Open Collections

UBC Theses and Dissertations

UBC Theses Logo

UBC Theses and Dissertations

Characteristics and formation of the Jeronimo sedimentary rock-hosted disseminated gold deposit, Atacama… Gale, Vanessa Genevieve 1999

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

Item Metadata

Download

Media
831-ubc_2000-0067.pdf [ 10.8MB ]
Metadata
JSON: 831-1.0089378.json
JSON-LD: 831-1.0089378-ld.json
RDF/XML (Pretty): 831-1.0089378-rdf.xml
RDF/JSON: 831-1.0089378-rdf.json
Turtle: 831-1.0089378-turtle.txt
N-Triples: 831-1.0089378-rdf-ntriples.txt
Original Record: 831-1.0089378-source.json
Full Text
831-1.0089378-fulltext.txt
Citation
831-1.0089378.ris

Full Text

CHARACTERISTICS AND FORMATION OF THE JERONIMO SEDIMENTARY ROCK-HOSTED DISSEMINATED GOLD DEPOSIT, ATACAMA REGION, CHILE  by VANESSA GENEVIEVE GALE B.Sc, Dalhousie University, 1997 A THESIS SUBMITTED IN PARTIAL FULFILLMENT OF THE REQUIREMENTS FOR THE DEGREE OF MASTER OF SCIENCE in THE FACULTY OF GRADUATE STUDIES (Department of Earth and Ocean Sciences) We accept this thesis as conforming to the required standard  THE UNIVERSITY OF BRITISH COLUMBIA October 1999 © Vanessa Genevieve Gale, 1999  U B C Special Collections - Thesis Authorisation Form  http://www.library.ubc.ca/spcoll/thesauth.html  In p r e s e n t i n g t h i s t h e s i s i n p a r t i a l f u l f i l m e n t o f the requirements f o r an advanced degree a t the U n i v e r s i t y o f B r i t i s h Columbia, I agree t h a t the L i b r a r y s h a l l make i t f r e e l y a v a i l a b l e f o r r e f e r e n c e and study. I f u r t h e r agree t h a t p e r m i s s i o n f o r e x t e n s i v e c o p y i n g o f t h i s t h e s i s f o r s c h o l a r l y purposes may be g r a n t e d by the head o f my department o r by h i s o r h e r r e p r e s e n t a t i v e s . I t i s u n d e r s t o o d t h a t copying or p u b l i c a t i o n of t h i s t h e s i s f o r f i n a n c i a l gain s h a l l not be a l l o w e d without my w r i t t e n p e r m i s s i o n .  Department o f  Bgrfln Cj AJ Q(e£tn Su<ZrU*^  The U n i v e r s i t y o f B r i t i s h Columbia Vancouver, Canada Date  l of l  12/14/99 8:35 A M  Abstract  The Jeronimo sedimentary rock-hosted disseminated gold deposit is located within the Potrerillos district of the Atacama region of northern Chile, to the east of the Potrerillos-El Hueso porphyry Cu-Mo-(Au) system of Eocene to Oligocene age. Jeronimo occurs within a zone of thrust and reverse faulting and folding associated with middle Tertiary compressional deformation along the Domeyko fault system. The deposit contains a resource of 16.6 million tonnes of gold averageing 4.80 grams/tonne. Jeronimo is manto-shaped, and is on average six metres in thickness. In plan view, it is elliptical and measures approximately 2.0 by 1.3 km. It is separated into two zones known as upper and lower Jeronimo, which are offset by subvertical normal faults. Jeronimo is stratabound, occurring within bioclastic limestone lithologies of the Jurassic El Asientos Formation limestone sequence. Ore distribution is also structurally-controlled, as mineralization is focussed in sub-vertical fractures and joints within the bioclastic units. Alteration phases include (1) strong, pervasive, replacement-style silicification; (2) carbonate, mainly restricted to vugs, consisting of manganese carbonate (rhodochrosite and kutnohorite), in the centre of the ore body, and calcite-dolomite on the margins; (3) argillization consisting of widespread disseminated and veinlet illite, and vug-filling kaolinite in the centre of the deposit. Other common alteration minerals include apatite, rutile, monazite and barite. The ore mineral suite at Jeronimo consists of pyrite, arsenopyrite, sphalerite, lead sulphosalts, orpiment, and realgar, with minor coloradoite, altaite, cinnabar and cassiterite. Jeronimo is characterized by the geochemical signature As-Mn-Zn-Pb-Ag-Hg. Gold is present as generally sub-micron-sized grains, ranging from 140 nm to 1.13 um, that are encapsulated in pyrite, arsenopyrite, quartz and realgar and also occur within vugs in the silicified matrix. Lead isotope results of the main stage sulphide and sulphosalt minerals ( Pb/ Pb: 18.564 to 18.644; Pb/ Pb: 15.592 to 15.662; and Pb/ Pb: 38.536 to 38.638) indicate that lead in the ore fluids was dominantly derived from a magmatic source, with some input from a more radiogenic source - igneous Carboniferous to Triassic basement rocks and/or the overlying Jurassic limestone and sandstone. The lead isotope composition of late orpiment indicates that lead in the late fluids was derived mainly from the Jurassic sedimentary host rocks. Carbon and oxygen isotope compositions of ore zone rhodochrosite and kutnohorite, ranging from8 0 of 16.65 to 22.52% (VSMOW) and 8 C of -2.84 to -1.3% (PDB), suggest contributions from both magmatic and Jurassic limestone wallrock sources. 206  207  204  208  204  204  18  13  0  0  The Jeronimo deposit is a gold-rich carbonate replacement deposit (CRD) showing distinct differences from sediment-hosted (Carlin-type) gold deposits. The critical features that define the CRD style of mineralization at Jeronimo include enrichment in base metals and manganese, carbonate alteration, the presence of native gold as visible grains, occurrence in a district containing porphyry and related styles of mineralziation, and evidence for a magmatic contribution to metals and hydrothermal fluids. A direct connection between Jeronimo and the Potrerillos porphyry Cu-Mo-(Au) deposit 4.3 kilometres to the west-northwest has not been established, but may be suggested by the existence of alteration and minor mineralization extending west from Jeronimo in favourable structures and lithologies. Alternatively, Jeronimo may be related to another concealed magmatic-hydrothermal center at depth.  Table of Contents  Abstract Table of contents  ii iii  List of figures  v  List of tables  vi  List of plates  vi  Acknowledgements Chapter 1: Introduction 1.1 Purpose 1.2 Methods 1.3 Thesis organization 1.4 Exploration history in the Potrerillos district 1.5 Previous work  viii 1 4 5 6 6 7  Chapter 2: Geological Framework 2.1 Regional geolgy 2.1.1 Pre-Jurassic record 2.1.2 Jurassic to early Cretaceous record 2.1.3 Middle Cretaceous to Cenozoic record 2.2 Geology of the Potrerillos district 2.3 Mineralization in the Potrerillos district 2.3.1 Potrerillos porphyry Cu-Mo-(Au) deposit 2.3.2 El Hueso sedimentary and volcanic rock-hosted disseminated gold deposit 2.3.3 Jeronimo sedimentary rock-hosted disseminated gold deposit 2.3.3.1 Sedimentary host rocks 2.3.3.2 Structural setting 2.4 Relationships among mineralized centres in the Potrerillos district  9 9 9 10 10 11 15 15 17 20 20 23 25  Chapter 3: Jeronimo Gold Deposit 3.1 Stratigraphic and structural controls on ore distribution 3.1.1 Lithologic control 3.1.2 Structural control 3.2 Hypogene alteration 3.2.1 Silicification 3.2.1.1 Distribution 3.2.1.2 Characteristics 3.2.1.3 Mineralogy  26 26 .26 42 48 48 48 49 55  iv 3.2.2 Carbonatization 3.2.2.1 Distribution 3.2.2.2 Characteristics 3.2.2.3 Carbonate mineral identification and composition 3.2.2.4 Mineralogy 3.2.3 Argillization 3.2.3.1 Distribution and characteristics 3.2.3.2 Mineralogy 3.2.3.2a Kaolinite 3.2.3.2b mite 3.2.3.3 Spatial variation in clay mineral distribution 3.2.4 Organic matter 3.2.5 Hypogene alteration relationships 3.3 Ore mineralogy 3.3.1 Mineral characteristics 3.3.2 Modes of gold occurrence 3.4 Geochemistry 3.5 Rare minerals 3.5.1 Aluminum-phosphate-sulphate minerals 3.5.2 Rare earth element-bearing calcite 3.5.3 Manganese sulphide minerals 3.6 Supergene weathering 3.7 Interpretation of deposit mineralogy and geochemistry 3.8 Summary  58 58 60 67 72 73 73 76 78 84 87 90 90 92 93 < • 106 113 124 124 126 128 131 133 136  Chapter 4: Isotopic studies of the Jeronimo deposit 4.1 Oxygen and carbon isotopes 4.1.1 Sample preparation, analysis and results 4.1.2 Interpretation 4.2 Lead isotopes 4.2.1 Lead isotope systematics 4.2.2 Sampling and analysis 4.2.3 Results.... 4.2.4 Interpretation 4.3 Conclusions  138 138 139 144 152 153 155 157 161 164  Chapter 5: Interpretation 5.1 Summary of Jeronimo 5.1.1 Strati graphic and structural controls 5.1.2 Alteration mineralogy and distribution 5.1.3 Mineralization 5.1.4 Isotopic compositions 5.2 Comparison with Carlin and gold-rich carbonate replacement deposits 5.3 Metallogenic relationships in the Potrerillos district 5.4 Genetic model for Jeronimo 5.4.1 Timing of Jeronimo ore deposition  166 166 166 168 169 170 171 174 178 178  V  5.4.2 Hydrothermal fluid origins 5.4.3 Fluid pathways 5.4.4 Depositional history 5.5 Concluding statement  179 180 181 184  References Appendix A: Appendix B: Appendix C: Appendix D: Appendix E:  187 Carbonate electron microprobe data Clay separation method Ore zone geochemical data Manganese sulphide mineral electron microprobe data Existing lead isotope data for the Potrerillos district  194 201 203 219 222  List of Figures 1.1 1.2 2.1 2.2  Map of Chile with location of Potrerillos district Geological map of the Potrerillos district Geological map of the Jeronimo and El Hueso region Stratigraphic column of the Mesozoic sedimentary formations in the vicinity of the Jeronimo deposit 3.1 Plan view of Jeronimo sulphide body, with GT contours, drill hole locations examined in this study, and cross-section lines 3.2 Diamond drill hole logs of ore zone intercepts used in this study 3.3a,b,c Cross sections through the Jeronimo sulphide body 3.4 Plan view of carbonate mineral distribution 3.5 Representative carbonate mineral X-ray diffraction spectra 3.6a,b Ternary plots of calcite-group and dolomite-group carbonate mineral electron microprobe analyses 3.7 Representative clay mineral X-ray diffraction spectra 3.8 Short-wave infrared reflectance spectra for kaolinite and dickite 3.9 Plan view of clay mineral distribution 3.10 Variation of selected elements across ore intercepts with respect to gold 3.11 Distribution of manganese carbonate minerals 4.1 Mineral 8 0 and 8 C isotopic compositions 4.2 Dominant carbon species as a function of pH and temperature 4.3a,b H 0 and H C0 8 O and 8 C compositions of fluids in equilibrium with host rock and ore zone carbonate minerals 4.4 Uranogenic Pb diagram 4.5 Thorogenic Pb diagram 5.1 Cross section through the Jeronimo and El Hueso deposits 5.2 Model of Jeronimo deposit formation 18  13  ls  2  2  2 3 12 22 27 34 43 59 68 71 79 82 88 117 129 145 148  13  3  150 159 160 175 183  vi  List of Tables 2.1 2.2 2.3 3.1 3.2 3.3 3.4 3.5 3.6 3.7 3.8 4.1 4.2 4.3 4.4 4.5 4.6 4.7 4.8 4.9 5.1 5.2  Paleocene to Eocene deformational events in the Potrerillos district Ar/ Ar results of Marsh (1997) Characteristics of the El Asientos formation units Example X-ray diffraction peaks for rhodochrosite and kutnohorite Example X-ray diffraction peaks for illite and kaolinite Empirically derived values of °A20 for known percentages of illite Calculated values of °A20 and Kubler indices for Jeronimo clay samples Characteristics of gold in the Jeronimo deposit Average geochemical values for Jeronimo ore Correlation of selected elements with respect to gold Minerals responsible for geochemical trends Sample characteristics of ore zone manganese carbonate minerals and 5 0 and 8 C results Sample characteristics of host rock calcium carbonate minerals and 5 0 and 8 C results Fractionation factors used for calculation of fluid 8 O and 8 C compositions Calculated Ff 0 and H C0 8 0 and 8 C compositions 8 0 and 8 C values of potential source reservoirs for Jeronimo fluids Sample characteristics of ore minerals analyzed for lead isotope composition Sample characteristics of carbonate and feldspar analyzed for lead isotope composition Ore mineral lead isotope results Carbonate and feldspar lead isotope results Characteristics of Carlin and carbonate replacement deposits Distinguishing features of Carlin and carbonate replacement deposits  40  39  18  I3  18  13  ls  18  2  18  2  13  13  3  13  15 20 21 69 80 86 86 114 115 116 123 140 141 149 151 152 156 156 162 162 173 185  List of Plates 3.1 Hand sample of bioclastic limestone, unit C 3.2 Thin section photo of bioclastic limestone, unit B 3.3 Thin section photo of bioclastic limestone, unit B 3.4 Thin section photo of sandy limestone 3.5 Hand sample photo of strongly silicified ore 3.6 Thin section photo of pervasive silicification 3.7 SEM BSE image of quartz veinlet 3.8 Thin section photo of silicified bioclasts 3.9 Thin section photo of silicified bioclasts 3.10 SEM BSE image of illite, apatite and lead sulphosalt in vug 3.11 Thin section photo of intergrown rutile and pyrite 3.12 Hand sample photo of disseminated kutnohorite  29 29 32 32 51 51 52 54 54 56 57 62  vii 3.13 3.14 3.15 3.16 3.17 3.18 3.19 3.20 3.21 3.22 3.23 3.24 3.25 3.26 3.27 3.28 3.29 3.30 3.31 3.32 3.33 3.34 3.35 3.36 3.37 3.38 3.39 3.40 3.41 3.42  Hand sample photo of rhodochrosite and orpiment veining Thin section photo of bioclast replaced by quartz and manganoan calcite Thin section photo of disseminated manganese carbonate mineralization SEM BSE image of rhodochrosite and kutnohorite aggregate Hand sample photo of pervasive argillization Thin section photo of patchy clay in quartz matrix Thin section photo of irregular clay veinlet SEM BSE image of kaolinite in vugs in quartz and carbonate aggregates SEM BSE image of intergrown illite and apatite Thin section photo of fine grained black alteration product Thin section photo of pyrite and rutile in diffuse vein SEM BSE image of arsenopyrite inclusions in vuggy pyrite SEM BSE image of arsenopyrite rim on pyrite grain SEM BSE image of diffuse arsenopyrite, rutile and pyrite vein Hand sample photo of ore with disseminated sphalerite and realgar veining Thin section photo of diffuse sphalerite and pyrite vein SEM BSE image of lead sulphosalt and realgar in vugs in quartz Hand sample photo of orpiment, realgar and calcite mineralization SEM BSE image of realgar, lead sulphosalt and rutile aggregate SEM BSE image of gold grain in realgar SEM BSE image of gold grain in illite-filled vug in quartz SEM BSE image of gold grain in vug in quartz SEM BSE image of gold grain encapsulated in quartz SEM BSE image of gold grains as inclusions in pyrite SEM BSE image of gold grain as inclusion in arsenopyrite SEM BSE image of gold grain in vug in manganese oxide mineral SEM BSE image of aluminum-phosphate-sulphate mineral SEM BSE image of rare earth element-bearing calcite Hand sample photo of alabandite aggregates Hand sample photo of oxidized ore  62 65 65 66 75 75 77 83 88 91 94 95 96 97 100 100 102 103 104 105 107 108 109 110 Ill 112 125 127 130 132  viii  Acknowledgements Funding for this project was provided by the National Science and Engineering Research Council, in the form of a P G S A scholarship, and by Homestake M i n i n g Inc. who funded analyses, room and board in Chile and airfares. I would like to thank my advisor Dr. John Thompson for his enthusiasm, guidance and lightheartedness. A l l those committee members and other providers of scholarly advice, including J i m Mortensen, Greg Dipple, D i c k Tosdal, Anne Thompson, Steve Rowins, Jim Lang and T i m Baker deserve a round of applause. I would like to thank Arne Toma and Marc Baker for handling my frequent computer crises and Bryon Cranston for being so helpful Thanks to all of the Homestake geologists at Agua de la Falda, including George Shroer, Devin Denboer, Benjamin Sanfurgo and John James, who found time for me in their hectic schedules and who made life in Potrerillos that much more interesting. Thanks to N i c k Callan for making me smile in the Andes in winter. I greatly appreciated the interest and efforts of other Homestake geologists including Richard Bedell, Dave Hendricks, B i l l Wright, R o n Parratt and Nathan Brewer. Cheers to James M c D o n a l d for thinking up the idea in the first place! A special, warm thank you to the bottom-dwellers Laurel Basciano, Brent Nassichuk and Pier Pufahl for making the basement the place to be. I never knew procrastination could be quite so creative. Thanks to Cari Deyell for being so sweet and always knowing just what to say. T o everyone else who made life more joyful for me at U B C , including M i k e St.Pierre, T o m M i l n e , E m i l y Chastain, Jen Hobday, Patrick Williston, Susannah Price, R i c h Harris, the Damn Bastards, the Stress Tensors and the Lunch Club, I salute you! Final words of appreciation for my family, especially my mum, for putting up with me and to A l a n , for making me laugh during the last dark days of thesis creation.  1 Chapter 1: Introduction  The Potrerillos district (26° 30' S, 69° 24' W), located in the Atacama region of northern Chile (Figure 1.1), hosts several types of mineral deposits, including porphyry copper, skarn and disseminated gold. One deposit within the district, the Jeronimo sedimentary rock-hosted disseminated gold deposit, discovered in 1995, is currently in production by Agua de la Falda S.A. (ADLF), a joint venture between Minera Homestake Chile S.A. and the Corporation National del Cobre (CODELCO). Jeronimo is situated in the footwall of a major thrust fault, the Potrerillos Mine fault (PMF), which is also known as the Silica Roja fault in the Jeronimo area (Figure 1.2). About 1.5 kilometres to the west-northwest, in the hangingwall of the PMF, is the El Hueso sedimentary and volcanic rock-hosted disseminated gold deposit. Approximately 4.3 kilometres west-northwest of Jeronimo, also in the hangingwall of the PMF, is the Potrerillos porphyry Cu-Mo-(Au) deposit. As of December 31, 1998, Jeronimo contained a resource of 16,600,000 tonnes averaging 4.80 grams/tonne gold. The deposit, which is stratabound and consists of replacement-style mineralization, is hosted by a sequence of Jurassic carbonate rocks and occurs preferentially in bioclastic limestone beds. Jeronimo hosts zones of both oxide and sulphide mineralization. Gold in the sulphide zone occurs as generally sub-micronsized, anhedral grains that are present either as inclusions in pyrite, arsenopyrite, realgar and quartz or within vugs in the quartz matrix.  PERU  20° Quebrada Blanca Collahuasi El Abra Chuquicamata  BOLIVIA La Escondida  •25  c  El Salvador  — Potrerillos COPIAPO  30°  M5  •  /-sr  ?  1 H u e s o  Jeronimo  Miocene Maricunga Belt gold-rich porphyry and epithermal deposits  C  •40°  ARGENTINA D  Cities Late Eocene to Early Oligocene porphyry copper deposits  45°  Potrerillos porphyry copper deposit, El Hueso disseminated gold deposit, Jeronimo disseminated gold deposit of the Potrerillos district  Figure 1.1 Map of Chile showing the location of the Potrerillos, El Hueso and Jeronimo deposits of the Potrerillos district. Also shown for reference are the Miocene Maricunga gold belt and late Eocene to early Oligocene porphyry copper deposits, located mainly along the Domeyko fault zone.  3  • 7,O7O,00ON 2630'S -  69"15'W Strike and dip of bedding or flow foliation Normal fault, bar on downthrown side Thrust fault, barbs on upper plate Potrerillos Mine fault, barbs on upper plate Steep reverse fault, barbs on hangingwall  Granodiorite porphyry (unmincralized) Granodiorite porphyry (mineralized) Microdioritc - fine-grained -hornblendc-pyroxene-biotite dioritc  t  N  Dacitic breccia, welded ash tuff, dacitic lava flow Andcsitc breccia and flows, andesitic sedimentary rocks, and welded rhyolitic tuff Agua Helada and Pcdcrnalcs formations - nonmarine sandstone, marine bioclastic limestone and marine calcareous sandstone El Asientos and Montandon formations - calcareous siltstonc, sandstone, shale, oolitic limestone and gyspum  Kilometres  Triassic granodiorite of the Pcdcrnalcs Complex - coarse-grained granodiorite to quartz monzonite; La Tabla Formation - rhyolitic tuffs and breccias, and andesitic dikes and flows  Figure 1.2 Generalized geologic map of the Potrerillos district showing the location of the Potrerillos porphyry Cu-Mo-(Au) deposit, El Hueso disseminated Au deposit and Jeronimo disseminated Au deposit. ModifiedfromMarsh (1997).  4  1.1 Purpose The goals of this study are: (1) to characterize and classify the style of mineralization at the Jeronimo deposit; (2) to determine its conditions of formation; and (3) to define its probable relationships to the other styles of mineralization in the Potrerillos district. Further understanding of the genesis of the Jeronimo deposit will aid in the identification of new exploration targets in the Potrerillos district. Mineralization at Jeronimo is hosted by discrete bioclastic limestone horizons and therefore may be defined as stratabound. Two types of sedimentary rock-hosted deposits most fit the characteristics of Jeronimo, the Carlin-type deposits common to the Great Basin region of the United States and intrusion-related carbonate replacement deposits. Carlin-type deposits are replacement-style disseminated gold deposits that are lithologically and structurally controlled, with mineralization that is concentrated where faults intersect favourable horizons, typically silty limestones and calcareous siltstones. Alteration types in Carlin-type deposits consist of decarbonatization, silicification and argillization, and mineral assemblages commonly include pyrite, orpiment, realgar, arsenopyrite, stibnite, cinnabar and barite. Carlin-type deposits are characterized by the occurrence of gold as sub-micron-sized inclusions or structurally bound ions within arsenian pyrite. Although almost all Carlin-type deposits either host or occur proximally to intrusive rocks, genetic relations between the two are not clear. Carbonate replacement deposits (CRDs) are manto-shaped and, like Carlin-type deposits, are lithologically and structurally controlled, occurring mainly in silty and sandy limestones. They show similar alteration types as Carlin-type deposits, including decalcification, silicification and argillization. Unlike Carlin-type deposits, spatial  5  relations, fluid inclusion compositions and isotopic evidence show that CRDs are genetically related to intrusions. They commonly occur in association with other intrusion-related deposits, such as porphyry and skarn deposits (Sillitoe, 1991a). Although CRDs are normally exploited for base metals, a variety of CRD is gold-rich, with only minor base metal enrichment. Gold rich-CRDs are mineralogically variable but may host a sulphide mineral assemblage including pyrite, arsenopyrite, sphalerite, galena, chalcopyrite, marcasite, pyrrhotite and are enriched in As-Sb-Pb-Zn ± Ag-Mn-TlTe-Cu-Hg-Bi-Sn.  1.2 Methods At the mine site, 19 drill core intercepts of the Jeronimo ore zone were logged and sampled. Additionally, portions of 6 drill holes extending up to 1.7 kilometres away from the western margin of the deposit were logged and sampled to assess the distal effects of Jeronimo mineralization. Mineral identification and examination of textures were performed at the University of British Columbia with the use of a transmitted and reflected light microscope, X-ray diffractometer and scanning electron microscope with an energy dispersive spectrometer. Mineral compositions were quantitatively determined using an electron microprobe with a wavelength dispersive spectrometer. Reconnaissance isotopic studies were performed in order to characterize the sources of the Jeronimo hydrothermal fluids. The carbon and oxygen isotopic compositions of ore zone and host rock carbonate minerals were measured at the Stable Isotopes Lab at Queen's University, Kingston, Ontario. The lead isotope compositions of  6  ore zone sulphide minerals, host rock carbonates, and feldspar from a local porphyry intrusion were measured at the Geochronology Lab at the University of British Columbia. Although a geochronological framework has been established for magmatism in the Potrerillos District, mainly through A r / A r dating by Marsh (1997)', the age of three 4 0  3 9  felsic porphyry intrusions in proximity to the Jeronimo deposit were determined in the present study based on analyses of zircon by the U-Pb method.  1.3 Thesis organization Following chapter one, which continues with summaries of the exploration history and previous work conducted in the Potrerillos district, chapter two presents the regional, district and local geologic setting of the Jeronimo deposit. Chapter three describes in detail the characteristics of the Jeronimo deposit, including lithologic and structural controls, alteration types and distribution, sulphide mineral characteristics, modes of gold occurrence and trends in geochemistry. Chapter four presents the results of the stable light isotope and lead isotope studies. Chapter five summarizes the Jeronimo characteristics and compares Jeronimo to Carlin-type deposits and CRDs. It also discusses district-scale metallogenic associations, suggests a model of ore formation for Jeronimo, and presents the conclusions of the study.  1.4 Exploration History in the Potrerillos District The earliest recorded mining activity in the Potrerillos district dates back to 1894, when local miners worked the seams and veins of the Potrerillos Cu-Mo-(Au) porphyry deposit, hosted in the Cobre stock (Figure 1.2) (Olson, 1983). Large-scale mining of this  7  deposit by the Andes Copper Company, a subsidiary of the Anaconda Copper M i n i n g Company, took place from 1926 to 1959, during which 169 million metric tonnes of sulphide ore with a mean grade of 1.1 % copper and 39.2 million metric tonnes of oxidized ore with a mean grade of 1.07 % copper were extracted (Reyes, 1981; cited in Olson, 1983). B y the early 1930s, the Andes Copper Company began using silicified volcanic rocks from a location about 2.8 kilometres to the east-southeast as flux for the Potrerillos smelter (Hernandez et al., 1991). High gold values in the smelter product were eventually traced to the silica flux and gold exploration was begun by C O D E L C O in 1982, leading to the discovery of the E l Hueso sediment- and volcanic-hosted disseminated gold deposit (Hernandez et al., 1991). Estimates of initial reserves for open pit mining were 16 million metric tonnes of oxidized ore averaging 1.68 grams/tonne gold and 7.8grams/tonne silver. E l Hueso was mined from 1986 to 1995 by Minera Homestake Chile S.A. under lease from C O D E L C O . In an effort to replace the depleting reserves of the E l Hueso mine, exploration by Homestake geologists in 1994 led to the discovery of the strata-bound Jeronimo oxide body, which is expressed at surface about 1.5 kilometres to the east-southeast of E l Hueso. Further drilling-based exploration led to the discovery of the Jeronimo sulphide body in 1995. Jeronimo is currently being exploited by a joint Homestake and C O D E L C O company known as Agua de la Falda, S.A ( A D L F ) .  1.5 Previous Work The stratigraphic sequences in the Potrerillos district were first named and described by Harrington (1961) and have been re-defined by Perez (1977). The first  8  regional map of the district was produced by Frutos and Tobar in 1973. The geology of the district was remapped and interpreted by Olson (1983), whose study remains the most comprehensive of the area. Dilles (1995) described the stratigraphy of the Jurassic carbonate units in proximity to Jeronimo and subdivided the Montandon and E l Asientos formations, the latter of which hosts the Jeronimo deposit, into units A through G . These subdivisions are currently used by Homestake geologists. Lewis (1996) mapped the stratigraphic sequences, intrusive stocks and structural features of the Jeronimo area. Elgueta and Fuentes (1997) examined the stratigraphy of the E l Asientos Formation and determined that it was deposited within a back arc or intra-arc basin with facies typical of epicontinental open sea and mixed clastic-carbonate platform environments. Marsh (1997) established a geochronological framework for mineralization within the district by conducting A r / A r studies on the E l Hueso and Potrerillos deposits and on porphyries 4 0  3 9  in the area. He determined that mineralization at E l Hueso was not related to the emplacement of the Cobre stock, instead suggesting that it was associated with a more proximal, concealed stock. The surficial geology of the A D L F property has been most recently mapped by Callan (1998). Petrographic studies of the Jeronimo ore and host rocks were performed by Honea (1997) and Wilson (1997). The most recent study of Jeronimo was performed by Caddey (1999), who mapped the structural geology in the first underground exposure of the Jeronimo sulphide body in a tunnel opened in late 1998.  9 Chapter 2: Geological Framework  Several geologic events controlled the occurrences and styles of mineralization in the Potrerillos district. Late Mesozoic basin opening resulted in the deposition of thick limestone sequences which host several centres of mineralization, including Jeronimo, El Hueso and skarn associated with the Potrerillos porphyry deposit. Early Tertiary compression caused folding, and thrust and reverse faulting of the limestone and overlying volcanic units and laid the groundwork that would largely control subsequent fluid flow in the district. This chapter briefly summarizes the regional geological features that influenced the formation of the Jeronimo deposit and discusses the geology and relationships between mineralized centres in the Potrerillos district. For comprehensive accounts of Andean geologic history the reader is referred to Mpodozis and Ramos (1990) and Davidson and Mpodozis (1991).  2.1 Regional geology 2.1.1 Pre-Jurassic record Chile's pre-Jurassic basement is comprised of accreted Precambrian to early Paleozoic allochthonous and displaced terranes, and autochthonous late Paleozoic terranes of the Gondwana continent (Mpodozis and Ramos, 1990). In northern Chile, exposed basement consists of medium- to high-grade metamorphic rocks intruded by upper Paleozoic and Triassic plutons and locally overlain by Triassic silicic volcanic rocks (Cornejo et al., 1993).  10  2.1.2 Jurassic to early Cretaceous record The tectonic regime governing the Andean margin during the Jurassic to early Cretaceous was extensional (Mpodozis and Ramos, 1990). A t this time, subduction of the A l u k plate resulted in the establishment of a calc-alkaline magmatic arc along the entire length of the margin. In northern Chile, subduction-related extension induced the opening of the Tarapaca back-arc basin, in which several thousand metres of carbonate and terrigenous sediments accumulated (Mpodozis and Ramos, 1990).  2.1.3 Middle Cretaceous to Cenozoic record Between 100 and 80 M a , the onset of sea floor spreading in the South Atlantic caused an increase in the rate of westward motion of the South American plate (Davidson and Mpodozis, 1991). This acceleration induced the closure, collapse and uplift of the Tarapaca basin, creating a morphogeological high referred to as the Domeyko protocordillera (Mpodozis and Ramos, 1990). Following a lull during the Late Cretaceous, calc-alkaline magmatism resumed during the Tertiary in northern Chile and recorded the eastward migration of the foci of arc magmatism (Davidson and Mpodozis, 1991). A t this time, oblique convergence induced strike-slip motion along the axis of the Domeyko protocordillera, forming the regional north-striking Domeyko fault system. Near the end of oblique convergence, during the late Eocene to early Oligocene, the large porphyry copper deposits of northern Chile were emplaced syntectonically along transtensional and transpressional structures linked with the Domeyko fault system (Davidson and Mpodozis, 1991; Tomlinson et al., 1994).  11  2.2 Geology of the Potrerillos district Pre-Jurassic basement assemblages in the Potrerillos district include Carboniferous to Triassic granite, diorite, tonalite and quartz syenite as well as metavolcanic rocks and minor quartz-muscovite schists (Olson, 1983). During the Jurassic, sedimentation in the Tarapaca back-arc basin resulted in the deposition of the Montandon Formation calcareous mudstones and limestones on the PreJurassic basement (Olson, 1983). This formation grades into the mixed carbonate-clastic sedimentary rocks of the E l Asientos Formation, which hosts the Jeronimo deposit, and is conformably overlain by the siltstones and limestones of the Cretaceous Pedernales Formation (Figure 2.1). The closure and uplift of the Tarapaca back-arc basin, which began in response to the middle Cretaceous shift from extensional to compressional tectonics, resulted in the formation of the regional Hornitos unconformity and locally exposed E l Asientos Formation rocks. During the Paleocene to Eocene, the Hornitos Formation andesitic lava flows, conglomerates, rhyolitic ash-flow tuffs and rhyolite flows were deposited unconformably over the E l Asientos and Pedernales formations (Olson, 1983). The source of these volcanic rocks has not been identified, however their tendency to thicken to the west suggests that their source may be in that direction (Olson, 1983). In the Potrerillos district, the Eocene to Oligocene transpression initiated sinistral, reverse movement on the major thrust known as the Potrerillos M i n e fault ( P M F ) (Figure 1.2). The Cobre porphyry, which hosts the Potrerillos porphyry C u - M o - ( A u ) deposit, was emplaced syntectonically with movement along the P M F (Tomlinson, 1994;  12  3 00C8St7 ^  13  Quaternary alluvium, mine dumps and leach pads |  Tertiary feldspar-biotite-quartz-hornblende porphyries Medium- to coarse-grained magnetite-bearing gabbro to diorite  Tertiary Hornitos Formation Aphanitic to plagioclase-porphyritic, massive to flow-banded, locally vesicular andesite flows Locally moderately to strongly welded, lapilli-ash tuffs with crystals of plagioclase and biotite and lithicfragmentsrangingfromlapilli-sized pumice to breccia-sized andesite Coarse, porphyritic andesite cobble-pebble clast conglomerate; thinly bedded epiclastic sandstones, siltstones, and pebbly to tuffaceous wackes  Cretaceous Pedernales Formation Interbedded fossiliferous, bioclastic and oolitic limestones, sandy limestones and calcareous shale  Jurassic El Asientos Formation Jag  Thin to lenticular, bedded marls, carbonaceous limestones andfine-grainedsandy limestones  Jaf  Coarsely fossiliferous limestone; thin to medium bedded oolitic limestone withfinefossil hash  Jae  Fine- to coarse-grained sandy limestones and calcarenites with intercalated bioclastic, dolomitic and oolitic limestones  Jad  Medium- to coarse-grained quartz arenite  Jac  Q  Coarse, black biosparite with continuous fossil hash and local coquina horizons; main ore host  Jab  Thinly and wavy bedded,fine-grainedsandy limestone with chert matrix, nodules and bands Jurassicfine-grainedplagioclase-porphyritic basaltic andesite sills and possible flows in unit Jag  ^  Hornitos unconformity  y  ^^**» Property boundary  ' High-angle normal fault  .