, breccias and vugs formed through dissolution of the host car-bonates can be seen in all three subdivisions (Velasco et al., 2003). All of the ore styles consist of fine grained sphalerite layers that are locally colloform and contain grains of dolomite and quartz (Velasco et al., 2003) and typical MVT textures such as: crackle breccias, ore matrix breccias, trash zone accu-mulations and ‘snow on the roof’ textures as described by Leach and Sangster (1993). Four main ore textures have been described at Reocín by Velasco et al. (2003): 1) colloform, layered sphalerite with variation in colour and grain size, 2) fine to coarse grained disseminated to massive sphalerite that replaces the host dolostone, 3) botryoidal sphalerite growths, and 4) rare, coarse grained sphalerite infilling late veins. Stages 1 and 2 are thought to have been contemporaneous with stages 3 and 4 rep-resenting later depositional stages (Velasco et al., 2003).The Reocín orebodies exhibit only slight deformation and faults are not commonly seen in the mine. There are three main fault systems as described by Velasco et al., (2003): 1) an east-west striking, sub-vertical or north-dipping fault system that is directly related to mineralization; proximal to these faults Zn and Pb grade increases as does the dolomitic alteration. These features suggest that the migration of hydrothermal fluids responsible for mineralization were controlled by this fault system, 2) a north-east-southwest striking, subvertical-dipping fault system that encloses the major part of the deposit, and 3) a third fault system that is barren of mineralization.0361-0128/ParagenesisTwo stages of dolomitization are recognized at the Reocín deposit (Figure 2.2). The first stage is the pervasive dolomitization of the primary limestone host rocks, which was controlled by faulting and de-stroyed all fossils and original textures (Velasco et al., 2003). The second stage occurred after erosion 22(Velasco et al., 2003) and was controlled by karstic cavities; it is this stage that was accompanied by ore deposition (Velasco et al., 2003). Fluid inclusion temperatures indicate temperatures of ore deposition were less than 100°C and C and O isotopic data suggest both dolomitization events occurred under similar conditions (basinal brines; Velasco et al., 2003).The paragenetic sequence of ore minerals (from most to least abundant) is relatively simple and in-cludes: sphalerite, wurtzite, galena, marcasite, pyrite with dolomite and rare calcite as the gangue minerals (Velasco et al., 2003). The paragenetic sequence consists of main stage, rhythmically depos-ited colloform sphalerite-wurtzite and sphalerite layers of variable grain size accompanied by ferroan-dolomite and a second stage that recrystallizes the main stage as rhythmically banded sphalerite and open space stalactitic growths (Velasco et al., 2003).Primary LimestonePervasive Dolomite (D1a&b)Saddle Dolomite (D2)Ferroan Dolomite (D3)Sparry, Coarse Dolomite (D4)Sphalerite 1-WurziteSphalerite 2Relative TimingMain Mineralization EventPhase Relative TimingErosionMarcasitePyriteGalenaFigure 2.2. Schematic diagram to show the paragenetic sequence of events at the Reocín Mine. As described by Velasco et al. (2003).Carbonate TexturesMany of the dolomitized rocks are easily distinguishable from the primary limestone host rocks due to their dominantly beige colour, relatively uniform grain size and presence of vugs, veins and abundant cavities (Velasco et al., 2003). The lateral extent of dolomitization was controlled by the lithology, per-meability and porosity of the differing carbonate host rocks it was overprinting; dolomite therefore only developed in highly permeable limestone, near to syn-sedimentary faults where primary porosity could be enhanced (Velasco et al., 2003).Different generations of carbonates at Reocín have been studied through optical microscopy, cath-odoluminescence and back-scattered electron microscopy by Velasco et al. (2003). These authors dis-covered four specific types of dolomite that formed during the two stages of dolomitization described above; D1 dolomite formed during the first stage of dolomitization, D2 dolomite formed during the second stage of dolomitzation and D3 and D4 dolomites formed during the waning of the second stage 23of dolomitization (Velasco et al., 2003). The four dolomites are described according to Velasco et al. (2003) in more detail below:D1 DolomiteThis dolomite can be split into two subtypes depending on the limestone host that was altered. The first is a grey dolostone (D1a) formed from a dark grey marly limestone (U4) that has retained most of the fossils and primary textures present in the original host rock. These rocks consist of medium to coarse grained dolomite crystals that are subhedral with straight and curved intercrystalline boundar-ies that form idiotopic to xenotopic mosaics (Sibley and Gregg, 1987). The dolomite crystals exhibit growth zoning defined by cloudy rhombs with clear outer rims. The second form of dolomite (D1b) is the most pervasive in the mine area and is formed from dolomitization of the ore-hosting fossiliferous limestone (U5). D1b dolomite is massive and homogeneous, coarse grained and has a characteristic beige colour with local vuggy porosity. The dolomite crystals form xenotopic to hypidotopic mosaics of anhedral to subhedral crystals with common cloudy cores and abundant allochem ghosts.D2 DolomiteThe D2 dolomite is a coarsely crystalline, white saddle dolomite formed by the replacement of porous dolostone. These zones of iron-poor dolostone may extend several metres inside the D1b dolomite.D3 DolomiteThe iron-rich, ore-bearing D3 dolomite characteristically occurs as fine to medium grained crystals that form xenomorphic to hypidiomorphic mosaics and is often associated with sphalerite. The dolomite grains lack cloudy centres, but commonly contain intergranular clays, organic matter, quartz, sphaler-ite and other detrital grains.D4 DolomiteThe final stage, D4, is a white, sparry, coarse grained dolomite similar to saddle dolomite or white sparry dolomite observed in other carbonate-hosted Zn-Pb deposits.C and O IsotopesC and O isotopes were recorded by Velasco et al. (2003) from roughly 80 representative carbonate samples from all stages of dolomite (D1 through to D4) as well as unaltered limestone. δ13C values are recorded with respect to PBD and δ18O values are recorded with respect to SMOW. In general, δ13C values range from -3.20 ‰ to 2.20 ‰ for all dolomite samples with the δ18O values showing a tighter range of 22.30 ‰  to 24.50 ‰ (Figure 2.3). The data from Velasco et al. (2003) are discussed in more detail below.24 -4.00-3.00-2.00-1.000.001.002.003.004.0020.00 22.00 24.00 26.00 28.00 30.00d13C_VPDB ‰ d180_VSMOW ‰ C and O Data for Host Rocks at the Reocín Mine, Spain UnalteredLimestonesD1a DolomitesD1b DolomitesD2 Dolomites (pre-ore)D3 Dolomites (syn-ore)D4 Dolomites(post-ore)Figure 2.3. Graph of C and O isotopic data at the Reocín Mine. Data from Velasco et al. (2003).Unaltered LimestonesThese limestones have δ13C of 2.00 ‰ to 3.50 ‰ and δ18O values of 25.20 ‰ to 28.60 ‰ and are representative of marine limestone (Veizer and Hoefs, 1976; Land, 1980) and Lower Cretaceous car-bonates in other parts of the Basque-Cantabrian basin (Velasco et al., 1994; Bustillo and Ordoñez, 1995; Fernández-Martínez and Velasco, 1995; Rosales, 1995). This range in both δ13C and δ18O values is thought to be due to changes in composition according to stratigraphic position of the limestone host rocks (Velasco et al., 2003).D1 DolomiteBoth D1a and D1b dolomite show similar isotopic compositions with δ13C values between 2.20 ‰ to -2.80 ‰ and δ18O values between 24.50 ‰ to 22.40 ‰. These values are lower than those for typical marine limestone (Velasco et al., 2003).D2, D3 and D4 DolomiteThe isotopic composition of pre-, syn- and postore dolomite (D2, D3 and D4 dolomite respectively) shows a similar range to D1 dolomite. However, there is a general trend to more depleted δ13C values with increasing iron content. This trend has been attributed to the incorporation of isotopically light CO2 from late hydrothermal fluids carrying base metals (Velasco et al., 2003). δ18O values for these do-lomites are also depleted and are thought to represent the equilibration of dolomite with isotopically 25light fluids or at elevated temperatures (Velasco et al., 2003).The broad range of δ13C values compared to the tighter range of δ18O values is thought to be due to the mixing of basinal pore fluids with fluids containing hydrocarbons (Velasco et al., 2003). The wider range of δ13C values may also be due to the presence of different fluids or carbonate species during the dissolution of the carbonate host rock or from changes in pH, temperature and/or oxygen fugacity (Velasco et al., 2003). However, the relatively narrow range of δ18O values suggest that all dolomite phases formed under similar pH, temperature and oxygen fugacity conditions (Velasco et al., 2003).SummaryFrom the results of Velasco et al. (2003) and other authors who have worked on Reocín, it has been shown that the deposit formed when base metal-rich brines migrated into a well-developed hydro-thermal karst network. The main sulphides that precipitated were sphalerite and galena, associated with iron-rich dolomite (D3), followed by a white saddle dolomite and marcasite. Mineral textures, fluid inclusion studies and C and O isotope data all point towards ore deposition due to fluid mixing of a deep, metalliferous, basinal brine and reduced sulphur under supersaturated conditions.Two main phases of dolomitization have been described by Velasco et al. (2003); a primary, pervasive phase (D1) that was controlled by faulting and limestone permeability and a secondary phase associ-ated with sulphide deposition (D2 through to D4). Both phases of dolomitization were strongly con-trolled by fault zones and host rock permeability, which provided the pathways for ascending metal-rich, basinal fluids. As such, mineralization mainly follows permeable or karstified beds acting as traps for mineralizing fluids and potentially hydrocarbons (Velasco et al., 2003).Pine Point, CanadaIntroductionThe Middle Devonian Pine Point region is an extensive mining camp hosted in the Devonian Presqu’ile barrier reef that included roughly 40 mineral deposits and produced over 58 Mt of Zn-Pb ore (Gleeson and Turner, 2007). These Middle Devonian carbonate rocks crop out south of the Great Slave Lake and extend west into the Western Canada Sedimentary Basin in northeastern British Columbia (BC; Figure 2.4).  Regional GeologyThe Precambrian basement (Figure 2.5) in the Pine Point region comprises granites, gneisses and quartzites (Norris, 1965).  An unconformity separates these basement rocks from an overlying pack-age of Ordovician or older siliciclastic, carbonate and clastic rocks, which are themselves unconform-ably overlain by Lower to Middle Devonian carbonates and evaporites (Gleeson and Turner, 2007). This Lower to Middle Devonian package of rocks is overlain by the Middle Devonian Presqu’ile barrier, 26Figure 2.4. Location of the Pine Point region and regional geology of the Presqu’ile barrier and Western Canada Sedi-mentary Basin. Modified from Qing and Mountjoy, (1994; a.) Deposit geology of the Pine Point region. Modified from Rhodes et al. (1984; b.).CORDILLERAN OROGENEROSIONAL EDGEPeace RiverArchGreat SlaveLakeMcDonald FaultBritishColumbiaPine PointLandNearshorePresqu’ile barrierOpen-marine platformRestricted intra-platformbasin-semirestricted platformSub-basins with reef swarmsSaskatchewanNorthwestTerritoriesElk Point Basin0 400kmN110°00’ W60°00’ Na. HighwayGreat Slave Lake0 10kmNWatt Mountain FormationWindy Point FormationMuskeg FormationSulphur Point FormationPine Point FormationKeg River FormationDawson LandingSulphur PointPresqu’ile PointPresqu’ile Dolomite and Zn-Pb Mineralizationb. 114°00’ W61°00’ N27Figure 2.5. Schematic stratigraphic log through the sedimentary sequence in the Presqu’ile barrier. Modified from Qing and Mountjoy (1994).PrecambrianbasementOrdovicianLower DeovianMiddle DevonianSulphur PointFormationWatt MountainFormationPresqu’ile barrierMuskegFmWindy Pt FmGranites, gneisses and quartzitesSiliciclastic, carbonate and clastic rocksCarbonates and evaporitesFossiliferous wackestones and packstonesKeg RiverFormationPine PointFormation Packstones, oatstones and medium grained dolomiteDolomitized packstones, grainstones and oatstonesMain host of ore mineralizationWindy Pt: Marine shales and mudstonesMuskeg Pt: Interbedded anhydrite and ne grained dolomiteSilty, pyritic shales, limestones and dolostonessub-Watt Mountain unconformity28which is host to the Pine Point ore deposits (Gleeson and Turner, 2007).The Middle Devonian Presqu’ile barrier is a 400 km long carbonate reef complex that extends from the Northwest Territories into northeastern British Columbia (Qing, 1998). The development of this barrier reef complex restricted seawater circulation in the adjacent Elk Point basin during the Middle Devonian and as such led to the deposition of evaporites and carbonates south of the barrier (Elk Point evaporate basin), and the deposition of typical marine sediments north of the barrier (Macken-zie shale basin; Qing, 1998). A thick package of shale overlies the Presqu’ile barrier and acts as a cap, which aided in MVT-style mineralization (Gleeson and Turner, 2007). The Pine Point region is located in the eastern section of this barrier, where a range of dolomite rocks host the mineralization (Qing, 1998). A major basement fault that can be traced from the Canadian Shield into the subsurface of the Western Canada Sedimentary Basin; the McDonald Fault, is thought to have been the main conduit for fluids that caused dolomitization and mineralization in this area (Qing, 1998).Deposit GeologySkall (1975) and Rhodes et al. (1984) divided the Presqu’ile barrier in the Pine Point region into four distinct stratigraphic units (Figure 2.5): 1) a basinal, regionally extensive shallow water carbonate plat-form (the Keg River Formation). This formation comprises fossiliferous wackestones and packstones, 2) the Presqu’ile carbonate barrier, which can be further subdivided into the Pine Point Formation and the overlying Sulphur Point Formation. The Pine Point Formation comprises shallow subtidal facies of packstones and floatstones followed by medium grained dolomite. The Sulphur Point Formation com-prises reef and shallow subtidal packstones, grainstones and floatstones that have been pervasively dolomitized to form the coarse grained