••'*'  Low-angle reverse fault  ^  r 1  Fold axis  o f  m a  e d  Approximate limits of sulphide and oxide ore bodies (>1 gram/tonne Au)  y High angle reverse fault Fault trace  PP ^ ° Cross section line L i m i t  "  Strike and dip of bedding  14  Mpodozis et al., 1994). The emplacement and mineralization of the Potrerillos porphyry has been dated at 35.6-35.9 M a by the A r / A r method (Marsh, 1997). More than ten 3 9  4 0  granodioritic, quartz monzodioritic to tonalitic plagioclase-porphyritic intrusions were emplaced in the Potrerillos district from the late Eocene to early Oligocene, between 40.8 and 32.6 M a (Marsh, 1997). Three of the porphyries dated by Marsh (1997), which are proximal to the Jeronimo deposit and occur within the E l Hueso open pit (Figure 2.1), were also dated in the present study by the U - P b analysis of zircon at the Geochronology Lab of the University of British Columbia. Results from two of the porphyries, Silica North and East, are in agreement with the results of Marsh (1997) and yield ages of ca. 40 M a and ca. 41 M a , respectively. The date obtained by Marsh (1997) for the third porphyry, Silica South, of 47.03 ± 0.07 M a was discordant and unreliable. U - P b analysis of zircon in this porphyry yields a more reasonable age of 39.2 ± 0.2 M a and indicates that the emplacement of this porphyry is similar in age to the others in the vicinity. In the Potrerillos district, Eocene to Oligocene transpression linked with deformation along the Domeyko fault system, was manifested as asymmetric folding and thrust and reverse faulting (Tomlinson et al., 1994). Caddey (1999) has subdivided this period of deformation into three phases, summarized in Table 2.1, which are marked by the progressive clockwise rotation of the principal stress direction (Oi). Deformation during phases D l a and D l b was brittle-ductile and initially involved the formation of asymmetric folds that, with progressive deformation, were sheared along their flat limbs to form thrust faults and along their steep limbs and axial planes to form reverse faults. Continued compression caused drag folding between thrust planes and resulted in the formation of an imbricated fold-thrust package. B y D i e , the principal stress direction  15  ( G i ) was oriented N 6 0 W - S 6 0 E and deformation during this phase was brittle. In the Potrerillos district, this phase was marked by the formation of conjugate shears, extension and release fractures related to extension in the direction of G . 3  Table 2.1 Summary of late Paleocene to Early Oligocene deformational events in the Potrerillos district. Summarized from Caddey (1999).  Dla Dlb  Principal stress direction E-W N73W-S73E  Die  N60W-S60E  Phase  Resultant structures Brittle-ductile deformation consisting of initial asymmetric folding. Progressive compression causing thrust faulting along flat fold limbs and reverse faulting along steep fold limbs and fold axial planes. Subsequent drag folding related to thrusting and formation of imbricated fold/thrust belt. In both events, o is vertical. Brittle deformation with extension along r j which is now N 3 0 W - S 3 0 E . Formation of conjugate shears, extension and release fractures. Local reactivation of thrust faults with cm-scale offset. Simultaneous with porphyry stock emplacement in the Potrerillos district and mineralization at Potrerillos and Jeronimo. 3  3  2.3 Mineralization in the Potrerillos district Three mines, shown in Figure 1.2, have operated in the Potrerillos district thus far: (1) the Potrerillos porphyry C u - M o - ( A u ) porphyry deposit hosted in the Cobre intrusion, (2) the E l Hueso sedimentary and volcanic rock-hosted disseminated gold deposit, and (3) the Jeronimo sedimentary rock-hosted disseminated gold deposit, which is currently in production.  .2.3.1 Potrerillos porphyry Cu-Mo-(Au) deposit The Potrerillos porphyry C u - M o - ( A u ) deposit is hosted by the Cobre porphyry, a 1.5 by 5 kilometre Oligocene plagioclase-, biotite- and hornblende-porphyritic quartz  16  monzonite that lies approximately 4.3 kilometres to the west-northwest of the Jeronimo deposit (Figure 1.2). Within the porphyry, 2 to 6 millimetre phenocrysts are set in a finegrained groundmass of quartz, potassium feldspar and plagioclase. Magmatic hornblende from the Cobre porphyry yields an age of 35.87 ± 0.21 M a whereas the alteration minerals sericite and orthoclase yield ages of 35.64 ± 0.03 M a and 35.06 ± 0.22 M a , respectively (Marsh, 1997). The Cobre porphyry is truncated at depth by the west-dipping P M F . Deep drilling in 1980 by C O D E L C O located the roots of the Cobre porphyry approximately 1.5 k m to the west of its displaced top (Marsh, 1997). The presence of fault-parallel quartz veins and associated sericitic alteration suggests that thrusting and mineralization were synchronous (Sillitoe, 1997). Marsh (1997) distinguished three phases of hypogene alteration in the Potrerillos porphyry deposit. A n early phase of potassic alteration involved the replacement of plagioclase by hydrothermal orthoclase, hornblende by biotite and rutile, and the formation of mosaic-textured quartz veinlets hosting chalcopyrite, bornite and magnetite. Subsequent sericitic alteration resulted in the replacement of plagioclase, hornblende, biotite and hydrothermal orthoclase by coarsegrained sericite. V e i n minerals deposited during sericitic alteration include quartz, anhydrite, ankerite, pyrite, Au-bearing pyrite, chalcopyrite, tennantite, sphalerite and galena. Finally, propylitic alteration affected the interior of the Cobre stock and involved the deposition of epidote, chlorite, calcite and albite. Olson (1983) described the contact metamorphism and skarn formation in the Montandon, E l Asientos and Pedernales Formations intruded by the Cobre porphyry. Sandstone units are recrystallized and contain disseminated pyrite and actinolite up to  17 800 metres from the contact, becoming thermally metamorphosed with calcite-anhydritepyrite-chalcopyrite veins within 100 metres of the contact. Marls and impure carbonates are thermally metamorphosed and host disseminated pyrite and sphalerite up to 1500 metres away from the contact, but exhibit skarn alteration, pyroxene-quartz and calciteanhydrite hornfels with disseminated pyrite and garnet, within 450 metres. Pure limestones are silicified up to 900 metres away from the contact. This effect is most prevalent in coarser-grained bioclastic lithologies. A t the margins of the plutons, mantoshaped bodies hosting copper mineralization occur in the pure limestone horizons of the E l Asientos Formation.  2.3.2 EI Hueso disseminated sedimentary and volcanic rock-hosted gold deposit The E l Hueso gold deposit is located approximately 1.5 kilometres to the westnorthwest of the Jeronimo deposit (Figure 1.2). The replacement-style mineralization at E l Hueso is focussed in steep, east-west trending faults and fans out into the relatively porous, gently west-dipping bioclastic limestone horizons of the upper E l Asientos Formation (unit G) and tuffs of the Tertiary Hornitos Formation. The steep faults are interpreted as tear faults that accommodated differential motion of thrust sheet segments during the early Tertiary compression (Tomlinson, 1994). Following mineralization, normal dip-slip offset occurred along one of the steep east-west faults, known as the Principal fault, resulting in the lateral juxtaposition of the Jurassic and Tertiary host rocks. L i k e the Potrerillos porphyry C u - M o - ( A u ) deposit, E l Hueso is truncated at depth by the P M F , however Marsh (1997) did not observe evidence of syntectonic  18  mineralization at E l Hueso and concluded that mineralization occurred prior to movement along the P M F . Prior to mineralization, the E l Hueso area was intruded by at least nine small Eocene plagioclase-biotite-quartz-hornblende-porphyritic stocks and dykes. Magmatic biotite from one porphyry yielded a A r / A r age of 40.54 ± 0.19 M a (Marsh, 1997; 4 0  3 9  Table 2.2). The distribution of the gold-bearing mantos, which formed after the emplacement of the stocks and dykes, is interrupted where the bioclastic horizons are intruded by the relatively impermeable intrusive rocks. M i n o r gold mineralization occurs at the porphyry margins, where they are cross-cut by gold-bearing quartz veins. Sericite associated with one such gold-bearing vein was dated at 40.25 ± 0.05 M a by the 4 0  A r / A r method (Marsh, 1997). 3 9  Four types of alteration are present at E l Hueso. Silicification occurs dominantly by pervasive replacement within the steep faults and favourable manto-shaped bodies, and as less common veinlets. In relatively low porosity siltstones and more strongly welded tuffs, a spotted or banded texture, called 'moteada', formed locally and consists of green-gray siliceous material and white illitic material forming either as 0.5 to 10 millimeter spots or as the matrix to the spots, or in irregular alternating bands. A post-ore stage of advanced argillic alteration, apparently unrelated to gold deposition, is recorded by the deposition of coarse-grained pink alunite, pyrophyllite, zunyite, dickite, woodehousite, rutile, diaspore and quartz veins along steep faults cross-cutting one of the porphyries and pyroclastic rocks (Marsh, 1997). Coarse-grained alunite from this stage of hypogene alteration was dated at 36.23 ± 0.07 M a by the A r / A r method (Marsh, 4 0  3 9  1997), which is roughly coincident with Cobre porphyry emplacement and mineralization  19  of the Potrerillos deposit. Late stage supergene alteration resulted in the leaching of most of the sulphide minerals and in the deposition of jarosite, chalcanthite, gypsum, scorodite, mansfeldite, chrysocolla, kaolinite, chalcocite, covellite, native copper, delafossite, acanthite and fine-grained alunite. The dominant preserved hypogene sulphide minerals in the deposit are pyrite, chalcopyrite, bornite and pyrrhotite. The majority of the gold at E l Hueso is located within the siliceous replacement bodies, where it is present as one to three micron-sized grains of native gold, or as alloys with C u , within quartz-pyrite veinlets, with chalcopyrite, bornite, galena, sphalerite, jamesonite, stannite, possible wittchenite, monazite and apatite (Marsh, 1997). M i n o r gold also occurs as disseminated, native grains within the moteada-textured zones. The results of the A r / A r dating study by Marsh (1997) indicate that E l Hueso 4 0  3 0  mineralization preceded the formation of the Potrerillos C u - M o - ( A u ) porphyry by approximately 4.5 M a , precluding the possibility that E l Hueso is a distal product of mineralization at Potrerillos. Rather, Marsh (1997) suggested that the source of E l Hueso mineralization is a more proximal, as yet unidentified pluton. In contrast, advanced argillic alteration in E l Hueso, which occurred at 36.23 ± 0.07 M a , is reasonably similar in age to Potrerillos mineralization (35.6-35.9 M a ) and may indicate that the advanced argillic alteration represents the remnants of an eroded, district-wide lithocap originating at the Potrerillos porphyry (Sillitoe, 1997).  20 Table 2.2 Selected Ar/ Ar results of Marsh (1997) for geologic events in the Potrerillos district.  Time (Ma) 6.3-25.1 35.6-35.9  36.2 40.2-40.8  Event Precipitation of fine-grained alunite and jarosite during supergene alteration of the E l Hueso deposit. Emplacement and cooling of the Cobre quartz monzodiorite porphyry stock and porphyry copper mineralization. G o l d mineralization in vicinity of Cobre stock. Hypogene alunite-pyrophyllite-diaspore-zunyite-dickite-rutile-pyrite alteration lacking significant gold mineralization at the E l Hueso deposit. Emplacement, gold mineralization, and cooling of porphyry stocks on the north and east sides of the E l Hueso deposit. G o l d mineralization accompanying illitization, sericitization, and silicification of the sedimentary and volcaniclastic host rocks.  2.3.3 Jeronimo sediment-hosted disseminated gold deposit The Jeronimo deposit may be divided into two distinct parts, separated by the west-northwest-dipping Jeronimo thrust fault (Figure 2.1). In the footwall to the east, the deposit is exposed at the surface and has been intensely altered by supergene oxidation. This portion of Jeronimo has been in production since 1994. T o the west, in the hangingwall, the sulphide-bearing portion of Jeronimo, which is the subject of this study, occurs at depths ranging from approximately 100 to 600 metres. In late 1998, a tunnel was constructed into the sulphide body and exploitation of this portion of Jeronimo is scheduled to begin shortly. The Jeronimo ore will be bio-oxidized, then treated with cyanide to recover the gold.  2.3.3.1 Host sedimentary rocks The Jeronimo deposit occurs within the Jurassic Lower E l Asientos Formation, which has been subdivided into units B through F (Table 2.3, Figure 2.2) by Dilles  21  Table 2.3 Characteristics of the early to middle Jurassic Montandon and El Asientos formations in the vicinity of the Jeronimo deposit. Summarized from Lewis (1996), Elgueta and Fuentes (1997), Olson (1983).  Formation  Unit  Thickness  Upper E l Asientos Fm.  G  -75 m  Lower E l Asientos Fm.  F  20-30 m  E  100-110 m  D  12-14 m  Barrier  C  8-12 m  Barrier  B  60 m  Epicontinental open sea  A  410 m  Montandon Fm.  Environment of deposition Epicontinental open sea  Inner platform Lagoon  Lithologies Thinly- to medium-bedded, calcareous mudstone; fine-grained, sandy fossiliferous limestones interbedded andesite and basaltic andesite flows. Thickly-bedded to massive fossiliferous limestone and oolitic limestone. Interstratified limestone; calcareous quartzfeldspar siltstone and sandstone; oolitic calcarenite; dolomitic limestone. Locally fossiliferous. M e d i u m - to thickly-bedded, medium-grained, calcareous, quartz-rich sandstone with intercalated coquina layers. M e d i u m - to thickly-bedded, medium- to coarse-grained calcareous packstone, grainstone and coquina; fossiliferous sandy limestone horizons. Thinly bedded, calcareous sandstone and siltstones with minor chert nodules; sparse intervals of fossiliferous sandy limestone. Calcareous mudstone with chert nodules and minor coquina layers.  22  J v v V v v v  V  Tertiary Hornitos Formation  v  v  v  Medium-bedded feldspathic siltstone; thicklybedded to massive fossiliferous, oolitic limestone layers, interstratified basaltic andesite flows.  Hornitos Unconformity,  Cretaceous Pedernales Formation  Jurassic Upper Asientos Formation 100 m  Aphanitic to plagioclase-porphyritic andesite flows; andesitic to rhyolitic lapilli-ash tuffs; epiclastic sandstones, siltstones and wackes; coarse andesite cobble-pebble conglomerates.  UnitG  Tr  1  75 m  UnitF  Thinly- to medium-bedded, calcareous mudstone; fine-grained, sandy fossiliferous limestone; interbedded andesite and basaltic andesite flows  Thickly-bedded to massive fossiliferous limestone and oolitic limestone.  25 m  50 m UnitE  Interstratified limestone; calcareous quartz-feldspar siltstone and sandstone; oolitic calcarenite; dolomitic limestone. Locally fossiliferous.  105 m  0  Jurassic Lower Asientos Formation  i  UnitD 13 m UnitC 10 m UnitB 60 m  Unit A 410m  TV".  i  Medium- to thickly-bedded, medium-grained calcareous, quartz-rich sandstone with intercalated coquina horizons. Medium- to thickly-bedded, medium- to coarsegrained calcareous coquina, packstone and grainstone; fossiliferous sandy limestone horizons. Thinly-bedded, calcareous sandstone and siltstone with chert nodules; sparse intervals of fossiliferous sandy limestone. Calcareous mudstone with chert nodules and minor coquina layers.  Jurassic Montandon Formation  Pre-Jurassic Unconformity  Granodiorite to quartz monzonite; rhyolite tuffs and breccias; andesite dykes and flows.  Permian, Carboniferous and Triassic basement Figure 2.2 Generalized stratigraphic section of pre-Jurassic, Jurassic, Cretaceous and Tertiary rock units in the Potrerillos district (modified from Lewis, 1996 and Marsh, 1997).  23 (1995). Unit A is part of the underlying Montandon Formation, and unit G has been assigned to the Upper E l Asientos Formation. A s stated above, thick sedimentary sequences were deposited in the Tarapaca back-arc basin during the Jurassic to Early Cretaceous. In the Potrerillos district, the basin existed from Pleinsbachian to Berriasian time and was the site of an epicontinental sea (Elgueta and Fuentes, 1997). Sedimentation along the basin axis was primarily deep water and mixed clastic-carbonate platform sedimentation occurred on the basin margins. According to Elgueta and Fuentes (1997), two transgressive-regressive cycles are recorded in the Mesozoic sedimentary units of the Potrerillos district. The first cycle is represented by the Montandon and Lower E l Asientos formations, which were deposited following the initial subsidence of the basin. The second cycle is marked by renewed extension and subsidence and is recorded by the deposition of the Upper E l Asientos and Pedernales formations.  2.3.3.2 Structural setting Unlike the Potrerillos porphyry and E l Hueso deposits, the Jeronimo deposit is located entirely within the footwall of the P M F . In the Jeronimo area, the P M F is known as the Silica Roja fault and it is recognized as the fault that places the Mesozoic sedimentary sequence over the Tertiary volcanic rocks (Figure 2.1). A s stated above, the west-dipping Jeronimo thrust fault separates the Jeronimo sulphide body in the hangingwall from the oxidized portion of the deposit in the footwall (Figure 2.1). G o l d mineralization has not been observed within the plane of the Jeronimo fault (Caddey, 1999).  24  Step-wise vertical offset of the horizons hosting the Jeronimo sulphide body, which separates Upper and Lower Jeronimo by up to 400 metres, occurred by steep faulting ranging in strike from north-south to east-west. According to Homestake geologists, these faults, which are known as the Polvorines and Mapuche faults, formed during basin growth in the Jurassic to early Cretaceous. This interpretation is based on timing relationships, as the normal faults must be older than the early Tertiary thrust faults that cross-cut them, suggesting that they are Mesozoic in age and existed when the Tarapaca basin was active. The normal faults generally do not host significant gold, except for localized zones of gold enrichment (e.g. 2.9 g/t). Underground mapping of a recently opened tunnel through the Jeronimo sulphide ore by Caddey (1999) has shown that, within the bioclastic horizons, mineralization is hosted in several fracture sets. Caddey (1999) interpreted these fractures to have formed as the early Tertiary deformational event drew to a close, in response to lateral extension perpendicular to N 6 0 W - S 6 0 E compression. The bioclastic horizons are preferentially fractured relative to the finer-grained lithologies, possibly in response to the greater competency of the bioclastic limestone lithology (Caddey, 1999). Several steep faults with centimetre-scale normal dip-slip offset, hosting calcite and arsenic sulphide minerals, were observed to cross-cut the C , D and E units (Caddey, 1999). Caddey (1999) proposed that these faults also formed at the end of early Tertiary compression and may have been conduits for fluid flow to the bioclastic horizons. Unfortunately, lack of exposure in the tunnel below the mineralized horizon precludes the observation of these faults and any contained mineralization at depth.  25  2.4 Relationships among mineralized centres in the Potrerillos district G o l d deposition at E l Hueso during the late Eocene (40.25 M a ; Marsh, 1997) appears to be the oldest identified occurrence of mineralization in the Potrerillos District. It was probably produced from magmatic-hydrothermal fluids exsolved from a local, unidentified porphyry (Marsh, 1997). Advanced argillic alteration at the deposit, dated at 36.23 ± 0.07 M a , must be related to a younger magmatic event, possibly the intrusion and mineralization of the Cobre porphyry, dated at 35.6-35.9 M a . A l o n g with the majority of northern Chile's large porphyry copper deposits, the emplacement of the early Oligocene Cobre porphyry and mineralization of the Potrerillos porphyry C u - M o - ( A u ) deposit occurred during a major compressional event, related to deformation within the Domeyko fault system. According to Caddey (1999), mineralization at Jeronimo also occurred during this compressional event, raising the possibility that Jeronimo mineralization may be related to magmatism at Potrerillos or at a more proximal, as yet unidentified, intrusion.  26  Chapter 3: Jeronimo Gold Deposit  3.1 Stratigraphic and structural controls on ore distribution The stratabound Jeronimo deposit dips gently to the northwest, following stratigraphy, and is on average 6 metres thick. In plan view, the deposit is amoeboidshaped and covers and area of approximately 2 by 1.5 kilometres (Figure 3.1). The sulphide body is subdivided into two zones referred to as upper and lower Jeronimo, separated by the Mapuche and Polvorines steep normal faults (Figure 3.1). Homestake geologists have plotted contours of grams/tonne of gold multiplied by thickness in metres (known as G T contours) for Jeronimo and have distinguished zones of <10 g/t-m, 10-50 g/t-m and >50 g/t-m. These contours, shown in Figure 3.1, emphasize the irregularity of gold distribution within the deposit.  3.1.1 Lithologic control The lithologies that most consistently host ore in the Jeronimo deposit are bioclastic limestones that comprise unit C and portions of unit B . T w o types of bioclastic limestones are recognized. The dominant ore host is a very poorly sorted bioclastic limestone with large, typically 8 by 4 mm, pelecypod and gastropod shells, referred to as a coquina by Homestake geologists (Plate 3.1). The coquina generally consists of 30% bioclasts, 50% coated carbonate grains and 20% calcite spar cement, with up to 4%  28  Plate 3.1 Hand sample photo of bioclastic limestone from unit C , dominant ore-hosting lithology. Sample DDH69-156.72.  Plate 3.2 Thin section photo of bioclastic limestone with pelecypod shells comprised of blocky calcite and dark brown, rounded, coated carbonate grains in a sparry calcite matrix. F i e l d of view is 5.1 mm. Taken under transmitted light with crossed nichols. Sample DDH69-156.72.  29  30 subangular detrital quartz grains, typically 0.3 by 0.2 mm (Plate 3.2). Pelecypod shells comprise the dominant bioclast type and are composed of blocky recrystallized calcite, commonly with patches of chalcedonic quartz. Other bioclast types in the coquina consist of symmetrically coiled cephalopod shells, generally 2 m m in diameter, composed of fine-grained recrystallized calcite; rod-shaped fossils, possibly echinoderm spines that are composed of chalcedonic quartz and are on average 10 by 0.1 mm; and, occasional spherical, sub-rounded bryozoan mat fragments that are generally 0.5 by 0.5 mm and are composed of fine-grained calcite. The spherical to elliptical coated carbonate grains, possibly pellets, range in size from 0.2 to 0.5 mm, are subangular to rounded and consist of fine-grained calcite. The grains are normally rimmed or wholly replaced by amorphous dark material, possibly fine-grained pyrite, organic material, micrite or clay. They are generally enclosed by the spar cement and are not in grain to grain contact. The secondary ore host is a bioclastic limestone referred to as a fossiliferous sandy limestone by Homestake geologists. It is moderately sorted and generally consists of 70% bioclasts, 10% intraclasts and 20% spar cement, with occasional fine-grained detrital quartz (Plate 3.3). Bryozoan mat fragments ranging from 0.5 to 0.7 mm comprise the dominant bioclast type and are generally subangular to subrounded, spherical, and have a porous texture. S E M - E D S analysis indicates that some bryozoan mat pores are filled by illite, while others remain as void space. The bryozoan mat fragments are often enclosed by an optically continuous calcite rim cement. Other bioclast types include grains consisting of fine-grained calcite, of similar external morphology to the bryozoan  31  Plate 3.3 Thin section photo of bioclastic limestone of unit B , with bryozoan mat fragments and occasional quartz grains in a sparry calcite matrix. Field of view is 2.5 mm. Taken under transmitted light with crossed nichols. Sample DM3-560.00.  Plate 3.4 Thin section photo of sandy limestone, consisting of quartz grains and calcite allochems in a matrix of chalcedonic quartz, clay minerals and sparry calcite cement. Poor ore host. Field of view is 5.1 mm. Taken under transmitted light with crossed nichols. Sample DDH91-562.80.  32  33  clasts, but without internal structure; occasional elongate chambered clasts composed of fine-grained calcite, possibly cephalopod fragments; and subangular, typically 0.8 by 0.4 m m coral fragments. Rounded intraclasts are generally 0.8 by 0.6 mm and consist of small bioclasts in light brown micritic matrices. Lithologies which do not host ore are poorly to moderately sorted, matrix supported sandy limestones. Within units B and D , these consist of 10 to 50% quartz grains and up to 70% calcite allochems, with occasional small bioclasts, typically 0.7 by 0.2 mm (Plate 3.4). Quartz grains range from very fine to medium sand (0.1 to 0.4 mm) and are round to subangular. Subrounded, brown-stained calcite allochems also range from 0.1 to 0.4 m m and do not posses internal structure. The matrix of the sandy limestones may consist of spar or chalcedonic quartz and includes up to 5% disseminated kaolinite or illite. The diamond drill core logs D D H 91 and 93 (Figure 3.2) illustrate the influence of lithology on ore distribution. Elevated gold grades occur mainly within the bioclastic limestone and coquina of unit C and the bioclastic limestone of unit B . Most logs in Figure 3.2 show that gold grade decreases strongly where sandy limestone units occur interlayered with the bioclastic limestone in unit B (e.g. D D H 72, D D H 100 - Upper Jeronimo intercept, Figure 3.2). The depositional and diagenetic properties of the bioclastic limestones of units C and B which account for the preferential migration of ore fluids through them, as opposed to the sandy limestones, include (1) the overall coarser grain size; (2) the absence of clay particles; (3) the irregularity of bioclast shape, which influences grain packing; (4) the greater intraparticle porosity, resulting from the porous texture of the bryozoan mat  Figure 3.2 Logs of ore zone diamond drill hole intercepts examined in this study. Au grade was determined by fire-assay at ALS Geolabs, Copiapo. Decalcification, argillization and silicification were assessed on a relative basis and were ranked from 0 to 4. Percentages of pyrite, calcite veining and late arsenic sulphide mineralization were visually estimated. Lithologies Shelly bioclastic limestone Bioclastic limestone Coarse-grained sandy limestone Fine-grained sandy limestone Coarse-grained calcareous sandstone Fine-grained calcareous limestone Sandy oolitic limestone Zone of faulting or brecciation  35 DDH59  J? <3°  » ^  4  ?  /  #  0 1 2 3 4 5 0 1 2 3 4 0 1 2 3 4 0 1—I 2 I3 1_ 4 0  670-  4  8 12 0  4  8  12 0  675680685/—\  •3 u Q  690695700A ' 1  705710-  •''iv'  715720.  DDH 72  » ^/ 0123450 715A' 1  720-  725-  •a  730-  CL.  VM" .-.••rV'  .'.••T¥'.  735-  740-4^  745-  750.  • •'iV-'.  V 2  e,v  4 0  2  y^  Jt  4/  <y  ^ ~ 4 0  2  <^  4  0 2 4 6 0 1 2 3  2  4  6  DDH 73  DDH83  <r 0 1 2 3 0  # 2  4 0  2  4 0  ^ 2  4 0  3  530  535 •  4 540  •a o.  u Q  545  550  |,vj  555  6  0  2  4  0  8  16  DDH86 xV  0 2 4 6 0  2  / J  &  Ni^'  4  0  2  4 0  2  40  2.5  50  2.5  50  20  40  2.5  5  535  540  •'iV'  P  545  550  555 A . '' ' 1  560  565  DDH91  0 6 560  565  570  575 - 1 * ^ * 1  580 A  585  590  iV ^  12 18  0  ^  2  4  0  «r 2  4 0  ^ 2  4 0  2  4 0  DDH  96  Upper Jeronimo  #  y 05  275  280  285  8  290-1^4  3 a.  Q  295 300  305 •" • . ' ,v  310  315  H P  10 15 0.  2  4  0  y 2  4  J i ff 0  2  0  4  0  5 0  . 1 2  DDH 96 Lower Jeronimo (1st intersection)  ^ 0  3  60  2  4  0  ^ 2  4 0  r 2  415 - * • •  D 420  ~1 425 —  ^  ^  -a CL T¥  C  430  435  V.  B  440  DDH 96 Lower Jeronimo (2nd intersection)  470  475  480  4  485  490  4  495  4  500  39  4  0  1  2  3  0  1  2  3  DDH 97 Upper Jeronimo V  0  265  270  20  40  0  2  4  0  2  4 0  2  4 0 4 8  12 16 0  2  4  H  275  280  H~  285 H  •*v. •'•T¥'  290 i.*--: 295  DDH 97 Lower Jeronimo ••5-  0 8  400  410 J  415  IS3  425 J  430  • '.' y" ;1  AOA*,  435 1  440  16 24  0  2  4 0  2  4 0  2  4  0  4  8 I  405 J  420  O?  w.-  12 I  0  3  DDH 100 Upper Jeronimo  DDH 100 Lower Jeronimo  4  350  •• .'•  • "  0  5  2  4  0  , v :  • JI . • TIT.  355  4  •  . -'V.' 360  .• 'A ' 1  ''T¥.-;. • '•'T¥;'  B  365  •B Q. Q 370  • '/T¥V.  • -i¥  • '•' T V ;  375  380  VI,,,  4  0  42  clasts; and (5) the greater moldic porosity of the coquina, caused by the preferential dissolution of shell clasts. Although the poor sorting of the coquina may have decreased its porosity, its effect appears to be compensated by the strong irregularity of the bioclast shapes, which causes the coquina to be very loosely packed. L o w primary porosity of the more silicate-rich sandy limestones may have caused them to act as relatively impermeable barriers, further enhancing fluid flow in the bioclastic limestone horizons. The timing of deposition of the calcite spar cement in the bioclastic limestones is difficult to determine. If this cement is a late phase, intergranular porosity may have been very high during ore fluid migration. Alternatively, it is possible that the spar cement was partially to entirely present prior to mineralization and that porosity and permeability were enhanced during mineralization by dissolution of the cement.  3.1.2 Structural control In the Jeronimo area, the late Mesozoic north to northwest-striking Polvorines and east to northeast-striking Mapuche normal fault systems (Figure 2.1) link together to form a semi-circle, downdropping the host horizons to the south, west and east by up to 400 metres in a series of steps (Figures 3.3a, b and c). Assay results of core from angled diamond drill holes that have pierced the subvertical Mapuche fault indicate that gold grade is locally elevated within the fault zone (e.g. 2.9 g/t; Figure 3.2, D D H 9 6 , Lower Jeronimo, 1 intersection). It is possible that the Jeronimo ore fluids used these faults as st  conduits, however the absence of significant overall mineralization within the fault zones has led Homestake geologists to conclude that they did not act as the principle  Figure 3.3 a,b,c Three cross-sections through the Jeronimo sulphide body, showing drill holelocations, major lithologies and structures. Geology interpreted by Homestake geologists.  44  A.  460885, 7070310  ' 7069060 4  A  6  2  3  9  5  A'  4000 i —  3900  Tertiary Hornitos Formation  Thrust fault  W] Volcanic conglomerate, epiclastic sandstone, —' andesitic to rhyolitic tuffs, andesite flows and breccias —  Hornitos unconformity  Jurassic Lower El Asientos Formation | HH  U n i t F : Fossiliferous limestone, oolitic limestone Unit E : Limestone, calcareous sandstone and siltstone Unit D : Calcareous sandstone  \//\ I  Normal fault  Unit C: Bioclastic limestone I Unit B : Calcareous sandstone and siltstone, fossiliferous sandy limestone  Jurassic Montandon Formation Unit A : Calcareous mudstone, minor bioclastic limestone Chloritized augite porphyry of unknown age  DDH 62 Diamond drill hole. Arrow projects into page, cross projects out o f page.  45  B.  460935,  461775,  7069590  7070280  B'  B  4000  Tertiary Hornitos Formation y  V  Aphanitic to plagioclase-porphyritic andesite flows; andesitic to rhyolitic lapilli-ash tuffs; epiclastic sandstones, siltstones and wackes; coarse andesite cobble-pebble conglomerates.  