Presqu’ile dolomite, 3) a back-barrier evaporite sequence (the Muskeg Formation) that consists of interbedded anhydrite and fine grained dolomites, and 4) a fore-barrier sequence of marine shales and mudstones (the Buffalo River and Windy Point formations; Rhodes et al., 1984). Following the deposition of these four units, the Presqu’ile barrier was subaeri-ally exposed, which resulted in the formation of the sub-Watt Mountain unconformity. This unconfor-mity is overlain by the Watt Mountain Formation, a series of green, silty, pyritic shales, limestones and dolostones (Rhodes et al., 1984).The Presqu’ile dolomite in the Sulphur Point Formation is the host to most of the mineralization in the Pine Point region. The dolomite consists of coarsely crystalline and saddle dolomites that host galena, sphalerite, marcasite and pyrite mineralization (Adams et al., 2000). The orebodies themselves are restricted to paleokarst networks that are primarily located subparallel to the McDonald Fault (Adams et al., 2000). ParagenesisIn the Pine Point region there are three stages of dolomitzation (Figure 2.6) that have been recognized by Qing and Mountjoy (1990); dolomite is observed pre-, syn- and post-mineralization (Gleeson and Turner, 2007). The first stage of dolomitization is a fine grained dolomite found in the back-barrier 29Muskeg Formation, the second stage of dolomitization is a medium grained dolomite found in the lower Pine Point Formation, and the third stage of dolomitization is a coarse grained dolomite found in the Sulphur Point Formation (Qing and Mountjoy, 1990). This coarse grained dolomite also extends into the overlying sub-Watt Mountain unconformity and the Watt Mountain Formation (Qing and Mountjoy, 1990). The third stage of dolomitization also corresponds to the formation of the Presqu’ile dolomite, which hosts the majority of MVT-style mineralization in this region and is comprised of both saddle and non-saddle coarse grained dolomite (Gleeson and Turner, 2007). Post-mineralization at Pine Point, there is further occurrence of saddle and non-saddle coarse dolomite as well as late blocky calcite and bitumen (Gleeson and Turner, 2007).Fine Grained DolomiteMedium Grained DolomiteCoarse Grained DolomiteSaddle Dolomite 1Non-Saddle Dolomite 1Saddle Dolomite 2Non-Saddle Dolomite 2Relative TimingMain Mineralization EventPhase Relative TimingLate Blocky CalciteBitumenSulphidesFigure 2.6. Schematic diagram to show the paragenetic sequence of events at the Pine Point Mine. As described by Qing and Mountjoy (1990) and Gleeson and Turner (2007).Carbonate TexturesIn the Pine Point region there are well preserved occurrences of diagenetic features in the carbon-ate rocks of the Presqu’ile barrier (Qing 1998). These features are thought to represent submarine, subaerial and subsurface environments (Qing, 1998). Submarine diagenetic features include: micrite envelopes, micrite, syntaxial cement, microspar cement, fibrous cement, and fine grained dolomite (Qing, 1998). Subaerial diagenetic features include localized pendant cements and dissolution and brecciation (Qing, 1998). Subsurface diagenetic features display the most significant alteration and in-clude blocky, sparry, coarse and saddle dolomites (Qing, 1998). As mentioned above, these subsurface features have been split into three stages of dolomitization by Qing and Mountjoy (1990). Fine DolomiteThe fine grained dolomite predates mineralization and is found in the Muskeg Formation (Gleeson and Turner, 2007). This dolomite is composed of subhedral to euhedral crystals and is interbedded with anhydrite of the evaporitic Elk Point basin (Qing, 1998). Fossils and primary sedimentary structures are often well-preserved with fractures and vugs being infilled by anhydrite (Qing, 1998).30Medium DolomiteThe medium grained dolomite also predates mineralization (Gleeson and Turner, 2007). This stage of dolomitization is found in the lower Pine Point Formation and is the most abundant phase within the region (Qing, 1998). The medium dolomite is composed of anhedral to subhedral crystals with well-defined grain boundaries (Qing, 1998). Fossils and sedimentary textures are generally recognizable, but are not as clear as in the fine grained dolomite, and fractures and vugs are filled by late-stage saddle dolomite (Qing, 1998).Coarse DolomiteThe coarse grained dolomite is associated with mineralization and comprises the Presqu’ile dolomite (Gleeson and Turner, 2007). This dolomite phase may be coarse saddle or non-saddle dolomite that tends to have anhedral to subhedral grains with indistinct grain boundaries; this dolomite commonly displays undulatory extinction (Qing, 1998).C and O IsotopesQing (1998) carried out C and O isotope analyses on the fine, medium and coarse grained dolomite phases in the Pine Point region. δ13C values are recorded with respect to PBD and δ18O values are recorded with respect to SMOW. This author’s work gave δ13C values of 0.27 ‰ to 1.70 ‰ and δ18O values of 29.27 ‰ to 27.00 ‰ for the fine grained dolomite, δ13C values of 0.60 ‰ to 2.50 ‰ and δ18O values of 27.11 ‰ to 21.23 ‰ for the medium grained dolomite and δ13C values of 1.80 ‰ to 1.90 ‰ and δ18O values of 27.00 ‰ to 26.59 ‰ for the coarse grained, mineralized dolomite (Figure 2.7). The gradual depletion in δ18O values has been attributed to increased fluid:rock interaction along the Presqu’ile barrier (Adams et al., 2000). 0.000.200.400.600.801.001.201.401.601.802.0020.00 22.00 24.00 26.00 28.00 30.00d13C_VPDB ‰ d180_VSMOW ‰ C and O Data for Host Rocks at the Pine Point Mine, Canada Fine DolomitesMediumDolomitesCoarse Dolomites(mineralized)Figure 2.7. Graph of C and O isotopic data at the Pine Point Mine. Data from Qing (1998).31SummaryFrom the work of several authors it has been shown that MVT-style mineralization in the Pine Point region is concentrated within the Presqu’ile dolomite, a coarse grained saddle or non-saddle dolo-mite from the Sulphur Point Formation (Qing and Mountjoy, 1990). Two further dolomite phases have been described by Qing and Mountjoy (1990); a fine grained and a medium grained dolomite, which are both pre-mineralization. Post-mineralization phases include further saddle and non-saddle coarse grained dolomite, blocky calcite and bitumen (Gleeson and Turner, 2007).It is thought that the McDonald Fault acted as a conduit for dolomitizing and mineralizing fluids to form the Pine Point deposits with the mixing of two brines (a warm barrier fluid and a cooler supersa-line brine) being the trigger for sulphide mineralization (Adams et al., 2000). Furthermore, the gradual depletion in δ18O values has been attributed to increased fluid:rock interaction along the Presqu’ile barrier (Adams et al., 2000).Lennard Shelf, Canning Basin, AustraliaIntroductionThe Devonian Lennard Shelf forms the northern margin of the Canning Basin where it overlaps Pre-cambrian basement of the Kimberley Block in northwest Australia (Figure 2.8). Mineral exploration in the area has led to the discovery of many MVT deposits hosted in the Devonian reef complexes exposed along the shelf (Wallace et al., 2002). Four Zn-Pb deposits have been economic in the eastern part of the Lennard Shelf; Cadjebut, Goongewa, Pillara and Kapok. Furthermore, in the south of the basin, a deep Zn-Pb deposit known as Admiral Bay has been discovered, but is yet to be economic due to its depth (Wallace et al., 2002). Regional GeologyThe formation of the Canning Basin was initiated in the Ordovician and by the Early Devonian, uplifting resulted in the formation of several discrete sub-basins (Arne et al., 1989). This was followed by sedi-mentation, which was concentrated in the northern part of the basin, from the Middle Devonian to the Early Carboniferous (Arne et al., 1989). A thick sequence accumulated in the Fitzroy Trough whilst Devonian reef complexes formed along the adjacent Lennard Shelf (Figure 2.9; Playford, 1980). The Lennard Shelf is separated from the Fitzroy Trough by a series of south-dipping northwest-southeast trending listric normal faults (Wallace et al., 2002).  A major period of uplift and exposure along the Lennard Shelf was followed by sedimentation of the lowermost Grant Group during the Late Carbonif-erous (Figure 2.9; Arne et al., 1989). Sedimentation continued throughout the basin during the Perm-ian with minor periods of uplift and erosion. The last tectonic event to affect the Lennard Shelf region of the Canning Basin involved uplift and erosion during the Early Jurassic; the Fitzroy Movement (Arne et al., 1989). As such, Jurassic and Cretaceous sediments are only preserved in the southern portion of 32Figure 2.8. Location of the Lennard Shelf and regional geology of the Canning Basin region, northwest Australia. Modified from Arne et al. (1989), originally after Playford (1981).Figure 2.9. Schematic stratigraphic log through the sedimentary sequence of the Lennard Shelf. Modified from Nicoll et al. (1993) and Tompkins et al. (1994).Post DevonianDevonian-CarboniferousConglomeratesPrecambrianKimberley Landmass300kmNCanning Basin18°00’S125°00’EFitzroy TroughPinnacle FaultLennard ShelfNapier RangePillara RangeEmanuel RangePillaraGoonwegaCadjebut/KapokAdmiral BayKimberley BlockFrasnianPlatform FaciesFamennian Platform FaciesFamennian-FrasnianMarginal Slope and Basin FaciesCanningBasinShale and siltstonesAbundantly fossiliferous shales and siltstonesShale and siltstones with limestone nodulesDolomitic siltstonesSiltstone and dolomitic siltstonesLimestone and vuggy dolomiteLimestoneDolomitic limestonesMain host of Cadjebut ore mineralizationPrecambrianMarbooFormationOrdovicianTremodocianEmanuelMarbooLowermemberMiddlememberUppermemberGap CreekLowermemberUpper memberDevonian FamennianFrasnianPillaraNullaraPillaraFormationNullaraFormationCarboniferousAndersonGrantAndersonFormationGrantFormation Dolomitic limestones33the basin. Finally, intrusion of leucite lamprophyre dykes took place in the Miocene (Arne et al., 1989).The Devonian reef complex along the Lennard Shelf is roughly 50 km wide by 350 km long. Two ma-jor depositional cycles separated by an unconformity have been recognized in this area by Playford (1984); the Frasnian Pillara cycle and the Famennian Nullara cycle (Figure 2.9). The Pillara cycle de-veloped above Ordovician and Precambrian basement rocks and is characterized by low-relief up-right reef deposits (limestones and vuggy dolomites; Playford, 1984). The Nullara cycle developed above the Frasnian platform created by the Pillara cycle and is characterized by advancing reef-margin platforms (limestones; Playford, 1984). The Famennian rocks of the Nullara cycle have been largely removed through erosion, but significant thicknesses of these platforms are exposed in the Napier Ranges (Tompkins et al., 1994).The structural history of the Canning Basin has essentially produced two major subsidence and sub-sequent burial events (Late Devonian to Early Carboniferous and Permian to Mesozoic) and two uplift and subsequent erosional events (Middle to Late Carboniferous and Late Mesozoic to Tertiary) on the Lennard Shelf (Wallace et al., 2002).Deposit GeologyThe Lennard Shelf carbonates range in age from Givetian to Famennian with the best developed reef complexes being Frasnian to Famennian in age (Wallace et al., 2002). In most areas, the Lennard Shelf is underlain by Precambrian metamorphic basement rocks, however in the eastern section of the shelf the Ordovician Emanuel and Gap Creek formations can be found beneath the Devonian sequence (Wallace et al., 2002).The eastern Lennard Shelf has been the most prominent in terms of Zn-Pb mineralization, hosting the most economic deposits to date with faults and fractures being the host of many of the MVT deposits in this region (e.g., the Pillara deposit; Wallace et al., 2002). By contrast, the Cadjebut deposit (the largest resource currently identified in the area) is characteristically stratabound due to a stratigraphic and evaporitic control on sulphide mineralization (Wallace et al., 2002). The main orebody at Cadjebut is banded, whereas the less significant orebody is hosted in a breccia (Wallace et al., 2002). Generally, Zn-Pb mineralization in this region occurs at the boundary between dolomitized and undolomitized carbonates and so suggests a structural and lithological control on permeability as the main factor influencing mineralization (Arne et al., 1989). Compactional flow from the Fitzroy Trough is thought to be the driving force of metalliferous brines and hydrocarbons into the Lennard Shelf (Wallace et al., 2002). Sulphides are observed depositing as cavity-filling cements in large hydrothermal karst systems (Wallace et al., 2002). Fluid inclusion studies conducted in this region give a range of ore-forming tem-peratures from 45°C to 115°C and suggest a very saline Na-Ca-Cl bearing brine (Lambert et al., 1984; Etminan and Lambert, 1986).34ParagenesisMcManus and Wallace (1992) suggested due to carbonate cement stratigraphic relationships and host rock primary porosity that there was a single, widespread event of sulphide precipitation during early burial and diagenesis of the Lennard Shelf in the Late Devonian to Early Carboniferous. Sulphides de-posited included: pyrite, sphalerite, galena and marcasite (Wallace et al., 2002). This mineralization age has since been confirmed by Rb-Sr dating of sphalerite and U-Pb dating of ore-stage calcite (Chris-tensen et al., 1995; Brannon et al., 1996).For the purpose of dolomite parageneis, this study will concentrate on the well-researched and most significant deposit to date; the Cadjebut deposit (Figure 2.10). The earliest dolomitization event within the Cadjebut region is referred to as dolomite I (Tompkins et al., 1994). This dolomite comprises 40 % of the total dolomite observed within the area and is characterized by dolomite cement overgrowths of carbonate strata (Tompkins et al., 1994). Dolomite II forms 55 % of the total dolomite observed within the Cadjebut area and pervasively replaces carbonate rocks of the Emanuel Range platform. It is characterized by a vuggy dolomite appearance (Tompkins et al., 1994). Dolomite III is a third phase of replacement dolomite, and is characterized by turbid cores and less turbid rims (Tompkins et al., 1994). Dolomite III along with dolomite II saddle cements are observed along the Cadjebut fault in the area of the mine (Tompkins et al., 1994). It is this saddle dolomite cement and accompanying calcite cements that are closely associated with sulphide deposition (Tompkins et al., 1994). The association of ore deposition with the late-stage calcite cements suggests that mineralization postdates dolomiti-zation and as such must be constrained to later stages of platform burial in the middle to late Carbon-iferous (Tompkins et al., 1994). Dolomite IDolomite IIDolomite IIIDolomite II CementDolomite III CementSulphidesRelative TimingMain Mineralization EventPhase Relative TimingFigure 2.10. Schematic diagram to show the paragenetic sequence of events at the Cadjebut deposit, Lennard Shelf, Can-ning Basin. As described by Tompkins et al. (1994).Carbonate TexturesIn general, potential carbonate host rocks for Zn-Pb mineralization on the Lennard Shelf are described as coarsely crystalline dolomites that comprise roughly 10 % of the Devonian carbonates in this region (Wallace et al., 2002). Specifically in the Cadjebut deposit, dolomite I differs from dolomites II and III as it preserves textural details of the previous host rocks’ matrix and grains (Tompkins et al., 1994). By contrast, dolomites II and III destroy all original textures (Tompkins et al., 1994). Dolomite II consists of 35fine to medium grained anhedral to subhedral crystals that are turbid. Dolomite III consists of turbid cores with less turbid rims and has a more ferroan composition (Tompkins et al., 1994). As mentioned above, it is the saddle dolomite cements that are associated with mineralization (Tomp-kins et al., 1994). Dolomite cements II and III overgrow the sulphides and are then themselves over-grown by late-stage calcite cements (Tompkins et al., 1994). There are two dominant calcite cements at the Cadjebut deposit. The first is a polygranular, cloudy calcite and the second is a clear, blocky cement (Tompkins et al., 1994). The second, clear, blocky calcite destroys all previous textures and all secondary porosity (Tompkins et al., 1994).C and O IsotopesAs with dolomite paragensis and carbonate textures, this study will focus on the most researched Cad-jebut deposit for C and O isotopic signatures (Figure 2.11). Tompkins et al. (1994) conducted a stable isotope study of several stages of dolomitization and cementation at the Cadjebut deposit, the results of which are discussed further below. δ13C values are recorded with respect to PBD and δ18O values are recorded with respect to SMOW. 0.001.002.003.004.005.006.0020.00 25.00 30.00 35.00d13C_VPDB ‰ d180_VSMOW ‰ C and O Data for Host Rocks at the Cadjebut Deposit, Lennard Shelf Region, Canning Basin, Australia Dolomite IDolomite IIDolomite IICement(mineralized)Dolomite IIIDolomite IIICement(mineralized)Figure 2.11. Graph of C and O isotopic data at the Cadjebut deposit, Lennard Shelf, Canning Basin. Data from Tompkins et al. (1994).The mean δ18O values for dolomite I are significantly heavier than those for dolomites II and III, where-as the δ13C values are slightly lighter; dolomite I has a δ13C range of 0.10 ‰ to 1.90 ‰ and a δ18O range of 29.79 ‰ to 32.26 ‰. Dolomite II has a δ13C range of 1.00 ‰ to 3.50 ‰ and a δ18O range of 29.37 ‰ to 26.02 ‰ with dolomite II cement showing a δ13C value of 3.00 ‰ and a δ18O value of 27.57 ‰. 36Dolomite III has a δ13C range of 1.00 ‰ to 5.00 ‰ and a δ18O range of 27.31 ‰ to 24.73 ‰ with do-lomite III cement showing a δ13C range of 0.50 ‰ to 5.00 ‰ and a δ18O range of 27.83 ‰ to 24.73 ‰ (Tompkins et al., 1994).The ferroan nature of dolomite III and dolomite III cement and the lighter δ18O signatures compared to dolomites I and II are thought to indicate that the source of dolomitizing fluids was not well mixed, potentially pointing towards subsurface fluids derived from different stratigraphic horizons (Tompkins et al., 1994). Furthermore, the lighter δ18O signature of dolomite III and dolomite III cement suggests a slightly elevated temperature of formation (Tompkins et al., 1994). This may be explained by a great-er burial depth of formation, or formation from isotopically lighter fluids than for dolomites I and II (Tompkins et al., 1994).SummaryFrom the work of several authors in the Lennard Shelf region, is has been shown that many MVT sys-tems have been discovered along the Devonian reef complexes in this area. In particular, four main Zn-Pb deposits have been discovered; Cadjebut, Pillara, Goongewa and Kapok with Cadjebut being the most economically significant to date.According to stratigraphic relationships and host rock primary porosity studied by McManus and Wal-lace (1992), there has been one major mineralization event within the Lennard Shelf. The timing of which has been confirmed as Late Devonian to Early Carboniferous by Rb-Sr and U-Pb dating by Chris-tensen et al. (1995) and Brannon et al. (1996).Within the Cadjebut Mine area, mineralization is accompanied by dolomite III cements that are char-acterized by a saddle dolomite texture and late-stage calcite cements (Tompkins et al., 1994). Dolomite III is pre-dated by dolomites I and II that are not associated with mineralization. Dolomite I preserves original textures, whereas dolomites II and III destroy all primary features (Tompkins et al., 1994). In general, however, potential carbonate host rocks for Zn-Pb mineralization on the Lennard Shelf are described as coarsely crystalline dolomites that comprise roughly 10 % of the Devonian carbonates in this region (Wallace et al., 2002). As with other MVT systems, dolomites associated with mineralization tend to show depleted δ18O signatures, which in this case is thought to be due to higher temperatures of formation, or formation from isotopically lighter fluids (Tompkins et al., 1994).Nanisivik, CanadaIntroductionThe Mesoproterozoic Nanisivik Mine is located on the Borden Peninsula in northern Baffin Island (Fig-ure 2.12). Mineralization in this region is related to the 300 km long, northwest trending Milne Inlet graben. Nanisivik was a major carbonate-hosted Zn-Pb deposit that was mined from 1976 to 2002 producing 19 Mt grading at roughly 10 % Zn and Pb (Turner, 2011). 37Figure 2.12. Location of the Nanisivik Mine and regional geology of the Borden basin and Milne Inlet region, northern Baffin Island. Modified from Arne et al. (1991).Figure 2.13. Schematic stratigraphic log through the sedimentary se-quence in the Bylot Supergroup of the Borden basin. Modified from Sherlock et al. (2004).ArcheanMary River GroupMiddle - Upper ProterozoicBylot SupergroupPaleozoicAdmirality and Brodeur groups1000kmN72°00’N73°00’N84°00’WNanisivikBorden Peninsula Milne Inlet78°00’WMilne Inlet grabenBaffin IslandNanisivikUndierentiated gneisses and plutonic rocksBasaltic ow-sills and siliciclastic sedimentary rocksQuartz areniteSiltstones and shalesDolomitic mudstones and boundstonesMain host of ore mineralizationShales and silty mudstonesSubarkose to shalesDolomitic sandstone, dolomite and dolostonesDolomitic limestonesArcheanMary RiverGroupMesoproterozoicBylot SupergroupEqalulik GroupMary RiverNauyatAdams SoundArctic BayUluksan GroupSociety ClisVictor BayNunatsiaq GroupStrathconaSoundPaleozoicAdmiralityBrodeurCambrianOrdovicianSilurian38Regional GeologyThe Nanisivik Mine is hosted in Mesoproterozoic sedimentary rocks from the Bylot Supergroup, which unconformably overlies an Archean gneissic basement (Sherlock et al., 2004). Rifting of this basement during the late Mesoproterozoic created a series of extensional rift basins in the eastern Arctic; the Borden, Fury and Hecla, Ashton and Hunting, and Thule basins that are thought to represent erosional remnants of a previously larger basin (Fahrig et al., 1981; Jackson and Iannelli, 1981). The Bylot Super-group (Figure 2.13), which is where most of the mineralization in this region is found, was deposited within the Borden basin. The basin comprises three troughs with the Milne Inlet graben being both the largest and the only one to contain sulphide mineralization (Turner, 2011).The Bylot Supergroup can be split into lower volcanic and siliciclastic rocks (the Eqalulik Group), a carbonate-dominated group (the Uluksan Group) and an upper siliciclastic and carbonate-dominated group (the Nunatsiaq Group; Jackson and Iannelli, 1981). Additionally, the Supergroup is crosscut by diabase dykes from the extensive dyke swarm of the Franklin igneous event (Fahrig et al., 1971). Paleozoic rocks of the Arctic platform overlie the Bylot Supergroup and comprise the Cambrian to Ordovician Admiralty Group and the Ordovician to Silurian Brodeur Group (Sherlock et al., 2004).In the Nanisivik region, the Mesoproterozoic host rocks have been gently deformed into large-scale open folds (interlimb angles of 160° to 170°) with north-trending axial traces. However, the overlying Paleozoic rocks are flat-lying and undeformed (Sherlock et al., 2004). Deposit GeologyIn the Nanisivik region the Eqalulik, Uluksan and Nunatsiaq groups can be further subdivided (Figure 2.13); the Eqalulik Group is split into the Adams Sound and Arctic Bay formations, the Uluksan Group is split into the Society Cliffs and Victor Bay formations and the Nunatsiaq Group is split into the Strathco-na Sound Formation and the Nunatsiaq Group (Sherlock et al., 2004). Zn-Pb mineralization at Nanisivik is hosted in the upper Proterozoic dolostones of the Society Cliffs Formation, which is locally capped by shales of the Victor Bay Formation (Uluksan Group; Arne et al., 1991). The sulphide ore bodies contain abundant pyrite, sphalerite, galena and pyrite pseudomorphs after marcasite with sparry dolomite be-ing the main gangue mineral creating a well-banded ore texture thought to be due to the progressive replacement of the carbonate host rock (Arne et al., 1991). There are two major ore bodies at Nanisivik; the Main orebody and the Area 14 orebody.  Other sulphide bodies consist of barren pyrite (Arne et al., 1991). The Main orebody is composed predomi-nately of pyrite, but contains economic quantities of sphalerite and galena. Sparry dolomite and late quartz crystals lining open vugs are the two main gangue minerals (Arne et al., 1991). Textural and mineralogical variations enable the subdivision of the Main orebody into four horizontal ore zones; the western, central, eastern and shale zones. Furthermore, the upper lens of the orebody may be split into six texturally distinct, laterally extensive mine units (Arne et al., 1991).The Main orebody is a horizontal, elongate mass approximately 3 km long, 130 to 200 m wide and 10 39to 20 m thick (Turner, 2011). The upper surface of this orebody is very flat with its top being roughly 100 m below the contact with the overlying Victor Bay Formation shale; this unconformity is thought to have acted as a stratigraphic trap for both gas and mineralizing fluid (Turner, 2011). It has been sug-gested that local proximity to the unconformity, stratigraphic level relative to the unconformity, prox-imity to underlying shale and allochthonous material and proximity to faulting are the main controls on mineralization at Nanisivik rather than the composition of the primary host rock (Turner, 2011). The flat and extensive nature of the orebody is taken as evidence that sulphide mineralization occurred after the Mesoproterozoic host rocks had undergone gentle deformation into large-scale open folds (Sherlock et al., 2004). Fluid inclusion studies at Nanisivik indicate that mineralization took place at a temperature range of 165°C to 210°C from a basinal brine containing 20 to 30 wt. % NaCl (Arne et al., 1991). ParagenesisThe age of the orebody at Nanisivik is controversial, with results ranging from Mesoproterozoic to Or-dovician (Turner, 2011). As a result of this, the timing of regional mineralization remains unclear. How-ever, textural and geometric relationships between mineralized and unmineralized host rocks provide constraints on the relative timing of sulphide deposition (Turner, 2011). Numerous sulphide showings are related to the dykes of the Franklin igneous event and so suggest that these fracture systems acted as fluid pathways (Turner, 2011); however, it remains unclear whether the Nanisivik region records a single mineralizing event, or whether the history of the Borden basin is more complex (Turner, 2011).Within the Society Cliffs Formation, the sulphides clearly crosscut stratigraphy. As mentioned previ-ously, this suggests that sulphide deposition occurred after folding and tilting of the host rocks (Sher-lock et al., 2004). In terms of sulphide mineralization, the general paragenetic sequence (Figure 2.14) for individual sulphide bands is early sphalerite, followed by pyrite and late galena with sparry dolo-mite gangue (Arne et al., 1991). In addition to a sulphide paragenetic sequence, the carbonate phases may be split (Figure 2.14) into the early Society Cliffs Formation dolostones followed by an ore stage sparry dolomite interbanded with sulphide mineralization and finally a late-stage sparry dolomite in-filling vugs (Ghazban et al., 1991).Figure 2.14. Schematic diagram to show the paragenetic sequence of events at the Nanisivik Mine. As described by Turner (2011).Society Clis DolostoneSparry DolomiteLate CarbonateSphaleritePyriteGalenaRelative TimingMain Mineralization EventPhase Relative Timing40Carbonate TexturesAs mentioned above the Main orebody has been divided into a western, a central and an eastern zone, as well as a shale zone. Furthermore, a series of six lateral units have been defined based on ore texture and mineralogy in the eastern and central upper lens (Arne et al., 1991). Units 1, 3 and 5 are similar and are characterized by horizontally banded ore and sparry dolomite (Arne et al., 1991). Unit 2 is comprised of well banded ore that is commonly inclined at a high angle compared to units 1, 3 and 5 (Arne et al., 1991). Unit 4 consists of poorly banded ore in which the sulphide layers have been disturbed to give the ore a ‘shredded’ appearance (Arne et al., 1991). Finally, in unit 6 the ore has been further disturbed to give a chaotic appearance (Arne et al., 1991).In terms of carbonates, there is not an extensive range in the textures observed at Nanisivik, with the main textural unit being a sparry dolomite associated with mineralization. Units 1, 3 and 5 show evidence of dolostone host rock replacement characterized by the progressive recrystallization of do-lomite and by the precipitation of fine grained pyrite (Arne et al., 1991). Units 1, 3 and 5 also exhibit coarsely banded pyrite-sparry dolomite-sphalerite layers with sparry dolomite comprising the main gangue mineral in these layers (Arne et al., 1991). The repetitively banded ore more characteristic of units 2 and 4 are composed of sphalerite-pyrite or sphalerite-sparry dolomite layers, again with sparry dolomite comprising the main gangue mineral (Arne et al., 1991). By contrast, unit 6 exhibits a