Jurassic El Asientos Formation Jaf  | Coarsely fossiliferous limestone; thin to medium bedded oolitic limestone with fine fossil hash  Jae |HJ Fine- to coarse-grained sandy limestones andcalcarenites with intercalated bioclastic, dolomitic andoolitic limestones Jad  7 Thrust fault Hornitos unconformity  Normal fault  Medium- to coarse-grained quartz arenite  Jac  7~A Coarse, black biosparite with continuous fossil hash and local coquina horizons; main ore host  Jab  j Thinly and wavy bedded,fine-grainedsandy limestone with chert matrix, nodules and bands Jurassic Montandon Formation  Jaa  Zone of grossular-wollastonitealbite skarn  Calcareous mudstone, minor bioclastic limestone Chloritized augite porphyry of unknown age  DDH 62 Diamond drill hole  46  c. 461340,  462250,  7068970  7069855 C  Tertiary Hornitos Formation Volcanic conglomerate, epiclastic sandstone, andesitic to rhyolitic tuffs, andesite flows and breccias Thrust fault  Jurassic Lower E l Asientos Formation I Unit F: Fossiliferous limestone, oolitic limestone  „ — — Hornitos unconformity  | Unit E : Limestone, calcareous sandstone and siltstone Unit D : Calcareous sandstone  t  Normal fault  Unit C : Bioclastic limestone Unit B : Calcareous sandstone and siltstone, fossiliferous sandy limestone  DDH  62  Diamond drill hole Jurassic Montandon Formation Unit A : Calcareous mudstone, minor bioclastic limestone •*+*  Chloritized augite porphyryof unknown age  47  conduits. In drill core, these fault zones consist of gougey, black rubble. They may have been poor conduits at the time of ore deposition due to the clay-rich nature of the fault zones, which is not conducive to fluid flow. Regardless, they influence the shape of the Jeronimo body as they have segmented and offset the host horizons. The early Tertiary Jeronimo thrust fault (Figure 2.1) separates the Jeronimo sulphide ore, located in the hangingwall, from the oxidized ore located in the footwall. Although the thrust fault is barren, it may have been active during mineralization (Caddey, 1999). During mapping of a recently opened tunnel through the sulphide zone, Caddey (1999) identified a set of sub-vertical normal faults (described in section 2.3.3.2), which cross-cut the ore-hosting C-unit and overlying D and E units, as well as the thrust faults and host calcite and arsenic sulphide minerals. The normal faults, which repeat at intervals of 120 to 140 metres, range in strike from N 5 0 W to N 6 5 W and have centimetrescale offsets. Caddey (1999) also observed that the bioclastic limestone horizons are preferentially fractured and jointed. These structures host manganese carbonate, quartz, pyrite and arsenic sulphide minerals and strike, in order of decreasing abundance, N 6 0 W , N 3 0 E , N - S and N 8 0 W . Caddey (1999) interpreted that (1) the faults, fractures and joints formed during a period of regional relaxation following the early Tertiary compressive event; (2) ore deposition was simultaneous with the formation of these structures; and (3) the faults acted as conduits, channeling the ascending fluids to the relatively porous and permeable bioclastic horizons. Due to a lack of deeper exposure in the tunnel, it was not possible to  48 examine the faults below the mineralized horizon in order to ascertain whether they display evidence of fluid flow at depth.  3.2 Hypogene alteration Three hypogene alteration assemblages occur in the Jeronimo deposit. Within the ore zone, the absence of primary carbonate indicates that complete decalcification of the bioclastic limestones has occurred. The limestones have been replaced by hydrothermal quartz, carbonate and clay minerals that formed during pervasive silicification, and localized carbonatization and argillization. This section discusses the characteristics and relations of these three phases of hypogene alteration within the deposit. The majority of the ore and gangue minerals within the Jeronimo deposit are restricted to very fine grain sizes, typically ranging from several hundred urn to less than one pxn, necessitating the use of an X-ray diffractometer ( X R D ) and scanning electron microscope ( S E M ) in order to identify the minerals and examine their textures and relationships. The X R D used in this study is a Siemens Diffraktometer D5000. The S E M is a Philips X L 3 0 electron microscope equipped with a Princeton Gamma-Tech energy dispersive system (EDS).  3.2.1 Silicification 3.2.1.1 Distribution Secondary quartz is the most abundant ore zone mineral, composing on average 50% of the ore. Silicification intensity varies considerably, with secondary quartz ranging from 5 to 90%. During core logging, silicification intensity was qualitatively  49  assessed and assigned a numerical value according to the ease with which the core was scratched by steel. In some drill hole intercepts (e.g. D D H 59, Figure 3.2), zones of strong silicification overlap with gold mineralization, however in most cases, silicification intensity increases only sporadically within the ore zone and also outside of it (e.g. D D H 72 and 91).  3.2.1.2 Characteristics In hand sample, strongly silicified ore is light to dark blue-grey, vuggy and can not be scratched by steel (Plate 3.5). In thin section, silicification is characterized by the pervasive occurrence of massive, vuggy aggregates of fine-grained to chalcedonic, anhedral, interlocked quartz grains. Quartz grains increase in size and become more euhedral towards vugs, and quartz prisms, generally 0.1 by 0.025 mm, with triangular terminations, commonly project into vugs (Plate 3.6). Veins of quartz are rare within the deposit and those observed in this study are irregular and discontinuous, ranging from 25 u,m to 0.5 mm in width, and host pyrite, rutile and arsenopyrite (Plate 3.7). The secondary quartz commonly preserves fossil shapes and textures. Pelecypod clasts, generally ranging from 2 by 0.5 mm to 10 by 1 mm, are entirely replaced either by equant, interlocking quartz grains or by elongate, anhedral fibrous quartz, and commonly host micron-sized carbonate inclusions (Plate 3.8). The porous texture of the bryozoan mats is also locally preserved (Plate 3.9). Silicified fossils and grains of pyrite may be rimmed by fibrous, radiating quartz. The centres of the silicified pelecypod clasts are typically vuggy, with drusy quartz, and may be filled by compact aggregates of carbonate minerals. Some silicified bioclasts are cross-cut by carbonate veinlets.  50  Plate 3.5 Hand sample photo of vuggy, blue-gray zone of strong, pervasive silicification in a bioclastic limestone. Samples DDH91-572.9 (upper) and DDH91-574.5 (lower).  Plate 3.6 Thin section photo of pervasively silicified ore. V u g in centre is lined with drusy quartz and filled by manganese carbonate. Brown wisps in matrix are illite. Field of view is 1.25 mm. Taken under transmitted light with crossed nichols. Sample DDH100-188.53.  5]  Plate 3.7 Scanning electron microscope back-scattered electron image o f rare quartz veinlet cross-cutting a calcite vein with rhodochrosite inclusions. The matrix hosts disseminated pyritewith arsenopyrite inclusions. Sample DDH91-574.39.  53  Plate 3.8 Thin section photo of silicified pelecypod clasts in matrix of fine-grained quartz and clay. Field of view is 5.1 mm. Taken under transmitted light with crossed nichols. Sample DDH97-273.06.  Plate 3.9 Thin section photo of silicified bryozoan mat clasts with internal structure preserved by secondary quartz. Irregular veinlet of manganese carbonate extending from lower left-corner to upper centre of photo. Field of view is 1.25 mm. Taken under transmitted light with crossed nichols. Sample DDH100-369.15.  54  55  3.2.1.3 Mineralogy Minerals that occur as grains intergrown with the vuggy matrix-forming quartz include pyrite, arsenopyrite, sphalerite, apatite, rutile and barite. The characteristics of the sulphide minerals are discussed in section 3.3.1. Subhedral to anhedral, equant apatite grains are typically 20 p m but range from 5 to 140 pm. Apatite is normally disseminated within the vuggy quartz matrix and may occur within large vugs (Plate 3.10). It is commonly intergrown with rutile and pyrite, and is locally intergrown with arsenopyrite. In one sample, subhedral, 20 p m rutile grains are enclosed by massive apatite in clusters 400 um wide. Locally, apatite occurs as (1) tabular, disseminated grains ranging from 60 by 25 p m to 215 by 65 um; (2) poikilitic aggregates up to 65 by 40 pm; and (3) as fine-grained crystals with pyrite in diffuse, irregular veinlets, on average 60 p m wide, that cross-cut the silicified matrix. Although the majority of apatite grains have hydrothermal textures, some composite, well-rounded grains resemble detrital apatite intraclasts present in the unaltered host rock. The apatite intraclasts, typically 100 pm, are a minor but ubiquitous constituent of the host rocks. Rutile grains are subhedral to anhedral, equant to weakly tabular, and are typically 16 by 12 um, but range from 6 by 6 pm to 20 by 10 pm. They occur mainly as disseminated grains in the quartz matrix and are strongly associated with pyrite (Plate 3.11). The two minerals occur together in diffuse aggregates and veinlets. Sporadically, rutile is present in the host rocks, where it is present as disseminated, weakly tabular, typically 80 by 35 um grains.  Plate 3.10 Scanning electron microscope back-scattered electron image of illite-filled vug in a pervasively silicified matrix with abundant disseminated pyrite. Tabular, subhedral apatite and aggregates o f acicular Pb-sulphosalt grains are present within the vug and at vug margins. Sample D D H 5 9 - 6 9 1 . 7 8 .  Plate 3.11 Thin section photo ot" blocky brown rutile grains intergrown with pyrite in vuggy. quartz-rich matrix. Field of view is 0.625 mm. Taken under transmitted, plain polarized light. Sample DDH73-278.32.  Barite occurs as both large, 300 by 20 urn, subhedral, tabular grains in diffuse aggregates within the quartz matrix, locally intergrown with pyrite, and as disseminated micron-sized grains. Barite also tends to form as a vug-filling mineral, usually occurring in vugs in the manganese carbonate minerals with kaolinite and quartz. In one sample, poikilitic barite hosts inclusions of anhedral galena, subhedral pyrite and subhedral to euhedral, tabular apatite, in a 325 u m diffuse vein that cross-cuts the vuggy quartz matrix. Micron-sized grains of La-, Ce- and Nd-bearing monazite commonly occur with illite in vugs in quartz, or as inclusions in quartz grains and aggregates. The grains are anhedral and are spherical to elliptical, ranging up to 2 by 1 um in size.  3.2.2  Carbonatization  Primary or diagenetic carbonate minerals within the E l Asientos Formation host rocks consist of nearly pure calcite and dolomite. E D S spectra indicate that Fe and, to a lesser extent, M n are trace constituents of these carbonate minerals. In contrast, E D S spectra of ore zone carbonate minerals show that they are strongly enriched in M n and are locally enriched in Fe.  3.2.2.1 Distribution Rhodochrosite and kutnohorite (see section 3.2.2.3 for carbonate mineral determination) are the volumetrically dominant carbonate minerals within an elliptical, roughly north-west trending zone overlapping the centre of the ore body (Figure 3.4). Manganese carbonate mineralization is not exclusive to any of the zones defined by the  59  60  G T contours. Hydrothermally-deposited calcite is the dominant carbonate mineral in a zone at least 200 metres wide enveloping the central manganese carbonate zone. It occurs as the dominant carbonate mineral only in G T zones of less than 50 g/t-m. Outside of the ore zone, late calcite veinlets and fracture coatings, ranging from <1 mm to 5 mm, are present throughout each of the units of the E l Asientos Formation, typically comprising between 1 and 3% of the rock.. Visual estimation of the abundance of these calcite veinlets across ore zone intercepts suggests there is an inverse relationship between gold grade and calcite veining. In particular, logs D D H 59, 72 and 91 illustrate this negative correlation (Figure 3.2).  3.2.2.2 Characteristics Typically, manganese carbonate minerals comprise between 0 and 20% of the ore and are fine-grained and difficult to observe in hand sample. In certain samples, manganese carbonate minerals, dominantly rhodochrosite, are present in quantities between 40 and 70%. They most commonly occur as very fine-grained white to light grey crystals, occasionally displaying a pinkish hue. Manganese carbonate is either disseminated and intergrown with grey quartz grains, giving the ore a salt and pepper appearance in hand sample, or occurs in aggregates up to 3 millimetres in size (Plate 3.12). Sporadically, light pink rhodochrosite occurs as thin, typically 4 millimetre, irregular veinlets cross-cutting the grey-white quartz and manganese carbonate matrix (Plate 3.13). Inspection of thin sections by optical microscopy and S E M revealed that the manganese carbonate minerals most commonly occur as subhedral, aggregated grains  61  Plate 3.12 Hand sample photo of ore with strong manganese carbonate mineralization. Kutnohorite occurs as disseminated grains and in fine-grained aggregates. Sample DDH100-175.23.  Plate 3.13 Hand sample photo of pink rhodochrosite veining. Manganese carbonate veins are cross-cut by discontinuous, late orpiment veinlets. Sample DDH86-547.61.  62  63  rimming or filling vugs within the silicified matrix or within the vuggy centres of silicified bioclasts (plates 3.6, 3.14). The manganese carbonate minerals also occur as fine-grained massive aggregates (Plate 3.15). Locally, small quartz grains, typically 0.1 millimetre, occur within vugs in the manganese carbonate aggregates. The manganese carbonate minerals also form irregular veinlets that cross-cut the quartz matrix, ranging in width from 25 to 175 um (Plate 3.9). Rhodochrosite constitutes the dominant manganese carbonate mineral and is strongly intergrown with kutnohorite and manganoan dolomite (Plate 3.16). In most cases, kutnohorite deposition appears to post-date rhodochrosite, as kutnohorite often occurs as subhedral, typically 15 um grains lining the margins of rhodochrosite aggregates. However, kutnohorite also occurs as distinct aggregates, veins and disseminated grains, and locally hosts rhodochrosite inclusions, indicating that the precipitation of these two minerals overlapped. Kutnohorite may be intergrown with manganoan dolomite in aggregates of anhedral to subhedral grains, generally 35 um, that locally host subhedral, typically 10 by 5 um rhodochrosite inclusions. Pure calcite, the dominant carbonate mineral in the zone enveloping the manganese carbonate zone, commonly occurs (1) filling open space at the centre of silicified pelecypod clasts, (2) as micron-sized inclusions in silicified bioclasts, and (3) in late, irregular veinlets intergrown with orpiment, realgar and minor cinnabar. One sample near the eastern edge of the deposit ( D D H 69-161.09) contains a silicified pelecypod clast with open-growth, pure calcite at its core, hosting small (~3 um) subhedral inclusions of rhodochrosite and the manganese sulphide mineral alabandite (MnS).  64  Plate 3.14 Thin section photo of silicified bioclast with centre of manganoan calcite. Primary fibrous texture of bioclast is preserved by quartz and carbonate. Field of view is 5.1 mm. Taken under transmitted light with crossed nichols. Sample DDH69-161.09.  Plate 3.15 Thin section photo of disseminated manganese carbonate mineralization. Occasional detrital quartz grains. Field of view is 2.5 mm. Taken under transmitted light with crossed nichols. Sample DDH91-568.89.  65  Plate 3.16 Scanning electron microscope back-scattered electron image o f a rhodochrosite-kutnohorite aggregate, with intergrown quartz and kaolinitefilled vugs. Sample D D H 3 5 - 4 0 2 . 5 9 .  67  3.2.2.3 Carbonate m i n e r a l identification a n d composition According to their structure, carbonate minerals may be grouped as calcite-, dolomite-, and aragonite-type (Chang et al., 1996). A s extensive solid solution is possible between the cations of carbonate minerals within these groups, it is necessary to determine both the structure type and composition of a carbonate mineral in order to identify it. The Jeronimo carbonate minerals were identified through a combination of X Ray Diffraction ( X R D ) and Electron Probe Micro-Analysis ( E P M A ) . X R D analyses of the carbonate minerals were performed using a Siemens Diffraktometer D5000 at 40.0 k V and 30.0 m A using C u K-alpha radiation and scanned at 0.04°20 intervals for 2.0 seconds between 3 and 60 °28.  Within the Jeronimo carbonate minerals, the  considerable solid solution of the cations M n , F e , C a 2 +  2+  2 +  and M g  2 +  resulted in shifts in  the positions of the X R D peaks, rendering the identification of specific carbonate minerals by X R D alone unreliable. X R D spectra were only used with certainty to distinguish between carbonate minerals of the different structural groups, not between compositional end members within a single structural group. The spectra indicated that the Jeronimo carbonate assemblage consists of calcite- and dolomite-group carbonate minerals, with calcite-group minerals comprising the most abundant phase. Representative examples of the carbonate mineral diffraction patterns obtained in this study are presented in Figure 3.5. Diffraction peaks for rhodochrosite, kutnohorite and and magnesian kutnohorite are presented for reference in Table 3.1.  68  Rhodochrosite Sample: DDH86-547.61 2.856 Rhodochrosite 104  o O  20  30  40  Degrees two-theta  B.  Kutnohorite Sample: DDH91-562.80  2.918 Kutnohorite 104  1000  10  30  40  50  60  Degrees two-theta Figure 3.5 Representative X-ray diffraction spectra of (A) rhodochrosite and (B) kutnohorite.  70  69 Table 3.1 Sample diffraction peaks for the calcite- and dolomite-group manganese carbonate minerals, rhodochrosite and kutnohorite, for 3-60 °29. From the Mineral Powder Diffraction File Data Book (1980), published by the Joint Committee on Powder Diffraction Standards (USA).  Rhodochrosite Hkl 012 104 110 113 202 024 018 116 211 122  I/Ii  35 100 20 25 25 12 30 35 2 14  Kutnohorite d(A) 3.66 2.84 2.39 2.172 2.000 1.829 1.770 1.763 1.556 1.533  Hkl 101 012 104 006 015 110 113 021 107 024 018 009 211 122 212  I/Ii  6 20 100 6 4 14 20 4 20 10 25 30 6 4 4  Kutnohorite (magnesian) d(A) 4.27 3.75 2.94 2.73 2.59 2.44 2.23 2.10 2.04 1.876 1.837 1.814 1.588 1.566 1.540  hkl 003 012 104 006 015 110 113 202, 107 024 018 116 009 211 122,027  I/Ii  2 8 100 2 2 6 10 6 4 10 12 10 2 4 2  d(A) 5.41 3.73 2.91 2.701 2.564 2.423 2.209 2.031 1.862 1.823 1.804 1.800 1.578 1.556 1.512  E P M A of the carbonate minerals was performed on a fully automated C A M E C A S X - 5 0 microprobe, operating in the wavelength-dispersion mode, with the following operating conditions: excitation voltage 15 k V , beam current 10 n A , peak count time 20 s, background count time 10 s, beam diameter 10 pm. F o r the elements sought, the following standards, X-ray lines and crystals were used: C a , natural calcite, C a K a , P E T ; Fe, natural siderite, F e K a , L I F ; M n , natural rhodonite, M n K a , L I E ; M g , natural dolomite, MgKoc, T A P . Data reduction was done with the " P A P " (|)(pZ) method (Pouchou and Pichoir, 1985). Oxygen was calculated stoichiometrically on the basis of six oxygen atoms per formula unit and carbon was calculated by difference. Average mineral formulas were determined from analyses with C atoms = 2.00 ± 0.04. Samples  70  selected for E P M A were enriched in M n or Fe as detected by S E M - E D S analysis. Microprobe data obtained for the carbonate minerals is presented in Appendix A . The calcite-group minerals rhodochrosite, siderite and calcite were identified by E P M A . They are defined as having mole fractions of M n C 0 , F e C 0 3  3  and C a C 0  3  of >  50%, respectively, according to the guidelines presented by N i c k e l (1992). Twenty-three rhodochrosite analyses give an average composition of (Mno.7oCao.iiFeo.o8Mgo.o2)C0 that ranges from (Mn .87Cao.iiMgo.oiFe .oi)C0 0  0  3  3  to (Mno.68Ca .2oFeo.iiMgo.oi)C0 . One 0  3  analysis of manganoan calcite with the formula (Cao.9iMn .o7Feo.o2)C0 was also 0  3  obtained. A n average of eight siderite analyses yields the composition (Fe .7oMno.24Ca .o4Mgo.o2)C0 , with a range of (Feo.82Mno.iiCa .o5Mgo.o2)C0 0  0  3  0  3  to  (Feo.5oMno. 5Cao.o5Mgo.oi)C0 . A ternary plot (Figure 3.6a) of the calcite-group carbonate 3  3  minerals indicates that (1) the compositions of carbonates in individual samples is distinct and relatively homogeneous; and (2) there is complete solid solution between siderite and rhodochrosite, consistent with the findings of Essene (1983). The dolomite-group minerals kutnohorite ( C a M n ( C 0 ) 2 ) and dolomite were 3  identified by E P M A . Minerals of the dolomite-group have the generalized formula A B ( C 0 ) 2 , where the A site is occupied by C a and the B site may be occupied by M g , 3  M n , Fe and excess C a (Frondel and Bauer, 1955). In this study, kutnohorite and dolomite are defined as dolomite-group minerals where the B site is occupied by greater than 50% M n or M g , respectively, according to the guidelines established by N i c k e l (1992). A n average of three kutnohorite analyses, with stoichiometrically calculated C of 2.00 ±  71  A.  86-547.61 91-572.28 96-299.45 96-493.69 96- 494.75 97- 414.39 100-355.81  CaO  FeO MnO  B.  + D D H 91-572.28 • D D H 96-299.45 v D D H 97-414.39 o D D H 100-175.23 • D D H 100-355.81  MgO  FeO  Figure 3.6 (A) Variation in CaO, MnO and FeO endmember proportions of calcite-group carbonate minerals for analyses with mol % MgO < 5; (B) variation in MgO, MnO and FeO endmember proportions of dolomite-group carbonate minerals for analysis with mol% CaO of40-60.  72  0.04, yields a composition of Ca(Mno.54Cao.i9Mgo.i5Feo.i2)(C03) . A single analysis of 2  manganoan dolomite yields the formula Ca(Mgo.57Mno.39Feo.o3Ca .oi)(C03)2. 0  A ternary.plot of the B site cations (Figure 3.6b) in the dolomite-group carbonate minerals for analyses with a stoichiometrically calculated value of C of 2.00 ± 0.14 indicates that (1) there is extensive solid solution between kutnohorite and dolomite, as shown by Essene (1983), and (2) two samples host kutnohorite that is enriched in Fe. The latter samples occur in the same drill hole (DDH100), but in offset segments of the C-horizon.  3.2.2.4 Mineralogy The manganese carbonate aggregates host inclusions of anhedral to subhedral pyrite, which average 5 um in size and are occasionally arsenopyrite-bearing; anhedral, typically 80 by 35 u m sphalerite; and tabular inclusions of rutile, typically 60 by 20 p m in size. Micron-sized, anhedral grains of L a - , Ce- and Nd-bearing monazite and barite occur as inclusions and grains within vugs in the manganese carbonate aggregates. Kaolinite or orpiment and realgar locally fill vugs in the manganese carbonate aggregates. In several samples from the centre of the deposit, rhodochrosite hosts 1 to 15 (xm, elongate and anhedral, but occasionally lath-shaped inclusions of rhodonite. Figure 3.4 illustrates the restriction of rhodonite to the centre of the ore body. Siderite occurs in one sample slightly northwest of the centre of the deposit ( D D H 96-493.69). It fills vugs in the quartz matrix and occurs in veinlets with organic matter consisting of carbon and sulphur (possibly bitumen) and arsenopyrite-bearing  73 pyrite. Siderite hosts sub-micron-sized, anhedal inclusions of cassiterite and an unidentified chrome-bearing mineral.  3.2.3  Argillization  Clay minerals are present within most Jeronimo ore samples, typically comprising less than 10% but locally forming up to 60% of the sample. The identification of the clay minerals is important for the characterization of the hydrothermal fluids, as clay minerals may be used to determine fluid temperature and p H (see Reyes, 1990). A s the clay minerals are very fine-grained and could not be identified with certainty by optical microscopy, it was necessary to identify them by X R D . The samples were analyzed using a Siemens Diffraktometer  D5000 at 40.0 k V and 30.0 m A using C u  K-alpha radiation and scanned at 0.02°20 intervals for 1.25 seconds between 3 and 30 °20.  3.2.3.1 Distribution and characteristics  In hand sample, the strongly argillized ore is medium to dark grey-brown and is easily scratched by steel (Plate 3.17). During core logging, the degree of argillization was qualitatively assessed according to the softness of the core. Generally, there is a positive correlation of gold grade with argillization that is well illustrated by the drill core logs of D D H 73 and 91 (Figure 3.2). In thin section, clay mineral aggregates occur (1) in patches, typically 2 m m wide, disseminated throughout the matrix, (2) filling vugs within aggregates of carbonate and quartz (Plate 3.10), (3) intergrown with fine-grained matrix quartz pervasively replacing  74  Plate 3.17 Hand sample photo of pervasively argillized ore with white to light gray kaolinite in vugs. Sample DDH97-285.5.  Plate 3.18 Thin section photo of discontinuous, irregular bands of clay within a finegrained, vuggy, quartz-rich matrix. Field of view is 1.25 mm. Photo taken under transmitted light with crossed nichols. Sample DDH96-493.69.  75  76 the matrix, (4) as discontinuous, anastomosing bands within the matrix (Plate 3.18), and (5) as rare, irregular, discontinuous veinlets, on average 120 u m wide, cross-cutting the fine-grained quartz matrix (Plate 3.19).  3.2.3.2 M i n e r a l o g y The identification of the clay minerals in the Jeronimo deposit required the physical separation of the clay-size fraction (see Appendix B ) for two reasons. Firstly, the clay content in the Jeronimo ore is locally minimal (e.g. 2%) and bulk rock X R D patterns contained too many overlapping diffraction peaks for reliable evaluation of clay mineral peaks. Secondly, the identification of clay minerals by X R D is normally made by analyzing a sample of oriented clay particles. Such samples are useful for clay mineral identification as they enhance reflections of the basal (00/) planes, for which the lattice spacings are a diagnostic characteristic of each clay mineral. The presence of nonplatey minerals, such as quartz and pyrite, disturbs the preferred orientation of the clay platelets and results in broader diffraction peaks. A s clay minerals are normally less than 2 u,m in diameter, this size fraction was selected for clay mineral identification. Unfortunately, due to the fine-grained nature of Jeronimo mineralization, considerable quartz, carbonate and pyrite are also included in the under 2 um size fraction and it was not possible to make pure clay separates. Nonetheless, a kaolin-group mineral and illite were clearly identified according to X R D patterns of their basal reflections. Their presence was confirmed through diagnostic tests consisting of heating and ethylene glycolation, in addition to qualitative compositional analysis by E D S . Figure 3.7 shows  Plate 3.19 Thin section photo of irregular clay veinlet cross-cutting a matrix of finegrained quartz. Field of view is 0.625 mm. Taken under transmitted light with crossed nichols. Sample DDH86-562.S3.  78  representative diffraction spectra for illite and kaolinite from the Jeronimo deposit and Table 3.2 lists sample diffraction peaks of kaolinite and two polytypes of illite for reference.  3.2.3.2a Kaolinite The presence of a kaolin-group mineral, the most common of which are the polymorphs kaolininte and dickite ( A l S i 4 O i ( O H ) ) , was indicated on X R D spectra by 4  0  8  001 reflections ranging from 7.12 to 7.22 A and 002 reflections ranging from 3.58 to 3.59 A (Figure 3.7). Although kaolin- and chlorite-group minerals differ greatly in terms of composition and structure, there is strong overlap in the positions of their basal diffraction peaks. It is therefore necessary to heat samples bearing kaolin- or chloritegroup minerals for one hour at 550°C to distinguish between the two. This process induces the dehydroxylation of the hydroxide sheet in chlorite and results in an increase in intensity of the chlorite 001 reflection at about 14.2 A and a weakening of the 002, 003 and 004 reflections (Moore and Reynolds, 1997). The heat treatment causes kaolin-group minerals to become amorphous and eliminates their reflections from the spectra. Figure 3.7a displays X R D spectra of an air-dried and heat-treated sample and shows that, with heating, the 7.19 A peak is nearly eliminated and that there is no growth of a peak near 14.2 A , suggesting the mineral is of the kaolin-group. The presence of a kaolin mineral in this sample was confirmed by qualitative E D S analysis, indicating that it is constituted by A l , S i and O.  A.  Kaolinite Sample: DDH96-431.37  3.346 Quartz 101, i l l i t e 006  o U  Untreated  Heated  15  20  Degrees two-theta  Figure 3.7 Representative X-ray diffraction spectra of (A) untreated and heated kaolinite, and (b) untreated and glycolated illite.  80  Table 3.2 Examples o f diffraction peaks for kaolinite and two polytypes o f illite taken from M i n e r a l Powder Diffraction F i l e Data B o o k (1980), published by the Joint Committee on Powder Diffraction Standards ( U S A ) .  Kaolinite (IT) hkl d(A) I/Ii 001 100 7.17 020 35 4.478  Illite (1M) hkl I/Ii 001 80 002 80  110  60  4.366  020  60  4.52  111  35  4.186  80  3.63  111  14  4.44  111  45  4.139  112 003, 022  100 3.35  113  8  3.89  021  40  3.847  112  80  023  12  3.72  021 002  35 80  3.745 3.579  114 006  14 100  3.46 3.34  111 111  5 35  3.420 3.376  114  16  3.20  112  20  3.155  112  20  3.107  d(A) 10.0 5.03  3.10  Illite (2M,) hkl I/It 002 90 004 50 110, 020 16  d(A)  10.02 5.02 4.48  81  The X R D patterns of kaolinite and its high temperature polymorph dickite overlap strongly and are difficult to distinguish. T o discriminate between kaolinite or dickite, short-wave infrared (SWIR) reflectance spectroscopy was performed on kaolin-bearing samples using a P I M A (Portable Infrared Mineral Analyser) at Petrascience Consultants, Vancouver. S W I R detects the vibrational energy of molecular bonds within molecules that have bending and stretching modes within the 1.3 to 2.5 um region of the electromagnetic spectrum and is sensitive to water, hydroxide, ammonia and carbonate molecules. The interaction of the infrared radiation with the molecular bonds of these molecules results in absorption of the incident radiation at specific wavelength values that are characteristic for each molecular bond. The resultant troughs on S W I R spectra, which display the intensity of the returned radiation over a range of wavelengths, are known as absorption features and produce patterns that are distinct for individual mineral phases. Kaolinite and dickite are distinguished on S W I R spectra by their distinctive doublets at 1.4 and 2.2 um. Figure 3.8 displays the SWER. spectra of three Jeronimo samples bearing kaolin-group minerals and includes typical S W I R spectra of kaolinite and dickite for comparison. The asymmetry of the 2.2 u m doublet in the Jeronimo spectra, a diagnostic feature of kaolinite S W I R spectra, suggests that the kaolin-group minerals are kaolinite. However, the width of the 1.4 um doublet may indicate that minor dickite, which has a wider 1.4 um doublet than kaolinite, is also present. Inspection of kaolinite-bearing samples by S E M indicates that it tends to occur as monomineralic aggregates filling vugs within both the manganese and calcium carbonate and quartz aggregates and appears to have formed as a late phase (Plate 3.20).  82  1.4  1.6  1.8  2.0  2.2  2.4  Wavelength (um)  Figure 3.8 Short wave infrared reflectance spectra of three kaolin-group mineral-bearing Jeronimo ore samples and typical spectra of kaolinite and dickite shown for comparison. The strongly asymmetric doublets of the Jeronimo samples at 1.4 and 2.2 um are most similar to those of the kaolinite spectrum, although the width of the 1.4 um peak may indicate that a minor amount of dickite is present.  Plate 3.20 Scanning electron microscope back-scattered electron image o f kaolinite in vugs in rhodochrosite and quartz. Sample DDH69-161.09.  3.2.3.2b Illite The occurrence of illite, ( K , H 0 ) ( A 1 , M g , Fe) (Si, A l ) O [ ( O H ) , H 0 ] , was 3  2  4  1 0  2  2  determined from X R D spectra of glycolated samples with peaks ranging from 9.89 to 10.18 A (Figure 3.7b). Structurally, illite is similar to muscovite but contains more S i , M g and H 0 and less tetrahedral A l and interlayer K (Moore and Reynolds, 1997). Illite 2  particles are normally interlayered with an expandable clay, most commonly smectite. Illite may be considered a transitional mineral phase between the end members smectite and muscovite (Moore and Reynolds, 1997). With increasing temperature, smectite, (Ca, N a ) . ( A l , F e , F e , M g ) . ( A l , S i ) ( O H ) « n H 0 , loses interlayer water molecules and 2 +  0  3  3 +  2  3  4  2  2  exchanges C a ions for A l and K , forming illite. Provided the components A l and K are available from fluids or the host rock, the amount of illite in an illite-smectite mixture increases with temperature until pure illite is formed. W i t h increasing temperature, illite progresses through a series of polymorphs until it attains the crystal structure and composition of muscovite (Duba and Williams-Jones, 1983). To evaluate the presence and degree of smectite interlayering in the Jeronimo ore, the illite-bearing samples were glycolated. During glycolation, smectite expands parallel to the C-axis due to the replacement of interlayer cations by the polar organic compound ethylene glycol, while structural integrity is maintained along the A and B axes (Moore and Reynolds, 1997). With ethylene glycolation, the smectite 001 peak shifts from 15 to 16.9 A . Glycolation of the Jeronimo clay-size separates resulted in small shifts ranging from 0.01 to 1.11 A of the illite 001/ 002 peak towards higher values of °20 (Figure 3.7b).  