core of white, sparry dolomite surrounded by dark sphalerite and then sphalerite intergrown with pyrite (Arne et al., 1991). The presence of sparry dolomite interbanded with pyrite is thought to indicate that the ore forming fluid was around a pH of less than 5 as carbonate minerals are extremely pH sensitive (Arne et al., 1991)C and O IsotopesGhazban et al. (1991) recorded C and O isotope data from a range of carbonate rocks at Nanisivik. These authors’ work analyzed samples of white sparry dolomite associated with sulphides, late car-bonate infilling vugs and host dolostone (Figure 2.15). δ13C values are recorded with respect to PBD and δ18O values are recorded with respect to SMOW.Host Dolostoneδ13C and δ18O values for the host dolostone samples range from 2.70 ‰ to 3.00 ‰ and 28.86 ‰ to 28.34 ‰ respectively which, as expected, sit in the range of Proterozoic marine carbonates (Ghazban et al., 1991).White Sparry Dolomite Associated with Sulphidesδ13C and δ18O values for the white sparry dolomite associated with sulphides samples range from 2.00 ‰ to -6.50 ‰ and 23.29 ‰ to 15.04 ‰ respectively (Ghazban et al., 1991). Two individual bands of 41sparry dolomite were analyzed to show that despite their centimetre thickness there is significant variation across the mineralized bands; more depleted δ13C values tend to be associated with more enriched δ18O values (Ghazban et al., 1991). This variation is thought to be due to local reduction of sulphate by hydrocarbons in the mineralizing fluid. These hydrocarbons may have potentially migrated into the ore-depositional site along with the mineralizing fluid from the Arctic shale beneath the Soci-ety Cliffs Formation (Ghazban et al., 1991).Late Carbonateδ13C and δ18O values for the late carbonate samples range from -3.00 ‰ to -5.00 ‰ and 18.45 ‰ to 15.66 ‰ respectively (Ghazban et al., 1991).The general results from this study show that both the sparry dolomite and late carbonates are isoto-pically depleted compared to the host rock and as such, isotopic equilibrium between these phases was not achieved (Ghazban et al., 1991). Furthermore, the ore forming fluids must have acquired a de-pleted δ13C signature from sources other than the host rock (Ghazban et al., 1991). The wide variation in δ18O values has been attributed to changes in temperature during carbonate precipitation (Ghazban et al., 1991). -10.00-8.00-6.00-4.00-2.000.002.004.0010.00 15.00 20.00 25.00 30.00d13C_VPDB ‰ d180_VSMOW ‰ C and O Data for Host Rocks at the Nanisivik Mine, Canada Host DolostoneSparry Dolomite(mineralized)Late Calcite VeinsFigure 2.15. Graph of C and O isotopic data at the Nanisivik Mine. Data from Ghazban et al. (1991).SummaryFrom the results of previous authors who have worked on the Nanisivik deposit, it has been shown that the Main orebody can be characterized by mineralogical and textural zonation and there is abun-42dant evidence of repetitive replacement of the dolostone host rock by banded sulphide and sparry dolomite ores. The Area 14 orebody is similar to the Main orebody, but displays simpler zonation. Isotopic signatures are consistent with other MVT-style deposits in that the ore associated and late-stage carbonates are significantly depleted with regards to the host dolostone. These variations have been attributed to local reduction of sulphate by hydrocarbons in the mineralizing fluid and changes in temperature during carbonate precipitation (Ghazban et al., 1991).A final point to consider is that Zn-Pb mineralization at the Nanisivik deposit may differ from ‘typical’ carbonate-hosted Zn-Pb deposits because its Mesoproterozoic age falls outside the general Phanero-zoic age of many MVT systems (Leach et al., 2010). Zn-Pb mineralization in Precambrian rocks is not common due to the low preservation potential of such systems (Leach et al., 2010). As such, many characteristics of the Nanisivik deposit may not necessarily correspond to younger Zn-Pb systems. This is because in the Proterozoic there were no skeletal particles in the host carbonates and so primary carbonate sediment mineralogy was dominated by seawater composition leading to less permeability produced by mineral-driven diagenesis (Turner, 2011). Furthermore, early dolomitization was com-mon in Proterozoic rocks, which would have led to increased stability and so less potential for later permeability (Turner et al., 2011).DiscussionThe four MVT Zn-Pb systems considered in this study; Reocín, Spain (Early Cretaceous), Pine Point, Canada (Middle Devonian), Lennard Shelf, Canning Basin, Australia (Devonian) and Nanisivik, Canada (Mesoproterozoic) were chosen due to their range in geographic location and age, but more impor-tantly because they show minimal deformation. This lack of deformation, which is usually the main factor in the destruction of such deposits, has allowed for the characterization of a detailed parage-netic sequence for each of these systems. Because of this detailed characterization, despite the wide range in age and location of each deposit, there are some significant similarities that have been sum-marized in Table 2.0.In summary, all four deposits exhibit extensive dolomitization associated with mineralization. This do-lomitization can commonly be split into differing stages based on textural characteristics and is com-monly followed by a late-stage carbonate; late-stage calcite in Pine Point, Lennard Shelf and Nanisivik and late-stage dolomite in Reocín. The dolomite texture most commonly associated with mineraliza-tion is a saddle dolomite; however, sparry dolomite is the host to most mineralization in the Nanisivik deposit. Karstic networks are common and control mineralization, except in the case of Nanisivik. As mentioned above, this may be a function of the age of the Nanisivik deposit and so the difference in primary composition of host carbonates compared to younger, Phanerozoic deposits. Finally, there appears to be a general depletion in δ18O values from least altered through to most altered and miner-alized host rocks in all four deposits. This is however not reflected in δ13C values, with only the Reocín deposit showing depletion in δ13C and the Lennard Shelf deposits actually showing enrichment in δ13C.43Table 2.0. Comparison table of the four highlighted deposits in this study. Data has been taken from all of the authors referenced above.Deposit Production Karstic Dolomitization Stages of dolomitization Dolomite textures associated with mineralizationδ13C values for least altered host rocksδ18O values for least altered host rocksδ13C values for altered host rocksδ18O values for altered host rocksδ13C values for mineralized rocksδ18O values for mineralized rocksDepletion1. Medium to coarse grained dolomite, pervasive, destroys all primary textures2. Saddle dolomite associated with mineralization and 3. Iron-rich dolomite associated with sphalerite4. Late stage, white, sparry dolomite1. Fine grained dolomite, preserves primary textures2. Medium grained dolomite3. Coarse grained dolomite associated with mineralization4. Late stage, blocky calcite1. Fine grained dolomite, preserves primary textures2. Fine to medium grained, vuggy dolomite, destroys primary textures. Associated with a saddle dolomite cement intergrown with mineralization3. Iron-rich dolomite. Associated with a saddle dolomite cement intergrown with mineralization4. Late stage, blocky calcite1. At least one stage of sparry dolomite2. Late calciteGeneral depletion in O, but enrichment in CGeneral depletion in O29.37 to 24.73 ‰N/A1.80 to 1.90 ‰0.50 to 5.00 ‰2.00 to 6.50 ‰27.00 to 26.59 ‰27.73 to 24.73 ‰23.79 to 15.04 ‰0.50 to 5.00 ‰N/AYesYesSaddle and non-saddleSaddleSparry0.10 to 1.90 ‰2.70 to 4.00 ‰29.27 to 27.00 ‰29.79 to 32.26 ‰28.86 to 28.34 ‰24.50 to 22.40 ‰ 24.50 to 22.40 ‰General depletion in both C and O associated with increase in iron content of dolomite2.20 to -2.80 ‰Yes 0.27 to -1.70 ‰ 27.11 to 21.23 ‰ General depletion in OYes Saddle and crackle breccia 2.00 to 2.50 ‰ 25.20 to 28.60 ‰ 2.20 to -2.80 ‰0.60 to 2.50 ‰Reocin, SpainPine Point, CanadaCadjebut deposit, Lennard Shelf, Canning Basin AustraliaNanasivik, CanadaYesYesYesNo62 Mt at 8.70 wt. % Zn and 1.00 wt. % Pb52 Mt at 7.00 wt. % Zn and 3.40 wt. % Pb16 Mt at 8.90 wt. % Zn and 5.00 wt. % Pb19 Mt at 10 wt. % Zn and Pb44Saddle DolomiteThe term saddle (or baroque) dolomite refers to sparry, commonly milky-white, gray or brown dolo-mite crystals with distinctly warped faces and lattices, undulose extinction, abundant fluid inclusions, ferroan composition and intercrystalline variations in trace element abundances (Radke and Mathis, 1980). Crystals are commonly coarse, millimetre-sized with curved faces and cleavages (Flügel, 2010); these curved faces result from the enrichment of Ca towards the crystal edges leading to lattice expan-sion (Searl, 1989). This type of dolomite commonly occurs as cement in moldic or vuggy carbonates and less commonly as massive replacement of carbonate host rocks associated with hydrocarbons or base metal mineralization (in particular, MVT deposits; Radke and Mathis, 1980). Despite its relatively common occurrence, the origin of saddle dolomite is still poorly understood; however, the presence of saddle dolomite does point towards certain characteristic environmental conditions. Almost all saddle dolomite has been interpreted as having formed from brines with salinities 2 to 6 times higher than that of seawater, and at temperatures between 60°C and 150°C (Radke and Mathis, 1980). These conditions tend to point towards deep burial and hydrothermal environments or by-products of ther-mochemical sulphate reduction (Machel, 1987); however, other environments of formation have been suggested. Collins and Smith (1975) suggested that saddle dolomite forms from seawater when a change in sea level causes the strata in question to return below the ocean surface. Assereto and Folk (1980) suggested that saddle dolomite forms after an influx of freshwater into a hypersaline environ-ment. Morrow et al. (1986) suggested that saddle dolomite forms during any of the three following scenarios: 1) freshwater mixing, 2) thermal convection of brines that dissolve underlying evaporites, or 3) a regional groundwater flow system. Despite multiple authors citing evidence for saddle dolomite formation in ways other than deep burial and hydrothermal activity, thus far no conclusive evidence for a low-temperature, low-salinity origin of saddle dolomite has been proved (Machel, 1987).C and O IsotopesIn most sites δ18O is a good tracer for MVT-style mineralization as all four deposits show a depletion in δ18O values from least altered to most altered and mineralized host rocks. δ13C by contrast does not show this depletion across all four deposits, and in the case of the Cadjebut deposit in the Lennard Shelf region there is actually an enrichment in δ13C values from least altered to most altered and min-eralized host rocks (Figure 2.16).ReocínThe Reocín deposit is the only deposit of the four that shows a depletion in both δ18O values and δ13C values from the least altered to most altered host rocks (Figure 2.3). The depletion in δ13C has been attributed to the incorporation of isotopically light CO2 from late hydrothermal fluids carrying base metals (Velasco et al., 2003). The depletion in δ18O is thought to represent the equilibrium of dolo-mite with isotopically light fluids or due to elevated temperatures of formation (Velasco et al., 2003). There is also a significantly broader range of δ13C values compared to δ18O values. This broader range 45is thought to represent the presence of different fluids or carbonate species during the dissolution of carbonate host rock or from changes in pH, temperature and/or oxygen fugacity (Velasco et al., 2003). The relatively narrower range of δ18O values is thought to indicate that all dolomite phases formed un-der similar pH, temperature and oxygen fugacity conditions (Velasco et al., 2003), which would point towards the presence of different fluids or carbonate species during the dissolution of carbonate host rock as the reason for a broader δ13C signature.   -10.00-8.00-6.00-4.00-2.000.002.004.006.0010.00 15.00 20.00 25.00 30.00 35.00d13C_VPDB ‰ d180_VSMOW ‰ Evaluation of C and O Data for all Four MVT Systems Reocín, SpainPine Point, CanadaCadjebut, Canning Basin,AustraliaNanisivik, CanadaFigure 2.16. Comparison graph of C and O isotopic data from all four key MVT systems. Data from Velasco et al. (2003), Qing (1998), Tompkins et al. (1994) and Ghazban et al. (1991).Pine PointThe Pine Point deposit shows a general depletion in δ18O values from the least altered to most altered host rocks (Figure 2.7). This depletion has been attributed to increased fluid:rock interaction along the Presqu’ile barrier (Adams et al., 2000).Lennard ShelfThe Cadjebut deposit in the Lennard Shelf region shows a general depletion in δ18O values from the least altered to most altered host rocks as expected (Figure 2.11), but interestingly also shows en-richment in δ13C values. The enrichment in δ13C is not well explained; however, the depletion in δ18O values is attributed to elevated temperatures of dolomite formation caused by greater burial depths (Tompkins et al., 1994).46NanisivikThe Nanisivik deposit also shows a general depletion in δ18O values (Figure 2.15) and this has once again been attributed to changes in temperature during carbonate precipitation (Ghazban et al., 1991).In summary, δ18O looks to be a good tracer for MVT-style mineralization as it shows a general deple-tion from least altered to most altered and mineralized host rocks. This depletion has been attributed to more fluid:rock interaction within mineralized regions and so equilibrium between dolomite and isotopically lighter fluids, but more importantly due to elevated temperatures of dolomite formation caused by greater burial depths or interaction with hotter mineralizing fluids. By contract, δ13C does not show the same depletion trend consistently. Instead, δ13C values across all four deposits remain unchanged or become enriched or depleted from the least altered to most altered host rocks. The majority of hydrothermal fluids have much larger quantities of H2O compared to carbon-bearing species such as CO2 or CH4 (Barker et al., 2013). As such, oxygen isotopes will show greater degrees of isotopic alteration relative to carbon isotopes in the same rock affected by the same fluid (Barker et al., 2013). This therefore explains why δ18O is a more useful indicator of MVT-style mineralization compared to δ13C, which may be less readily altered by the mineralized fluid.In addition to the characteristic similarities highlighted in Table 2.0, there are also several structural controls that are common across all four deposits. Despite the fact that each deposit exhibits minimal deformation, they all have at least one significant fault that has been interpreted to be the pathway for mineralizing fluids to reach the host rocks. The Reocín and Pine Point deposits in particular are both associated with one major fault system (Velasco et al., 2003; Qing, 1998), the Lennard Shelf deposits, not including the Cadjebut deposit, are mainly hosted within fault and fracture networks (e.g., the Pil-lara deposit; Wallace et al., 2002) and mineralization at the Nanisivik deposit is thought to have been facilitated by the proximity to an overlying unconformity and faulting (Turner, 2011).ConclusionFour specific MVT Zn-Pb systems; Reocín, Spain (Early Cretaceous), Pine Point, Canada (Middle De-vonian), Lennard Shelf, Canning Basin, Australia (Devonian) and Nanisivik, Canada (Mesoproterozoic) were studied to give a global context of carbonate-hosted Zn-Pb systems. These four deposits were chosen due to their range in geographic location and age, but more importantly because they show minimal deformation. This lack of deformation, which is usually the main factor in the destruction of such deposits, has allowed for the characterization of a detailed paragenetic sequence for each of these systems.In environments that satisfy one or more of the 14 MVT-style characteristics listed in Chapter 1 (Leach and Sangster 1993; Leach et al., 2005) there are several, more detailed, characteristics defined through this study which may aid in narrowing down the location of MVT-style mineralization.  Some of the key features observed in all four deposits, and summarized in Table 2.0, include: 1) extensive dolomitiza-tion, which is commonly characterized by coarse grained dolomite that destroys all primary features, 472) at least one stage of dolomitization separately defined by textural characteristics, 3) evidence of karstic networks, and 4) late-stage carbonates, most commonly blocky calcite. The mineralized zones in particular, according to the four deposits studied above, will commonly exhibit saddle dolomite (sparry, commonly milky-white, gray or brown dolomite crystals with distinctly warped faces and lat-tices, undulose extinction, abundant fluid inclusions, ferroan composition and intercrystalline varia-tions in trace element abundances; Radke and Mathis, 1980) and a general depletion in δ18O values from the least altered through to the most altered and mineralized host rocks.In addition to extensive dolomitization, saddle dolomite and depleted δ18O values, there are also sev-eral structural controls that are common across all four of the deposits. They all have at least one significant fault that has been interpreted to be the pathway for mineralizing fluids to reach the host rocks. As such, exploration for MVT-style mineralization should not only focus on passive margin en-vironments with extensive dolomitization and the presence of saddle dolomite in mineralized zones, but also on the proximity to local structures which may have acted as pathways for mineralizing fluids.It is anticipated that the variables associated with the specific textural and isotopic signatures observed in each of the four deposits studied could be used as vectoring tools and so aid in the exploration of similar carbonate-hosted Zn-Pb deposits. Additionally, now that a global context of MVT systems has been presented, this study is able to focus on the four specific carbonate-hosted Zn-Pb systems within the Kootenay Arc.48Chapter 3: Salmo-type Carbonate-Hosted Zn-Pb Deposits: Jersey and H.B.IntroductionCarbonate-hosted Zn-Pb deposits in southeastern British Columbia (BC) and northeastern Washington are hosted within the Kootenay Arc; a lower to mid-Paleozoic curving belt of complexly deformed sedi-mentary, volcanic and metamorphic rocks (Fyles, 1967) that represent the suture between the ances-tral North American margin and the Intermontane superterrane (Colpron and Price, 1995). The North American part of the central Kootenay Arc comprises a thick clastic assemblage of Neoproterozoic Windermere Supergroup, upon which the Lower Cambrian Reno Formation, the Cambrian Laib Forma-tion (containing the Truman, Reeves and Emerald members), the Middle Cambrian Nelway Formation and the Ordovician Active Formation were deposited (Figure 1.4). The Kootenay Arc Zn-Pb deposits are hosted in either the Reeves member or the Nelway Formation and may be characterized as Salmo-type or Metaline-type based on the age of their host rocks and the degree of deformation (Fyles, 1970). The Jersey and H.B. Zn-Pb deposits in southeastern BC (Fig-ure 1.5) are characterized as Salmo-type deposits (Fyles, 1970). Mineralization from both deposits is hosted in the Reeves member dolomite and is comprised of stratiform lenses of pyrite, sphalerite and galena, with galena dominating at the H.B. Mine. Proximity to Mesozoic intrusions has led to late-stage skarn mineralization at both of these mines; particularly Jersey. The two intrusions in close proximity to Jersey and H.B. are the mid-Cretaceous Dodger and Emerald biotite granite stocks (Cathro and Lefebure, 2000). Tungsten skarn mineralization at the Jersey Mine occurs as disseminated scheelite (Ca(WO4))  with minor powellite (Ca(MoO4)) and trace wolframite ((Fe,Mn)WO4) and molybdenite (MoS2; Fyles and Hewlett, 1959). Skarn mineraliza-tion replaces the Reeves member limestone and Truman member argillite forming a green and brown banded skarn containing diopside, garnet and calcite (Cathro and Lefebure, 2000).Jersey Deposit GeologyThe Jersey deposit is located at the summit between Sheep and Lost creeks, roughly 11 km southeast of the town of Salmo (Figure 1.5). Between 1949 and 1970, whilst the mine was active, 6.4 Mt of ore was processed at 2.60 % Zn and 1.04 % Pb. Measured and indicated ore reserves as of April 1st, 1965 are reported at 671 075 tonnes grading 4.10 % Zn and 1.20 % Pb (MINFILE 082FSW009, BC Geological Survey, 2016). The deposit is underlain by rocks of the Laib Formation (Figure 3.0): 1) the Truman member, made up of interbedded thin grey and white, locally dolomitic limestone, 2) the Emerald member, a black argillite unit, 3) the Reeves member, which in this area consists of limestone and dolomite and, 4) the upper Laib Formation, made up of green phyllites and micaceous quartzites. These sedimentary rocks were intruded by small plugs, dikes and sills of Cretaceous granite and those that are in contact with the granitic bodies are typically skarnified, resulting in a variety of recrystallized coarse-grained marble 49to garnet-pyroxene–bearing skarns (Fyles and Hewlett, 1959).Sheep CreekAspen CreekSalmo River0 2kmNH.B.Jersey Lost CreekPlutonSheep Creek StockSalmo StockSalmo River AnticlineSalmo117°30’ W49°20’ NCretaceous IntrusionsLower to Middle Jurassic Elise FormationOrdovician Active FormationLower Cambrian Truman andEmerald membersLower Cambrian ReevesmemberFigure 3.0. Deposit geology of the Jersey and H.B. deposits, southeastern British Columbia. Modified from Giroux and Grunenberg (2010).The Zn-Pb deposits at the Jersey Mine are hosted by fine grained, poorly layered to massive dolomite of the Reeves member (Fyles and Hewlett, 1959; Simandl and Paradis, 2009). This mineralization occurs near to the base of the Reeves member and forms stratiform lenses of pyrite, sphalerite and galena in dolomitized zones (Fyles and Hewlett, 1959; Simandl and Paradis, 2009). Five ore bands, ranging in thickness from 0.3 to 9.0 m have been identified. These bands in order of stratigraphic sequence are:1. upper lead band,2. upper zinc band,3. middle zinc band,4. lower zinc band, and505. lower lead bandThe Truman member of the Laib Formation forms the mine footwall rocks. It consists of dense, red-dish green skarns and a brown argillite that hosts W and Mo mineralization (Fyles and Hewlett, 1959; Giroux and Grunenberg, 2010). Several zones of significant and locally very different mineralization have been identified. Historically, mined areas produced Zn, Pb and W, with other areas recognized for their high Mo, Au, Bi, As, Cu, Ag, Cd and Ba concentrations (Fyles and Hewlett, 1959; Giroux and Grunenberg, 2010).The Jersey deposit rocks were deformed by three phases of folding. Within the mine area, structure is dominated by the major north-northeast–trending Salmo River anticline. The deposit mineralization is associated with the east limb of this anticline (Fyles and Hewlett, 1959). H.B. Deposit GeologyThe H.B. deposit is located on Aspen Creek, a tributary of Sheep Creek, 8 km southeast of Salmo (Fig-ure 1.5). Between 1912 and 1978, whilst the mine was active, 6.7 Mt of ore was recovered at 4.10 % Zn and 0.10 % Pb. Measured and indicated reserves as of December 31st, 1978 are reported at 36 287 tonnes grading 4.10 % Zn and 0.10 % Pb (MINFILE 082FSW004, BC Geological Survey, 2016). The deposit is underlain by the Reeves member limestone and the Lower to Middle Ordovician Ac-tive Formation (Figure 3.0). These units are in contact with each other along a fault, with the Active Formation rocks overthrust from the east over the Reeves member rocks (Fyles and Hewlett, 1959). There are two distinct calcareous layers of the Reeves member, an upper unit about 110 m thick that is separated from a lower 12 m member by 15 to 30 m of micaceous, brown, limey argillite. The H.B. orebodies occur roughly 100 m to the west of the thrust fault.The mineralized zones are in dolomitized limestone of the Reeves member and contain banded galena, sphalerite, pyrite and pyrrhotite similar to that at the Jersey deposit, except that Pb dominates (Giroux and Grunenberg, 2010).In the vicinity of the H.B. Mine, the beds are folded into a broad synclinorium, and the limestone layers in the mine are on the west limb of this structure. The principal ore zones consist of three steeply-dipping, parallel zones lying approximately side-by-side and extending as pencil-like shoots for about 900 m along the south plunge of the controlling structures (Fyles and Hewlett, 1959; Giroux and Grunenberg, 2010).A suite of regional rocks from the Kootenay Arc was collected to define the background stable isotopic signature of the carbonate alteration footprint in this region. In addition to regional samples, mine specific samples were collected to define the carbonate alteration footprint of Zn-Pb mineralization. Studies of C and O isotopes characterize the isotopic footprint of the background Kootenay Arc rocks and the Zn-Pb mineralization hosted within this region and, when used in conjunction with other geo-logical data, place better constraints on the fluid components and fluid source responsible for these 51Zn-Pb deposits. In turn, this information will improve the current deposit model for these poorly un-derstood systems and aid in exploration for similar deposits.SamplingSampling within southeastern BC and northeastern Washington concentrated on collecting samples for C and O isotope analysis. These samples included drill core and surface grab samples of mineralized and non-mineralized rocks in order to define the C and O isotopic signature of background and Zn-Pb mineralization within this region. Appendix B provides a complete list of all the samples taken for this M.Sc. thesis.RegionalA regional sample set was collected to define the C and O isotopic signature of host rocks distal to min-eralized areas and so provide a background host rock signature (Figure 3.1). These regional samples were taken as a roadside transect from east to west across this area of southeastern BC and northeast-ern Washington. Samples were taken of every different carbonate lithology exposed along the road, with several small transects taken at each sample stop. In total, 84 regional samples were collected, incorporating samples from the Reeves member dolomites and limestones, the Nelway Formation dolomites and limestones and the Active Formation calcareous shales.JerseySample collection at the Jersey Mine (Figure 3.1) concentrated on two drill holes, an underground transect and a stratigraphic transect through an open pit. Twenty samples were taken from drill hole JS07-17 (0 – 254 feet) and 25 samples were taken from drill hole JS07-08 (0 – 432 feet). These two drill holes displayed all of the significant lithologies present at the Jersey Mine and crosscut horizons of Zn-Pb mineralization; this enabled the sampling of mineralized and non-mineralized host rocks. A small, east-west, 9 sample transect was conducted across the mine in the underground workings to intersect the high-grade lower Pb band described above. In addition to drill hole and underground sampling, a 22 sample transect was conducted from the base to the top of the B Zone Pit. Again, this pit enabled the sampling of mineralized and non-mineralized host rocks in stratigraphic section.H.B.Sample collection at the H.B. Mine (Figure 3.1) concentrated on two drill holes. Eighteen samples were taken from drill hole HB1008 (0 – 70 feet) and 11 samples were taken from HB1005 (0 – 119 feet). These two drill holes displayed all of the significant lithologies present at the H.B. Mine and crosscut horizons of Zn-Pb mineralization; allowing for the sampling of mineralized and non-mineralized host rock.52Figure 3.1. Sampling location map on regional geology. Modified from Wheeler and McFeely (1991); Paradis et al. (2009).SalmoStockWallack CreekStockLost Creek            PlutonSpirit BatholithPend Oreille RiverSlate Creek ThrustFlume Creek FaultSheep Creek AnticlineLeadpoint FaultBlack Canyon ThrustColumbia ThrustSalmo Valley FaultBlack Blu ThrustLeadpoint FaultSheep Creek AnticlineSalmo River AnticlineSalmo River AnticlineBritish ColumbiaWashingtonIdaho49°00’ N117°30’ W 117°00’ W49°00’ NCretaceous IntrusionsLower to Middle Jurassic Elise FormationDevonian to Lower Triassic Charbonneau Creek AssemblageOrdovician Active FormationMiddle Cambrian Nelway FormationLower Cambrian to Ordovician Upper Laib FormationLower Cambrian Laib FormationLower Cambrian Reno FormationLower Cambrian Quartzite Range FormationLower Cambrian and Proterozoic Three Sisters FormationLower Cambrian and Proterozoic Monk FormationPend Oreille MineRemacMineJerseyMineH.B.MineRegional SamplesJersey SamplesH.B. SamplesRemac Samples0 5kmNINSULAR BELTCOAST BELTINTERMONTANE BELTOMINECA BELTFORELAND BELTBRITISHCOLUMBIAWASHINGTONALBERTAIDAHOVancouverSeattle SpokaneNelsonKelownaPrince GeorgeMetalineFallsSalmo53PetrographyPetrographic studies were undertaken to define the different carbonate phases within the region; specifically at the Jersey Mine. In addition, petrographic studies combined with observations made in the field and using cold-cathode cathodoluminescence described below contribute to a paragenesis of events for the Jersey deposit (Figure 3.2).PhaseLimestoneFine Grained DolomiteMedium Grained DolomiteCoarse Grained DolomiteCalciteRelative TimingPyrite 1Pyrite 2Calc-SilicatesSphaleriteCalcite CementMain Mineralization EventFigure 3.2. Paragenetic diagram to highlight the sequence of events at the Jersey Mine.Six samples were taken of each significant lithology at the Jersey Mine as this deposit exhibited the best available core and is very similar to the H.B. deposit in terms of location, lithology and mineralogy. Petrographic methods are outlined in Appendix A.Carbonate UnitsFour carbonate units were defined during petrographic observations: limestone, fine grained dolo-mite, medium grained (pervasive) dolomite and coarse grained dolomite. In addition to these carbon-ate rocks, two calcite vein-types were also observed: coarse sparry calcite and coarse sparry calcite + sulphides. The limestone unit is recognized by coarse grained, euhedral calcite grains with excellent twinning. By contrast, the dolomite units are defined as fine, medium or coarse grained, subhedral to anhedral dolomite with excellent twinning. These dolomites may be composed purely of medium grained (pervasive) dolomite, or occur as banded or zebra dolomites with alternating sections of fine, medium or coarse grains; these banded or zebra dolomites often display an oxidized appearance. Para-genetically, the fine grained dolomite occurred first and was subsequently overprinted by the medium grained dolomite. These two dolomites are followed by the final dolomite phase; the coarse grained dolomite.  