85 In this study, the percentage of illite, or illite crystallinity, was assessed using three methods. In the first method, described by Moore and Reynolds (1997), the positions of the illite 001/002 and 002/003 peaks are used to assess the percentage of illite. The value of °A29, equal to the difference in the illite 001/002 and 002/003 peak positions (for C u K a ) , has been empirically derived for a range of percentages of illite in illite-smectite mixtures (Table 3.3). Values of °A20 for 19 Jeronimo glycolated clay-size separates (Table 3.4) show limited variation, ranging only from 8.25 to 9.0 °A20, which corresponds to values of 90-100% illite for all but one sample which contains 80-90% illite. These values indicate that the amount of smectite interlayering is consistently low. The second method used to assess the degree of smectite interlayering involves the measurement of the Kiibler index, which is defined as the width of a peak at half of its height above background. Kubler indices are most effectively used to assess relative differences within a sample set as their measurement varies between laboratories, according to machine conditions and sample preparation techniques (Robinson et al., 1990). According to Eberl and Velde (1989), the following two properties control the width of illite peaks, (1) the X-ray scattering domain size, or the size and distribution of coherent illite particles, and (2) the degree of structural distortion, which is controlled by the amount and distribution of expandable layers present between illite particles. For 19 Jeronimo clay-size fraction samples, the Kubler index, measured on the illite 001/002 peak of glycolated samples, varies from 0.36 to 1.09 °20 (Table 3.4), although systematic spatial variation in the value of the Kubler Index is not apparent throughout the deposit. In fact, the Kubler index shows large local variations (e.g. 0.43  86 °28 variation between two samples 6.23 m apart in DDH53) and can not be used to define broad spatial trends in illite domain size.  Table 3.3 Empirically derived values of °A29 for known percentages of illite in illite-smectite mixtures (from Moore and Reynolds, 1997). 001/002 002/003 % Illite Reichweite d(A) d(A) °26 °29 °A29 5.49 10 0 8.58 10.31 5.61 15.80 15.88 5.68 20 0 8.67 10.20 5.58 5.94 0 8.77 10.09 5.53 16.03 30 6.16 40 0 8.89 9.95 5.50 16.11 5.44 6.52 9.77 16.29 50 0 9.05 7.01 9.22 5.34 16.60 60 1 9.59 7.38 1 9.40 9.41 5.28 16.79 70 17.05 7.88 9.64 9.17 5.20 80 1 8.38 9.82 5.10 17.39 90 3 9.01  Table 3.4 Calculated values of °A29, corresponding illite percentages, and Kubler indices measured on glycolated clay-size separates of Jeronimo ore zone samples. Kubler Index % illite Sample 001/002 092/003 °A29 (°29) (°29) (°29) 90-190 0.36 17.55 8.99 DDH39-398.97 8.65 9.42 90-109 17.55 8.99 DDH53-636.10 8.65 9.85 8.70 99-199 8.80 17.59 DDH53-642.33 9.79 8.79 99-109 DDH59-691.78 8.70 17.49 1.99 8.59 99-109 DDH59-695.16 8.75 17.25 99-199 9.97 8.80 17.55 8.75 DDH59-703.50 99-199 9.73 17.50 8.69 DDH72-745.31 8.90 9.79 8.55 99-199 DDH73-284.16 8.80 17.35 99-199 1.99 8.85 17.25 8.49 DDH73-285.92 9.55 99-199 8.65 17.55 8.99 DDH83-542.60 89-99 9.97 DDH86-545.32 9.00 17.25 8.25 1.99 8.55 99-190 DDH86-562.83 8.80 17.35 99-190 9.85 17.59 8.79 DDH91-568.89 8.80 9.85 90-199 8.80 17.49 8.69 DDH91-574.39 99-199 9.69 DDH96-431.37 8.75 17.59 8.75 99-199 9.67 DDH96-493.69 8.65 17.65 9.09 99-199 9.85 8.75 17.55 8.89 DDH97-285.78 99-199 9.73 DDH97-414.39 8.75 17.55 8.89 8.79 99-199 1.93 DDH100-175.23 8.80 17.59  87  S W I R was also used to assess variation in illite composition in the Jeronimo claysize separates. Post and Noble (1993) have shown that the 2.2 um A l - O H absorption feature on S W I R illite spectra shifts from 2.202-2.210 um for illites of low aluminum content to 2.192-2.198 um for those with a higher aluminum content. For all of the Jeronimo ore zone samples, the position of the A l - O H trough varied only from 2.200 to 2.212 um, indicating that there is no significant change in the aluminum content of the ore zone illites and that it is consistently low. However, the presence of fine-grained pyrite in the clay separates caused many of the S W I R spectra to be noisy and the exact position of the A l - O H absorption feature was difficult to assess. Inspection by S E M indicates that illite occurs preferentially in patches around the rims of fossils and grains of quartz, carbonate and pyrite, and finely disseminated within the vuggy silicified matrix. It is often stained a brown-yellow colour or dusted by veryfine grained black material, possibly an oxide or organic matter. Ulite-filled vugs may host grains of pyrite, arsenopyrite, galena, sphalerite, coloradoite, cassiterite, rutile, gold, barite, monazite-(La) and apatite. It also commonly occurs as a very fine-grained mixture with apatite and pyrite (Plate 3.21).  3.2.3.3 Spatial variation in clay mineral distribution Overlapping the centre of the deposit, an elliptical zone spanning a distance of approximately 1.0 by 0.5 km consists of ore bearing both of the clay minerals kaolinite and illite (Figure 3.9). This zone also hosts highly variable quantities of clay minerals, ranging from 1 to 30%. Outside of the zone, the clay assemblage consists only of illite, which occurs in more restricted volumes of between 2 and 10%. The origin of clay in  Plate 3.21 Scanning electron microscope back-scattered electron image o f band o f fine-grained illite and apatite with framboidal pyrite, cross-cutting matrix o f vuggy quartz and pyrite. Sample DDH69-161.09.  89  90 the host rocks - primary, diagenetic or hydrothermal - is difficult to determine, but hydrothermal processes have likely partially altered most units throughout the area.  3.2.4 Organic matter Organic matter consisting of C and S was identified by E D S analysis in several polished sections. In one sample, it is present as fine-grained amorphous material in typically 20 urn, irregular veinlets with siderite and pyrite with arsenopyrite inclusions. Silicified bioclasts are commonly coated by very fine-grained, dark gray, nonreflective material (Plate 3.22). This material may consist of organic matter but may also be a fine-grained oxide.  3.2.5 Hypogene alteration relationships The preservation of internal and external bioclast structures by secondary quartz indicates that decalcification and replacement of the E l Asientos limestone by silica occurred simultaneously. However, the presence of vugs lined with drusy quartz in the siliceous matrix also implies that localized bulk dissolution and removal of the host rock occurred. The presence of the manganese carbonate minerals in vugs within the siliceous matrix and bioclast centres, along with the sporadic occurrence of manganese carbonate veinlets cross-cutting quartz aggregates indicates that manganese carbonate mineralization followed silicification. The localized occurrence of quartz grains within vugs in the manganese carbonate aggregates implies that minor silicification also followed manganese carbonatization. The intergrowth of the calcium and manganese  93  Plate 3.22 Thin section photo of fine-grained black material, possible organic matter or a fine-grained oxide mineral, rimming silicified bioclasts. Field of view is 5.1 mm. Taken under transmitted light with crossed nichols. Sample DDH97-421.84.  92 carbonate minerals and the similarity in their mode of occurrence implies that the two were deposited simultaneously. The manganese carbonate minerals, which dominate in a zone overlapping the centre of the deposit, form the more proximal phase, while calcium carbonate minerals comprise the more distal carbonate phase. V e i n and fracture infills of calcite, often associated with orpiment, realgar and cinnabar, appear to be a late phase and show no relation to manganese carbonate mineralization. The occurrence of illite patches and finely dispersed illite within the siliceous matrix throughout the entire deposit, suggests that illite deposition may have occurred during silicification. However, the presence of illite in vugs in quartz and carbonate aggregates indicates that its depositional also followed both of these phases of alteration. The deposition of kaolinite, uniquely as a vug-filling mineral within quartz and manganese carbonate aggregates, undoubtedly followed both silicification and manganese carbonatization. It is not possible to determine with certainty whether kaolinite, which is restricted to the centre of the deposit, is of supergene or hypogene origin, however the absence of dickite indicates that the kaolin-group mineral formed at relatively low temperatures.  3.3 Ore mineralogy The majority of the ore minerals within the Jeronimo deposit are restricted to very fine grain sizes, typically ranging from several hundred um to less than one pm, necessitating the use of a S E M to identify the ore minerals and examine their textures and relationships.  93  3.3.1 Mineral characteristics The major sulphide and sulphosalt minerals in the Jeronimo ore deposit are pyrite, arsenopyrite, sphalerite, orpiment, realgar, galena and one or more lead sulphosalts, including probable twinnite or guettardite. Pyrite is ubiquitous throughout the deposit and forms very fine-grained, euhedral to anhedral, commonly vuggy crystals, normally ranging from 0.5 m m to 2 |J,m, but typically 20 um. It occurs mainly as disseminated grains within quartz and carbonate aggregates and is intergrown with illite in vugs. Pyrite forms rare, diffuse, discontinuous veinlets (typically 60 to 100 um) that cross-cut the silicified matrix, and bands that concentrically rim silicified bioclasts. In these instances, pyrite grains are typically adjacent to or intergrown with subhedral rutile (Plate 3.23). Pyrite occurs locally in isolated framboids, typically 10 to 15 (am in diameter, composed of aggregates of micronsized, rounded, spherical pyrite grains or as clusters of framboids typically 80 by 60 um wide. The framboids occur within quartz and manganese carbonate aggregates and are also commonly associated with rutile. Arsenopyrite normally occurs as (1) inclusions and veinlets filling fractures in vuggy pyrite and lining pyrite rims (plates 3.24 and 3.25), (2) as subhedral to euhedral rhomb-shaped, typically 50 by 35 um, acicular grains, ranging from 20 by 4 um to 100 by 10 um, disseminated in massive quartz and manganese carbonate (Plate 3.26), and (3) as small, anhedral to subhedral inclusions or grains within vugs. Pyrite hosting arsenopyrite inclusions, veinlets and rims occurs as disseminated grains in rhodochrosite, kutnohorite, siderite and quartz. The equant, anhedral arsenopyrite inclusions, ranging from less than 1 um to 20 um, tend to fill vugs within pyrite and are normally  94  Plate 3.23 T h i n section photo of pyrite (beige) and rutile (gray) in diffuse band w i t h i n the quartz matrix. F i e l d o f v i e w is 0.625 m m . T a k e n under reflected light. S a m p l e DDH96-493.69.  100 Lim Plate 3.24 Scanning electron microscope back-scattered electron image o f arsenopyrite inclusions in porous pyrite grains intergrown with siderite. Sample D D H 9 6 - 4 9 3 . 6 9 .  96  Plate 3.26 Scanning electron microscope back-scattered electron image o f a diffuse band o f arsenopyrite, pyrite and rutile in a fine-grained quartz-illite matrix. Sample DDH93-533.07.  98 concentrated near or at the rims of the pyrite grains. Straight and irregular fractures in pyrite, locally occurring along cleavage planes, are filled by sub-micron-wide, discontinuous arsenopyrite veinlets. Arsenopyrite tends to rim only small portions of the pyrite outer surface in discontinuous layers typically 1 um wide and 6 p m long. Arsenopyrite grains also occur filling vugs in a quartz veinlet cross-cutting a manganese carbonate aggregate. The larger, disseminated, rhomb-shaped arsenopyrite grains host subhedral to anhedral, equant, typically 12 p m inclusions of pyrite and rutile and also locally occur in embayed contacts with rutile and apatite (Plate 3.26). In two samples, nickeliferous arsenopyrite occurs as anhedral, round, micron-sized grains in vugs in quartz and rhodochrosite. Disseminated sphalerite grains are anhedral to subhedral and occur dominantly within manganese carbonate aggregates. They range in size from sub-micron-sized equant, anhedral grains to subhedral, tabular, typically 80 by 35 um grains. In several samples, generally 35 p m sphalerite grains form 4 mm wide diffuse bands, with pyrite and minor apatite, cross-cutting the silicified matrix (plates 3.27 and 3.28). Sphalerite also locally occurs as micron-sized to 100 pm grains in vugs in quartz and carbonate aggregates. Galena is normally micron-sized and anhedral, occurring within vugs in quartz and manganese carbonate. It occurs locally as larger, typically 40 um, subhedral, cubic inclusions in quartz aggregates. Galena grains are locally intergrown with or occur as inclusions in lead sulphosalt grains. In two samples, galena is intergrown with an aluminum-phosphate-sulphate (APS) mineral, described in section 3.5.1, in a vug in quartz.  99  Plate 3.27 Hand sample photo of silicified ore with sphalerite (beige) in diffuse patches cross-cut by fine realgar veining. Sample DDH97-273.06.  Plate 3.28 Thin section photo of band of sphalerite (light gray) and pyrite (white) crosscutting quartz matrix. Field of view is 5.1 mm. Taken under reflected light. Sample DDH97-287.25.  100  101  One or more lead sulphosalt minerals are present mainly as tabular to acicular grains, ranging from 8 by 1 um to 45 by 6 um, in vugs in quartz (Plate 3.29). The lead sulphosalt grains are locally intergrown with barite, galena and the A P S mineral, and occur as inclusions in compact aggregates of orpiment filling vugs in quartz. S E M - E D S analyses indicate that the lead sulphosalt grains range from antimony- to arsenic-rich. A s at least six lead sulphosalt minerals are known with extensive solid solution between antimony and arsenic (Wuensch, 1974), it is not possible to say which or how many of the lead sulphosalt minerals are present without electron microprobe analyses and the fine grain size of the majority of the lead sulphosalt grains precluded this. However, one sample contained aggregates of acicular lead sulphosalt grains up to 450 by 300 urn in size, which were large enough to physically separate and analyze by X R D . The X R D spectrum indicated that the mineral may be twinnite (Pb(Sb, A s ) S ) or guettardite 2  4  (Pb (Sb, A s ) S ) . 9  1 6  3 3  Orpiment and realgar normally occur in late veinlets with calcite and minor, typically 10 um cinnabar grains (Plate 3.30). The arsenic sulphides also occur as vugfilling minerals in manganese carbonate aggregates. Locally, realgar forms large, typically 1 cm, vug-filling aggregates hosting inclusions of subhedral lead sulphosalt(s), sphalerite, cinnabar, apatite, coloradoite (HgTe), arsenolite (AsO), and gold (plates 3.31 and 3.32). The deposit hosts a number of minerals that occur locally as sub-micron-sized grains. Typically 600 nm grains of cassiterite are present as rounded inclusions in rhodochrosite, kutnohorite, siderite, quartz and pyrite, and locally occur in illite-filled vugs. Rare stibnite is present as anhedral, sub-micron-sized inclusions in rhodochrosite.  Plate 3.29 Scanning electron microscope back-scattered electron image o f acicular Pb-sulphosalt grains and blocky realgar in vugs in a quartz matrix. Sample DDH53-641.17.  103  Plate 3.30 Hand sample photo of late orpiment. realgar and calcite mineralization. Sample DDH86-545.72.  Plate 3.31 Scanning electron microscope back-scattered electron image o f a realgar, Pb-sulphosalt and rutile aggregate in a quartz matrix. The realgar has locally altered to arsenolite. Sample DDH97-273.06.  Plate 3.32 Scanning electron microscope back-scattered electron image of a gold grain encapsulated i n realgar that is locally altered to arsenolite. Sample DDH97-273.06.  106  A copper-silver-antimony sulphosalt grain, probably freibergite [(Ag, C u , Fe)] (Sb, 2  A s ^ S n ) ] , occurs as a micron-sized inclusion in quartz. T w o telluride minerals are present locally, coloradoite (HgTe) and altaite (PbTe) and occur as micron-sized inclusions in quartz. S E M - E D S analysis indicates that A g and B i are the main constituents of micron-sized grains in vugs in quartz. It is possible that the metals are present as native elements, however the grains were too small to obtain pure E D S analyses. Similarly, Cr-rich micron-sized grains are present as inclusions in siderite, however they also occur too finely to determine the mineral species.  3.3.2 Modes of gold occurrence Fifteen gold grains were located in six polished sections by S E M . G o l d in the Jeronimo deposit occurs as anhedral, sub-micron to micron-sized free grains that are not related to a specific mineral phase and instead occur in association with a variety of minerals. Table 3.5 describes the size, shape and occurrence of the gold grains. E D S analyses indicated that minor quantities of silver are contained within the gold, however they did not meet the requirement of 30-45% A g designated to constitute electrum ( A u A g ) according to Ramdohr (1980). G o l d grains most frequently occur at the linings of vugs in quartz, which may be filled by fine-grained illite (Plate 3.33) or exist as open space (Plate 3.34). T w o gold grains are encapsulated by quartz (Plate 3.35). G o l d grains are also present as inclusions in grains of pyrite and in arsenopyrite inclusions in pyrite that are disseminated in the vuggy quartz matrix (plates 3.36 and 3.37). In one sample, gold grains occur as inclusions within large realgar aggregates (Plate 3.32). G o l d grains were also observed in one sample that has undergone supergene oxidation (Plate 3.38), discussed in section 3.6.  Plate 3.33 Scanning electron microscope back-scattered electron image o f gold grain at margin o f illite-filled vug in quartz. Sample DDH97-421.84.  Plate 3.34 Scanning electron microscope back-scattered electron image of gold grain in open vug in quartz. Quartz also hosts framboidal pyrite. Sample DDH97-378.78.  109  Plate 3.35 Scanning electron microscope back-scattered electron image o f a gold grain encapsulated in quartz intergrown with kutnohorite. One grain of altaite (PbTe) also present in quartz. Sample DDH100-354.86.  110  Pyrite f  Arsenopyrite  6  Jim m\\m  ,J|  *^^  ^^Bfe  Gold •  •  V  5 Lim  Plate 3.36 Scanning electron microscope back-scattered electron image o f two gold grains encapsulated in pyrite. Pyrite also hosts arsenopyrite inclusions. Sample DDH96-493.69.  Plate 3.37 Scanning electron microscope back-scattered electron image o f a gold grain encapsulated in arsenopyrite located at the rim o f a pyrite grain. Sample DDH96-493.69.  Plate 3.38 Scanning electron microscope back-scattered electron image o f gold grain in vug in amorphous manganese oxide in oxidized sample. Sample DDH100-188.53.  113  In at least two circumstances, gold grains occur in proximity to grains of the telluride minerals coloradoite (HgTe) and altaite (PbTe). Although gold was not detected by E D S analysis in the arsenopyrite inclusions, veinlets and rims associated with pyrite, it is possible that gold is present in concentrations below the E D S detection limit. Analysis by more sensitive techniques such as secondary ion mass spectrometry could be used to test whether gold is present either structurally bound or as very fine grained (i.e. < 0.1 urn) particles within the sulphides, in addition to occurring as free grains (e.g. Cabri et al., 1989).  3.4 Geochemistry Homestake geologists have sampled mineralized drill hole intercepts for the analysis of the following suite of elements: A g , A l , B a , B e , B i , C a , C d , C o , C r , C u , Fe, Ga, H g , K , L a , M g , M n , M o , N a , N i , P, Pb, Sb, Sc, Sr, T i , Te, U , V , W , Z n . A t the mine site, H Q and NQ-sized drill core was cut in half lengthwise with a rock saw and half of the core was sent to the Chemex Laboratory in Reno. The samples were digested using a nitric acid-aqua regia leach procedure, which only partially digests the elements A l , B a , Be, C a , Cr, G a , M g , K , Sc, N a , Sr, TI, T i , W and analyses of these elements are only semi-quantitative. The resultant solutions were measured by inductively-coupled plasma atomic emission spectroscopy, with the exception of H g , which was measured by cold vapour atomic absorption spectroscopy. The geochemical data used for interpretation in the present study were restricted to the diamond drill holes logged by the author and are presented in Appendix C .  114  Table 3,5 Descriptions of gold occurrences in the Jeronimo deposit.  Sample DDH96493.69  Size 930 x 660 nm (Figure 3.37)  Shape Anhedral, rounded, earshaped  Occurrence Inclusion in subhedral, tabular, 14.75 x 6.25 pm arsenopyrite grain hosted in pyrite grain disseminated in vuggy quartz matrix  288 x 250 nm  Anhedral, spherical  Inclusion in anhedral, 4.9 x 1.6 p m pyrite grain disseminated in vuggy quartz  863 x 735 nm  Anhedral, rounded, tadpole-shaped  Inclusion in subhedral, 25 x 13.5 p m arsenopyrite grain at edge of large pyrite aggregate  713 x 368 nm (Figure 3.36)  Anhedral, rounded, elliptical Anhedral, rounded, spherical Anhedral, elongate, irregular Anhedral, elongate  A s inclusion in anhedral, 55 x 30 p m pyrite grain in vuggy quartz matrix  441 x 375 nm (Figure 3.36) DDH97273.06  1.13 x 0.477 pm 962 x 385 nm (Figure 3.32)  DDH97378.78  DDH97421.84 DDH100188.53  DDH100354.86  538 x 201 nm  Anhedral, tadpole-shaped  183 x 143 nm (Figure 3.34) 340 x 128 nm  Anhedral, spherical Anhedral, elliptical  1.02 x 1.00 pm (Figure 3.33) 210 x 177 nm  Anhedral, equant, with scalloped edges  298 x 255 nm (Figure 3.38) 1.89x0.816 pm (Figure 3.35) 251 x 246 nm  Anhedral, spherical Anhedral, elliptical  Anhedral, elliptical  Anhedral, spherical  As inclusion in anhedral, 55 x 30 um pyrite grain in vuggy quartz matrix Inclusion in 2 x 1.5 cm massive realgar aggregate in quartz matrix, in proximity to coloradoite. Within pocket of 3.9 x 2.7 p m octahedral arsenolite grains in 2 x 1.5 cm massive realgar aggregate in quartz matrix A t edge of 11.5 p m pyrite grain within vug in quartz At edge of vug in quartz Within vug in quartz, surrounded by finegrained illite and sub-micron-sized anhedral pyrite grains At boundary of illite-filled vug in quartz  Within a vug in a Manganese-oxide mineral which has filled a vug in quartz, within 5 pm of 2 sub-micron-sized cassiterite grains Within a vug in a Manganese-oxide mineral which has filled a vug in quartz A s inclusions encapsulated in quartz, in proximity to sub-micron sized grains of altaite and monazite  115  Trends of the elements B e , B i , C d , G a , N a , Sc, T i , U and W were not considered as their concentrations are mainly at detection limits. Polished section observation by S E M indicated that micron-sized fragments of a C u - N i - Z n alloy are present in vugs in some samples. These fragments must originate from contamination by machinery during the sampling or sample preparation processes. Contamination may have occurred at the mine site during diamond drilling or during the halfing of the core with a rock saw, or at U B C during polished section preparation. Powdered samples may also be contaminated during the process of pulverization by the addition of Fe, C o , C r , C u , M o , M n , N i and V from iron and steel equipment (Fletcher, 1981). A s a result of these potential sources of contamination, trends in the elements C u , N i , C o , C r , M o and V were not considered in this study. Trends in the elements Z n , Fe and M n were still considered as the minerals sphalerite, pyrite and carbonate are present in such high quantities that they are likely to overwhelm any potential effects of contamination. The average concentrations of elements in ore grade samples from the Jeronimo intercepts logged in this study are presented in Table 3.6. The Jeronimo ore is characterized by high concentrations of A s , M n , Z n , Pb, Sb, A g and H g .  Table 3.6 Elemental abundance in Jeronimo ore with gold greater than ore equal to 1 g/t based on geochemistry only from holes logged in this study (140 samples).  Element A g (ppm) A l (%) A s (%) B a (ppm) C a (%) C u (ppm) Fe (%) H g (ppm)  Average 9.28 0.41 0.94 25.68 2.53 9.96 4.15 10.51  Range 0.1-347 0.06 - 2.29 0.02 - 42.4 0.5-159 0.02-15 0.5-142 0.65 - 10 0.08 - 100  Element K(%) L a (ppm) M g (%) M n (ppm) N i (ppm) Pb (ppm) Sb (ppm) Z n (ppm)  Average 0.14 12.16 0.49 26956 13.11 2132 145.46 5439  Range 0.005 -0.81 0.5 - 57 0.005 - 2.96 35 - 257,400 0.5 - 40 1 - 53,900 2 - 2000 6 - 97,000  116  The correlation of elements with gold can be visually assessed by comparing element distributions across the ore intercepts. The iogs presented in Figure 3.2 reveal that gold distribution is often saddle shaped with an upper and lower peak. The absence of mineralization between the peaks corresponds to a decrease in host rock grain size, most likely in response to lowered porosity and permeability. Figure 3.10 shows the variation in the distribution of the elements A g , A s , B a , C a , Fe, H g , K , M n , Te, P, Pb, Sb and Z n across four intercepts in drill holes D D H 59, 86, 91 and 96, which are located in the centre and margins of the ore zone (Figure 3.1). Table 3.7 summarizes the behaviour of these elements across the ore intercepts with respect to gold. It is apparent from different drill core intercepts that individual elements may have multiple relationships with gold.  Table 3.7 Summary of element behaviour with respect to gold in the ore zone intercepts of DDH 59, 86, 91 and 96.  Elements that correlate with overall gold distribution Elements that correlate with one gold peak when gold distribution is saddle-shaped Elements that correlate negatively with gold Elements that show no downhole correlation with gold in the drill holes studied  A s , H g , Te, Pb, Sb, Fe, M n , (K, P, Zn) A g , A s , H g , M n , Pb and Sb Ba, C a , M n , (Mg) Al, La  Elements that show good overall correlation with gold are A s , H g , Te, Pb, Sb, Fe, M n , while K , P and Z n correlate weakly with gold. In cases where the distribution of gold is saddle-shaped, the elements A g , A s , H g , M n , Pb and Sb tend to correlate strongly with the upper gold peak and show only weak correlation with the lower peak. Elements that consistently form as envelopes to the gold mineralization are B a and Ca. The  117  Figure 3.10 Concentrations of selected elements across ore zone intercepts of diamond drill holes DDH59, DDH86, DDH 91 and DDH 96 (Lower Jeronimo, 2nd intercept). Geochemical data are listed in Appendix C. On each plot, the concentration of the selected element is represented by a solid line while gold concentration, which is included for reference, is shown as a dotted line.  Au (ppm) 5  Au (ppm)  Au (ppm) 10  4  530  8  DDH 86  12  470  DDH 96  540 480 550  490  560  720  570 0  100  200  300  Ag (ppm)  400  0.0  Au (ppm)  2.5  5.0  Ag (ppm)  500  0.0  0.3  0.9  1.2  Au (ppm)  Au (ppm) 53<?r  0.6  Ag (ppm)  «  4  47<?,  *  5  It  DDH 96  540  48 550:  \  490 560  !  "3L  cr  As (ppm) 720  0  16000  32000  570,  25  As (wt %)  As (ppm)  10  0 560  4  8  0  0  BTj  Too  Hg (ppm)  150  590  5000  16  8  470  12  480  490  580  2  3000  As (ppm) Au (ppm)  12  570  7  500 0 1000  Au (ppm)  Au (ppm) 5  50  0  2  4  Hg (ppm)  6  0  2  4  Hg (ppm)  6  119  Au (ppm) 5  A u (ppm) 10  4  0 530  8  12  540  550  560  100  Ba  200  570  (ppm)  A u (ppm) 5  Ba 10  530  250 (ppm)  500  Ba  A u (ppm) 4  )  0  8  470  DDH 86  540  *  100 (ppm)  A u (ppm) 4 8  1  200  12  DDH 96  480  550 490  5~  560  10  20  570  J  10  Ca (wt%)  0  560  0  4  8  12  470  0  4  8  480  580  590  8  490  14 Mn (wt%)  28  12  A u (ppm)  16  570  Mn (wt%)  4  Ca (wt%)  A u (ppm) 10  r_  500  20  Ca (wt%)  A u (ppm)  5  •  1  500 £  j M  j n  (  w t  o  / o )  12  Au (ppm)  Au (ppm)  Pb  0  5  560  670  Au (ppm) 16  DDH 91  4  470  8  12  DDH 96  680 570  480  580  490  690  4500  9000  590  75  Pb (ppm)  6 7 0  o  150  l)  0  530  20  30  40  Au (ppm)  Au (ppm)  5  10  Pb (ppm)  Pb (ppm)  Au (ppm) Sb  JT  t—r  50(  4  470  0  4  8  12  5  10  15  480  350  Sb (ppm)  700  250  Au (ppm) Z  n  500  Sb (ppm)  Sb (ppm)  Au (ppm)  670 °  470  4  8  12  480  30000  Zn (ppm)  60000  J  0  25  Zn (ppm)  50  Au  Au (ppm) 670  0  5  10  0 4 7 0  (ppm)  4  8  12  DDH 96  1  480  •""I 490  500 0  Fe (wt%)  670  5  2  3  4  5  F e (Wt %)  Au (ppm) 0  1  10  680  530  0  Au (ppm) 4  540  690 550  560  570  150  0 670  300  TI (ppm)  TI (ppm)  Au (ppm)  Au (ppm)  5  600  Au (ppm) K  470°  4  8  12  680  690  1200  P (ppm)  2400  2500  P (ppm)  5000  0.125 K (wt%)  0.250  122  relationship of Mn to gold varies spatially within the deposit. At the centre of the deposit (DDH 96), Mn overlaps strongly with gold, however in more distal intercepts it either correlates well only with the upper gold peak or forms an envelope around the ore zone. Positive correlation of Pb and Sb suggests that geochemical trends of these elements are dominantly controlled by the occurrence of the lead sulphosalt mineral(s). Similarly, Hg and Te correlate well, indicating that the occurrence of coloradoite dominantly influences trends in these elements. Inspection of ore sample polished thin sections by optical microscopy and SEM revealed which minerals have produced the observed geochemical trends (Table 3.8). The relationships summarized in Table 3.7 suggest that gold mineralization is related to pyrite, arsenopyrite, coloradoite, lead sulphosalt(s), orpiment, realgar and manganese carbonate minerals. This correlation does not necessarily imply that there is a genetic relation between gold and these minerals. For example, the arsenic sulphide minerals are recognized to be a late phase, either reflecting changing conditions in the latter stages of mineralization or the passage of a second fluid through the ore-bearing units.  123 Table 3.8 Minerals responsible for trends in specific elements within ore intercepts and their typical mode of occurrence. Element Mineral Occurrence in Jeronimo Native silver? Micron-sized grains in vugs in quartz. Also, gold grains Ag host minor silver. Al Illite Illite is disseminated and in patches throughout the Kaolinite silicified matrix. Kaolinite fills vugs in carbonate and quartz aggregates. Orpiment Arsenopyrite occurs dominantly as disseminated, 50 by As Realgar 35 um grains. Realgar and orpiment occur in late Arsenopyrite veinlets and vug infills, locally hosting gold. See below lead sulphosalt(s) for lead-sulphosalt occurrence. Ba Barite Typically 300 by 20 p m grains in diffuse aggregates within the quartz matrix. Carbonate Calcium carbonate minerals fill vugs in quartz mainly at Ca the margins of the deposit and occur in late veinlets with orpiment and realgar. Fe Pyrite Ubiquitous, disseminated 20 um grains. A l s o in diffuse veinlets with rutile. Cinnabar Cinnabar occurs as 10 p m grains in late orpiment, Hg Coloradoite realgar, calcite veinlets. Coloradoite occurs as micron(HgTe) sized inclusions in quartz. Disseminated and in patches throughout the silicified K Illite matrix. Monazite occurs as micron-sized grains in vugs in Monazite La quartz. REE-calcite occurs as aggregates of dendritic REE-calcite crystals in the silicified matrix. M g is present in calcium-rich carbonate minerals that fill Carbonate Mg vugs in quartz mainly at the margins of the deposit. Manganese carbonate minerals fill vugs in quartz mainly Mn Carbonate at the centre of the deposit. P Apatite Typically 20 um grains are disseminated and line vugs in quartz. Also in diffuse aggregates with pyrite and rutile. Lead sulphosalt(s) Lead sulphosalt grains occur in vugs as isolated and Pb Galena aggregated grains. Galena is micron-sized to 40 u m in vugs. See above for lead sulphosalt occurrence. Stibnite Sb Lead sulphosalt(s) Stibnite occurs as micron-sized inclusions in rhodochrosite. Micron-sized inclusions in quartz. Te Coloradoite (HgTe) Altai te (PbTe) Zn Sphalerite Disseminated 80 by 35 p m grains. A l s o in diffuse bands with pyrite and apatite  124  3.5 Rare minerals The Jeronimo ore deposit hosts several uncommon minerals. The following section describes the compositional and textural characteristics of three of these minerals and discusses the significance of their occurrences in the deposit.  3.5.1 Aluminum-phosphate-sulphate minerals Minerals of the aluminum-phosphate-sulphate ( A P S ) group are isostructural with alunite ( K A 1 ( S 0 ) ( 0 H ) ( 5 ) , however the S 0 " molecule is partially to entirely substituted 2  3  4  2  4  by P 0 " . The resultant excess negative charge is compensated by the coupled 3  4  substitution of the divalent or trivalent C a , S r , B a , P b 2 +  2+  2 +  2 +  or C e  3 +  cations for the  univalent K cation (Stoffregen and Alpers, 1987). +  E D S analyses indicated that an A P S mineral is present in three samples within strongly silicified ore from the centre of the Jeronimo deposit. This ore contains little or no carbonate and hosts minor illite in pockets in the silicified matrix. The mineral is intermediate between the svanbergite ( S r A l ( P 0 ) ( S 0 ) ( O H ) ) and woodhousite 3  4  4  6  ( C a A l ( P 0 ) ( S 0 ) ( O F f ) ) end members, and is accompanied by minor substitution of B a 3  4  4  6  for Sr and C a . The A P S grains tend to form as aggregates of bladed crystals (typically 20 by 1 pm) within vugs in the quartz matrix (Plate 3.39). They are locally intergrown with anhedral, sub-micron-sized grains of galena, an acicular lead sulphosalt mineral and finegrained illite. Apatite, which is present throughout the ore zone as a minor constituent, is absent from the APS-bearing samples. According to Stoffregen and Alpers (1987), A P S minerals may form as a result of the dissolution of apatite, as in the case of the  Plate 3.39 Scanning electron microscope back-scattered electron image o f acicular aluminum-phospate-sulphate mineral with illite in vugs in quartz. Sample DDH97-421.84.  