In addition, several samples display trace abundances of fine grained calc-silicates.Banded or zebra dolomites are thought to represent the transition of fine grained to medium grained or medium grained to coarse grained dolomite, with coarse grained dolomites forming due to re-54crystallization. Recrystallization textures such as coarse grained marbles and calcite veins and small occurrences of calc-silicate minerals (specifically talc and/or muscovite) in several of the samples are thought to be indicators of contact metamorphism from Mesozoic plutons proximal to the Jersey Mine. In addition to petrographic observations of Jersey samples, hand sample description of regional, Jer-sey and H.B. samples led to the identification of three distinct vein-types: coarse sparry calcite, coarse sparry calcite + sulphides and tabular calcite veins. The two sparry calcite vein-types have been re-crystallized by contact metamorphism resulting from the emplacement of proximal Mesozoic plutons. Due to recrystallization and a lack of cross-cutting relationships, it is difficult to determine an accurate paragenesis of these vein-types. The coarse sparry calcite + sulphide veins must have been syn-min-eralization based on their sulphide association. The remaining coarse sparry calcite veins that are not associated with sulphides may have formed before, during or after the sulphide associated veins were emplaced. The tabular calcite veins are thought to form late as they display very planar, straight veins that cross-cut the host rock. However, only one tabular calcite vein (HB1008_43B) is recognized in the samples collected so an accurate paragenesis cannot be completed. MineralizationSulphide mineralization from the Jersey Mine is strongly deformed by three phases of folding. This mineralization is most commonly associated with, and forms bands through, the medium and coarse grained dolomites and the coarse calcite veins (Figure 3.3). The main mineralization observed in thin section from these samples is sphalerite and pyrite. There are two pyrite phases; one disseminated throughout early limestone and dolomite and a second as inclusions in later sphalerite grains (Figure 3.3). Sphalerite mineralization is banded throughout the early dolomite phases, but is most strongly associated with coarse calcite veins. More detailed descriptions of these thin sections can be found in Appendix C. 0.5 mmFigure 3.3. Photographs to display sulphide associations at the Jersey Mine. Plane polarized image of sulphide banding throughout coarse grained calcite and dolomite (left). Reflected light (FOV = 1 mm) image of pyrite grains forming as inclusions and along grain boundaries of pre-existing sphalerite (right).55Cold-Cathode CathodoluminescenceIn addition to petrographic studies, cold-cathode cathodoluminescence (cold-cathode CL) analysis was undertaken to further define the different carbonate phases. CL imaging is useful as it provides data that is otherwise not readily visible such as information regarding diagenetic change, cementation or porosity loss (Boggs and Krinsley, 2006). Cold-cathode CL methods are outlined in Appendix A.The luminescence characteristics of carbonate minerals are primarily controlled by the abundance of Mn, rare earth elements and Fe (Boggs and Krinsley, 2006). Mn2+ and rare earth elements are the most important activator elements, whereas Fe2+ acts as a luminescence quencher (Boggs and Krin-sley, 2006). A general consensus is that the CL brightness is a function of the concentrations of Mn2+ activators and Fe2+ quenchers (Boggs and Krinsley, 2006).Two polished thin sections (JS07-08_180 and JS07-17_55.5) and a slab from regional sample NC29 were analyzed by cold-cathode CL. These samples were chosen due to their abundance of calcite (JS07-17_55.5 and NC29 represent the Reeves member limestone and JS07-08_180 represents a min-eralized calcite vein).ResultsAll samples displayed a dull-orange fluorescence under cold-cathode CL regardless of whether the phase fluorescing was the Reeves member limestone (JS07-17_55.5 and NC29) or a sparry calcite vein associated with sulphides (JS07-08_180). In addition to the background dull-orange fluorescence, there is at least one other phase of calcite represented by a brighter orange fluorescence (Figure 3.4c and d). This brighter fluorescence represents a phase of fine grained interstitial calcite cement, in some cases observed to rim sphalerite grains and infill fractures and twinning of the sparry calcite vein. a. b.c. d.1 mmFigure 3.4. Cold-cathode CL images for samples JS07-08_180 and NC29 (FOV = 2 mm). Background dull-orange fluores-cence of the Reeves member limestone (a.). Plane polarized image of sample JS07-08_180 displaying sulphides hosted in a sparry calcite vein (b.). Background dull-orange fluorescence of sparry calcite with region of brighter fluorescence out-lining sulphides and infilling twinning (c.). Background dull-orange fluorescence of sparry calcite with region of brighter fluorescence infilling fracture (d.). 56The dull-orange fluorescence is thought to represent the relatively Mn2+-depleted background com-position of all limestone and calcite samples, as the two limestone samples (JS07-17_55.5 and NC29) show no variation in this fluorescence (Figure 3.4a; Boggs and Krinsley, 2006). The brighter orange fluorescence associated with sulphides is thought to represent a calcite phase that must have a higher concentration of Mn2+ (Boggs and Krinsley, 2006). Carbonate Staining and UV Light AnalysisCarbonate staining and ultraviolet (UV) light analysis also aided in the determination of different car-bonate phases. Offcuts from the six samples sent for thin section production were analyzed through carbonate staining and UV light. Only samples from Jersey were analyzed as the Jersey and H.B. de-posits are very similar in terms of location, lithology and mineralogy. Carbonate staining and UV light analysis methods are outlined in Appendix A.Carbonate StainingThe identification of different carbonate minerals is enabled by carbonate staining (Hitzman, 1999). The easiest method of staining hand samples and drill core uses two diluted hydrochloric acid (HCl) solutions; one containing alizarin red S and one containing potassium ferricyanide (Hitzman, 1999). These solutions allow for the rapid identification of calcite, ferroan calcite, ferroan dolomite and rho-dochrosite (Hitzman, 1999). Table 3.0 indicates different carbonate phases and their associated aliza-rin red S and potassium ferricyanide stains as defined by Hitzman (1999).Carbonate Mineral Formula Alizarin red S Potassium FerricyanideCalcite (CaCO3) Pink to red UnstainedAragonite (CaCO3) Pink to red UnstainedFerroan Calcite (Ca,Fe)CO3 Pink to pale pink Pale to deep blueDolomite (CaMg(CO3)2) Unstained UnstainedFerroan Dolomite (Ca(Mg,Fe)CO3)2) Unstained Pale to deep turquoiseSiderite (FeCO3) Unstained UnstainedRhodochrosite (MnCO3) Unstained Very pale brownMagnesite (MgCO3) Unstained UnstainedWitherite (BaCO3) Red UnstainedCerussite (PbCO3) Mauve UnstainedTable 3.0. Carbonate phases and their associated alizarin red S and potassium ferricyanide stains. Modified from Hitzman (1999).Alizarin Red SThe alizarin red S solution stains calcite and ferroan calcite a pink to red colour, whereas dolomite will not stain (Hitzman, 1999). However, Warne (1962) suggests that ferroan dolomite could obtain a lilac stain.57Potassium FerricyanideThe potassium ferricyanide solution stains ferroan calcite a pale to deep blue colour and ferroan dolo-mite a pale to deep turquoise colour, whereas non-ferroan calcite and dolomite will not stain (Hitzman, 1999). ResultsThe staining results for these six samples are summarized in Table 3.1 and photographs are displayed in Figure 3.5. All samples, except JS07-17_172, displayed a light red to red stain using the alizarin red S solution and no stain using the potassium ferricyanide stain. Sample JS07-17_172 displayed no stain with either solution. A light red to red stain with alizarin red S and no stain with potassium ferricyanide indicates a calcite/aragonite composition according to Hitzman (1999), whereas no stain with either solution indicates a pure dolomite composition (Hitzman, 1999).Sample ID Hole ID Depth (ft) Alizarin Red Potassium Ferricyanide Potential Carbonate PhaseJS07-17 23 JS07-17 23 Red No stain Calcite/aragoniteJS07-17 55.5 JS07-17 55.5 Red No stain Calcite/aragoniteJS07-17 120 JS07-17 120 Light red No stain Calcite/aragoniteJS07-17 172 JS07-17 172 No stain No stain DolomiteJS07-08 100 JS07-08 100 Red No stain Calcite/aragoniteJS07-08 180 JS07-08 180 Red No stain Calcite/aragoniteTable 3.1. Alizarin red S and potassium ferricyanide staining results for the Jersey Mine.No Stain Alizarin red S Potassium FerricyanideSample IDJS07-17_23JS07-08_100Figure 3.5. Selection of images to show results of the staining on Jersey samples. JS07-17_23 and JS07-08_100; typical light red to red stain.Despite these staining results, samples JS07-17_23, JS07-17_120, JS07-08_100 and JS07-08_180 were recorded as dolomite based on hand sample description, HCl acid tests and thin section petrography. These samples must therefore have a higher calcite/aragonite content than previously considered to 58explain their red stain using the alizarin red S solution, or the staining solution was affected by another factor. For the purpose of C and O isotope analysis, all samples recorded as dolomite based on hand sample description, HCl acid tests and thin section petrography remain characterized as dolomite ir-respective of carbonate staining results.UV Light AnalysisUV fluorescence is a form of luminescence that represents the ability of a mineral or element to emit light when excited by UV light (Rost, 1992). Fluorescence occurs when impurities or activators are present within the mineral in question. These activators are usually metal cations: W, Mo, Pb, B, Ti, Mn, U and Cr (Marfunin, 1979). Rare earth elements may also contribute to fluorescence as well as the presence of organic matter, whereas (as in cold-cathode CL) elements such as Fe act as fluorescence quenchers (Marfunin, 1979). Many minerals will fluoresce one colour under short-wave UV light (100 – 280 nm) and another under long-wave UV light (315 – 400 nm). Analysis of fluorescence in carbonate rocks has been used extensively to aid in the determination of carbonate mineralogy, hidden microfab-rics and organic components (Flügel, 2010). ResultsThe UV fluorescence results for these six samples are summarized in Table 3.2 and photographs are displayed in Figure 3.6. All samples, except JS07-08_100 and JS07-08_180, displayed no fluorescence under UV light. Samples JS07-08_100 and JS07-08_180 displayed a pale yellow to orange fluorescence under UV light; both samples displayed this fluorescence in association with the sulphides (sphalerite and pyrite). Sample ID Hole ID Depth (ft) Fluorescence Potential CauseJS07-17 23 JS07-17 23 No fluorescenceJS07-17 55.5 JS07-17 55.5 No fluorescenceJS07-17 120 JS07-17 120 No fluorescenceJS07-17 172 JS07-17 172 No fluorescenceJS07-08 100 JS07-08 100 Pale yellow disseminated throughout sulphides Pb/MnJS07-08 180 JS07-08 180 Distinct yellow and orange throughout sulphides Pb/MnTable 3.2. UV light analysis results for the Jersey Mine.59Visible Light Ultraviolet LightSample IDJS07-08_100JS07-08_180Figure 3.6. Selection of images to show results of the UV light analysis on Jersey samples. JS07-08_100 and JS07-08_180; yellow to orange fluorescence.No fluorescence under UV light suggests that activator elements, such as metal cations, were not present within most samples (Marfunin, 1979). The presence of Fe in these samples could act as a fluorescence quencher, however as none of these samples stained under a potassium ferricyanide solution this indicates that they did not have a ferroan composition. A yellow to orange fluorescence is associated with Pb2+ or Mn2+ replacing Ca2+ (Marfunin, 1979). In this case the fluorescence observed in samples JS07-08_100 and JS07-08_180 may be attributed to the presence of either Mn2+ or Pb2+. The presence of Mn2+ surrounding sulphides and infilling fractures was observed in cold-cathode CL analy-sis; the fluorescence observed in UV light may also be a reflection of this relationship. Additionally, Pb2+ is a possible activator due to the association of this fluorescence with sulphides and so perhaps very fine grained cerussite (PbCO3) that is not recognized in hand sample or thin section.C and O Isotope AnalysisApproximately 200 samples from this field area were analyzed for C and O isotopes. These samples were collected to define the isotopic footprint of alteration associated with Zn-Pb mineralization host-ed within this portion of the Kootenay Arc. C and O isotope analysis methods are outlined in Appendix A and a table of C and O data, quality assurance and quality control (QAQC) can be found in Appendix E.All isotopic values are given in per mil (‰) with δ18O values recorded with respect to Vienna Standard Mean Ocean Water (VSMOW) and δ13C values recorded with respect to Vienna Pee Dee Belemnite (VPDB).RegionalFrom the 84 samples taken from regional roadside outcrops, only 70 returned meaningful δ13C and δ18O values. The remaining 14 samples did not contain enough carbonate to produce reliable δ13C and 60δ18O values.The following results present all data from regional samples. Figure 3.7 displays the δ13C and δ18O values of all regional samples from BC and Washington. Samples from BC and Washington display δ18O values as low as 6.41 ‰ and 8.67 ‰ with respect to VSMOW and δ13C values as low as -3.12 ‰ and -2.40 ‰ with respect to VPDB. The data from BC shows a greater spread of isotopic values, with some samples even being enriched in δ13C and δ18O compared to other samples. Furthermore, there is a cluster of data points from BC that is extremely depleted in δ13C (up to -15.38 ‰) compared to the other samples.    