Summitville gold-copper deposit in Colorado, where primary apatite in the host rock and hydrothermal apatite in clay-rich alteration zones is absent from APS-bearing quartz kaolinite and quartz-alunite alteration zones. Stoffregen and Alpers (1987) determined that acidic fluids with high total dissolved phosphate contents are required to precipitate woodhousite. They found that A P S minerals associated with mineral deposits, such as porphyry-Cu, epithermal gold and polymetallic vein deposits were restricted to zones of advanced argillic alteration. In the Jeronimo deposit, the intergrowth of the A P S mineral with illite suggests that it is associated with argillic alteration.  3.5.2 Rare earth element-bearing calcite Rare earth element (REE)-bearing calcite is present in three strongly silicified samples at the centre and margins of the ore zone (Figure 3.4). The samples also host calcium and manganese carbonate minerals, illite and kaolinite in vugs. E D S analyses indicated that the calcite hosts Y and the light rare earth elements L a , N d and Ce. They occur in the vuggy quartz matrix as dendritic crystals, typically 4 by 0.5 urn, in aggregates that range in size from 5 by 15 p m to 20 by 25 p m (Plate 3.40). Locally, the REE-bearing calcite is intergrown with titanite. One aggregate occurs within a vein of manganese-bearing dolomite and pyrite. The significance of the REE-bearing calcite is discussed in Section 3.7.  20 Lim  Plate 3 . 4 0 Scanning electron microscope back-scattered electron image of an aggregate of rare earth element-bearing calcite grains in a vuggy quartz matrix. Sample D D H 6 9 - 1 6 1 . 0 9 .  128 3.5.3 Manganese sulphides Alabandite (MnS) and hauerite  (MnS2),  two rare sulphides of manganese, are  present as vein-forming minerals within the sandy bioclastic limestones of unit E outside of the Jeronimo ore zone (Figure 3.11). Alabandite is also present within one ore zone sample as small inclusions in calcite. The vein-forming alabandite and hauerite occur in intensely hydrobrecciated, argillized and silicified host rock with abundant quartz, gypsum and calcite veining. The strongly coloured, forest green alabandite forms anhedral, subrounded, typically 0.2 mm crystals in discontinuous, 1 to 5 mm veins. Alabandite veins, which brecciate quartz veins and illitized host rock, are cross-cut by veins of dark red hauerite varying from 0.1 to 0.5 mm in width (Plate 3.41). Hauerite also occurs as isolated veinlets, up to 2 mm wide cross-cutting the illitized matrix. The alabandite and hauerite veins are brecciated and cross-cut by gypsum veins. Although hauerite veins often cross-cut alabandite veins, hauerite grains also occur as inclusions within alabandite, suggesting that the crystallization of these two minerals overlapped temporally. Sulphide minerals that occur in association with the vein alabandite and hauerite include pyrite, Mn-bearing sphalerite, galena and oldhamite (CaS). EDS analysis indicated that the Mn-bearing sphalerite hosts minor indium. Other minerals associated with the vein Mn-sulphides include barite, apatite, monazite, rutile, and lead and copper sulphosalts. In one ore zone sample (DDH69-161.09), fine-grained alabandite occurs as anhedral, rounded, typically 5 um inclusions in coarse-grained calcite filling the vuggy centres of silicified bioclasts. The calcite also hosts subhedral, typically 10 um pyrite  129  x  c .a u  O  _  BO g  >. 5 X  o  O J3  o  2 .2 8  J3  J3  " S3 "ea  U  C  c o N o  f l  -a —  O  4J  G  on 3u  -  •S  w  cfl  60 i  E § E 2  JS •& E  » S  _C c fl  .3 —  Cfl  u  ta c o £ c o  c  o 1c  oo  c o  —  u i o n  3  O  GJ  C  S  Q  S"  '3  O -j o JS o o •-  •K Cfl  3  3 H  o  p o «  5 "2  » S u u u ca  c2 g O  M  o  o o  O  in  «r~j ON —  4.  l ia  t i  (fl  S  1 -s«  o  III D. Q  f  fe  .a  o  EA  -M  — —  5  E—  ^ o  8  5o  ca A a. 00  g a —, a  00  ' cj o  73  IS  c  tL, T3  3 O  «  <D ti cr  § 2  —  •s o -a  <2  r-> u  —i  i  & •= T3 on  H  s  to  9 —  u Q  A  c3 2  O  -  E  t oo E 3 3 O u cj o 33 o f j >,  ^ -  -  Cfl  ft « 2?  ts « c o oo  iS  (5 c  ^  O  0 E  c  o  £  E  O O  —I  3  §  O  C  O O  O  CJ  •a £ 3  —  JS  O  w  JU  a e  51 «s  o  s  3  — . a  —  O  3  a  S "o  s a « 00  l  cfl  cfl  g  c  1  i  2  • | „-  o .S o £ U tu o  5 *  1  o  2 u BSD  c a  E  «  3  ^  3  •— JS / —™  130  Plate 3.41  H a n d sample photo ot" alabandite aggregates and dark gray s i l i c i f i e d matrix  brecciated by g y p s u m . F i e l d o f v i e w is 20 c m . S a m p l e D M 3 - 3 6 0 . 0 0 .  131  inclusions and anhedral, typically 14 p m inclusions of rhodochrosite that often occur in contact with the alabandite inclusions. Electron-probe micro-analyses of alabandite and hauerite from three drill core samples were done on a fully-automated C A M E C A S X - 5 0 microprobe, operating in the wavelength-dispersion mode, with the following operating conditions: excitation voltage 20 k V , beam current 20 n A , peak count time 20 s, background current time 10 s, beam diameter 5 pm. Data reduction was done with the " P A P " (|)(pZ) method (Pouchou and Pichoir, 1985). For the elements sought, the following standards, X-ray lines and crystals were used: M n , natural rhodonite, MnKoc, L I E ; S, natural pyrite, S K a , P E T ; Fe, natural pyrite, F e K a , LLF; C a , natural diopside, C a K a , P E T ; A s , natural tennantite, A s K a , T A P ; C d , synthetic C d metal, CdKcc, P E T ; Sn, synthetic Sn metal, S n K a , P E T ; Pb, natural galena, P E T . The elements C u , N i , Sb and Zn were sought but not found. The results of 17 analyses of alabandite and 16 analyses of hauerite, presented in Appendix D , indicate that there is limited substitution of Fe, C u , Z n and Pb in alabandite. The hauerite analyses are among the purest ever reported, with trace substitution (< 0.1%) of Fe, N i , Z n a n d Sn.  3.6 Supergene weathering Local oxidation of the upper Jeronimo sulphide body has occurred. One sample inspected in this study consists of a matrix of vuggy quartz with large, typically 600 by 200 pm, irregular vugs lined with drusy quartz prisms. The vugs are filled mainly by an amorphous, porous manganese oxide with inclusions of an Fe-oxide, while others are filled by calcite or illite (Plate 3.42). The manganese oxide hosts micron-sized  132  Plate 3.42 Hand sample photo of oxidized Jeronimo ore. Vuggy. dark, gray silicified matrix with red brown manganese and iron minerals. Sample D D H 100-188.53.  133 inclusions of gold, cassiterite and Ni-bearing mineral (possibly Ni-oxide). Vugs in quartz matrix of this sample host pyrite, arsenopyrite, cassiterite, barite and monazite. The presence of gold grains within vugs in the manganese oxide mineral, which was most likely produced by the oxidization of the manganese carbonate minerals (Zeegers and Leduc, 1991), indicates that the process of oxidation has locally remobilized gold in the deposit.  3.7 Interpretation of deposit mineralogy and geochemistry Several tentative conclusions regarding fluid conditions and depositional history can be drawn from the mineralogy and mineral relations at Jeronimo. Decalcification and simultaneous silicification of the host limestone most likely occurred due to the cooling of the hydrothermal fluid. A t temperatures below 300°C and at a constant value of / € 0 , 2  calcite solubility increases and quartz solubility decreases as temperature decreases (Fournier, 1985a). The replacement of limestone by quartz may take place by the slow cooling of a fluid, at near neutral p H , as long as boiling does not occur (Fournier, 1985b). The process of boiling results in the loss of volatiles, such as C O 2 , and induces calcite precipitation as the following reaction is driven to the left (Fournier, 1985b). CaC0 ( ) + C 0 3  S  2 ( g )  + H 0 <-> C a 2  2 + ( a q )  + 2HC0 " 3  ( a q )  The precipitation of calcite in a hydrothermal system is largely controlled b y / C 0 , which 2  varies mainly by the processes of boiling, as mentioned above, and fluid mixing (Simmons and Christenson, 1994). In a hydrothermal system, sudden decreases in fluid pressure, which may induce boiling, result when ascending fluids encounter open spaces, such as vugs and fractures (Giere, 1996). The tendency of the Jeronimo carbonate  134  minerals to occur within vugs suggests that their deposition may have resulted from the process of boiling, however, Simmons and Christenson (1994) noted that, in the Broadlands-Ohaki geothermal system, calcite that precipitated in response to boiling has a distinctive platey habit, which is not present at Jeronimo. The compositional zonation of the Jeronimo carbonate minerals reflects the M n / C a ratio in the hydrothermal fluid. This ratio must have decreased as the fluid spread toward the margins of the deposit, possibly due to interaction with the host limestone. The presence of illite and kaolinite within the deposit indicates that A l and K must have been transported to the system, and were possibly acquired from the alteration of detrital feldspar in units below the ore-hosting horizon. In currently active geothermal systems in the Philippines (see Reyes, 1990), illite deposition occurs from fluids of near neutral p H over the temperature range of 230 to 320°C. This range suggests that illite was precipitated during one of the main stages of alteration in the deposit. In contrast, kaolinite in the Philippine geothermal systems is deposited from acidic fluids with temperatures ranging from ambient to 120°C. It is possible that kaolinite was deposited from cooler, more acidic, oxidizing supergene fluids and may not be related to ore formation. A c i d i c fluids are required to transport above average abundances of R E E s (Giere, 1996). The presence of monazite and REE-bearing calcite within Jeronimo ore implies that the fluids that deposited these minerals were acidic. R E E s are often transported with elements such as T i , which may account for the abundance of rutile within the deposit (Giere, 1996). The passage of acidic fluids is also invoked by the presence of the A P S  135  mineral. L i k e kaolinite, the A P S grains, which occur in vugs, are a late mineral phase and may have been deposited due to interactions with supergene fluids. In hydrothermal fluids at near neutral p H , bisulphide complexing plays an important role in gold transport (Seward, 1991). Te may also act as a ligand, forming complexes with gold such as Au(Te2)2 \ Au(TeS)2 ", A u ( T e ) H S " or AuTe2 " (Seward, 2  3  2  3  v  2  1991). The presence of coloradoite and altaite in proximity to gold grains suggests that some gold may have been transported in Te complexes. G o l d deposition may have occurred by several processes including boiling, fluid-rock interaction and fluid mixing (Seward, 1991). The occurrence of free grains of gold in association with a variety of minerals within the Jeronimo deposit contrasts with the typical mode of occurrence of gold in Carlin systems, where it normally occurs as sub-micron-sized inclusions (Au°) or structurally bound ions ( A u ) in arsenian pyrite, as in the case of the T w i n Creeks 1  deposit, Nevada, where gold is structurally bound and occurs as submicroscopic inclusions of native gold in arsenian pyrite (Simon et al., 1999). The occurrence of gold in Jeronimo is more similar to carbonate replacement deposits, in which gold occurs as free grains (Section 5.2). The geochemical characteristics of the Jeronimo deposit can be compared to those of a Carlin-type deposit using the data of Radtke (1985) for the average composition of ore grade samples from the main pit of the Carlin deposit, Nevada. The elements M n and Pb are more than 100 times greater in Jeronimo than at Carlin, while Z n is more than 25 times greater, and A g and A s are more than an order of magnitude greater. Concentrations of the elements H g , Sb and C u are similar in both deposits, while B a is more than an order of magnitude greater at Carlin than in Jeronimo. It is more difficult to  136 compare element concentrations in Jeronimo with those of gold-rich carbonate replacement deposits (CRDs), as the geochemical signatures of these deposits are highly variable.  3.8 Summary The Jeronimo ore is hosted within coarse-grained bioclastic limestones with low contents of sand and clay. Enhanced porosity and permeability resulting from the large grain size, irregularity in grain shape, lack of material clogging pores and tendency to fracture caused the ore fluids to focus within these horizons. The general paucity of mineralization in the Mesozoic normal faults and early Tertiary thrust faults and the presence of mineralization in the latest steep faults and fractures, suggests that the ore fluids were introduced towards the end of the early Tertiary compressional deformation event. Four stages of hypogene alteration are recorded in the Jeronimo ore zone. Initial decalcification of the host rocks and simultaneous pervasive, replacement-style silicification most likely resulted from the slow cooling of the hydrothermal fluids. The presence of drusy quartz-lined vugs indicates that bulk dissolution and removal of the limestone matrix occurred locally. Silicification was accompanied by the precipitation of pyrite, rutile, apatite, arsenopyrite, barite and monazite. Carbonatization occurred in two stages. The first stage involved the precipitation of manganese and calcium carbonate minerals in vugs in the silicified matrix, probably as a result of boiling or fluid mixing. Manganese carbonate comprises the dominant phase at the centre of the deposit, while calcium carbonate is the more distal dominant phase.  137 This stage of carbonatization was accompanied by the deposition of pyrite, sphalerite, rutile, monazite and barite. A late stage of carbonate deposition is recorded by veinlets and fracture-infills of calcite, orpiment, realgar and cinnabar. T w o stages of argillization are recorded in the Jeronimo ore. Illite occurs both disseminated and in patches within the quartz matrix, and was possibly deposited during silicification, but also occurs in vugs in quartz and carbonate aggregates, indicating that it was also precipitated subsequent to silicification and carbonatization. It is accompanied by pyrite and apatite deposition and occurs throughout the entire deposit. The occurrence of kaolinite is restricted to vugs in carbonate and quartz aggregates. Kaolinite-filled vugs are monomineralic and occur only in the centre of the deposit. Kaolinite may have been deposited during supergene alteration. The major sulphide minerals in the Jeronimo deposit include pyrite, arsenopyrite, sphalerite, orpiment, realgar, galena and one or more lead sulphosalt minerals. G o l d in the deposit occurs as generally micron-sized grains of native gold, most commonly at vug margins in quartz, but also encapsulated by quartz, pyrite, arsenopyrite, realgar and a manganese oxide mineral. Jeronimo is characterized by high concentrations of A g , A s , H g , M n , Pb, Sb and Z n .  138  Chapter 4: Isotopic studies of the Jeronimo deposit  T w o reconnaissance isotopic studies were performed on the Jeronimo deposit in order to identify and assess the influence of potential source reservoirs of the hydrothermal fluid(s) from which the ore zone minerals were deposited. The first study involved the measurement of the C and O isotopic compositions of ore zone manganese carbonate minerals and host rock limestone. The second study consisted of analyzing the lead isotope composition of a suite of ore zone sulphide minerals, host rock limestone, and feldspar from a local porphyry stock. Following brief introductions to the use of these isotopic systems, the methods, sampling and analysis techniques, results, and interpretations are presented for each study.  4.1 Oxygen and carbon isotopes The relative differences in mass of the isotopes of the light elements H , C , O and S are great enough to cause fractionation during physico-chemical reactions (Rollinson, 1993). This tendency to fractionate has led to the development of geologic reservoirs with distinct light element isotopic compositions. A s these light elements are the dominant constituents of hydrothermal fluids and are important components of many gangue minerals, analysis of the isotopic composition of gangue minerals may be used, under certain conditions, to identify hydrothermal fluid sources. In this reconnaissance study, the C and O isotope compositions of hydrothermally-precipitated rhodochrosite and kutnohorite from the Jeronimo ore zone, and calcite and dolomite from the E l Asientos Formation host limestone, were measured in order to identify the source or sources of the hydrothermal fluid that deposited the  manganese carbonate minerals. Carbonate mineral isotopic compositions have been used in the study of several types of sediment-hosted mineral deposits, including Carlin-type and carbonate replacement deposits. Results from the Alligator Ridge, Carlin, Cortez, Vantage, Post and Betze Carlin-type deposits, summarized in Arehart (1996), generally indicate that the oxygen isotope composition of carbonate minerals systematically becomes lighter with increasing hydrothermal alteration, reflecting both the mixing of a sedimentary carbonate isotopic signature with that of an isotopically lighter hydrothermal fluid and the decrease in fluid temperature away from the centre of the system, which promotes the precipitation of isotopically heavier hydrothermal carbonates. Isotopic studies of several carbonate replacement deposits (e.g. Santa Eulalia, M e x i c o , Megaw, 1987; E l Mochito, Honduras, Vasquez et al., 1998), show that carbon and oxygen isotopic compositions are lowered in carbonate replacement deposit host rocks, possibly due to the addition of C and O from ore fluids or other sources.  4.1.1 Sample preparation, analysis and results Six ore zone rock samples bearing manganese carbonate minerals and four E l Asientos Formation limestone host rock samples were selected for analysis and are described in tables 4.1 and 4.2. The ore zone samples were disaggregated by hand by crushing in a porcelain mortar and pestle and manganese carbonate-rich rock fragments and manganese carbonate grains were picked under a binocular microscope. B u l k samples were taken of the host rock and crushed in a porcelain mortar and pestle.  CD  CO  CD  '3 Nl E t;  5 ^ S° 2o  •s °  c  o  .-  Cu  ,  CD OO  03  _  Q  00  X  £ 03  C O X>  u  O  <S  §  a  CD  M  03  03  Cu  a. E 03  oo  OO  03  oo C  o  o  w  O- :  60  C3  2  U  E  -2 Cu E  CD i-i  $  0 3 ^2 CO "O  12 CD  C  co »  E  s o  m oo  as OS OS  in  E  co  CM  O  u  12  vq cn  00  VO  d  CO  I? O  Q  to  CO  oq rn  m  oo  ON  in  CN  od  roo  vo as  > O CQ co  CM  m oo VD  r-  , U PQ fc Q  vo  co  co  in  £  t-i  C  00 CN  03  o  o  "o  Q  U  Q  00  s 03  o in rN  ffi as Q >n  03 r-H  ON  ^  oo cn  5  i  03 03  03  00  o  •a «  >>  03  03  g>§1  3  §  £ "S  c  g1s  a. .E  &  E -a p c  o  03  ON  £ V £ o O T3  "•4—»  O X> ^  co  0  0  as rn o< od ^ O m VO Q m  x> 00 rN d  VO m  co OO  TJ  CD  J=  oo 3  CO  -a c  T3  ir  x  (U PH  <D  c  4-H  -4-J  O 'C co O (U JC  Hi  •c  00  > -4—)  00  c  CO  *-> 03  00  Q  2M  60 00 X  >-, § 00  £  o c  -1—»  3  CO  E x  o  H 3  PH  '1  ca,  l ca a  'I 'I  3 >  ' CO CO CCJ  U  N  ca 3  ca  o N t; •e ca  la  o 60  oo  od  o  oo S  00 tD X 00| 00 . X) 3 CO  I "O  ca  ca  6  <D » tD  •c O X! O  £> H I E ET X  0  00  c cd  ^ o <u VH > X o  I  co T3  ca  o  oo  i n  CN  vd  i n  N _>,  "o  to  C  ca  CN  s O  U  o  cn Ov  u  13  T3  o  O  2  i n  VO  vd  >  CN Ov O CN  CN  cn  Ov  od  CN vn CN CN  > n  od  CO  _« ' "H.  o mQ  c n  oo c n  e oo o lb N  Ov vo Ov  CN  vo i n  CN  CN  O l Ov  o  CNI  cn  Ov O CN  T-M  od  o  hrU PQ o  CN  Q u ta 7 b P H  Ov  OV  VO VO  c n  Ov  CN  00 CN  c  o  X  JD  O oo CN  H, _a> JO  ca  H  ca I CO  •<* m i  vo 00  VO i n i r-H  Ov  ca oo CN CN r i n i  Ov  X) 00 CN CN r -  ^H  OV  til ^  c « -9  ca  o (L)  3  > >. &  3  -a  O r-  oo  ta  O -C co O <D X  12 o  00 3 >  M-H  -a §1 CO  <D  CO co  o  E  CO  3  O  l-c  CO  o  c  ca - a 00 « tU 00 <u x ca >^ tU ca x o 00 00 cU >  ca x  lH  H-H  4-H  *  PH  (D  •c £ o E o ca  tU X  o  c  X co  P3 tU  <u  CO  c n i n  Ov Ov CN VO Ov Ov  CN  oo  > n  i n  r--  i n  t  c n  o o  o o  o  143  Fifty milligram samples were submitted of both carbonate types. One sample duplicate, made from different parts of the same sample and processed separately, was included for both the ore zone and host samples to assess sample heterogeneity. X R D analyses of the samples indicated that they also contained minor quartz, kaolinite and illite. The carbon and oxygen isotope compositions of the samples were determined at the Queen's University Department of Geological Sciences Stable Isotopes Lab following the method of M c C r e a (1950) in which carbonate is dissolved in 100% phosphoric acid. Host samples were dissolved in the following three steps: (1) 4 hours at 25°C to dissolve the calcite fraction, (2) 24 hours at 25°C to remove the dolomite and calcite fraction, and (3) 24 hours at 50°C to dissolve the remaining dolomite. The manganese carbonate samples were dissolved for 8 days at 50°C. The 5 C and 5 0 of the emitted C 0 gas 1 3  1 8  2  were measured simultaneously using a Finnigan Mat 252 mass spectrometer. Analytical reproducibility of  1 3  C and 5 O values is between 0.1 and 0.2 %c. Standards included in l s  analyzed batches have a reproducibility of 0.03 to 0.05 %o. Isotopic compositions are reported as per mil deviations from a standard. Using 8 notation, C and O isotopic compositions are expressed as: 5 ' C %o = C / ' C ( s a m p l e ) - C/ C(standard) C / C (standard)  x 1000  5 ' Q %c = Q/' Q(sample) - Q/ 0(standard) 0 / 0 (standard)  x 1000  3  13  2  1 3  8  18  13  6  1 8  12  1 2  18  16  l 6  The measurements were reported relative to a belemnite sample from the Cretaceous Peedee Formation, South Carolina (PDB) for C and to Vienna standard mean ocean water ( V S M O W ) and P D B for O. Values of 5 0 are discussed relative to V S M O W . 1 8  144  The carbonate mineral isotopic compositions are presented in tables 4.1 and 4.2 and are displayed graphically in Figure 4.1. The ore zone minerals range in 8 0 values 1 8  from 17.97 to 22.52% and in S C values from -2.84 to -1.3%o, while the E l Asientos I 3  0  Formation host rocks range in 5 0 values from 13.13 to 23.87%o and in 5 1 8  1 3  C values from  -1.83 to-1.42% . 0  The ore zone group sample duplicate pair shows very little variation in 5 O , with l s  a difference of 0.44% , while 5 o  1 3  C varies more widely by 0.69%o. The pair of laboratory  duplicate samples within this group differs by only 0.13%o in 5 0 and 0.03%c in 5 C . 1 8  The host rock sample duplicate pair varies in 8 0 by 0.06 %c and 8 1 8  1 3  1 3  C by 0.18 %o for  calcite and in 5 0 by 2.06% and 5 C by 0.61% for dolomite. 1 8  I 3  0  0  4.1.2 Interpretation Interpretation of mineral stable isotope data is based on the assumption that the minerals analyzed have not undergone retrograde isotopic exchange and that their present values are those preserved from the time of deposition. O f the host rock samples, only two (DDH20181-459.15 and DM3-560.00) possess isotopic compositions within the field of Jurassic carbonates ( 5 0 = 23 to 29 %o and 5 1 8  1 3  C = -2 to 3 %c) defined by Veizer and  Ffoefs (1976). Figure 4.1 shows that the remaining host rock samples ( 8 0 = 12 to 19 %o 1 8  and 8  1 3  C = -2 to 1 %o) have undergone exchange with a fluid isotopically lighter in O.  The alteration of these host rock samples is discussed in Section 5.3.  145  30  Composition o f Jurassic carbonates (Veizer and Hoefs, 1976)  \  Host rock samples with typical marine limestone isotopic signature  \ v"  *  •  J  Host rock samples that have exchanged with isotopically light fluid  • Jeronimo ore zone carbonate o E l Asientos host limestone i  -4  1  -  1  1  3  -  2  -  1  1  0  1  1  1  2  per m i l 6 C ( P D B ) 13  Figure 4.1 M i n e r a l isotopic compositions for Jeronimo ore zone hydrothermal carbonate minerals and sedimentary carbonate minerals in the E l Asientos Formation host rocks. O n l y two host rock samples plot within the field o f Jurassic marine limestone defined by Veizer and Hoefs (1976), shown in gray. Other host rock samples have exchanged with isotopically lighter fluids from an unidentified source.  146  A s the ore zone carbonate minerals were precipitated directly from a hydrothermal fluid, it is useful to calculate the isotopic composition of a fluid in equilibrium with the minerals in order to identify possible fluid sources (see Ohmoto, 1986; Taylor, 1987). T o calculate the isotopic composition of a fluid in equilibrium with a mineral, it is assumed that (1) the fractionation factors between the mineral and fluid are accurate, and that (2) the mineral and hydrothermal fluid were in isotopic equilibrium during deposition. Depending on the isotopes considered, several characteristics of the fluid must be known. F l u i d temperature affects the fractionation of both oxygen and carbon isotopes (Ohmoto, 1986). A decrease in fluid temperature results in the precipitation of minerals with higher 8 0 and 5 C values. The 5 1 8  1 3  1 3  C of a fluid is also  influenced by the types and proportions of the carbon species present. Depending on oxygen fugacity, carbon may be in oxidized, neutral or reduced form, for example, C  4 +  in  carbonates, C ° in graphite and diamond, and C " in organic compounds. There are large 4  kinetic isotope effects associated with changes in the redox state of an element, such that the heavy isotopes are preferentially concentrated in elements with higher oxidation states. In the case of C , 5 C onates>8 Cg phite>8 Corganic matter- A s carbon is present 13  13  carb  13  ra  within carbonate minerals in the Jeronimo ore zone, they are most likely in equilibrium with an oxidized C - b e a r i n g fluid. In an oxidized fluid, p H controls whether carbon is 4+  present as C 0 , H C 0 " , H C 0 or C 0 " .  A t temperatures less than 3 5 0 ° C , the dominant  2  2  3  2  3  3  carbon species are F f C 0 , H C 0 " and C 0 " (Ohmoto and Rye, 1979). 2  2  3  3  3  It was not possible to measure fluid temperature directly from fluid inclusion homogenization temperatures as the fluid inclusions in Jeronimo are small and vapourrich. Futhermore, carbonate geothermometry (see Anovitz and Essene, 1987) could not  147 be used to gain an estimate of temperature as calcite and dolomite from a single sample were in isotopic disequilibrium. Therefore, fluid isotopic compositions were calculated over a range of temperatures from 200 to 350°C. This range was chosen because (1) although not directly related to the carbonate minerals, illite intergrown with the vuggy quartz constrains fluid temperatures to between 230 and 320°C (Reyes, 1990); and (2) reported values of fluid inclusion homogenization temperatures from Carlin-type sedimentary rock-hosted disseminated gold deposits generally range from 175 to 300°C (Arehart, 1996) and from 200 to 500°C for carbonate replacement deposits (Haynes and Kesler, 1988; Titley, 1991; Lang and Baker, 1999; Megaw, 1998). F l u i d p H was estimated in order to determine which carbon species were present. According to Krauskopf and B i r d (1995), except for very localized, sporadic variations, the p H of common hydrothermal solutions varies by no more than two p H units away from neutrality. Note however, that neutral p H increases with temperature as the dissociation constant of water increases, for instance, reaching 5.5 at 275°C (Krauskopf and Bird, 1995). Taking this temperature as an average value for Jeronimo mineralization, the p H of the fluids that deposited the carbonates most likely ranged from 3.5 to 7.5. Figure 4.2 from Krauskopf and B i r d (1995) shows the dominance of the carbon species H C 0 , H C G y and C 0 " as a function of p H and temperature. The 2  2  3  3  diagram indicates that, over a range of temperatures from 0 to 300°C, the dominant carbon species is H2CO3 for the p H range stated above. This result is corroborated by Ohmoto (1986), who stated that, in most geological fluids above 100°C, the amount of HCO3" is negligible relative to the amount of H C 0 . Consequently, in this study, the 2  3  Figure 4.2 Dominant C-species as a function of temperature and pH at 1 bar. Redrawn from Krauskopf and Bird (1995).  149  fluid 8 C was calculated for carbonate minerals in equilibrium with the H 2 C O 3 species in 1 3  the fluid. Although some oxygen would be derived from H 2 C O 3 , the 5 0 of the fluid 1 8  would be dominated by H 0 . Therefore, the 8 0 of fluid was only calculated for 1 8  2  equilibria between the carbonate minerals and H 2 O . A s limited data are known for oxygen and carbon isotope fractionations between the fluids and manganese carbonate minerals, fractionation factors for dolomite, listed in Table 4.3, are used instead.  Table 4 . 3 Fractionation factors o f mineral-fluid pairs, taken from the database I S O F R A C , written by J. Martin.  Pair of substances (element)  Reference Ohmoto and Rye, 1979 Ohmoto and Rye, 1979 O ' N e i l et a l , 1969 Matthews and Katz, 1977  C a l c i t e - H C 0 (C) D o l o m i t e - H C 0 (C) C a l c i t e - H 0 (0) D o l o m i t e - H 0 (0) 2  3  2  3  2  2  Figure 4.3 shows the variation in the isotopic composition of H 0 and H C 0 2  2  3  species in fluids in equilibrium with the ore minerals and Jurassic limestone samples. Table 4.4 lists the results of the calculations of fluid isotopic compositions. For the fluid that deposited the ore zone manganese carbonate minerals, at the intermediate temperature of 275°C, the calculated fluid 5 0 2 o ranges from 10.5 to 16.3% , while the 1 8  H  S CH2C03 1 3  0  ranges from -8.0 to -6.5%o. A fluid in equilibrium with the Jurassic limestone  samples, at 275°C, has a 8 0 2 o ranging from 16.8 to 17.5%o and a S C 2C03 ranging 18  13  H  H  from -4.0 to -3.0% . o  In the case of the Jeronimo deposit, potential sources of oxygen in the fluid that deposited the ore zone carbonate minerals are magmatic and meteoric. The carbon-  Figure 4.3 Isotopic compositions of O in H 0 and C in f^CO, in equilibrium with the ore zone and host rock carbonate, calculated over a range of temperatures. Fractionation factors between rhodochrosite-fluid, kutnohorite-fluid and dolomitefluid pairs are calculated using dolomite-fluid equilibria of Matthews and Katz (1977) and Ohmoto and Rye (1979), and between calcite-fluid pairs using Ohmoto and Rye (1979) and O'Neil et al. (1969). 2  151 Table 4.4 Calculated isotopic composition of fluids in equilibrium with El Asientos Formation marine limestone and Jeronimo ore zone hydrothermal carbonate minerals. H 0 (O) H C O ( C ) 2  El Asientos host 20181-459.15 Limestone DM3-560.00 Ore zone 91-562.80 Mn-carbonates 100-175.23 86-547.61 91-572.28 91-572.28d 96.299.45 98-619.94 100-355.81  2  200°C  275°C  350°C  200°C  14.3 13.7 11.1 12.7 6.8 8.6 8.7 9.5 8.1 8.6  17.5 16.8 14.7 16.3 10.5 12.2 12.3 13.2 11.8 12.3  19.6 18.9 17.1  -1.1 -2.1 -5.4 -5.4 -5.2 -5.5 -5.5 -4.8 -6.2 -6.4  18.7 12.9 14.6 14.8 15.6 14.2 14.7  3  275 ° C -3.0 -4.0 -7.1 -7.1 -6.9 -7.2 -7.1 -6.5 -7.8 -8.0  350°C  -5.1 -6.0 -9.0 -9.0 -8.8 -9.1 -9.0 -8.4 -9.8 -9.9  bearing species in the fluid may have been acquired from magmatic fluids, dissolution of E l Asientos Formation host limestone, and decomposition of organic matter contained in the host rocks. Values of 5 0 18  and 8 C for material from these reservoirs are listed in 13  Table 4.5. At  275°C, 5  13  C 2 c o 3 of the ore fluids ranges from H  -8 to -6.5%o. These  values are  lower than the calculated composition of fluid in equilibrium with the host limestone (-3 to  -5%o)  and higher than organic carbon sources  (-25  to  -10%o). The ore  fluid 5 C 2C03 13  H  partially overlaps with the accepted range for magmatic C0 of -7 to -3%o (Ohmoto and 2  Rye, 1979)  (Table 4.5).  The available C isotope data do not point definitively to a fluid  source. The observed 8 C 2co3 of ore carbonate may reflect the mixed input of 13  H  magmatic volatiles and/or organic material with dissolved host rock limestone.  275°C range from 10.5  The values of 5 0 2 0 for the ore zone fluid at 18  H  These values are significantly heavier than those of meteoric fluid  (-10%o),  partially overlap with both the ranges for magmatic H 0 (7 to 13%o) 2  the dissolution of the host carbonate  (16  to  18%o). Although some  to  16.3%o.  however they  and O derived from component of  152  meteoric H 2 O may have been present, the 8 0 data for the hydrothermal fluid can be 1 8  explained solely by the mixing of magmatic H 0 with O derived from the dissolution of 2  the host limestone.  Table 4.5 Isotopic compositions of potential reservoirs for the Jeronimo fluid. Reference 5 0 (SMOW) 5 C (PDB) Magmatic H 0 Ohmoto (1986) 7-13 %o — Meteoric H 0 Ohmoto (1986) — 10 %0 -16-18 %o This study Carbonate species —3 to -5 %o derived from dissolution or decarbonatization of carbonate rock Ohmoto (1986) C 0 from -25 to-10%o decomposition of organic matter Ohmoto and Rye (1979) Magmatic C 0 -7 to -3 % 1 8  1 3  2  2  2  2  —  0  4.2 Lead isotopes The lead isotope composition of ore-related minerals has been widely used to constrain the source, or sources, of hydrothermal ore-forming fluids as well as the influence of the material through which the fluids were transported. The application of lead isotope analysis to the study of ore deposits is based on the assumption that the fluid that deposited the lead-bearing minerals also deposited other ore minerals, gold in this case. A n overview of lead isotopes and their applications to the study of mineral deposits is given in Tosdal et al. (in press). Lead isotopes have been used to identify the source reservoirs of a variety of mineral deposits including Mississippi valley-type (e.g. Kesler et al., 1994; H e y l et al., 1974), sedimentary-exhalative (Beaudoin, 1997), volcanogenic massive sulphide (Childe and Thompson, 1997), porphyry-related (e.g. Mukasa et al.,  153 1990) and epithermal precious metal deposits (e.g. Richards et al., 1991). Relatively few studies have been undertaken on the lead isotopes of sediment-hosted, disseminated gold deposits. However, in preliminary studies of the lead isotope composition of sulphide minerals from the Getchell, Betze-Post and Jerritt Canyon District Carlin-type deposits of Nevada, summarized in Tosdal et al. (1998), trends in the lead isotopes were used to assess potential source reservoirs of the ore fluids. Lead isotope studies of carbonate replacement deposits indicate that proximal ore fluids possessed a significant magmatic component (Megaw, 1998).  4.2.1 Lead isotope systematics The three radiogenic isotopes of lead are ZUD  2 0 6  Pb,  2 0 7  P b and  2 0 8  P b . The isotopes  P b and ' P b are the daughter products of the radioactive parent isotopes z u  respectively, while isotope,  2 0 4  2 0 8  P b is the product of the radioactive decay of  2 3 2  TJ and  U,  T h . A fourth lead  P b , is stable.  A t the time of the formation of the earth 4.6 b.y. ago, distribution of the elements TJ, T h and Pb was homogeneous (Zartman and Doe, 1981). Geologic processes such as partial melting, fractional crystallization, metamorphism and alteration have fractionated these elements differently forming reservoirs with distinct  U/  Pb and  Th/  U  ratios. The incompatible elements uranium and thorium normally exist in a tetravalent oxidation state and are close in ionic radius (0.97 and 1.02 A, respectively), causing them to behave similarly during most geologic processes. However, their behaviour differs under oxidizing conditions, as uranium may exist in a 6+ oxidation state and is highly soluble in aqueous fluids, while thorium remains tetravalent and is insoluble. Under most  154  conditions, lead is stable in a divalent oxidation state and has a significantly larger ionic radius (1.20 A) than its parent isotopes. Lead is able to form minerals such as galena and can substitute for many low valence metal cations in sulphide, silicate and carbonate minerals. These minerals accept very limited quantities of uranium and thorium and their initial lead isotopic composition remains relatively constant over time. Lead isotopes are normally displayed on two covariant diagrams. A s absolute abundances of the isotopes are difficult to determine, they are measured as ratios relative to  2 0 4  P b , since the value of this isotope is constant and does not change with time. The  uranogenic covariant diagram plots the daughter product of abundant lead daughter isotope,  2 0 6  Pb (  207  Pb/  204  P b vs.  diagram is similar but plots the daughter product of (  208  p b /  204  p b  y  s  206  p b /  204  2 0 6  Pb/  2 3 5  2 0 4  U decay against the most  P b ) . The thorogenic  T h decay against  Pb  p b )  A feature that is normally present on covariant lead isotope diagrams is a model growth curve. This curve represents the rate of change over time in the production of the daughter lead isotopes in the continental crust. The shape of growth curves depends on the initial ratios of U / P b and T h / U in a given reservoir. Such a curve is included on the covariant diagrams to indicate whether the source reservoir(s) was depleted or enriched in uranium or thorium relative to the curve. Growth curves are constructed from ore deposits with good age constraints and with hydrothermal systems that are equivalent to the average compositions of continental crust. The curve of Stacey and Kramers (1975), used in this study, is based on two stages of growth with distinct U / P b and T h / U ratios.  155  4.2.2 Sampling and analysis The samples analyzed for their common lead isotopic composition consist of a suite of ore-zone sulphide and sulphosalt minerals, in addition to five limestone samples from the E l Asientos Formation and one feldspar sample from the Bochinche porphyry. The limestone and feldspar were analyzed in order to assess the lead isotopic composition of the E l Asientos Formation host rocks and of the Eocene to Oligocene porphyries in the area, as these are potential source reservoirs for lead in the ore fluids. The lead isotope analyses were performed at the Geochronology Laboratory at the University of British Columbia. Samples bearing sphalerite, orpiment, realgar, lead sulphosalt, and feldspar, and limestone (described in tables 4.6 and 4.7) were crushed by hand and grains were handpicked to form 10 to 50 milligram samples. The samples were leached in dilute hydrochloric acid to remove surface contamination and subsequently dissolved in nitric acid. Galena-bearing samples were bulk leached in 2 N H C l . Although this leaching process dissolved all of the minerals in the galena-bearing samples, with the exception of quartz and arsenopyrite, the lead content of galena is high enough to overwhelm the addition of trace lead from other minerals. A l l samples were passed through ion exchange columns in hydrobromic acid. Lead was collected in hydrochloric acid and precipitated as lead chloride. Approximately 50 to 100 nanograms of the lead chloride as loaded on a rhenium filament using a silica gel-phosphoric acid mixture. Isotopic  156  Sample 86-546.58  Mineral Orpiment  Characteristics  93-527.33  Orpiment  A s irregular yellow aggregates, typically 400 pm, in veinlets with pyrite, calcite and minor cinnabar filling open space in quartz matrix.  93533.07a,b  Galena  Anhedral, disseminated galena, typically 5 um. Also bladed, subhedral, typically 175 by 25 u m arsenopyrite grains disseminated in vuggy quartz matrix, intergrown with apatite, pyrite and rutile. Sample repeat (first analysis subject to machine error).  97273.06a,b  Sphalerite  Amber-coloured, anhedral, typically 125 um grains, as inclusions in massive realgar and disseminated within the vuggy quartz matrix. Sample duplicates.  97273.06c  Lead sulphosalt  97273.06d  Realgar  Steel-gray, rod-shaped grains, typically 125 by 25 um, as inclusions in massive realgar aggregates. X R D analysis suggests twinnite or guettardite. Bright orange-red subhedral grains in aggregates hosting pyrite, sphalerite, apatite, the lead sulphosalt, coloradoite and gold, typically 1 cm.  97-277.17  Galena  Anhedral to subhedral, cubic grains, typically 45 um, filling vugs in quartz matrix, intergrown with pyrite and an A P S mineral.  97-282.89  Galena  Cubic to tabular, subhedral grains, typically 170 by 75 um, in poikilitic barite vein with pyrite and apatite, cross-cutting vuggy quartz matrix.  97-287.25  Sphalerite  Equant, subhedral, typically 20 um grains in dense band with pyrite and apatite, cross-cutting quartz. A l s o in vugs in quartz matrix.  A s irregular yellow aggregates typically lOOum, filling vugs in Mn-carbonate and quartz matrix.  Table 4.7 Carbonate and feldspar minerals analyzed for their common lead composition. Sample sites shown in Figure 3.11.  Sample DM1438.91 DM3360.40 DM3560.00a,b DDH91562.80 DDH20181 -459.15a,b Bochinche porphyry  Mineral Calcite Calcite/ Dolomite Calcite Dolomite Calcite Plagioclase  Sample characteristics Matrix of carbonate grains, with some dolomite, and chalcedonic quartz Calcite grains in sparry dolomite matrix Bioclastic limestone with calcite spar cement. Lab duplicates. Vein of dolomite cross-cutting fossiliferous, calcitic matrix Bioclastic limestone with sparry matrix, composed of <90% calcite. Lab duplicates. Approximately 5mm, subhedral to euhedral, tabular plagioclase phenocrysts, weakly argillized.  157  compositions were determined using a modified V G 5 4 R thermal ionization mass spectrometer. Potential sources of uncertainty during analysis include (1) thermal mass fractionation of the isotopes, as the light isotopes tend to volatilize first, and (2) uncertainties measuring the abundance of  2 0 4  P b , due to the low abundance of this isotope.  The measured ratios were corrected for instrumental mass fractionation of 0.12% per mass unit based on repeated measurements of the N . B . S . S R M 981 Standard Isotopic Reference Material. Reported uncertainties were obtained by propagating all mass fractionation and analytical errors through the calculation and are given at the 2o (95%) confidence level. The total measured procedural blank on the trace lead chemistry was 100-120 picograms. Sample precision was tested by analyzing a sample duplicate (97273.06) and a lab duplicate (DDH20181-459.15).  4.2.3 Results A significant amount of unpublished lead isotope data from the Potrerillos District were made available to the author by R . M . Tosdal (written communication, 1999). This data set, presented in Appendix E , includes analyses of (1) Carboniferous and Permian igneous rocks, (2) Triassic and older sedimentary rocks, (3) Jurassic andesitic rocks, (4) Jurassic limestone and sandstone, and (5) Eocene to Oligocene porphyry stocks. A s these rock types are all potential source reservoirs for lead in the Jeronimo minerals, their lead isotope compositions are shown as fields with which the Jeronimo data could be compared on the covariant diagrams. Lead isotope data were also available for sulphide minerals from the Potrerillos porphyry C u - M o - ( A u ) deposit ( R . M . -Tosdal, unpublished  158  data, reported in Marsh, 1997) and from the E l Hueso disseminated gold deposit ( R . M . Tosdal, unpublished data). These are included on the diagrams as fields for comparison. Of the seven samples of E l Asientos limestone analyzed in this study, only two overlapped the field of Jurassic limestone and sandstone defined by Tosdal (unpublished data) on the uranogenic diagram, while no samples plotted within this field on the thorogenic diagram (insets in figures 4.4 and 4.5). The Jurassic sedimentary rock field is defined by the analysis of eight Jurassic limestone and sandstone samples from the E l Asientos and Aguas Heladas formations in the Potrerillos district. The samples of the E l Asientos Formation analyzed in the present study are drill core samples of unmineralized limestone from units B , C and E , ranging in location from the western margin of Jeronimo up to 2.4 kilometres west. Too few samples were analyzed to assess systematic changes in the lead isotope composition between units. The results of the E l Asientos formation limestone lead isotope analyses performed in this study suggest that either the field of Jurassic sedimentary rocks for the Potrerillos district is more variable than is indicated by the data of Tosdal, or that the lead isotope compositions of the limestone samples analyzed in this study have been modified by alteration. However in the latter case, it is apparent that, given the relatively radiogenic compositions of the limestone, they were not affected by fluids associated with the comparatively non-radiogenic Potrerillos porphyry, E l Hueso or Jeronimo deposits. Likewise, the one sample of Oligocene porphyry analyzed in this study plots slightly outside of the field defined by the data of Tosdal. When considering the position of the data points for carbonate and feldspar, which host only trace lead, it is important to remember that significant uncertainty ellipses, shown in figures 4.4 and 4.5, are associated with these analyses.  159 —  CJ  CJ 3  O -3  -_)  LC  o  o •  cc  JS CC  N  CC  CL>  c  cj  CC  1 OJD STCL c3  +  ,CJ  *tH  b  ca cc CJ  CL*  p. L-  CJ  C UJ  O O o X  _: CJ  c o  c  cc N  o  'ga is CC  u o 3  a  •  o  v  —'  TC  cc  _CL  3 CJ  09 l_ eo CC J 3 > S E3 u o  a>  Av  cn  u a o  CJ  V  c in r-  rve  *-• 09  CC  CL  ,  •—-  CJ  JC  •fi 2 CU  OU  \ \\ \!  3 CC  >-, CJ CJ  S  JO  CN OS  o B  11 CJ c/c SJ)- =  o CL.  ca a 2 ,5 <0 tc  CL O  9 £ £ c  O  —  CC CJ _ c  cs  Sr w 3 to  c c_ o  0)  <u 5 0 t- o | 'cc  o  ti  5 a. S 81  PH  PH o es  fe  ca •a BA  —i ,—  V  H ON •P 2  T3 CJ CL — I — O CC J = J3 O to  W  5 S o _  1I •5  C L2  LLC  c 5  rt  CL-O  C re  c  L  3  CJ  l e i  a o & a <H 3  CJ  53  CC — I  D  <o  a. a.  co  i  CL  ^  —  O  P  to O  £  b  |M  3  60 .  ro a.  ^5  160 T3 cu  to  3  cn  2 *- -S  •s s = "  c  c  on  N  O  CJ  u o  1«  cd  53  4}  co  3  cn  U.  irt  M P uu.3  i> g S-  "3 00  "eL  in  C  O  ffl >. w § C Je  cn  cd  « Q. 3 S  S & g E q o 3 o  Cd  u i)  -S  a, b C  •£ S .2 C c n cd  a ~ 3 M  O •= a  Er  > 2  S3  —  S .I-H  CU 'cn  OH  is Ui  l-i  cu  c  >% o Cu  0£  cu  o cfl  Cu  uu  161 T w o trends are apparent in the Jeronimo sulphide and sulphosalt mineral data, presented in tables 4.8 and 4.9, when they are plotted on the uranogenic and thorogenic diagrams (figures 4.4 and 4.5).  On both diagrams, analyses of the minerals realgar,  sphalerite, lead sulphosalt and galena align along a steep trend extending from the Eocene to Oligocene porphyry field toward the pre-Jurassic igneous and sedimentary rock fields, referred to jointly as the pre-Jurassic basement field. This trend is approximately parallel to the long axes of the uncertainty ellipses for the analyses of galena and of minerals with trace lead and may not define as much of a trend as suggested by the data points. Regardless, the average of this array lies between the Eocene to Oligocene porphyry and Pre-Jurassic basement fields. In contrast to the majority of the sulphide and sulphosalt analyses, late orpiment defines a separate trend. On the uranogenic diagram, the orpiment analyses plot between the Eocene to Oligocene porphyry field and the Jurassic sedimentary rock field. O n the thorogenic diagram, the orpiment trend line is enriched in  2 0 8  P b relative to the Eocene to  Oligocene porphyry field and can be traced back to the cluster of Jeronimo sulphide and sulphosalt analyses. A t present, the orpiment trend lines are defined by only two points and more analyses are required to substantiate these trends.  4.2.4 Interpretation Paragenetically, the sulphide and sulphosalt minerals that plot between the Eocene to Oligocene porphyry and the pre-Jurassic igneous and sedimentary rock fields appear to post-date the main phase of silicification, occuring either as veins or bands  162 Table 4.8 Common lead isotope results for Jeronimo ore zone sulphide and sulphosalt minerals. Sample source  Mineral  Quality of analysis  % error 2 0 6  p  b  /  2 0 4  p  b  (2a)  % error  % error 207p  b /  204p  b  (2a)  208  p b /  204  p  b  (2a)  DDH86546.58  Orpiment  Good  18.695  0.08  15.607  0.08  38.622  0.08  DDH93527.33  Orpiment  Good  19.119  0.06  15.656  0.05  38.687  0.06  DDH93533.07a  Galena  Good  17.401  0.10  15.433  0.09  36.918  0.11  DDH93533.07b  Galena  Good  18.644  0.02  15.653  0.02  38.632  0.02  DDH97273.06a  Sphalerite  Good  18.582  0.01  15.623  0.01  38.566  0.01  DDH97273.06b  Sphalerite  Good  18.564  0.01  15.592  0.01  38.589  0.01  DDH97273.06c  Lead sulphosalt  Good  18.579  0.01  15.619  0.01  38.554  0.01  DDH97273.06d  Realgar  Good  18.573  0.01  15.617  0.01  38.540  0.01  DDH97277.17  Galena  Good  18.614  0.06  15.662  0.06  38.638  0.06  DDH97282.89  Galena  Good  18.589  0.01  15.635  0.00  38.606  0.01  DDH97287.25  Sphalerite  Good  18.567  0.01  15.611  0.00  38.536  0.01  Table 4.9 Common lead isotope results for El Asientos Formation host rock carbonates and Bochinche porphyry feldspar. 207 204 % % *W Pb Quality Pb/ Pb % Mineral Sample error error of error source (2a) (2a) analysis (2a) 0.04 0.04 38.833 0.04 15.691 good 19.245 Calcite DM1438.91 0.06 0.05 38.237 15.541 fair 18.941 0.06 Calcite/ DM3Dolomite 360.40 0.43 0.42 38.791 0.43 15.632 poor 18.887 Calcite DM3560.00a 0.07 0.07 39.071 15.735 good 19.005 0.07 Calcite DM3560.00b 40.573 0.15 0.15 0.15 15.775 fair 22.306 DDH20181 Calcite -459.15 0.06 40.395 15.729 0.05 fair 22.163 0.06 DDH20181 Calcite -459.15 0.17 0.14 38.975 15.754 fair 19.173 0.16 Dolomite DDH91562.80 38.532 0.01 15.617 0.01 18.547 0.01 good Bochinche Plagioclase porphyry m  204  p b /  p b  U  4  163  cross cutting the quartz matrix or as minerals included in vugs within it. It is difficult to paragentically relate most of these sulphide and sulphosalt minerals directly to gold, with the exception of realgar in sample DDH97-273.06d, which encapsulates gold grains. The occurrence of realgar in this sample, as aggregates with pyrite, sphalerite, lead sulphosalt, coloradoite and gold, differs strongly with the dominant mode of occurrence of realgar in the deposit, in late veins with orpiment, calcite and cinnabar.  This textural difference,  combined with the difference in the lead isotope composition between the realgar and late orpiment samples, suggests that realgar may have been deposited during multiple stages of mineralization. Based on the trends in the geochemical data discussed in Section 3.4, gold correlates positively with two elements hosted in minerals analyzed for their lead isotope composition, Pb and Sb. This correlation may substantiate a genetic link between gold, galena and the lead sulphosalt(s), which would validate the application of the conclusions drawn from the lead isotope data of the sulphide minerals to gold paragenesis. The lead isotope compositions of these minerals appears to reflect the mixing of lead from the Eocene to Oligocene porphyries with lead from a more radiogenic source. T w o end-member cases may be considered for the latter source. The source may be the Jurassic E l Asientos Formation host rocks. In this case, the magmatic-hydrothermal ore fluids, during their passage through the E l Asientos formation from the porphyry to the site of ore deposition, would have acquired a component of the Jurassic host rock lead signature. The more radiogenic source may also be the pre-Jurassic basement igneous and sedimentary rocks. In this case, the source porphyry may have been emplaced into basement rocks and exsolved the ore-bearing fluids at this level, leading them to  164 exchange lead with the basement rock. The fluids may or may not have subsequently exchanged lead with the overlying Jurassic sedimentary rocks. A s the fields of the Jurassic sedimentary rocks and pre-Jurassic basement rocks overlap, it is also possible that both acted as sources of some lead in the Jeronimo ore fluids. The occurrence of orpiment contrasts with the other sulphide and sulphosalt minerals as it often forms in late veins with realgar, calcite and minor cinnabar.  The  orpiment signature probably reflects the input of a component of lead from the E l Asientos Formation host rocks. Lead from this source may have been added during mixing of meteoric water in equilibrium with the E l Asientos formation host rocks with the Jeronimo ore fluid.  4.3 Conclusions The 8 0 values of ore zone Mn-carbonate minerals suggest that the fluids from 1 8  which they were deposited possessed a mixed signature reflecting the input of magmatic volatiles and material derived from the dissolution of the host E l Asientos formation limestone. The 8 C values of these carbonates do not point to definitive sources, but do 1 3  not disagree with the oxygen isotope data. The 8 0 and 8 C values of minerals in the 1 8  1 3  pre-Jurassic basement were not tested in the study, however, they may also have influenced the ore zone carbonate isotopic composition. The common lead isotope data of main stage ore zone sulphide and sulphosalt minerals also reflect the mixing of lead derived from a magmatic source with lead derived from the Jurassic E l Asientos Formation limestone and/or pre-Jurassic basement  165 rocks. The common lead isotope composition of late orpiment indicates that it has derived a significant amount of lead from the host limestone. Although not definitive, the results obtained in these reconnaissance isotopic studies of the Jeronimo deposit strongly suggest that the hydrothermal fluids associated with the deposition of the manganese carbonate minerals and the sulphide-sulphosaltgold event were dominantly of magmatic origin. Although the genesis of some Carlintype deposits may involve some magmatic components, the strong magmatic affiliation of Jeronimo, suggested by the isotopic data, is more consistent with formation in a manner similar to carbonate replacement deposits.  166 Chapter 5: Interpretation  5.1 Summary of Jeronimo Jeronimo is a stratabound, disseminated gold deposit, hosted by Jurassic marine limestones. It is located in northern Chile, in the Potrerillos district, and lies between the Maricunga gold belt and Chilean porphyry copper belt. Thrust faults and folding in the Jeronimo area record early Tertiary compressive deformation, resulting from subductionrelated transpression related to deformation within the regional Domeyko fault system, along which Eocene to Oligocene porphyry copper deposits in northern Chile were emplaced.  5.1.1 Stratigraphic and structural controls The distribution of alteration and mineralization within Jeronimo was strongly controlled by lithology. The ore is dominantly hosted by coarse-grained bioclastic horizons with minor siliciclastic components. In contrast, fine-grained silty and sandy limestones, located above and below the rocks hosting the ore zone, and also interfingered with it, host minor to no gold. The bioclastic limestone horizons are favourable ore hosts because of their relatively high porosity and permeability, resulting from their (1) large grain size, (2) irregular grain shape, (3) lack of clay particles clogging pore throats, (4) high intraparticle porosity, and (5) high fracture-related secondary porosity. Several types of faults influenced the shape of the Jeronimo ore body.  Steep  normal faults segmented and offset portions of the ore-hosting horizon in a series of steps  167 with a cumulative offset of 400 metres. Although mineralization is only locally observed within the planes of these faults, they may have channeled some fluids to the bioclastic horizons. Late Paleocene to Early Oligocene subduction-related transpression resulted in basin closure and folding and thrust faulting of Mesozoic carbonate and Tertiary volcanic strata. According to Caddey (1999), at the end of this compressive event, fracturing perpendicular to the principal stress direction occurred preferentially in the bioclastic horizons, due to their high competence relative to the other limestone lithologies. These fractures, which presently host quartz, carbonate and sulphide mineralization, greatly increased the secondary porosity of the bioclastic horizons. Mineralized, steep, normal faults, observed to cross-cut the C-horizon and overlying D and E units may also have acted as feeders, channeling the ore fluids to the bioclastic horizons (Caddey, 1999). The relationship between thrust faulting and mineralization in the Potrerillos district is variable. The emplacement and mineralization of the Potrerillos porphyry were synchronous with movement along the Potrerillos M i n e fault (Tomlinson, 1994; Mpodozis et al., 1994). In contrast, movement along this fault cross-cuts and post-dates mineralization at E l Hueso (Marsh, 1997). It is possible that fluids were channeled towards the Jeronimo ore zone along the P M F (also known as the Silica Roja faults in the Jeronimo area), however the relationship between this thrust and Jeronimo mineralization remains uncertain.  168  5.1.2 Alteration mineralogy and distribution The earliest alteration events recorded by the Jeronimo ore are decalcification and silicification. Silicification occurs mainly as the pervasive replacement of limestone by fine-grained, locally chalcedonic quartz. The preservation of delicate bioclast structures by quartz indicates that decalcification and silicification were simultaneous. The presence of drusy quartz-lined vugs throughout the matrix and at the centres of some shell-shaped bioclasts shows that bulk dissolution and removal of the host limestone occurred locally. Although a wide range of minerals are present within vugs in quartz, a suite of minerals including pyrite, arsenopyrite, sphalerite, rutile, apatite, barite and micron-sized monazite occur as disseminated grains, partially encapsulated by the quartz matrix. Carbonatization within the Jeronimo deposit is marked by the precipitation of manganese and calcium carbonate minerals. The manganese carbonate minerals rhodochrosite and kutnohorite comprise the dominant carbonate mineral assemblage in an elliptical zone with an area of 1.5 by 0.8 kilometres, which trends northwest and overlaps the centre of the deposit. Rhodochrosite and kutnohorite are intergrown and occur as porous aggregates within vugs in the quartz matrix and within the drusy centres of silicified fossils. Manganese carbonate minerals encapsulate pyrite, sphalerite, rutile, rhodonite, and micron-sized barite and monazite, while vugs within the manganese carbonate aggregates are filled mainly by orpiment, realgar and kaolinite. R i m m i n g the manganese carbonate-dominated core of the deposit is a zone marked by calcium carbonate deposition. This stage of calcium carbonate mineralization involved the deposition of calcite and dolomite (locally manganoan) in the cores of silicified bioclasts  169 and in vugs in the silicified matrix. Late fractures and veinlets are filled by calcite, orpiment, realgar, and minor cinnabar. The Jeronimo clay mineral assemblage consists of illite and kaolinite. Illite is present in minor abundance (<10%) throughout the ore zone. It occurs mainly in patches within the vuggy quartz matrix, locally intergrown with apatite and pyrite, but is also present as finely disseminated particles and vug-filling aggregates that locally host gold grains. Illite crystallinity does not vary significantly throughout the deposit. Kaolinite distribution is restricted to a northwest trending zone, 1.0 by 0.5 kilometres in size, overlapping the centre of the deposit, where it is generally present in quantities up to 10 %. In contrast to illite, kaolinite occurs only as a vug-filling mineral in quartz and carbonate aggregates, and is not associated with other minerals. Kaolinite may be supergene in origin.  5.1.3 Mineralization Major (>1%) minerals in the Jeronimo ore suite consist of pyrite, arsenopyrite, sphalerite, galena, lead sulphosalt(s), realgar and orpiment. Typically, arsenopyrite and pyrite, commonly in association with rutile, are disseminated within the quartz matrix. Arsenopyrite also occurs as inclusions, veinlets and partial rims to pyrite grains. Sphalerite occurs mainly as inclusions in the manganese carbonate minerals and also in bands cross-cutting the quartz matrix with pyrite and apatite. Galena and lead sulphosalt grains occur dominantly in vugs in the quartz matrix. Realgar and orpiment are present mainly in late fractures and veinlets, normally with calcite and occasional grains of cinnabar. Minor, generally micron- to sub-micron-sized ore minerals including  170  cassiterite, coloradoite, altaite, stibnite and native silver (?) are present mainly as inclusions in quartz and carbonate minerals. G o l d occurs in native form generally as submicron-sized anhedral grains present as inclusions in pyrite, arsenopyrite, realgar and quartz or at vug margins in the quartz matrix. Geochemical data from the ore intercepts of the drill holes logged in this study indicate that Jeronimo ore is characterized by high concentrations of A s , M n , Z n , Pb, Sb, A g and H g . The behaviour of these elements with respect to gold varies widely between ore intercepts. In general, elements that correlate positively with gold include A s , H g , Te, Pb, Sb, Fe and M n .  5.1.4 Isotopic compositions Values of  5  1 8  0  ranging from  16.65  to  22.52%  0  ( V S M O W ) for manganese  carbonate minerals from the Jeronimo ore zone suggest that they may have been precipitated from magmatic fluids that had exchanged oxygen with the E l Asientos Formation host rocks. Whereas values of  5  1 3  C  of  -2.84  to  -1.30%o  (PDB) for the  manganese carbonates do not identify specific fluid sources, they are in agreement with the 8 0 results and suggest that the carbonate 8 C isotopic composition was derived 1 8  1 3  from the mixing of dissolved material from the E l Asientos limestone with magmatic volatiles and/or organic matter. The lead isotope composition of Jeronimo ore zone galena, sphalerite, lead sulphosalt and realgar is interpreted to represent lead derived from the Eocene to Oligocene porphyries mixed with a component of lead derived from the Jurassic E l Asientos Formation limestones and/or pre-Jurassic basement rocks. The lead isotope  171  values of orpiment, present in late veins with calcite and realgar, are closer in composition to lead derived from the E l Asientos formation host rocks. This late vein material may have been deposited from meteoric fluid in equilibrium with the host limestone.  5.2 Comparison with Carlin and gold-rich carbonate replacement deposits T w o deposit types show similarities to Jeronimo: (1) Carlin-type disseminated gold deposits and (2) gold-rich carbonate replacement deposits. The main characteristics of these deposit types are summarized in Table 5.1, which also includes the characteristics of the Jeronimo deposit for comparison. Berger and Bagby (1991) and Arehart (1996) have recently summarized Carlintype deposits. Ore deposition in this type of deposit is strongly influenced by lithology and structure. Preferential ore-hosting horizons are calcareous sandstones and siltstones, and mineralization typically fans out where steep faults intersect favourable horizons (e.g. Getchell, Joralemon, 1951). The dominant sulphide mineral assemblage in Carlintype deposits includes pyrite, arsenopyrite, realgar and orpiment and the deposits are characterized by the metal association A s - S b - H g - T l . Alteration types in Carlin systems consist mainly of decalcification, silicification and argillization, with the deposition of calcium carbonate minerals in some cases. A u in ore from unoxidized Carlin-type deposits occurs mainly in association with arsenian pyrite (e.g. T w i n Creeks, Nevada, Simon, 1999), which is normally present as distinct grains or rimming pyrite. G o l d may occur as (1) thin films coating arsenian pyrite, (2) sub-micron-sized inclusions of native A u in arsenian pyrite, and (3) as structurally-bound atoms of A u in arsenian pyrite.  172  Although intrusive units are almost always present within or in proximity to Carlindeposits, their involvement in mineralization is not clear. The intrusions may play passive roles, as interfaces along which fluids may ascend, or active roles providing heat, fluids, and/or metals to the deposits. In contrast, clear genetic links between igneous intrusive activity and the formation of carbonate replacement deposits (CRDs) can, in most cases, be made (Titley, 1991). C R D s , summarized in Megaw (1998), comprise a member of a zoned sequence of magmatic-hydrothermal deposits centered on intrusions, including porphyry, skarn and vein deposits (e.g. Ruth District, Nevada, Albino, 1995). C R D s occur as replacementstyle deposits in limestones. They are strongly lithologically controlled, preferring sandy and silty limestones, and normally possess manto shapes. Although typically mined for base metals, some C R D s are gold-rich and possess only minor base-metal enrichment (e.g. Star Pointer, Albino, 1995). Sillitoe and Bonham (1990) proposed that gold-rich C R D s , summarized in Sillitoe (1991a), form the most distal member of the sequence of intrusion-related deposits, occurring outward of the base-metal C R D s . The dominant sulphide mineral assemblage in gold-rich C R D s consists of pyrite, arsenopyrite, sphalerite, galena, chalcopyrite, marcasite and pyrrhotite and their geochemical signature, which is more variable than that of Carlin-type deposits, consists of As-Sb-Pb-Zn ± A g , M n , TI, Te, C u , H g , B i , Sn. L i k e Carlin-type deposits, alteration types in gold-rich C R D s consist of decalcification, silicification, argillization and carbonatization. In contrast to Carlin-type deposits, igneous intrusions are considered to be sources of heat, fluid and metals for C R D s , although they may also derive components from the host rock (Titley, 1993). G o l d in C R D s occurs as discreet particles, normally ranging from 1 to 100 pm,  173  that are either encapsulated in gangue and sulphide minerals or occur at mineral edges (e.g. Purisima Concepcion, Alvarez and Noble, 1988). Gold-rich C R D s are commonly enriched in M n (e.g. Purisima Concepcion, Alvarez and Noble, 1988; Cove, Emmons and Coyle, 1988; Star Pointer, Albino, 1995).  Table 5.1  Characteristics of Carlin-type and gold-rich carbonate replacement deposits, with the Jeronimo  deposit characteristics for comparison.  