Figure 3.7. Plot to display the δ13C and δ18O values of all regional samples. Figure 3.8 displays the δ13C and δ18O values of all regional samples from BC and Washington according to rock-type and Formation or member. In both BC and Washington samples, samples from the Nelway Formation display δ18O values as low as 6.41 ‰ relative to VSMOW. From these Nelway samples it is the banded and zebra dolomites, particularly in the BC samples, that represent the lowest δ18O val-ues. Samples from the Reeves member in BC display δ18O values as low as 8.53 ‰ relative to VSMOW however, a cluster of these Reeves samples also represent the enriched δ13C and δ18O isotopic signa-ture observed in Figure 3.7 and include intermediate and resilient dolomites (up to 24.81 ‰ δ18O and 1.41 ‰ δ13C). Sparry calcite vein samples from BC and Washington display δ18O values as low as 8.67 ‰ relative to VSMOW. Additionally, the extremely depleted δ13C values displayed in Figure 3.7 mostly represent phyllitic limestone and one carbonaceous slate sample.Non-mineralized and mineralized rocks from BC and Washington therefore display a fairly similar iso-topic signature; samples from both locations are depleted in δ18O (up to 15.00 ‰ and 12.00 ‰ re-spectively), but not depleted in δ13C relative to Cambrian marine carbonate (the supposed starting 61Figure 3.8. Plots to display the δ13C and δ18O values of all regional BC and Washington samples according to their rock-type and Formation. Regional BC data according to Formation (a). Regional Washington data according to Formation (b). Regional BC data according to rock-type (c). Regional Washington data according to rock-type (d). -20.00-15.00-10.00-5.000.005.000.00 10.00 20.00 30.00d13C_VPDB ‰ d180_VSMOW ‰ C and O Data for Regional Samples Collected in BC According to Formation Active FormationNelway FormationEmerald memberReeves memberCambrian MarineCarbonatea. -20.00-15.00-10.00-5.000.005.000.00 10.00 20.00 30.00d13C_VPDB ‰ d180_VSMOW ‰ C and O Data for Regional Samples Collected in BC According to Rock-Type Carbonaceous SlatePhyllitic LimestoneLimestoneBanded DolomiteZebra DolomitePervasive DolomiteIntermediate DolomiteResilient DolomiteSparry Calcite VeinCambrian MarineCarbonatec. -20.00-15.00-10.00-5.000.005.000.00 10.00 20.00 30.00d13C_VPDB ‰ d180_VSMOW ‰ C and O Data for Regional Samples Collected in WA According to Formation Nelway FormationJosephineMineralizationYellowhead 1MineralizationCambrian MarineCarbonateb. -20.00-15.00-10.00-5.000.005.000.00 10.00 20.00 30.00d13C_VPDB ‰ d180_VSMOW ‰ C and O Data for Regional Samples Collected in WA According to Rock-Type LimestoneFine Grained DolomiteJosephine MineralizationYellowhead 1MineralizationPervasive DolomiteZebra DolomiteTabular Calcite VeinSparry Calcite VeinSparry Calcite Vein +SulphidesCambrian MarineCarbonated. 62point for these Cambrian-Ordovician carbonate host rocks; Wadleigh and Veizer, 1992; Maloof et al., 2005; Saltzman, 2005; Dilliard et al., 2007; Prokoph et al., 2008). Data taken to form the Cambrian marine carbonate isotopic range has been collected from several authors: Wadleigh and Veizer (1992), Maloof et al. (2005), Saltzman (2005), Dilliard et al. (2007) and Prokoph et al. (2008). These authors present secular δ18O and δ13C data taken from Cambrian aged whole rock (limestone and dolomite) and fossil (belemnites, trilobites and brachiopods) samples that range from -5.00 ‰ to 5.00 ‰ with respect to VPDB δ13C and 20.00 ‰ to 28.00 ‰ with respect to VSMOW δ18O. These secular variations throughout the Cambrian are attributed to increased evaporation and precipitation rates owing to cooler or warmer climates for δ18O variations, and increased burial rates of isotopically light organic C for δ13C variations. Phases of increased δ18O values are attributed to cold climates, whereas a decrease in δ18O values is associated with warmer climates. Additionally, increased burial of organic C will lead to an increase in δ13C values. This Cambrian marine carbonate range therefore represents the isotopic signature of the most unaltered, background host rock from which all Cambrian carbonates will have formed, with variations from this range caused by fluid:rock interaction after deposition.From these mineralized and non-mineralized samples, it is the Nelway Formation samples that display the greatest depletion relative to Cambrian marine carbonate, in particular the banded and zebra do-lomites, which suggests that these rocks have experienced the greatest fluid:rock interaction, or have interacted with a fluid with the lowest isotopic composition. The samples that display a significant de-pletion in δ13C compared to all others were all collected from the Ordovician Active Formation; a unit composed of carbonaceous shales, slates and argillaceous limestone. Finally, the samples that display enrichment in δ13C and δ18O values are only found close to the Remac Mine, which will be discussed further in Chapter 4.In addition to the isotopic composition of the mineralized and non-mineralized host rocks, determina-tion of the isotopic composition of the mineralizing fluid is also important. The definition of the isoto-pic composition of the mineralizing fluid can place important constraints on the components, source and temperature of this fluid. The results described above have facilitated the calculation of δ18O of the fluid in equilibrium with a specific δ18O of the rock at a specific temperature. Essentially, these cal-culations are able to estimate the δ18O signature of the fluid responsible for causing the resultant δ18O signature of a specific sample. These equations have resulted from experimental data from O’Neil et al. (1969) for calcite samples and Böttcher (1994) for dolomite samples and are as follows:Calcite samples:  ∆18O Cal − H2O = 2.78 × 106T2 − 2.89      (1)Dolomite samples:  ∆18O Dol − H2O = 3.19 × 106T2 − 4.42        (2)These equations are based on the assumption that there is complete equilibration between the rock and the fluid and that the reaction between the two takes place within a closed system (O’Neil et al., 1969; Böttcher, 1994). Under the assumption that:        63    ∆A − B =  δ18OA −  δ18OB      (3)The above equations (one and two) can be rearranged to calculate the δ18O of the fluid as follows:Calcite samples:  δ18Ofluid =  δ18Orock − 2.78 × 106T2 + 2.89        (4) Dolomite samples:  δ18Ofluid =  δ18Orock − 3.19 × 106T2 + 4.42      (5)Where 2.78, 2.89, 3.19 and 4.42 are constants, δ18Orock is the δ18O value of a specific sample pro-duced by isotope analysis and T is temperature in Kelvin. These equations allow for the calculation of the δ18O of the fluid for a range of temperatures; in this case a range of 25°C to 250°C was used as this is a reasonable temperature range for Mississippi Valley-type (MVT) systems (Leach and Sangster, 1993; Leach et al., 2005). The calcite equation was used to calculate the δ18O of the fluid for all calcite vein, limestone and phyllitic limestone samples, whereas the dolomite equation was used to calculate the δ18O of the fluid for all dolomite samples.Figure 3.9 presents the results of calcite samples from regional BC outcrops. In this plot, the two most depleted calcite vein samples have been taken to represent the original fluid composition. As metioned previously, extensive recrystallization of all calcite veins through contact metamorphism has made the determination of a paragenesis of different vein types difficult; however, as these two represent the most depletion they have been used to represent the closest isotopic composition of the original fluid. Additionally, the three most enriched whole rock samples have been used to represent the most unaltered host rocks and so provide a background value from which all other samples have been altered. All samples that plot between these two endmembers (unaltered host rock and fluid) have been depleted relative to the background host rock values. Furthermore, assuming that the fluid had a seawater origin (δ18O = 0 ‰; Taylor, 1974; Shackleton and Kennett, 1975) as is acceptable for and is consistent with an MVT system (Anderson and Macqueen, 1982) allows for the calculation of a minimum temperature range for the system. In this case, a minimum temperature range for the sys-tem using the two most depleted calcite vein samples is approximately 150°C to 200°C as highlighted by the black arrow in Figure 3.9.Evaporation of this seawater-sourced fluid to form the saline brines commonly associated with MVT-style deposits as the ore forming fluids (Leach and Sangster, 1993; Leach et al., 2005) will affect the δ18O value of the fluid. Evaporation preferentially removes the lighter 16O isotope rendering the re-maining fluid relatively heavier in 18O. This in turn will increase the δ18O value of that fluid. In the case of evaporated seawater, this δ18O value will therefore increase above 0 ‰; by how much is dependent on the degree of evaporation. Consequently, the minimum temperature range of the system will also increase. Utilizing seawater as the mineralizing fluid is therefore one endmember scenario and only represents the very minimum temperature range of fluid:rock interaction in these MVT-style deposits. Despite this, the temperature range of 150°C to 200°C indicated in Figure 3.9 represents typical MVT deposit mineralization temperatures (Leach and Sangster, 1993; Leach et al., 2005).64Figure 3.9. Plot to display the δ18Ofluid and corresponding temperature for regional BC calcite samples.   65Using the most depleted and most enriched samples observed in Figure 3.9, polygons representing the background host rock value and fluid composition can be displayed on forthcoming plots. Figure 3.10 displays the remaining BC and Washington regional samples in relation to these two polygons. As with Figure 3.7, all remaining BC and Washington regional samples have a similar isotopic signature; all samples are depleted relative to the background host rock values. Furthermore, it is again the banded and zebra dolomites that display the most depletion and the limestone samples that display the least. Additionally, for regional Washington samples there is no distinction in isotopic signature between mineralized and non-mineralized samples; this will be discussed further in Chapter 5. Lastly, with the exception of one, all calcite vein samples (tabular, sparry or sparry + sulphides) plot within the main mass of altered rocks; -7.00 ‰ to -13.00 ‰ δ18O. These samples are thought to represent the original fluid that has been modified by fluid:rock reaction prior to vein formation.JerseyFrom the 76 samples taken at Jersey, only 60 returned meaningful δ13C and δ18O values. The remaining 16 samples did not contain enough carbonate to produce reliable δ13C and δ18O values.The following results present all data from Jersey samples. Figure 3.11 displays the δ13C and δ18O values from all non-mineralized and mineralized samples taken in BC; regional, Jersey Mine, H.B. Mine and Remac Mine. The Jersey isotopic values display a marked depletion in δ18O (up to 12.50 ‰); however, there is no distinction in isotopic signature between mineralized and non-mineralized samples. There is only a slight depletion in δ13C values (up to -2.45 ‰) and again, no distinction in isotopic signature between mineralized and non-mineralized samples. Figure 3.12a displays the δ13C and δ18O values from Jersey samples according to rock-type. As with the regional samples, the most depleted rock-types from the Jersey Mine are the banded and pervasive dolomites. However, as above, there is no distinction in the isotopic signature between mineralized and non-mineralized banded dolomite. The limestone samples show the least depletion (up to 16.51 ‰ δ18O and -1.97 ‰ δ13C).66Figure 3.10. Plots to display the δ18Ofluid and corresponding temperature for the remaining regional BC and Washington samples. Regional BC samples (a).      a. 67Figure 3.10 continued. Plots to display the δ18Ofluid and corresponding temperature for the remaining regional BC and Washington samples. Regional Washington samples (b).     