Carlin-type deposits  Morphology  Dominant host lithology Hypogene alteration assemblage  Major sulphide assemblage  Fault and lithologycontrolled; mineralization tends to spread out where conduits intersect favourable horizons Calcareous siltstones and sandstone Decalcification, silicification, argillization, Calcitization Pyrite, arsenopyrite, realgar, orpiment, stibnite, cinnabar  Geochemical signature  As-Sb-Hg-Tl  Mode of gold occurrence  Occurs mainly as thin films, very finegrained inclusions and structurally bound atoms in arsenian pyrite  Relationship to igneous intrusion  Role of igneous intrusions is uncertain  Gold-rich carbonate replacement deposits Strong lithological and fault-control; tend to be mantoshaped  Jeronimo  Silty and sandy limestones  Bioclastic limestones  Decalcification, silicification, argillization, Carbonatization Pyrite, arsenopyrite, sphalerite, galena, chalcopyrite, marcasite, pyrrhotite  Decalcification, silicification, argillization, carbonatization Pyrite, arsenopyrite, sphalerite, realgar, orpiment, galena, lead sulphosalt(s) As-Mn-Zn-Pb-SbAg-Hg  As-Sb-Pb-Zn±AgMn-Tl-Te-Cu-HgBi-Sn Occurs as grains of native gold, typically 10 to 100 pm, as inclusions or at grain boundaries Genetically related to igneous intrusions  Manto-shaped, strong lithological control due to increased porosity and permeability  A s grains of native gold, generally up to 1 u,m, as inclusions in sulphides and at vug boundaries Proximal to the Potrerillos porphyry which hosts the Cobre porphyry C u M o - ( A u ) deposit  174  5.3 Metallogenic relationships in the Potrerillos District In addition to Jeronimo, the Potrerillos district hosts two other mineral deposits, the Potrerillos porphyry C u - M o - ( A u ) deposit and the E l Hueso sedimentary and volcanic rock-hosted disseminated gold deposit. The Potrerillos deposit, hosted in the quartz monzodioritic Cobre porphyry, is located approximately 4.3 kilometres west-northwest of Jeronimo. The Cobre porphyry was emplaced into the Montandon, E l Asientos and Pedernales formations and produced a skarn zone extending up to 1.5 kilometres from the porphyry margin. The emplacement of the porphyry was syntectonic with movement along the Potrerillos M i n e fault (Tomlinson, 1994; Mpodozis et al., 1994). According to Sillitoe (1997), the presence of quartz veins associated with sericitic alteration, within the plane of the Potrerillos M i n e fault, which truncates the Cobre porphyry at a depth of about 600 metres, indicates that porphyry-style mineralization occurred synchronrously with thrust movement. The E l Hueso deposit is located at the surface approximately 1.5 kilometres to the west of Jeronimo. Mineralization in E l Hueso is replacement-style and is concentrated in steep, east-west trending faults and spreads out into the bioclastic limestones of the upper E l Asientos Formation and the tuffs of the Tertiary Hornitos formation. E l Hueso is also truncated by the Potrerillos M i n e fault, which intersects the surface at the eastern edge of the deposit (Figure 5.1). Marsh (1997) dated hydrothermal and magmatic minerals from Potrerillos and E l Hueso in order to determine the relationships among E l Hueso, Potrerillos and intrusions in the district. A n age of 40.2 to 40.8 M a for sericite associated with gold mineralization in a porphyry dyke in the E l Hueso open pit was interpreted by Marsh (1997) as the age of gold mineralization at E l Hueso. This age is significantly  175  (sarpui) U0IJBA3J3  176  older than the age of 35.6 to 35.9 M a obtained from magmatic and hydrothermal hornblende, biotite and sericite from the Potrerillos deposit and discounts a genetic relation between the two. In contrast, hypogene alunite in the E l Hueso deposit yields an age of 36.23 ± 0.07 M a , which is reasonably close to the age of Potrerillos mineralization and may represent the remnants of a district-wide lithocap, formed as the epithermal expression of Potrerillos mineralization. During the present study, an attempt was made to date mineralization in the Jeronimo deposit by the analysis of rutile by the U - P b method. Unfortunately rutile was not present in sufficient quantity or in large enough grains to obtain enough sample for dating by this method. The potential of a genetic link between Jeronimo and the Potrerillos porphyry was also tested in this study. It is possible that hydrothermal fluids originating at the Potrerillos porphyry ascended along conduits, such as (1) west-dipping thrust faults and associated tear faults, or (2) the relatively porous and permeable, gently west-northwestdipping bioclastic units. In the present study, the latter possibility was tested by logging six drill core intercepts of the C-horizon extending from the western margin of Jeronimo to a distance of 1.7 kilometres from the western limit of the deposit, shown in the eastwest cross-section in Figure 5.1, in order to assess variation in mineralogy and alteration in the C-horizon at an increasing distance from the deposit. Unfortunately, the distribution of drill holes in the area between the Homestake property boundary and the Potrerillos porphyry, a distance of approximately 1.5 kilometres, which was owned by C O D E L C O at the time of core logging, is sparse and access to existing cores and logs was limited.  177  Figure 5.1 shows that the most distal occurrence of the C-horizon observed in this study was in drill hole D M 3 . A t this location, the bioclastic limestone of the C-horizon is composed of a calcite matrix with illite- or kaolinite-filled vugs; disseminated pyrite grains, locally hosting thin growth layers of arsenian pyrite; and minor disseminated rutile and sphalerite. A limestone sample from this location was analyzed for carbon, oxygen and lead isotopic composition. The light isotope and lead analyses both indicated that limestone at this location possesses the typical signature of Jurassic carbonates according to the fields established by Veizer and Hoefs (1976) for the carbon and oxygen, and Tosdal (unpublished data) for lead, suggesting that, at this location, the calcite in the C-horizon had not exchanged with fluids of a different isotopic composition. A t the western margin of Jeronimo, the unmineralized C-horizon was examined in drill hole D D H 2 0 1 8 1 . The bioclastic limestone at this location consists of a calcite matrix with kaolinite-filled vugs, patchy quartz veining, and minor disseminated pyrite and rutile. The 8 C and 8 0 values for the C-horizon are again typical of Jurassic 1 3  1 8  carbonates according to Veizer and Hoefs (1976), however this sample is strongly enriched in radiogenic lead relative to other Jurassic marine carbonates in the Potrerillos district. With the exception of D D H 20181, the drill holes in the E - W section intersect a bedding-concordant, 2 to 18 metre thick zone of brecciation, lying either within or immediately below the C-horizon. This breccia zone consists of rubbly, strongly decalcified and weakly silicified fragments of limestone with up to 7 % pyrite. G o l d  178  grade is consistently higher in the breccia than in the more competent parts of the C horizon and increases towards the Jeronimo ore body (from <0.01 to 0.72 g/t Au). The presence of alteration minerals within the C-horizon up to a distance of at least 800 meters beyond the western edge of the Jeronimo ore zone and the occurrence of an altered zone of brecciation, which is semi-concordant to the C-horizon and shows a gradual decrease in gold grade away from the deposit margin, indicate that fluid flow has occurred along this plane. However, the isotopic composition of calcite within the C horizon indicates that this mineral has not exchanged with fluids bearing a magmatic component. The concentric elliptical patterns of zonation of the carbonate mineral types and clay minerals (sections 3.2.2.2 and 3.2.3.3), focussed on the centre of the Jeronimo ore body, do not argue for the lateral introduction of ore fluids to the C-horizon from the west. Instead, this pattern of zonation suggests that the fluids ascended along steep faults until they intersected segments of the C-horizon and spread laterally and concentrically.  5.4 Genetic model for Jeronimo 5.4.1 Timing of Jeronimo ore deposition Although it is possible that Jeronimo mineralization may have occurred prior to the offset of bioclastic horizon by the Mesozoic steep normal faults, it is unlikely considering the regional context of mineralization. According to Sillitoe (1991b), most of Chile's gold deposits formed after the onset of compression beginning in the Tertiary. This is the case for the other two major deposits in the Potrerillos district, the Potrerillos porphyry Cu-Mo-(Au) and El Hueso disseminated gold deposits, which formed during the Eocene and Oligocene. Furthermore, the presence of mineralization in faults and  179  fractures interpreted by Caddey (1999) to have formed during the waning stages of compression in the early Oligocene, simultaneous with Potrerillos porphyry C u - M o - ( A u ) mineralization and emplacement of most of northern Chile's porphyry copper deposits, suggests that Jeronimo mineralization did not occur until this time.  5.4.2 Hydrothermal fluid origins The lead isotope compositions of Jeronimo ore zone sulphide minerals indicate that lead hosted in these minerals was dominantly derived from a magmatic source. Some lead derived from the Jurassic host rocks and/or the dominantly granitic preJurassic basement was also present in the fluid. F r o m the lead isotope data, it is not possible to determine whether the porphyry was emplaced into the basement or overlying sedimentary rock, nor is it possible to identify a specific intrusion as the source of the fluid. A link can be made between the deposition of the sulphide minerals and gold based on the positive correlation of gold with lead and antimony in the geochemical data and on the occurrence of a gold grain in main stage realgar which was analyzed for its common lead composition. Likewise, the carbon and oxygen isotope composition of ore zone manganese carbonate minerals suggests that carbon and oxygen in the fluids from which they were deposited had a magmatic signature, mixed with dissolved material from the E l Asientos formation limestone.  180  5.4.3  Fluid  pathways  Potential conduits that may have channeled fluids to the site of Jeronimo ore deposition include one or a combination of the following. •  The Mesozoic steep normal faults. These locally host gold mineralization but are not strongly altered and do not display evidence of significant fluid flow.  •  The early Tertiary thrust faults. The Potrerillos M i n e fault was a focus of fluid flow during the formation of the Potrerillos porphyry C u - M o - ( A u ) deposit. In the case of Jeronimo, fluids may have traveled up the fault plane until they encountered a receptive unit or other structure through which the fluids could have flowed laterally to the site of ore deposition.  •  The bioclastic unit. F l u i d flow may have occurred within the gently west-northwestdipping, relatively porous and permeable bioclastic unit. This possibility is not supported by the carbon and oxygen isotopic composition of carbonate minerals within the bioclastic horizon distal to Jeronimo as these do not record the passage of fluids with a magmatic signature. Furthermore, the concentric patterns formed by the distribution of the carbonate and clay minerals argue against the lateral introduction of fluids to the site of Jeronimo mineralization.  •  The early Tertiary normal faults. Fluids could have flowed up the steep, normal faults observed to cross-cut the C , D and E horizons by Caddey (1999). Mineralization is present in these structure, however, to confirm their role as conduits, they need to be examined below the ore-hosting horizon in order to verify i f there is evidence of fluid flow at depth.  181  5.4.4  Depositional history  A t the site of deposition, cooling of the hydrothermal fluids resulted in the dissolution of the host limestone and simultaneous deposition of quartz, accompanied by the precipitation of pyrite, arsenopyrite, sphalerite, rutile, apatite, barite, monazite and possibly some illite. The presence of gold as inclusions in pyrite, arsenopyrite and quartz suggests that some gold deposition occurred at this time. The deposition of monazite and rutile during silicification suggests that the fluids were somewhat acidic. Local bulk dissolution of the limestone occurred during silicification, creating significant open space. Subsequently, carbonate minerals were deposited from fluids including components most likely derived from magmatic volatiles and dissolved E l Asientos formation limestone. The carbonate minerals were deposited in response to a decrease in /cc-2 which most likely occurred by either boiling or fluid mixing. The dominance of manganese carbonate minerals at the centre of the deposit implies that fluid M n / C a ratio decreased as the fluid spread away from the centre of the ore zone, possibly due to interaction with the limestone host. Pyrite, sphalerite, rutile, rhodonite and barite were also precipitated during this phase of carbonate deposition. Following carbonatization, the deposition of illite occurred at temperatures likely to be between 230-320°C from fluids of near neutral p H . Illite deposition was accompanied by the precipitation of pyrite and apatite. The latest events recorded in the Jeronimo ore are the deposition of kaolinite, A P S mineral precipitation, and localized supergene oxidation. Kaolinite was deposited from acidic fluids at approximately 120°C, possibly from oxidized supergene fluids. Acidic  182  fluids were also required to deposit the A P S mineral. The deposition of Mn-oxide observed in one sample from the sulphide body occurred by late oxidation of the manganese carbonate minerals and involved the remobilization of some gold. A model representing the potential components involved in the formation of the Jeronimo deposit is presented in Figure 5.2. The model shows that: •  The source intrusion(s) is emplaced into either the dominantly granitic pre-Jurassic basement or the Jurassic sedimentary rocks.  •  The potential paths of fluid flow (represented by arrows with curved tails) to the site of deposition may include one or more of the following conduits: (1) the late Mesozoic steep normal faults; (2) an early Tertiary thrust fault combined with a medium for the lateral transport of the fluids to the site of ore deposition, such as the bioclastic limestone; (3) the early Tertiary normal faults.  •  There is a concentric pattern of zonation, focussed on the centre of the deposit, based on carbonate mineral composition, with manganese carbonate minerals at the centre of the deposit and calcium carbonate minerals at the margins.  •  Late arsenic sulphide and calcite veins were deposited from fluids that had equilibrated with the E l Asientos formation host rocks.  •  Late kaolinite deposition was restricted to the centre of the ore body.  183  W  Figure 5.1 Model of the formation of the Jeronimo deposit. Curved arrows indicate potential paths of fluid flow to the site of mineralization, including (1) up late Mesozoic basin growth normal faults; (2) up early Tertiary normal faul ts; and (3) along early Tertiary thrust faults and within bioclastic units, from an intrusive to the west. The early Tertiary normal faults are observed only in tunnel exposures. The extent of their continuation below unit C is not known. Heavy arrows indicate paths of fluid flow required to account for the common Pb compositions of main stage sulphide minera Is. The Jeronimo ore zone is shown in gray, and is shaded according to the dominant carbonate composition.  184  5.5 Concluding statement The key features of the Jeronimo sediment-hosted disseminated gold deposit include the following: •  Jeronimo is a carbonate-hosted gold deposit on the eastern side of the Potrerillos-El Hueso porphyry C u - M o - ( A u ) system of Eocene to Oligocene age. The precise age of mineralization at Jeronimo has not been defined, but it is interpreted to be of similar age based on the regional setting, and the style and zoning of mineralization in the district.  •  Jeronimo is both lithologically and structurally controlled. Ore is hosted in bioclastic limestone horizons, due to their enhanced primary and secondary porosity and permeability. Jeronimo mineralization is present in steep Tertiary normal faults.  •  Alteration types recorded in the Jeronimo ore include decalcification, silicification, carbonatization and argillization. Jeronimo hosts manganese carbonate mineralization in a zone overlapping the centre of the deposit. The Jeronimo clay mineral assemblage consists of illite and kaolinite, the latter of which may be supergene in origin. Other alteration minerals that are characteristic of the Jeronimo ore are rutile, apatite, monazite, and barite.  •  The Jeronimo sulphide and sulphosalt mineral suite includes pyrite, arsenopyrite, sphalerite, galena, lead sulphosalt, orpiment, realgar. Coloradoite, altaite, cassiterite and stibnite occur as occasional micron-sized grains. High concentrations of A s , M n , Zn, Pb, Sb, A g and H g characterize the Jeronimo deposit. G o l d is present as generally sub-micron-sized grains of native gold encapsulated within pyrite, arsenopyrite, realgar and quartz, or within vugs in the quartz matrix.  185  •  The Jeronimo ore-forming fluids were derived from a magmatic source, but also incorporated material from the Jurassic sedimentary host rocks and/or the dominantly granitic pre-Jurassic basement.  Table 5.2 summarizes the distinguishing characteristics between the two deposit types. The presence of significant lead and zinc; the occurrence of gold as discrete, native grains; the presence of manganese mineralization; and the magmatic source of the Jeronimo fluids indicate that Jeronimo is more similar to gold-rich carbonate deposits than to Carlin-type deposits.  Table 5.2 Distinguishing characteristics of Carlin-type and gold-rich carbonate replacement deposits  Carlin-type deposits Paucity of base metals  In hypogene ore, gold occurs mainly as thin films, very fine-grained inclusions and structurally bound atoms in arsenian pyrite Sulphide assemblage is very fine grained, generally < 1 p m to 100 p m N o Mn-enrichment Uncertain relationship to igneous intrusions  Gold-rich carbonate replacement deposits Typically show some base metal enrichment. Sulphide assemblage includes sphalerite and galena G o l d typically occurs as free grains, ranging from 1 to 100 pm, as inclusions and at grains boundaries of sulphide and silicate minerals Sulphide minerals frequently > 100 p m Mn-enrichment is common (e.g. Cove, Purisima Concepcion, Star Pointer) Genetically related to igneous intrusions  The classification of Jeronimo as a gold-rich carbonate replacement deposit implies that a source intrusion must be present in the vicinity of the deposit. Whether this intrusion is the Cobre porphyry or another, unidentified porphyry is uncertain. In general, the tendency of thrust faults and bedding to dip shallowly to the west-northwest suggests that the source intrusion may be in that direction, such that exsolved fluids would migrate up hydraulic gradients. Given the occurrence of gold-rich C R D s as the  186  distal members in a zoned sequence of deposits, other intrusion-related deposits, such as base metal CRDs and skarn are potential exploration targets in the Potrerillos district.  187  References  Albino, G . V . 1995. Porphyry copper deposits of the Great Basin - Nevada, Utah and adjacent California. In Porphyry copper deposits of the American cordillera. Edited by F . W . Pierce and J.G. B o l m . Publication of the Arizona Geological Society, pp. 267-296. Alvarez, A . A . and Noble, D . C . 1988. Sedimentary rock-hosted disseminated precious metal mineralization at Purisima Concepcion, Yauricocha district, central Peru. Economic Geology, 83: 1368-1378. Bagby, W . C . and Berger, B . R . 1985. Geologic characteristics of sediment-hosted, disseminated precious-metal deposits in the western United States. In Geology and geochemistry of epithermal systems. Edited by Berger, B . R . and Bethke, P . M . Reviews in Economic Geology, 2: 169-202. Beaudoin, G . 1997. Proterozoic Pb isotope evolution in the Belt-Purcell Basin; constraints from syngenetic and epigenetic sulfide deposits. Economic Geology, 92: 343350. Blatt, H . , Middleton, G . and Murray, R. 1980. Origin of sedimentary rocks. Prentice-Hall, U S A , 782 pp. Cabri, L . J . , Chryssoulis, S.L., De Villiers, J.P.R., LaFlamme, J . H . G . and Buseck, P.R. 1989. The nature of "invisible" gold in arsenopyrite. Canadian Mineralogist, 27: 353-362. Caddey, S. 1999. Preliminary structural investigation, ore controls, and exploration implication of the A g u a de la Falda gold deposit, Potrerillos, Chile. Minera Homestake Chile S.A. internal report, 28 pp. Callan, N . 1998. Geologic map of the Agua de la Falda property, Potrerillos district, Atacama Region, Chile. Minera Homestake Chile S . A . internal report. Chang, L . L . Y . , Howie, R . A . and Zussman, J. 1996. Rock-Forming Minerals. Volume 5B: Non-silicates. Sulphates, Carbonates, Phosphates and Halides. Second edition. Longman Group Limited, England. Childe, F . C . and Thompson, J . F . H . 1997. Geological setting, U - P b geochronology, and radiogenic isotopic characteristics of the Permo-Triassic Kutcho assemblage, northcentral British Columbia. Canadian Journal of Earth Sciences, 34: 1310-1324. Choquette, P . W . and Pray, L . C . 1970. Geologic nomenclature and classification of pOrosity in sedimentary carbonates. The American Association of Petroleum Geologists Bulletin, 54: 207-244.  188  Cornejo, P., Mpodozis, C , Ramirez, C F . and Tomlinson, A.J.1993. Estudio geologico de la region de Potrerillos y E l Salvador (26°-27° Lat. S.): Santiago, Servicio Nacional de Geologia y M i n e r i a - C O D E L C O , Registered report IR-93-01, 258 pp. Davidson, J. and Mpodozis, C . 1991. Regional geologic setting of epithermal gold deposits, Chile. Economic Geology, 86: 1174-1186. Dilles, P . A . 1995. Geologic Mapping of the Manto A g u a de la Falda project area with technical and exploration recommendations, E l Hueso M i n e , Chile. Minera Homestake Chile S.A. internal report. Duba, D . and Williams-Jones, A . E . 1983. The application of illite crystallinity, organic matter reflectance, and isotopic techniques to mineral exploration: a case study in southwestern Gaspe, Quebec. Economic Geology, 78: 1350-1363. Eberl, D . D . and Velde, B . 1989. Beyond the Kubler index. Clay Minerals, 24: 571-577. Elgueta, S. and Fuentes, M . 1997. Sedimentological study of the Jurassic sequence of the area between quebradas E l Asiento and E l Hueso: implications for the exploration of gold ore bodies of Jeronimo type. Minera Homestake Chile S . A . internal report, 13 pp. Emmons, D . L and Coyle, R . D . 1988. Echo B a y details exploration activities at its Cove gold deposit in Nevada. M i n i n g Engineering, 40: 791-794. Essene, E . J . 1983. Solid solutions and solvi among metamorphic carbonates with applications to geologic thermobarometry. In Carbonates: Mineralogy and Chemistry. Edited by R . J . Reeder, Reviews in Mineralogy, 11, Mineralogical Society of America, Washington, D . C , pp. 77-96. Fletcher, W . K . 1981. Handbook of exploration geochemistry. Elsevier, the Netherlands, 255 pp. Fournier, R . O . 1985a. The behaviour of silica in hydrothermal solutions. In Geology and geochemistry of epithermal systems. Reviews in Economic Geology Volume 2. Edited by Berger, B . R . and Bethke, P . M . Society of Economic Geologists, pp. 45-62. Fournier, R . O . 1985b. Carbonate transport and deposition in the epithermal environment. In Geology and geochemistry of epithermal systems. Reviews in Economic Geology Volume 2. Edited by Berger, B . R . and Bethke, P . M . Society of Economic Geologists, pp. 63-72. Frondel, C . and Bauer, L . H . 1955. Kutnahorite: A manganese dolomite, C a M n ( C 0 ) . American Mineralogist, 40: 748-760. 3  2  Frutos, J. and Tobar, A . 1973. Unpublished map of the E l Salavador-Potrerillos region. Inst. Invest. Geol., Santiago, Chile.  189  Fu, M . , Changkakoti, A . , Krouse, H . R . , Gray, J. and Kwak, T . A . P . 1991. A n oxygen, hydrogen, sulfur and carbon isotope study of carbonate-replacement (skarn) tin deposits of the Dachang tin field, China. Economic Geology, 86: 1683-1703. Giere, R. 1996. Formation of rare earth minerals in hydrothermal systems. In Rare Earth Minerals. Edited by Jones, A . P . , W a l l , F . and Williams, C T . Chapman and H a l l , Great Britain, 372 pp. Harrington, 1961. Geology of parts of Antofagasta and Atacama provinces, northern Chile. Bulletin of the American Association of Petroleum Geologists, 45: 169-197. Haynes, F . M . and Kesler, S.E. 1988. Compositions and sources of mineralizing fluids for chimney and manto limestone-replacement ores in M e x i c o . Economic Geology, 83: 1985-1992. Hernandez, E . , Lazcano, A . and Carrasco, M . 1991. Geology of the E l Hueso gold deposit, Potrerillos district, Atacama, Chile. Minera Homestake Chile S . A . internal report, 20 pp. Heyl, A . V . , Landis, G . P . and R . E . Zartman. 1974. Isotopic evidence for the origin of Mississippi valley-type mineral deposits: a review. Economic Geology, 69: 992-1006. Honea, R . M . Petrology of the Jeronimo deposit. Minera Homesake Chile S . A . internal report, 203 pp. Joralemon, P. 1951. The occurrence of gold at the Getchell mine, Nevada. Economic Geology, 46: 267-310. Kesler, S.E., Cumming, G . L . , Krstic, D . and Appold, M . S . 1994. L e a d isotope geochemistry of Mississippi valley-type deposits of the southern Appalachians. Economic Geology, 89: 307-321. Lang, J.R. and Baker, T. 1999. The role of fluid immiscibility in some M e x i c a n skarn deposits. In Mineral Deposits: Processes to Processing. Edited by Stanley et al., Balkema, Rotterdam. Lewis, P. 1996. A g u a de la Falda/El Hueso Project, Potrerillos District, Chile. Minera Homestake Chile S . A . internal report, 67 pp. Marsh, T . M . 1997. Geochronology, thermochronology and isotope systematics of the C u A u and A u - A g deposits of the Potrerillos District, Atacama Region, Chile. Unpublished Ph.D. thesis, Stanford University, 341 pp. Matthews, A . and Katz, A . 1977. Oxygen isotope fractionation during the dolomitization of calcium carbonate. Geochimica et Cosmochimica Acta, 41: 1431-1438.  190  M c C r e a , J . M . 1950. O n the isotopic chemistry of carbonates and a paleotemperature scale. Journal of chemical physics, 18: 849-857. Megaw, P . K . M . 1998. Carbonate-hosted P b - Z n - A g - C u - A u replacement deposits: A n exploration perspective. In Mineralized intrusion-related skarn systems. Mineralogical Association of Canada Shortcourse Series, 26: 337-358. Megaw, P . K . M . 1987. Oxygen and carbon isotopic shifts between altered and unaltered limestone in the peripheries of the Santa Eulalia M i n i n g District, Chihuahua, M e x i c o . Geological Society of America Program Abstracts, 19: 769. Moore, D . M . and Reynolds, J.R.C. 1997. X-ray diffraction and the analysis of clay minerals, 2 edition. Oxford University Press, Oxford, 378 pp. n d  Mpodozis, C . and Ramos, V . 1990. The Andes of Chile and Argentina. Circum Pacific Council for Energy and Mineral Resources, Earth Science Series, 11: 59-90. Mpodozis, C , Tomlinson, A . J . and Cornejo, P. 1994. Acerca del control estructural de intrusivos eocenos y porfidos cupriferos en la region de Potrerillos-El Salvador. Septieme Congreso Geologico Chileno, Universidad de Conception, 2: 1596-1600. Mukasa, S.B., V i d a l , C . E . and Injoque-Espinoza, J. 1990. Pb isotope bearing on the metallogenesis of sulphide ore deposits in central and southern Peru. Economic Geology, 85: 1438-1446. Nickel, E . H . 1992. Solid solutions in mineral nomenclature. Canadian Mineralogist, 30: 231-234. Ohmoto, H . 1986. Stable isotope geochemistry of ore deposits. In Stable isotopes in high temperature geological processes. Edited by J.W. Valley et al. Mineralogical Association of Canada Reviews in Mineralogy, 16: 491-559. Ohmoto, H . and Rye, R.O. 1979. Isotopes of sulfur and carbon. In: Geochemistry of hydrothermal ore deposits. Second edition. Edited by Barnes, H . L . John Wiley & Sons. U S , pp. 509-567. Olson, S.F. 1983. Geology of the Potrerillos District, Atacama, Chile. Unpublished Ph.D. thesis, Stanford University, 190 pp. O ' N e i l , J.R., Clayton, R . N . and Mayeda, T . K . Oxygen isotope fractionation in divalent metal carbonates. The Journal of Chemical Physics, 51: 5547-5558. Perez, E . 1977. Estratigraphia de Quebrada Asientos, norte de Potrerillos. Inst. Invest. Geol., informe inedito, Santiago, 202 pp.  191 Pouchou, J.L. and Pichoir, F. 1985. PAP d)(pZ) procedure for improved quantitative microanalysis. Microbeam analysis, 1985: 104-106. Radtke, A.S. 1985. Geology of the Carlin gold deposit, Nevada: U.S. Geological Survey Professional Paper 1267, 124 pp. Ramdohr, P. 1980. The ore minerals and their intergrowths. Second edition. Pergamon Press, Germany, 1205 pp. Reyes, A.G. 1990. Petrology of the Philippine geothermal systems and the application of alteration mineralogy to their assessment. Journal of Volcanology and Geothermal Research, 43: 279-309. Richards, J.P., McCulloch, M.T., Chappell, B.W. and Kerrich, R. 1991. Sources of metals in the Porgera gold deposit, Papua New Guinea: Evidence from alteration, isotope and noble gas geochemistry. Geochimica et Cosmochimica Acta, 55: 565-580. Robinson, D., Warr, L.N., Bevins, R.E. 1990. The illite 'crystallinity' technique: a critical appraisal of its precision. Journal of Metamorphic Geology, 8: 333-344. Rollinson, H. 1993. Using geochemical data: evaluation, presentation, interpretation. Longman, United Kingdom. Roth, T. 1993. Geology, alteration and mineralization in the 21A zone, Eskay Creek, Northwestern British Columbia. Unpublished M.Sc. thesis, 230 pp. Rye, R.O. 1966. The carbon, hydrogen and oxygen composition of hydrothermal fluids responsible for the lead-zinc deposits at Providencia, Zacatecas, Mexico. Economic Geology, 61: 1379-1397. Seward, T.M. 1991. The hydrothermal geochemistry of gold. In Gold Metallogeny and Exploration. Edited by R.P. Foster. Blackie and Son, Glasgow, pp. 37-62. Simmons, S.F. and Christenson, B.W. 1994. Origins of calcite in a boiling geothermal system. American Journal of Science, 294: 361-400. Simon, G., Kesler, S.E. and Chryssoulis, S. 1999. Geochemistry and textures of goldbearing arsenian pyrite, Twin Creeks, Nevada: implications for deposition of gold in Carlin-type deposits. Economic Geology, 94: 405-422. Sillitoe, R.H. 1997. Comments on the Jeronimo gold deposit and its position in the Potrerillos district, Northern Chile. Minera Homestake Chile S.A. internal report, 11 pp. Sillitoe, R.H. 1991a. Intrusion-related gold deposits. In Gold Metallogeny and Exploration. Edited by R.P. Foster. Blackie and Son, Glasgow, pp. 165-209.  192  Sillitoe, R . H . 1991b. G o l d metallogeny of Chile - an introduction. Economic Geology, 86: 1187-1205. Stoffregen, R . E . and Alpers, C . N . 1987. Woodhousite and svanbergite in hydrothermal ore deposits: products of apatite destruction during advanced argillic alteration. Canadian Mineralogist, 25: 201-211. Taylor, B . E . 1987. Stable isotope geochemistry of ore-forming fluids. In Short course in stable isotope geochemistry. Edited by T . K . Kyser. Mineralogical Association of Canada Short Course Handbook, 13, pp.337-452. Titley, S.R. 1991. Characteristics of High-temperature, Carbonate-hosted Massive Sulphide Ores in the United States, M e x i c o and Peru. In Mineral Deposit Modelling. Edited by: R . V . Kirkham, W . D . Sinclair, R.I. Thorpe and J . M . Dunne. Geological Association of Canada Special Paper 40, pp. 585-614. Tomlinson, A . J . 1994. Relaciones entre el porfido cuprifero y la falla inversa de la mina de Potrerillos: un caso de intrusion sintectonica. Septieme Congreso Geologico Chileno, Universidad de Conception, 2: 1629-1633. Tomlinson, A . J . , Mpodozis, C , Cornejo, P., Ramirez, C.F., Dumitru, T. 1994. E l sistema de fallas sierra castillo-agua amarga: transpresion sinistral eocena en la Precordillera de Potrerillos-El Salvador. Septieme Congreso Geologico Chileno, Universidad de Conception, 2: 1459-1463. Tosdal, R . M . , Wooden, J.L. and Bouse, R . M . Pb isotopes, ore deposits and metallogenic terranes. In Reviews in Economic Geology. Society of Economic Geologists. In press. Tosdal, R . M . , Cline, J.S., Hofstra, A . H . , Peters, S.G., Wooden, J . L . and Young-Mitchell, M . N . 1998. M i x e d sources of Pb in sedimentary-rock-hosted A u deposits, northern Nevada. In Contributions to the gold metallogeny of northern Nevada. Edited by R . M . Tosdal. U S G S Open-File Report 98-338, pp. 223-233. Vasquez, R., Vennemann, T . W . , Kesler, S.E. and Russel, N . 1998. Carbon and oxygen isotopic haloes in the host limestone, E l Mochito Zn-Pb (Ag) skarn massive sulfide-oxide deposit, Honduras. Economic Geology, 93: 15-31. Veizer, J. and Hoefs, J. 1976. The nature of 0 / 0 and C / C secular trends in sedimentary carbonate rocks. Geochimica et Cosmochimica Acta, 40: 1387-1395. 1 8  1 6  1 3  1 2  Wilson, G . C . 1997. Altered porphyries and sedimentary rocks from the E l Hueso District, Chile. Minera Homestake Chile S . A . internal report, 60 pp. Wuensch, B . J . 1974. Sulfide crystal chemistry. In Short Course Notes Volume 1, Sulfide Mineralogy. Edited by P . H . Ribbe. Mineralogical Society of America, U S A , pp. 21-44.  