b. 68   Figure 3.11. Plots to display the δ13C and δ18O values of all BC regional, Jersey, H.B. and Remac samples. Non-mineral-ized samples (a). Mineralized samples (b).69   -20.00-15.00-10.00-5.000.005.000.00 10.00 20.00 30.00d13C_VPDB ‰ d180_VSMOW ‰ C and O Data for the Jersey Mine According to Rock-Type Banded LimestoneMineralized BandedLimestoneCoarse LimestoneBanded DolomiteMineralized BandedDolomitePervasive DolomiteSparry Calcite VeinSkarnCambrian Marine Carbonate-20.00-15.00-10.00-5.000.005.000.00 10.00 20.00 30.00d13C_VPDB ‰ d180_VSMOW ‰ C and O Data for the H.B. Mine According to Rock-Type Banded DolomiteMineralized BandedDolomiteTabular Calcite VeinSparry Calcite VeinSparry Calcite Vein +SulphidesCambrian Marine Carbonatea. b. Figure 3.12. Plots to display the δ13C and δ18O values of all Jersey and H.B. samples according to their rock-type. Jersey samples (a). H.B. samples (b).As with the regional samples described above, the mineralized and non-mineralized samples from Jersey display depletion in δ18O, but not in δ13C compared to Cambrian marine carbonate (up to 12.00 ‰ δ18O; Wadleigh and Veizer, 1992; Maloof et al., 2005; Saltzman, 2005; Dilliard et al., 2007; Prokoph et al., 2008). Furthermore, it is again the banded and pervasive dolomites from these samples that are 70the most depleted compared to Cambrian marine carbonate and so have likely undergone the most fluid:rock interaction, or have interacted with a fluid with the lowest isotopic composition (Wadleigh and Veizer, 1992; Maloof et al., 2005; Saltzman, 2005; Dilliard et al., 2007; Prokoph et al., 2008). In addition, as well as containing MVT-style Zn-Pb mineralization, the Jersey Mine also contains younger skarn mineralization associated with the nearby Cretaceous intrusions. These skarn samples are, for the most part, slightly depleted relative to Cambrian marine carbonate (up to 5.00 ‰ δ18O; Wadle-igh and Veizer, 1992; Maloof et al., 2005; Saltzman, 2005; Dilliard et al., 2007; Prokoph et al., 2008). However, there is one skarn sample that is extremely depleted and has a δ13C of -7.90 ‰ and a δ18O of 10.22 ‰.Utilizing equations four and five and the polygons of background host rock and original fluid as defined in Figure 3.9 has enabled the determination of δ18O of the fluid for Jersey samples; these results are displayed in Figure 3.13a and b. As above, the pervasive and banded dolomites are the most depleted with respect to the background host rock values and the limestones are the least. Again, there is no distinction in isotopic signature between mineralized and non-mineralized samples with the exception of the least depleted limestone samples; all mineralized samples are more depleted relative to these. Lastly, all of the sparry calcite vein samples plot within the main mass of altered rocks; -7.00 ‰ to -20.00 ‰ δ18O. These samples are thought to represent the original fluid that has been modified by fluid:rock reaction prior to vein formation.H.B.From the 29 samples taken at H.B., only 26 returned meaningful δ13C and δ18O values. The remaining three samples did not contain enough carbonate to produce reliable δ13C and δ18O values.The following results present all data from H.B. samples. Figure 3.11 displays the δ13C and δ18O values from all non-mineralized and mineralized samples taken in BC; regional, Jersey Mine, H.B. Mine and Remac Mine. The non-mineralized H.B. samples display only a slight depletion in δ18O (up to 21.20 ‰); however, mineralized H.B. samples display a marked depletion in δ18O (up to 15.26 ‰). In addition, there is only a slight depletion in δ13C for non-mineralized samples (up to -1.21 ‰); however, mineral-ized samples display a depletion of up to -2.85 ‰. Figure 3.12b displays the δ13C and δ18O values from H.B. samples according to rock-type. The most depleted rock-type from the H.B. Mine is the mineral-ized banded dolomite. Furthermore, there is a significant difference in isotopic signature between the mineralized and non-mineralized banded dolomites; mineralized and non-mineralized samples differ by up to 5.84 ‰ δ18O and 1.58 ‰ δ13C. Non-mineralized samples collected at H.B. therefore display no depletion in δ18O and no depletion in δ13C compared to Cambrian marine carbonate (Wadleigh and Veizer, 1992; Maloof et al., 2005; Saltzman, 2005; Dilliard et al., 2007; Prokoph et al., 2008) and so differ from the depletion observed in the regional and Jersey samples. Mineralized samples however, are significantly depleted relative to Cambrian marine carbonate and display similar depletions as observed in regional and Jersey samples (up to 15.26 ‰ δ18O; Wadleigh and Veizer, 1992; Maloof et al., 2005; Saltzman, 2005; Dilliard et al., 71Figure 3.13. Plots to display the δ18Ofluid and corresponding temperature for the Jersey and H.B. samples. Non-mineralized Jersey samples (a).   a. 72Figure 3.13 continued. Plots to display the δ18Ofluid and corresponding temperature for the Jersey and H.B. samples. Mineralized Jersey samples (b).   b. 73Figure 3.13 continued. Plots to display the δ18Ofluid and corresponding temperature for the Jersey and H.B. samples. H.B. samples (c).  c. 742007; Prokoph et al., 2008). Furthermore, it is again the banded dolomites from these samples that are the most depleted compared to Cambrian marine carbonate and so have likely undergone the most fluid:rock interaction, or have interacted with a fluid with the lowest isotopic composition (Wadleigh and Veizer, 1992; Maloof et al., 2005; Saltzman, 2005; Dilliard et al., 2007; Prokoph et al., 2008). The significant difference in isotopic values between non-mineralized and mineralized rocks collected at H.B. contrasts to the fairly similar isotopic values between non-mineralized and mineralized rock col-lected regionally and at the Jersey Mine. Utilization of equations four and five and the polygons of background host rock and original fluid as defined in Figure 3.9 has enabled the determination of δ18O of the fluid for H.B. samples; these results are displayed in Figure 3.13c. As above, the banded dolomites are the most depleted with respect to the background host rock values with the mineralized banded dolomites displaying further depletion. Lastly, all calcite vein samples (tabular, sparry or sparry + sulphides) plot within the main mass of al-tered rocks; -6.00 ‰ to -17.00 ‰ δ18O. These samples are thought to represent the original fluid that has been modified by fluid:rock reaction prior to vein formation.Distance from Mesozoic PlutonsAs the samples from this region were collected in close proximity to several Mesozoic plutons, a po-tential concern is that the isotopic signature of the fluid, and therefore of Zn-Pb mineralization, may have been affected. In particular, the Jersey and H.B. samples are locally within 0.5 to 3 km from these intrusions. Extensive evidence of recrystallization textures (coarse grained calcite veins and marbles) and small occurrences of calc-silicate minerals observed in hand sample and thin section indicate that these rocks have been affected by contact metamorphism. Furthermore, these plutons produce skarn mineralization at the Jersey Mine, which overprints the previous MVT-style mineralization. Figure 3.14 displays δ18O and δ13C values against distance from the proximal intrusions. In Figure 3.14a there are four samples with a depleted δ18O signature within 3 km of a pluton (up to 6.41 ‰). However, there is also a sample at a distance of 10 km that has a depleted δ18O signature of 9.98 ‰. The remaining samples plot in a constant range of 10.00 ‰ to 25.00 ‰ δ18O regardless of distance. In Figure 3.14b there is a cluster of strongly δ13C depleted samples at 2 to 3 km distance. However, these samples were collected from the Active Formation and represent those that also show a significant δ13C depletion in figures 3.7 and 3.8. The remaining samples again plot in a constant range of 2.00 ‰ to -4.00 ‰ δ13C regardless of distance. Therefore, there is no significant relationship between proximity to a particular pluton and isotopic value. From these data it can be assumed that although these rocks were affected by contact metamorphism generated by the emplacement of Mesozoic plutons, this metamorphic event did not affect the previous isotopic signature of MVT-style Zn-Pb mineralization. As such, this contact metamorphism must have taken place within a closed system with no external fluid interac-tion, at least in the areas examined in this study.75  0.005.0010.0015.0020.0025.0030.000.00 2.00 4.00 6.00 8.00 10.00 12.00d18O_VSMOW ‰ Distance km d18O and Distance from Mesozoic Plutons RegionalJerseyH.B.Remac-18.00-16.00-14.00-12.00-10.00-8.00-6.00-4.00-2.000.002.004.000.00 2.00 4.00 6.00 8.00 10.00 12.00d13C_PDB ‰  Distance km d13C and Distance from Mesozoic Plutons RegionalRegional ActiveFormationJerseyH.B.Remaca. b. Figure 3.14. Plots to display the δ18O and δ13C values of regional, Jersey, H.B. and Remac samples with distance to proximal Mesozoic intrusions. δ18O data (a). δ13C data (b).Vein-Wall Rock PairsFigure 3.15 displays the δ18O and δ13C values of the three distinct vein-types (coarse sparry calcite, coarse sparry calcite + sulphides and tabular calcite veins) and their associated wall rocks. Samples 76plotted (NC21B and C, NC23A and B, NC26B and C, NC27A and B, NC33B and C, HB1008_43A and B, HB1008_48A and B and HB1008_67A and B) are those where microsampling could drill carbonate material from vein and host rock within the same sample. There is no consistent relationship in δ18O values of vein-wall rock pairs; some veins are depleted relative to their wall rock and some are en-riched. However, there is a significant relationship in δ13C values of vein-wall rock pairs. All samples, with the exception of the one tabular calcite vein, display a depletion of roughly 1.00 ‰ to 2.00 ‰ δ13C between wall rock and vein. This suggests that the external fluid that altered the host rocks was isotopically light in δ13C and that the veins have therefore sampled the δ13C depleted wall rocks. As noted above, contact metamorphism generated by the emplacement of proximal Mesozoic plutons did not affect the isotopic signature of the samples and so it is unlikely that these vein-wall rock pairs have been altered. Therefore, these vein-wall rocks pairs represent the original fluid, which would have been isotopically light in δ13C.  -3.00-2.50-2.00-1.50-1.00-0.500.000.5010.00 15.00 20.00 25.00d13C_PDB ‰ d18O_VSMOW ‰ Vein - Wall Rock Pairs          Regional Wall Rock                 Regional Sparry Calcite         Vein         H.B. Wall Rock           H.B. Sparry Calcite Vein +        Sulphides                 H.B. Sparry Calcite Vein         H.B. Tabular Calcite VeinFigure 3.15. Plots to display the δ18O and δ13C values of vein-wall rock pairs.DiscussionApproximately 200 samples were collected from the Active Formation, Nelway Formation, Emerald and Reeves members (Figure 1.4) of the Kootenay Arc for C and O isotopic analysis. These rocks are host to several Zn-Pb deposits, have undergone three phases of folding and have a regional metamor-phic grade of lower greenschist facies (Fyles and Hewitt, 1959). The samples were collected from the Jersey and H.B. mines and from regional roadside outcrops. Mineralization in these rocks is commonly hosted in the Nelway Formation and Reeves member and consists of at least two phases of pyrite; an early diagenetic phase and a later phase of inclusions within coarse grained sphalerites, sphalerite 77and galena. This sulphide assemblage has been recognized through hand sample and thin section descriptions. Additionally, a yellow-orange UV fluorescence emitted from two samples (JS07-08_100 and JS07-08_180) may be attributed to the presence of fine grained cerussite (Pb2+) not visible in hand sample or thin section. This fluorescence may also be attributed to the presence of Mn2+, which has been observed to rim sulphide grains at the Jersey Mine. Sampling of the Active Formation, Nelway Formation, Emerald and Reeves members has enabled the analysis of mineralized and non-mineral-ized host rocks of the Kootenay Arc.Hand sample and thin section observations have enabled the description of six carbonate units: lime-stone, fine grained dolomite, medium grained (pervasive) dolomite, coarse grained dolomite, inter-mediate dolomite and resilient dolomite and three calcite vein-types: coarse sparry calcite, coarse sparry calcite + sulphides and tabular calcite veins. Additionally, cold-cathode CL analysis revealed the presence of a fourth calcite; a fine grained, relatively Mn-rich calcite cement that infills fractures and surrounds sulphide mineralization. The association of this Mn-rich calcite cement with Zn-Pb mineral-ization suggests that the mineralizing fluid had a higher Mn content than the host rock it was reacting with. Source of MnDissolved Mn can be introduced to seawater from a variety of sources: 1) terrigenous runoff, 2) fallout of atmospheric volcanic particles, 3) injection of hydrothermal fluids associated with mid-oceanic ridge volcanism, 4) diffusion from nearshore reduced sediments, 5) diffusion from pelagic sediments and 6) leaching from mafic to intermediate volcanic rocks (1 – 5; Bender, 1977; 6; Klinkhammer, 1977). From this range of different Mn sources, the most common is from seawater leaching Mn from mafic to in-termediate volcanic rocks (Klinkhammer, 1977). Mn is easily mobilized as Mn2+ under anoxic conditions (Hylander et al., 2000) and can easily substitute for Ca2+ or Mg2+ in carbonates. In addition, Mn may be soluble in CO2-rich, acidic, sulphate-bearing fluids (Corbett and Leach, 1997). It is therefore considered that the source of Mn in the Zn-Pb mineralizing fluid could have been from seawater either leaching Mn from proximal volcanic rocks (Lower to Middle Jurassic Elise Formation volcanics) or pelagic sedi-ments (Ordovician Active Formation shale). However, Paradis (2008) and Paradis et al. (2014) recently dated these Salmo-type deposits using Re-Os methods of the sulphide mineralization and produced a Devonian age, which precludes the Jurassic Elise Formation as the source of Mn.Several samples display evidence of contact metamorphism from the proximal Mesozoic intrusions; all of the sparry calcite veins and several limestone or marble samples have been recrystallized. The presence of calc-silicate minerals in thin sections JS07-08_180 and JS07-17_23 also indicate contact metamorphism. Furthermore, there is evidence of skarn mineralization at the Jersey Mine, which is at-tributed to the emplacement of these Mesozoic plutons. Despite this contact metamorphic overprint, there appears to be no isotopic overprint of the original Zn-Pb mineralization. Figure 3.14 displays the δ18O and δ13C values of these samples with distance from Mesozoic intrusions; there is no shift in isotopic signature with proximity to these plutons. As such, it is assumed that the original isotopic signature of Zn-Pb mineralization has been preserved and that any subsequent metamorphism was 78purely thermal and took place within a closed system with no further external fluid interaction, at least in the areas examined in this study.From the samples collected, limestones are the least depleted whereas samples of banded, zebra and pervasive dolomites are the most depleted relative to Cambrian marine carbonate (Wadleigh and Veizer, 1992; Maloof et al., 2005; Saltzman, 2005; Dilliard et al., 2007; Prokoph et al., 2008). Samples of the intermediate and resilient dolomites display an enrichment compared to other samples in both δ18O and δ13C; however, these carbonate units are only observed in close proximity to the Remac Mine which will be discussed in Chapter 4. Additionally, there is a cluster of analyses that are strongly depleted in δ13C; these samples are from the Ordovician Active Formation, which is a unit of carbo-naceous shales and phyllitic limestones. These samples, although taken in several different locations, represent a s