193 Zartman, R.E. and Doe, B.R. 1981. Plumbotectonics; the model. Tectonophysics, 75: 135-162. Zeegers, H. and Leduc, C. 1991. Geochemical exploration for gold in temperate, arid, semi-arid, and rain-forest terrains. In Gold Metallogeny and Exploration. Edited by Foster, R.P. Blackie and Sons Ltd, Great Britain, pp. 309-335.  Appendix A: Carbonate mineral electron microprobe data (see section 3.2.2.3 for operating conditions)  <g $2 o co 8 o cS SI 8  w  .  co  CM  ^ CM' §  ^  c\i d d  »  •i-  • io  ^i  CN Si—* ™ CM O §  i—*  1  •<* It co  LO  LO  c\i  d  ? <o <o g ^  O °> CM lO  1^  CO O CO  CO  co co O CM  ^  O  O O  i —  eg  d  T~  "Cf  "> §  °  LO  co ^  o  CM  d  3  S o m g o 1 0  £  O  i -  Lf) S  00 O  O)  og d  d  r  O  O)  O  i - o  O  P O  N ^i-  O  •<t CO CM  o o  i -  1-  T-  0)  o o  ^ ^.  co o oo  O  to  O  O  T-  IO  t\i d  3  LO o O TCM d  -  O O O  2  B  o  O  T  -  03  2  ^.M in '•^ ^ <°.  O  CD ^  ri ° 00 't ° T -  ,  r-  co «N  I°° co 5 o n  n  <°.  °  O  Tf  w  K  m  a  i -  O  co  -r-  °  «t  LO  . CO  co d  co  N  CO CO  oo  CD d  7 _  'Cf  CM  S{  T-  CO  T- o I-; O r d  CO ~. rn LO CO M °? i s CM CM CM LO CM  oo o  3  CO T-  C 2  o  d  +  0 LL  0 5  'cf  00 .  CO 0  3  ^  r\i  O O 03 O ^ O CM  O cn  T-  CO £ LO  CO .  oori  'cf 00  0  0  rf  O O cj d  T—  oo O  oV  d  CM CO CO CO 00 O O CO CO ni d d r 0  co  CD CD d  CO•  CO CO O CO CO CM  'Cf  -  . oo co  CM  CM 'Cf O O  T-  CO LO i LO O O CO CO CM O O T-  LO  d  CM 'cf CO d T—  IN d  CM  . LO f^  00 T O O CM d  CM  LO  O  O CM CM  'Cf  'Cf  CD 'cf CO CM CO 0 O "3: CM 1 01 d d r d  . . o5  co oo  O  .  -<cr  °> CM  ° 3 f  CO 'Cf O O  co  CO CD  co  O  o i d d r d  CO  w  r  o o co  cd  ? d o i ^  s  To*  CO  S3  O  1-  O d  O  O  CM CM CO 'cf O "cr Tc j d d r d  CM  N  T-  Tf  O  O o  co co  CM CO CO CM O CO CO i —  (\i O  d  r- o  T-  CO  LO 00  O  §  c\i o o T - ' o  CM O  O  00  'cf  oo  O  LO i -  co oo  o o  o  <M O  ° ri LO  15 - ^  ,CM  T-"  c\i  -r- CM  2  CO CO CO i -  00 O O  LO  CO "Cf  CM CO T O LO CO d r d  -r-'  r-~  d  co co co i -  O  o  K o  -  r-'  i -  o  CD  LO CO CM  LO i -  £>O0  CM O  3  O  i - IO  CM CM LO CO -i- i - LO i T- d  O O CM O  CO CO QQ CM OO .  ^  CM 00 O O  cvi d d  CNJ CO  o  T-  o o -i- o r-~ c\i d o ' T - o'  CM §  CO  CO  CM LO CO 00  1  CM O  co co co -*f CM  "-  d  d  r  CM 'cf  CO O CM CM LO 1 -  CM d  d  O  O  co co  d  N  CO O T-" d  O d  "cf  O  O  CO  c\i d d  T-  O  O  T - r-  o  io  T-  c  O ^ O ^  CO d  T-  n  d  CD CD £ CO ffl (O) Tl  d  S«  rf  . CO O) CD § o oi  o  (M  ^  LO csi LO  _  . co co  CSJ  T-  Osl  CM  co  M  SS  CM  § » <N  d d  rf CM O q i - rf CM d d  00  O)  T-  o  CD  T-  LO  d 03  ^  r -  d  rf  o co o  LO LO  d  T-  rf  £5  CO CM  co r»  r- $ CD ^ o LO O CD rf  \h ° o Oram O ^ O  rf O CM  CO O  T-  rf  d d  00 rf  CO 00 01 T O O CO LO  O  rf 0  CO T O rf  -i-  -i-  q d  T-  O  d  d  d  . (M  O  2  rf  r  n  C O Trf  2  d  LO CO T - LO t 0 O CO LO o  01 d  d r  d  CD CM i - CO O O CO LO  d d  co oo o o CM CM d d  LO  00 LO T-'  T-  o d  T-  o d  CO CM T— - i — O O O rf LO O  od  5 .  d d ^  d  CM O) LO O 0 iq  i -  CM  g  00  . co  O CO X CM CO  (? 01 LO O ' CM rf ^ T - rf  rt 05 ^ . rf CO  O ri ^  CO  5  ri  0 5  jCO  CM  N  CO  CM  fe d  oo r- I? LO co LO . o d CM g d  rf 0  2  O  CM  CM  T-  d d  O  r-~ o d  Ol CM CM CM 0  O ^ O  C)  O i - l"» O 01 d d T - d  CM  O O O  T—  CO 00 O  SI d°  O  T-  rf 0  CM  °  O  CJ CJ  o o  CO LO  q d  r -  LO CD rf O O T - CO O  o  O  d  O) CO rf O rf CO  CM O  0  .  01 d  rf O  00 hrf O  d T-" d  O  O m ^  CM CM 0) CM CO 0 O CO rf T -  D) CC CC  o  d o V  rf  LO CM 00 i — O) O O CO rf O  01 d  1 0  ° . CM  r? Oi rf  1^ T rf O  d d ^  o  o  h~ 1 rf q -r^ d  CM  r- d  s. % s a §  rf  T-  rf rf  o  r  rf  t- ^ 05 ^  T"  01 d d  O O rf CM d d  O  co  d oj  rf  . CO  CO  ^ g  CO  ™  •  CM  CM  1-  0  co in 6  LO CO O LO CM O O CO LO o  d  oo o d  CD CO 0) rf  q q CM d  LO CO CD cd d cri rf  co  i -  o  . rf rf" O  CO CO 05 - i -  co LO LO co  ^  O  o o  O CM  r>- o d i< LO LO  r-»  o  S  CM CO CO  CD  r ? 0  LO 00  CO  O  O  O d  O rf co o 01 O O T - ci  rf  OJ  CM  0  CM  CM T- Tf  CD  in  1^  oj N O i01 d o T -  CM LO CO O LO  d rf  O O TCM d d  . o g d  ffl  CD CD  o  CD CM CM rf  O  o co q 01 d d ^ d  CM O  o  i -  01 o d  O) i-~  r  D) ro C  O S O S  O  q d  197  S. CO CM g rf LO CD . CO O  .  o o  CO  o  IO  rr g co  LO oo  ob d d CM rf  . CM co LO  CM  O  N CM  CM  d  d  N  LO  d  T-  N  o  O  CM  CO  CM  d  d  O  T—  m co O  T -  o d  rf CO CD OO i O O CM LO o c\i d d T - d  CD LO oo rf CO rf O l< d  to CO O) CO T O O CM LO O c\i d d T - d  LO  V  O  CD CD O LO  o h-  ?  !D CO  =0  • rf  CM - & ~ CO LO 8 d  CO  LO O  CO O  CD CM  c\i d d  CO LO  TO  T-  d c o io  CO CO o o O O CO CD o  CM O* O*  LO  T-  O  I" o  CM m  LO CD rf LO LO 0 0 T  -  LO CD CO LO i O  <—I  r  IO  c\i d d  CO  r  O  T-  d  CD 00  CC  CO O § CO LO LO . CO  CO CM O • CD• CM rf  O  •  O) CO LO CO q q CM LO CM d o T ~  io d  o c <°,  CM ,_  rf CO I - CM q q c\i co c\i d d r s  LO  CO  _4-  «? S  rf CM  <q S  3 CM £ g CD CD CD CD CD  to w N m O i - LO T CM d d T—  <  T -  q d  _CD  CC  w o d  d  CM • CO  o's2 6  M O O O) T O O CO rf O CM d d T— d CD  O c g CM §  CD  co oo  rf O  CM O  CM CO  CD LO  1O  O  CM  d  d  T—  d  00  OJ It X!  >* TJ CD  O O O o> cc lO ^ O CM  a> rc O 2 O  +  CM OJ LL  co oo rf LO CM LO LO CD CO CM o cq T— o O d T - : d d rf  3 _o CO O  CD O  c  O rc cc cc  CD  oo CD rf oo 1— LO 1— CD 02 CD CM o d d LO rf  CD  k.  J J O O O Q  O cn cc c a; O 2 O 2 LL  CD (0 CC  o ^ o ^  CD it TJ  >i  -Q TJ 0)  JC CJ _o CC  O  198  r*  co  is CO CM rf rf LO  d co co  CM  CO CD o < rf LO ID LD ^ CM co rf CO "—  OJ  0  .-;  d  CM  k  O  oo  O  1  °  Q  OO LO «? UO « CM. O ._: CD CO o ° CO T °  rf rf  O  O  CO  CM  d  i—  rf  00 rf  O  d  •  CO  5  co  c g  I  rf rf  o o  ^ °  CO Is CO T-  °  CO CM rf T O O CO LO CM O i - O  O  o O  rf _ 9  5  2 *  53  rf  00  T-  i -  O  CO LO O  T~ T-  O CO  CM LO Is CM O •>LO CM O i - O  i i -  O CO  CM LO LO CO CM O -i- i - LO i O O CM O  co  ^  rf  CO LO CM O -i- i - LO O CM O  O  co o od  co  T-  CO ID LO T- CM O 1 - T - LO 1 CM O i - O O CO  £ °2 co"- O _ ^ ^ co ,A cd CM ^ cq CO rf "- oo ^ CO  "co  rf  „  CN 00 ° ° oolr? CM  o o  rf  O CM  9 §  m  CD CO O 00 CO O 00 O CM d o d o  rf rf rf  rf  O  O  CM  LO CD i - CO I l i - CO CO  i -  d d  CO CO CO CM O O CO LO CM O i - O  co  • n ™"  oo co co oo o  ~ N CD CO i ~  T-:  ^  T  o o  rf  1 0  00 CD to T CO. £  T-  "- rf  CD  ^ T o  00  d d  i—  CO CD OO CO O i - CM CM d T d  V ^ CO CD LO CO i —  CD CO LO  d  °2  CM £  T-  rf  Is CD N CM  ci £ co  0  CD O 00 CO O CO O  O CO  CO T-  CM  rf  CO ~ CO CM • CD• LO f-J CD is rf ° CO i -  rf rf  LO is CO ' . _ ; is C-5 CO  °  in °  o ^ °  CO CO ^ O CD ~ . T—  ^ co  .  m  ^ CO ^ LO . - : co i s ° CO " -  rf  O ^ °  in CM ^CM no U  IS is  rf  o o o LO o O O  iCO -  CM CO 00 O CO CM O i - O  o  00 CM CM OO O O CO  O O  rf  rf  c\i d  c\i d  i-  CO oo CO. . O CM CO LO O CD Q ^ CO T- °  •i rf  CM  O O O CM  rf  rf O  S  CO  rJoSSOO  T-  O  d 6  rf  O O  d d  r  fs  "  5 <§  rf  O O  oo o  rf  CO O CD CO -r-  cj d  CO  CO  rf  LO CD LO  d d  Is T- | s O O CO  J  co  cn  oo K-  00 O O CO CM CM ' CO i —  ^% 00 co  M  r?^  O  co  CO LO  „  co CO  rf •  CO «.  i s CO fv,  n  LO  T-  CM  rf  CO  0  T T^- S co CM  M  LO , N  rf  co co  CO  CO CM O O  00 00 i — CO iCM  CO CO CO o  O  d  d d  LO I - CO T s  rf  co q rf q 01 d T - d d 0  1-  O  1  LO CM O  CJ  m  CO  i -  CO CO  rf „ rf g5  i -  CM ^ :  <° CO  rf rf O  CM  CM LO CO T - CO  d  i r  r  d d  i i -  LO CM CO T-  CO  ^  LO CM O CO CM  d  o  d  CD  o c  co oo oo o co  CM co CO . LO K J O Tf ^ CO  co  VJ  O O O CD ca O O CM  i -  °  oo 5 0)  rs  CM C OO O i -  oo o  rf O  O O  °  in  l°>  CM wCD  rf  CO  ^ ° CO  a. £ £  d  rf  T-  00 O  CM  co o LO q 01 d r - d d 0  CD  =3 CM  O  CD 2  CO  ^  O ^  O O  O O  CD CO  2  O  O cu  O ) CO 2  O  c 2  CO  oCD  OO _^_ CD CO " O CM CO  ? 8  co  IV. L O ^"  co i LO  IV  co rf  CO 1  W  CM CM  £ g LO CM  CO T-  rCM  co CO  d CO  C\i CM  co ^ ^  LO O  rf i-  CM  O  T— TT-  O  O  |v rf CD T - 1— O CO 00 LO 1 CM  d  d  d  d  CD |v O O  CD LO O O CO o  CM  O  T-  O  CD L O CO CM  00  T—  CO O  LO CM  £  00 O  CM d  CO  co co  LO rf  co rf  ^~  LO  LO  rf  o  IV. d rf  LO  <°.  T—  C\J  °  O CO  FFL  0  ^ O T-  ^  CD «  <°  oo  c\i CM  '  0  CD  iv.  CM  rf  rf  T—  Lo  rf  CO  oi  OJ  CO  ^  O  r i  1—  CM CO  LO * |v CO  CD O  CM rf  CO  CM  O  rf rf LO IV rf rf T— 1  O)  °> CO  rf  CO CM  O  O  ri  O  LO O  00 CM  O T-  rf CM  O  CO i -  d  -i-  i-  O  O  00 00 O O  O  T-  CD CD  O O  CM  O  T-  O  O  CO CM O O  rf  LO CM CM T -  CO  cd  cg  d  d  O  co  rf  CO CM  C  O  LQ CM  CO CD CD d CM CM  rf  d  d  O CO rf  CM  CD CD rf CM IV CD  RO  . 02 rf  cd  d  r-~ J ^  co CM CD i OJ o o CO o LO r f LO IV d CO 1— r f  rf  CM d  ^  001  O  O  CM  O  ^  <°. CM CM  c? CO T—  O •  r  co  °  CN  d  CD  rf  O)  CD CO i - rf O O i - t- CO O CM  O  T-  rf  CO  rf  CM  O  T-  O  O  N M  O O  CO O  O  O  O  CD  °  co ri  ^ CD CM  CM  CO  d  CO  CD  o  CN rf  T-  IV. CD  _;  LO O  |v rf  CM O  CD CO  cj  d  r  d  CM q d  „  CD O iO CO i -  CO  £  c\i  d  d  CM  rf  O  |v CD r f i CM O LO O  c\i  d  T -  T—  rf  d  1q d  d  CD O  c  CD  |v CO iv LO O) CM 00 |v CO m  LO rf  T -  tv CO CM  • ^  rf  rf  |v 00  q  LO  CD  co  q  CM  d  d  d  d  CO  < ° CM. CD  LO  §  |v  CO  CM  CM O "  o  |v  o  O  i-  LO  o  d  d  , <]>  La T3  O 9 ? O cn CO C O 5 o CM  o 0  LL  +  O  CD CO 2  O  r- CM  £  CP  O O CD CO 2  O  O  OJ c0  r-  =  +  CM CD  CD 3  _o  CO  O  ^  o o 5 co - °°. co  o  CNJ  CO ts CM fs 05 CM m CM O CO CM CO  rf  oo  _ rf  CO LO  °  co ^  is  n l  cj  CO ^  . oo co d 0  oo  CM  6  r -  T-  LO  00 LO 00 O T-  00  o  o  is co rs co is  ^  CM  • O  LO OO  CM CO  00 CO  CO O T  1  co co f s  ^  CM CO  LO  co co |s CO CM fs  O  CO CM  d d  T-  i2r  CM CM  CO rf i - CM d d d  CO LO  d  d d  CD 00 T OO O T -  T-  T-  r  1  r  LO rf CM CO  d d ci  r  Is T— LO fs O) O i - CM LO d d d T00  CO CO CM CM 00 00 CO CM f s CO ^  oo •  1  CO CD CM fs 00 CO O i - CM LO  CO o 00 CO co • f,; rn' CO  rf  T-  rf  f ^ % 5 ih o o o O oi m c  1  CO CM T - 00 O i - CM LO  g  LO  T-  fs fs CO o O d  o  CM O  00 00 OO O  rf  ri cd  d  LO  CM.  • o co  co LO co co  i-  io o  „ « co  f- LO T - CO fs  00 CO  CO  « S co  LO  CO  '. CO  CO CM LO rf  d d  rf CO fs co o i - co i T— d o' o"  fs  •  m  0 0  o  CO CO O O0 O i r  d  d  %  + CD CO  o ^ o  =  CD  201  A p p e n d i x B : C l a y m i n e r a l sample preparation for X R D B . l Clay minera] separation The following method used to separate clay minerals. The procedure was outlined by Marumo (1992, as cited in Roth, 1993). Materials Required: 1. 2. 3. 4. 5. 6. 7. 8.  1 L glass beaker 500 ml glass beaker 4 to 8 plastic centrifuge tubes a centrifuge distilled water a long stirring rod sodium metaphosphate ( N a P 0 ) small beaker 3  Method: 1. Gently crush or crumble the sample and place in the 1 litre beaker. Ensure that the sample is not ground too finely, as coarser minerals might be crushed into the clay size fraction and w i l l not separate. Also, excessive grinding might damage clay minerals. 2. A d d approximately 100 ml of distilled water to the beaker to create a slurry and then progressively fill the beaker to about 1 litre with distilled water. 3. Let the sample settle for 10 to 15 minutes and observe: (a) If most of the clay material is still suspended, proceed to Step 5 below. (b) If the clay has flocculated and settled to the bottom, proceed with Step 4. 4. Proceed with this step only i f necessary, to minimize potential contamination and exchange of N a with the clays. (a) Prepare a small amount of dilute sodium metaphosphate solution by dissolving about A teaspoon of sodium metaphosphate in about 200 m l of distilled water. l  (b) A d d about 5 to 10 ml of the solution to your clay solution and stir. (c) A l l o w the sample to settle again for 10 to 15 minutes and observe as in Step 3. (d) If the clays continue to flocculate, repeat (b) and (c). Repeat this process until the clays remain relatively suspended after 15 minutes, then proceed to Step 5. 5. Let the clay solution settle for 3 hours.  202  6. Carefully pour off the top 1/3 of the settled solution into the clean 500 m l beaker. This fraction should contain only suspended clays. 7. Refill the 1 litre beaker with distilled water, stir and allow the sample to settle again for the next run. 8. Measure about 30 ml of the solution from the 500 ml beaker into each of the centrifuge tubes. Ensure that the centrifuge is precisely balanced (each tube should be weighed to ensure equal weight distribution in the centrifuge). 9. Centrifuge the sample for about 15 minutes at about 1,800 rpm to draw the clays to the bottom of the tubes. 10. Carefully pour off the excess fluid and examine the settled clay in the tube. If it appears that minerals other than clays are being separated (i.e. pyrite w i l l leave a thin black layer in the clay extract), then the settling time must be increased. 11. Repeat steps 8 to 10 until all the solution from the 500 ml beaker has been centrifuged. 12. Repeat steps 5 to 11 until enough material has collected in the centrifuge tubes for X R D analysis. Depending on the amount of clay extracted during each centrifuge sessions, this may take up to a week or more. 13. D r y the sample by scooping the wet, pasty clay out of the centrifuge and allowing it to air-dry in a clean glass dish or on weighing paper.  B.2 Preparation of oriented samples for X R D analysis Oriented samples of clay mineral separates for X-ray diffraction were prepared according to standard methods outlined in the Clay Identification Manual, complied by Dr. L . A . Groat, D r . W . C . Barnes and M r s . S. Horsky. The procedure is summarized below: 1. The clay separate was dispersed in a small beaker of distilled water. 2. The mixture was pipetted onto glass slides and allowed to air-dry overnight.  203  Appendix C: Geochemical data from selected diamond drill hole intercepts  o LO cn d r r  LO  CO  d  CO  47700 3900 3280  8430  8290 3920  CM  1790 10200  CD CM  1395  O c n  LO  CO  2120  r r  LO  co  d  LO  CO  o co  00  r r  |v  o  00  r r  CO  |v  d  LO  o o  LO  d  LO CO  o o  O J  O J  LO  CO CO  o o  O J  LO  1430  1420  o  d  r r  17300  LO  25000  1195  d  1655  |v  r r  r r |v  O J O J  d  O J 00 O J  LQ  o r r  o  CD  d  r r 00  LO  CD CO  o  r-  CM  LO  o  co  LO  o  CO  co d  4550  CO  5910  |v  o  CD  CO  1710  in  o  2030  cn  CD CO  d  r r  |v  CD  CM  co  CO CM  o  rv  r r  CM CM  00  LO  CM  o  CD  CM  co  CM  rf  CD  LO  CD  r r 10  CD IV  r r  r~-  O J  o  CO  o  co  r r  1450  r r  LO  00  CO LO  2250  CO  o o  CO  LO  d  1070  LO  CO  O J  LO  1030  41400  o  CO  CM  d  1180  co  CO  o  CM  rf  2970  rr o cn r r  O J  1225  1230 O O J  CO  22300  0.46  r r  14100  0.53  CO CD  IV  CM  49800  IV  CO  o  o  r r  rf  O J  LO  o  r r  rv  co d  LO  CO  |v  LO r r  CO CM  r r  r r  CD CD O J  d  CO  r r  LO  LO  oo  ppm  ppm  ppm  ppm  ppm  z  Q.  Sr  H  >  Zn  CO  Sc  CD  LO  ppm  C O  o  IO  r r  2420  LO  2  CM  o  CD  LO  1445  0.01 0.01 0.03  o  CD  LO  CO  LO CD  LO CO  co  r r  1030  00  o  00  o LO  LO  d  rv  CD IO  |v  d  1560  CM  r-  o  LO  CO  2470  |v  o  1495  co co o  cn o CD co  LO  1245  0.01 0.005 0.005 0.005 0.005 0.005 0.01 0.01 0.03 0.02  o  0.01  IV  r r |v  O J  ppm  I  LO CD  O  CM  Pb  O J  LO  0.28  0.08 0.03 0.02 0.02 0.05 0.04 0.09 0.04 0.03  0.06  0.13  0.08  3.24  71000 48400 15800  3.05 6.92  27200 3770  3.37 3.76  8310 20200  2.88 2.28  22100 28400  4.42 5.13  7960  2140  o  O  LO  CM  O CD CO  |v  Sb  ppm  O |v O J  o  LO CD  |v  CO  r--  o  ppm  ppm  Cu  LO CD  LO  O J  r-  ppm  ppm  o  LO CO  O  o LO cn o rf o CO co d  CO  ppm  ~s  Cr  CD  Co  |v  CO  Cd  rf  LO  O  CD CM  o r r co  Mo  co  o  O CO  tv  3580  CO CM  LO  CD  0.26  CO  o  LO  o  18400  o  o  o  0.59  CO  o  O J  co  1975  |v  o  LO  CO  LO  0.75  co  d  CD  ppm  co  LO  o  r r C J  Mn  CO CO  0.11  CO  LO  0.16  LO  d  0.08  C J O J  d  0.14  00  LO  0.11  CO CM  d  0.16  CO  LO  0.13  r r  CD  d  0.13  co  O J  IO  ppm  CO  d  La  CD CM  d  6.27  O J rf  LO  d  25900  CO  LO  d  7.05  O J  d  22800  |v rf  O J  6.01  CO  LO  LO  d  3.55  LO  d  LO  cd  25400 100000  CO CO  CO  1.36  CO  3020  CO  2.39  O J  O  3780  ^;  2.32  CD  CM  1800  rf  3.15  CO  CO CM  1990  O J  LO  1.74  rf  1.13  CO  LO  ppb  CO  Fe  0.35  14.5  0.11 0.11 0.21 0.16 0.17 0.33 0.33 0.35 0.35 0.55  O J  rf  ppm  LO  CM  0.25  0.27  0.44 10.25 10.6  0.25  As  14.35  ppm  C O  o co  0.25  ppm  < <  CD LO  0.25  0.23 0.08  ppm  11.85  0.25  3490  2.68 2.75 0.122 0.070  E  O  co  O J  LO  LO  CM  LO LO  o  rf  rf  O J  o  Au  680.12  o  LO  rf  o o  LO  CD  ci  CO  O J  CD  co  00  ci  o  O J  o  rf  CD  CD  LO  |v O J  LO  LT)  d  SO  From  LO  To  r-  CO 00  CO  rf  |v  CM CO  ppm  CD  rf  00  rf rf  Ba  00  CM  d  O O  cn  Ca  CO  0.32  CO  0.64  tv 00  CD  CO  1260  CO  00  C J  0.32  CD  CO  C O  rf  0.26  CO  O J  0.33  CD CD  CD  2.25  690.4 691.15 o  CO  co  2.47  CD  00  |v r r  0.35  CD CD  0.36  691.15 690.37 o  CM CD CO  29.4  0.16  CO  0.29  43.4 85.6  CO  4.21  CO CD CD  C O  23300  LO  d  CM CO  |v  2690  45.8  0.08  o  CD CD  O J  11200 rf  r r  CD CD  6190  LO  CO  O J  13300 16600 4940 6380 2990 23800 23100 30100 19500 29300  o  r~-  cn  CD  5810  0.06  27.8  1.79  68.6  0.06 0.14  rf  LO  rf  CO  |v  LO  OS  LO  r r Oi  rf  C O  LO  o  SO  o  CO CD CO  o>  CO  o  LO CD  CO  CO  o  CD CD  CD  LO CM  d  o  CD CO  co  O J  o co  d  3.11  |v  rf  o  CM  2.27  co  25.6  CO O) CO  0.12  rCD  o  O J  o  11.6  CD CD CD  CD  o  rf  C J  o  0.24  C O C n  O |v  17.6  O  00  O  O J rf  O J  OS  CD CD CD  E  Hole  |v  17.2  O  |v  0.36  o o  CM  CM  0.29  O J O |v  o  o o  0.39  iv  o  o  0.21  |v  o  8.82  CO  r-~  o  1.96  O J  0.81  O |v  O J  0.33  IV  0.56  o  1.25  CO  1.45  IV  O  SO  DDH59 DDH59 DDH59 DDH59 DDH59 DDH59 DDH59 DDH59 DDH59 DDH59 DDH59 DDH59 DDH59 DDH59 DDH59 DDH59 DDH59 DDH59 DDH59 DDH59  LO  iv  o  SO  DDH59  r r  CO  OJ CM  o  CO  CD OJ h-  CO IO  CD  CD  o  co  3830  o o  rf  CO  rf  o  CD  CO  3280  o  CO  o  in  is  o  OJ  co  OJ CO OJ  m d  rs  CD  rf  CO  LO  LO  LO  co  CO CD  CO  | s  CD CM  o rr  CD  o  CO  co  rf  12600  | s  CD  CO  m  in  O  o  d  o  in  CD CD  o  CO CM  O  O 00  CM  CD  CM  CD OJ  o  CD  CD  O  CD  co co  in  o  o  CD  d  LO  CM  rf  in  CO  ppm  ppm  ppm  ppm  ppm  iz  Q.  Sb  Sc  CO CM  2480  CO  1430  o  5520  CM LO  CD  ppm  CM  ppm  o co o  CO rs  Pb  OJ  ppm  CO CD  ppm  00  d  Mo  CD  in  CM  o  ppm  CD  CO  CO  CM  CD  in  m o d  CD  CM  CO  CM  o> CO  d  OJ  OJ  m co d rf  w  H  >  14700  o  o  OJ  o  05  49400  00  rf  3350  1200 1020  OJ  Mn  o>  CM CM  CD LO  CD  ppm  ^9  2  rf  CO  CD  CM  1  1370  0.69 0.24 1.06  1080  0.02  5190 LO | s  LO  0.61  22400  0.16  5060  0.02  CD  d  m d  o>  o  o  CO  LO  1100  0.61 0.19  0.99  0.32 0.24 0.08 0.08 0.07 0.14  0.06  0.14  0.97  CO  rf  CD  CD CD  CO OJ  CD  , _  00  1920 o  d  00  2640  CO  1210 1180 1260  0.42  in  CO  m d  ^9  2700 o  0.57  4320 5450 3300 101500 54000 189500 16500  0.88  1535  0.69 0.77  1115  1.62  0.16  00  CD  0.09  1220 1.15  3490  o co  rf  0.08  ppm  in  rs s  o  co  CO  |s  co d  o  0.08  ppm Co  X  in  CO  rs  rs  O  00  d  rf  in  d  o  CD  rf  CO  d  CM  o  O  in  in  rf  O  CD  CM  r-  rs  co  LO CO  o o  in  o  CO  in  1  CM OJ  rf  rr  in  o  0.08  ppm Cd  o  ^9  o  o co  0.11  CO  ppb  CD CM  ppm  co  Fe  o o  Cu  CD  d  OJ  La  Ba  Ca  LO  CO  1.53  < <  co  7160  ppm  ^9  CM  rf  rf  0.12 0.17 OJ  o  rs CM  in  CD  in  ppm  o  d  CM  rs  in  1075  1.27 2.36 2.53 3.08  6950 23100  3.04 3.42 2.49  7930  1.58 CO  3.04  3580 7110 20200  rf  o  CM CO  31100  CM  o  o  3.79  co  o  o  66100  d  d  co  CD  rf  o  LO  o o  CM  m d d  o o  in rf  CM  CD  CD  m  rf  in  d  rf  d  1.29  |s  CM  CM  ^9  0.41  CM  o o  CO  o  1200  05  o  0.21  rf  CO  1.41  rs  CM  co d  CO  o>  o C D O co L O co d co o  d  in  O  OJ  o  2.53  r--  o o  CM  d  o  ppm  CM  LO  CM CO  d  1.52  0.25 0.25  LO  0.25  LO  11.65  LO  4.22  rf  o  LO  rs  CO  iri  CM CD  5310  0.25 0.25  5.71  0.25  3.78 0.97  0.25 42.5 0.25  0.57 3.22  CM  CD  d  As  To  rf  LO  Ol  14.2  £  d  CD  ppm  723.75  E  LO  o co  CO  rf  2.99  co o o r- r - -  From  s  00  co  3.87  rs  o  o  LO  rf  ppm  | s  o  co  o  LO  CO  O 00  Au  723.75 720.35  CD  is  CO CO  CO  1.23  rf  r~-  |s  co  0.25  rr d  CD  O rf rf  co L O co d CM  0.25  co d  o  O  CM  0.25  d  LO  CM  oo cp  CO CO  0.25  d  CO  00  o  00  0.25  0.42 0.72 0.85  1155  0.53 0.18  0.21  1490 1780  0.35  2070  0.15 0.16 0.27  1000  0.26 0.32  CM CO | s  0.18  d  0.37  1615  0.53 3.93 0.34 0.18  0.36  281.08 279.9 278.89 277.67  4.030 4.630 0.210 0.04  728.5  0.04 0.07  CM rs CD  o  rs  LO  d  is  o  LO  CO  CO  LO  o co  o  00  o  CM CO LO  O  | s  LO  eg  o CO co  LO  CD  CO  CM  CM  o  o  CM  CO  CD  CO  o  o  co  d  CD  o  CD  rf  CO CM |s  CM CM r~  CO  o  o  CM CM  O I  o co  d  co  r-  o  d  h-  CM  rr  d  co  CM |s  CM  CD CM  LO  0.02  LO CO is  CM  2.330  741.42 rs  CO  2.490  rf  728.5  r--  rs  o  0.98  CO CO  in  cp  o rr  CD  0.56  740.1  5 is  745.6  279.9  DDH73  278.89  DDH73  277.67  DDH73  276.65  DDH73 DDH72 DDH72 DDH72 DDH72 DDH72 DDH72 DDH72 DDH72 DDH59 DDH59 DDH59 DDH59 DDH59 DDH72 DDH72 Hole  r-~  d  LO  so  rf rf  d  LO  co  so  LO  rf  LO CD  o  rs  OS  d  rf  OS  d  E E  ppm  ppm  ppm  < <  As  9.95  1.41  0.27  d  Au LO  rv  d  rr  282.85  o  CM CD CO  O CO  o  O  LO  LO  CO  ~S  ppm  ppm  ppm  Co  o  Cu  CM LO  51.5 rr rv  d  CD CO  rr  cq d CD  co  CO  o  o  rr rr  co  CM LO  CD  LO  cq o rr co  X  CO LO  o  o  CO  0 ^  CM  d  co o rv cn  ppm  ppm  ppm  z  o_  Sc  CO  LO  ppm  2  co r r  Pb  d CO  Sb  0.78 161500  LO  ppm  LO  ppm  CM CM  ppm  o LO CD  o rr cn LO CM  LO  CO  rr  LO  o  LO  LO LO |v  00  o CM cn CO CO  w H  CM rv  LO CO  o  CD LO CM  o oo oo |v  CM CD  o rr  CM CM  o  o LO  rr  CO  co o  rv  2550  CM CM  o  o  o o o  o co 14200  rv  o  CD CO CO  CD CO  CO  o  CD  rr  20400  d CD  r-  o  r-  CO  ppm  co CM  ppm  CD  >  Zn  r-  ppm  rv  1005  h-  2080  1860  LO  1150  CD  3210  d  CD  rr  2030  CD  3200  d  1940  CO  1790  d  1490  LO  1270 rr  4150  CO CO  2610  LO  CM  1390  d  1160  1990  rv  rr  rr  00  o CM  CM  CD  o co LO o co r r 00  CO  rr  co  rr  CD  CD  CO  rr  CO CM  o  CD CO  o  rr LO  CO  LO  LO  LO  00  CM  CO CM  LO  d  o co  CM  rr  rv  co  rr  o o  CM LO  CO  CO  CO  o  CO  co  d  LO  LO  rr  o CD 00  LO  o CD CO  LO CM CM CM  CO CM  CD CM  CO  CO CM CM CM  CO LO  o o O  CD  o LO  rr  o LO CO  o o rr  o CO  o  CM CD  O  rv  rT  oo  CM  rr  CD CM  o  o  CO  rv  rr  o  o  LO  rv  CM rv  16500  rv  LO CM  o  d  1390  0.18 49500  0.17 2260 rv  0.26  0.28 8700  4970 LO  2860  0.21  LO  rr  2310  0.29 1540  0.72  LO  3220  rr  14200  8230 CO  0.09  CO  0.15  co cq  0.06  1.93  CM  0.04  o co d  o  Mo  o  d  O  0.03  0.29  0.22 0.13  0.39  3.09 5410 0.13  1.55 6620  1.59 0.32  4.95 1980  2.78  0.08  3.95  5.86 40400 d  1180  2.68 4740 O  0.04  0.09  0.07  0.21  0.31  5.73  CM  42400  3.45 48300  LO  5.36  rr  o  73900  O  2110  CO CO  0.55  CD CM  O CO  CM  95700  CO CO CO  O  CM  4340  co  o  ppm  CM CM  O  Mn  LO  0.13  rCD  0.25  rT rr  0.08  O  O  00  0.19  CM  rr  o  0.81  O  rr  100000  CM CO  o o co  0.17  LO  rr  O  4.12  O CM  CD  3170  CO CD  4.96  LO  LO  9880  LO  rr  5.84  CO CO  6450  co  4.44  0.25  0.25 0.25  14.1 0.25  CM CM  6170  0.25  0.25  13.8 0.25  0.25 CO  0.25  d  rr  6.61  CD CM  0.25  CO CM  0.25  CO ID  LO  0.25  LO CM  rr  7590  9.11  3.01 31.5  LO  O CM  ppm  CM  ip  0.63  LO  LO  3.64  o  r-  CM  8420  rr CO  2.97  O  CM  7.43  co d CO  2.95  O  CM  0.87  O  rr  1.03  O CD  CD  0.25  O LO  rv  ppm  1.02 13500  0.94 1575  0.47 6890  1.54 43000  3730 O CD  15900  0.66 1010  11900 O  17600  1.53 20400  8950 O CM  1455  0.17 15100 o  1025  0.27 o  1330  0.24 o  0.33  0.18  5.94  o  LO  CO  La  CO  2.93  CO CM  d  o co  CO  CM LO  ppb  CM  oo d  o  LO  Fe  rr  d  Cd  rr  ppm  d  LO  0.79  CD  Ca  d  LO  Ba  co  rr  0.55  CD  LO  0.42  d  rr  0.67  LO CM  0.57  o rr rr  1325  LO  d  LO  d  2.29  545.72  0.09  0.09  0.01  545.15  544.4 545.15  544.4  543.3 543.43  d  d  0.43  0.38  2.15  1.78  1.29  LO  0.49  50.01  LO  0.01  LO  542.35  o co r r  1.89  rT  542.9  d  287.9  DDH86 DDH86  543.43  d  542.9 542.12 542.35  d  543.3  DDH83 DDH83 DDH83 DDH83 DDH83 DDH86 DDH86 DDH86 DDH86  d  286.92  DDH83  d  537.9  DDH73  d  286.92  286.26  DDH73 CM  To  286.26  285.65  DDH73  285.65  284.98  DDH73  CO  284.98  DDH73  CD LO  282.85  DDH73  rr  281.08  DDH73  d  From  LO  LO  00669  Hole  206  CD  CO  co  |v  LO  LO  CM CM  CO CD  o  o  CO rs  rf  CM  LO  rs  d  1820  oo co o  co o  CM  CM  00  rf  CM  rf  o  O  CM  CM CO  CM  CM  CD  CO CO  m in o  CM  rf rf  rf CD  CM  CM rf  CD  CM  CM CM CM  m rr  CM rf CM  CO  CM 00  o  CO  CO  CO CO  o m  rr  LO  in  o rr CD r f co  rr rr O rf  CO  o  CO  CD CO  O CM  CO rf rf  CO  O  CO  o  00  co o  co  o  CO CO  is  CO  o  rf 00  CM  CO  rs  CO  LO  o  CM  LO  m  00 CO  rf rf rf  o  CO  rr m  0-  in  o  o  in  rr  CD  O)  CM  o  o o rf co CM rr  d  CM CM  o  rf  in  o  o  co  I—  >  rf  in  O  in  LO  rr in co  CD  ppm  ppm  ppm Mn  Mo  ppm La  Fe  Ca  Hole  |  |s  co CO in  CNI  o co rr  CD CM  2050  co  co  1240  o  2680  1010  rf  ppm  o in co d co  CM  CNI CO  ppm  rs  CM  Zn  CD  |s  co o o  CO  in  LO  o  CM  rf  co  rf rf  CD  d  rf  r^  CO rs CO  m o  LO  o  CO  o  1320  oo  CO  2850  rs  CD  CO  ppm  in  o  ppm  rs  00  CM CO  ppm  CD  o  CM  in  ppm  m d  d  co  Sc  in  rf  Sb  2440  co o 00 oo  4630  0.19  rf  59900  0.35  CO CD  61900  0.34  rr o  d  52500  0.09  in  LO  in  14900  0.06  in  CD CO  1080  0.46  1695  0.45  2790  d  0.23  0.25  CNJ  o  o  3280  rf  O) rf  4190  co  CM  rf  CM 00  ppm  0.47  1445 1865  0.24  2040 2310  00  3220  0.25  o  CNJ  ppm  ppm  rf  CO CD  Pb  0.25  ppm  o  O  co  128000  LO  CM  0.07  o  O rf CD  42100  CO CO CM  CD  CO  0.14  o  CM  d  0.02  CM  CM  1290 227000  co  CO  d  o  rf  d  0.005  o  CM  LO  0.16  1.66  1060  1.47 1.32  1180  1.94 2.57  27200 61300 2210  1.42  41800  1.42 1.75  00  LO  LO  0.07 0.09 0.07 0.17 0.12 0.14 0.11 0.12 0.13 0.19  O) X  1.36  ^2  1090  LO  co  co  ppm  CNJ  rf  6200  o  o LO rr  m  2.45  o  CO  co  o co  2.42  in  co co  CO  6360  0.25 0.25  cd O CO  0.25  m  0.25 LO  rf  0.25  CNJ  0.25  |s  Co  CO  -  Cd  ppm Ba  0.61 0.29 0.38 0.68 0.34 0.19 0.46  o  in  CNI CO CNI  1.02  ppm As  o  LO  LO  0.72  63200 424000 147000  ppm  CNJ  LO  CO  1.27  14500 52500 90700  2.13 ppm Au  < <  0.35  4040 20200 37700  0.14 1.89 1.56 2.44  546.45 546.1 E  CO  0.09  CO  rf  0.12  O  co d  in  0.11  rf  LO  0.01  CD IO  LO  0.005  CO  CNJ  o  0.005  co  o  0.11  CO CM  is  o  in  d  3160  1.39 1.08  CO  0.53  1.28  co  rf  o  CO  d  2820  1.19  5  E  1.35  2320  2.41  in  CD  0.09  1.14  CD CM  |s  549.85  1390  12.95  co in d  CO  o  ppb  d  |s  o  ppm  iq  CNJ  o  Cu  LO  rr  CNJ rf  549.55 549.85  0.25  o in  0.25  O CNJ CO  o  548.6 549.05  13.5 in  10  CD  CD  548.6 549.05 549.55  0.25  o  CNI  o  547.95  0.25  CO  CO CM  547.95  14.05 LO  545.72  551.98  o m  CNJ  o  547.11  CO  CNJ  CD  547.11  rr  o m rr  o  546.45  rr  O LO LO  CO  546.1  O CO  d  |s  d  rf  o  C\i  CO  o rr  CNJ  o  d  CNI CNI  o  rf CNJ CNJ  o  o  CO  co  rf  O  o  CD  0.27  d  CNJ  CO  0.37  d  m  CNJ  0.76  LO LO  o  rf  CO  0.42  d  in  CO CO  co  0.34  d  LO LO  o  CNI  0.35  d  LO LO  d  From  550.3  LO LO  d  CM  14.2  co  0.08  rr  0.01  CO LO LO  0.01  LO LO LO  in  0.01  rT LO  d  0.01  CO LO LO  0.02  LO LO LO  d  0.13  rs LO LO  1.19  co o o NJ d Crr  550.3  d  0.08  CO LO LO  d  1.35  CO LO LO  LO  d  0.25  is LO LO  O O CO  0.05  OO LO LO  551.98  CO LO LO  To  DDH86  DDH86 DDH86 DDH86 DDH86 DDH86 DDH86 DDH86 DDH86 DDH86 DDH86 DDH86 DDH86 DDH86 DDH86 DDH86 DDH86 DDH86  207  562.65 563.55 3.53 cq d  560.5 562.35  560.5 562.35 562.65 3.19 d  7.91  0.25  E  ppm  ppm  ^?  ppm  < <  CO  Ba  Ca d  rf CO  d  o rO)  00  co  CM  rf  rf  CM  oo  o  CO  o  CO  LO CM  co  OJ CJ  o  |v  o  CD  CD  CO rf  rf CO CM  rf CO rf  cn CD  rf  CO  rf CO CO  co  CO  cn  CM CO  LO  LO  LO  LO  cf-  •-8  X  CO  2 CO  LO  LO  CO  cn  CO  LO  ppm  ppm  ppm  ppm  ppm  ppm  ppm  ppm  Mo  2  0.  Pb  Sb  Sc  Sr  1.22 16299  0.54 54200  14932 35700  15947  8025  CO  o  1705  CO rf  8930  d cn CO  4380  CM  rf  4820  d  o CM  cn  co cn m  CO  o CO iv  m CO  LO  cn CD CD  cn  d  o CO CD  o  rf O r-  cn  cn CO  LO  LO  d  LO  h-  d LO  LO  LO  CM  on rf  LO r-.  co o  OJ T  CD OJ iv  o CO CO OJ  LO CJ  CO  »o  rf rv  O  rf  CD  CO  CO CO  cn o co  CM  o  r~  "  CM  o  CM CM  in LO  CM  cn  '  in in  o co cn  CM OJ iv 00 rf CM  o  00 OJ  co  rv CO OJ  co  CD  cn cn CO lO CM  oi  CM  iq  CO  CD  m oj CO CO  rf  cn  CD  o in  co m CM  CD CM  CM CO  co co o o CM rf  o  d  OJ CM CO  CO CJ CO  CO  CO  LO  T  rf  CO  CM  co  09  LO  o co co  CO  CO rv CM  ppm  0.45 257400 119400 149800 o  |v  >  Zn  0.63 192600  iq rv  ppm  0.54 115400  LO  1640  1.49 54000  d  1760  18951  0.15  0.18  0.04  d  14300  5.06 5987 0.15  CD  7400  3.17 4693  d  ppm  3.92 6086  CO  Mn  1.37  CM CO  d  1.05  rf CO  0.97  0.16  0.23  0.58  0.23 o  2.28  0.42  4.03 7395  CO OJ  1.92  0.59  co cn d OJ  0.36  4.58 3978  o  0.11  0.11  3.94 2783  cn  0.17  0.07 LO  0.14  4.34  LO OJ  4305  CO  0.18  0.06 LO  0.08  co  0.05  cn  CO  4607  4.33  00  7.03  rv  4966  o  8.52  CO  7123  CO  5.82  LO  2350  co cn  1103  co  1.04  CO rv  d  1370  LO  0.97  rv LO  4570  2.42 2.15  4.89  co d  rf  OJ  2.07  00  LO  LO  38200  CO  ppm  d LO  La  CO rf  d  LO  3.24  CO CD  co  30600  d co  rf CJ  2.31  3.41  6.48  5.85 d  5.34  0.47 1847  0.56 1285  0.29 2514  0.59 3885  0.42 2587  0.61 3893  d  co  o  1.68  <N  d  CM  CO  18000  CO  d  LO  ppb  rf  rf  d  ppm  o  d  CM  Fe  ppm  6.55  rv  d  CJ  Cu  ppm  4.62  CD  6.48  o  1.31  LO  0.25  CO  0.25  1.45 4521 o  0.25  4139 ~  0.25  1905 1  0.25  1584 cn  9.61 d  1787  1765 co  1417  o  d  0666  CM  Cr  CD  d  Co  d CD  Cd  i—  LO  1.71  d  CJ  1.78  d  CD  5.26  d 0.98  d 1.09  d 1.47  d  0.84  d  0.88  d  0.15  cd d  2840  d  0.12  d  11800  573.35  572.9  572.45  571.9  571.5 03  2.98  570.55  572.9  DDH91  569.85 d  0.35  1.79  CO  d  0.25 1700  9800  11100  o co  As  569.5  568.7  568.3  567.5 CO  567.2  0.06  LO rr'  566.7  565.8  572.45  DDH91  571.9  DDH91  571.5  DDH91  570.9  DDH91  569.85  DDH91  1  569.5  DDH91  LO CO  d  0.21  6570  d  ppm  0.51  564.7  563.6  567.5  rv  cri d  0.16  1.31  o  0.59  £  d  ppm  in in  o  Au  OJ O rv  To  From  568.25 CM rr'  567.2  568.7  DDH91  CO CO-  566.7  566.2  564.7  DDH86 DDH86 DDH86 DDH86 DDH86 DDH86 DDH91 DDH91 DDH91 DDH91 DDH91  CO  0069  Hole  208  rf rv rf  O CM  cn  in CM  m cn CO CM rf  rv  m CD  o  CO CM  in  CM  cn o  rv  in  o  rv rf  in  r~  CM  co  CO  o  rv CO  o  CD CO  o  rf  -  CM  -  rf O  *-  LO  CM  rf  CM  o  CM  CM  OJ  CO  CD CM  CM  co  cn co  in  0.07 1.75  1403  LO  LO  CO  co  rs  3930  2.79 1.77  6937  2.73  4475  2.21  1618 3017  3.05  uo oi  1219  1.69  LO CM  CO  CD  CO  rf  io oi  00  m o  rf  uo m CM oi  LO  CD  CO  CO  UO  LO  CO  |s  CD  co  LO CM  LO  CM CO  o  rf  LO  CM OJ OJ  m  CD  LO  CO  CD  LO CO is  LO  rr  LO CM  LO  oi  OJ CO CM  00  00  rf  LO CM  LO CM  rr co  CD  00  co  LO  oi  CM OJ CM  00  o  LO  rf CD  m  CM CM  CD  CO  CD  CO  |S  rf OJ  oi  OJ in rr co oi  LO  d  co  CD  CD rs  00  |s  LO  CD  OJ rs  rf  o  uo d  |S  co co co  ppm  D)  io  ppm  iz 0.  co H  d  CD  r-  CD CD  co  LO  m  ppm  2  d  CM  CO CM  rf CO  00  CM CO rf  CD  CM CO  ppm  LO  0.64  LO  d  oi  >  Zn  d  oi  CD  ppm  CD  LO  •*9  oi oi  ppm  Fe  CO  CO  CD  Sc  ppm Cu  OJ  CD  LO uo oi oi  ppm  ppm  Co  Cr  3.51  ppm  Cd  X  IO  CO  Sb  -*9  CM  |s  IO  oi  ppm  0 ^  co LO co d  rs  Pb  CM  d  CD  CM  uo  0.54  0.07 0.08 0.06 0.13  0.17 0.13 0.08  3.53  9630  4.98  50000  1.74 2.48 2.64  CO  CO  CO  co  CD  ppm  CM  iq iq oi oi  O  ppm  o  d  rr  rs LO  rf  d rr  LO  oi oi  Mo  d  LO  CM  4.19  LO  rf  rr  rr  1347  o  d  CM  2.38  LO  00  0.92  CM  co  OJ  OJ  0.83  CD  iq rr oi CM  1896  CM  CD  LO  4354  rr  CM  1.36  CD CO  rf  rr CM  CD  CM  49800  CO  00  co  m  CO  CM  2.33  LO  d  1.63  rr cd d  00  2.23  d  d  rr d  rs  ppm  d  LO  rs  d  |s  CD  rs  CD s rf  0.33  d  oo iq d  LO  o  co  rf rs  0.47  d  CO  co  iq m oi CM  2E+05 183500 154000 164700  d  CD  CO  ppm  LO  CM  rf  r--  LO  Mn  d  CO CM  o  co  0.11  d  LO  d 0.07  LO CM  d  rf CD  0.14  LO  2.39  134.1 d  1.74  CM  0.11  Ba  Ca  0.29 0.71 0.33  LO LO  LO  0.19  As  CM  CO CO  0.23  ppm  < <  oo cd d  CM  0.06  ppm  0 ^  rr d  d  CO CM  co  0.09  ppm  ^9  CO  LO  CM  rs CM rf  co  0.09  0.32  1705  CO  d  d  CM  00  co  ppm  1912  rf  d  lO  d  8.62  13500  CD  CO  d  d  3.34  1018  CM  rs  0.31  1472 3444 8034  rr  co  co  CD CM  CD  CO  d  La  o  CM  CM  2.37  o  CM  1.84  o  co rr  o o co  3.51  CO  CD  d  d  rr co  d  CO  1.98  CO CD CM  o  CM  O  1.74  co co o co  CO  co  co  2.23  co o  LO rf  CO O CO  2.93  o  co  rf  CD  3.18  CO CO  co  CM  03  4.89  CM CO  CD LO  o  2406  o  CM  rf  CO  3.94  O) CO  rr  d  3693  CM  CM  6.05  rr  CO  o  rr  2254  CM LO  o  |s  3016  CO  rs  CM  2.62  0.37  CO  d  CO  CM  1028  0.58  1.12  h- o o co CD  0.45  0.51  d  CM rf  15.4  575.4 To  574.3 574.65 575.05 573.8 E  CO |s  ppm  N LO  LO CO |s  0.52  d  co  d  d  0.43  d  d  CD rf  0.37  0.14 0.22  0.16 0.07  is |s LO  0.56  0.12 0.005 0.04  577.8  co d  579.3  CO  CD is LO  E From  d CM  573.4  573.8  574.3 574.65 575.05  575.4  LO  d  LO |s  Au  r-  d  CD rs  0.29  583.1 582.5 581.9 580.1  582.5 580.1  581.9  rr d d  d  o  ppb  rr d  d  o  LO  CM  5.01  CO  CO  d  rr  6.84  CO CO  rr  CO  7.99  d  579.3 577.8 CO  0.43  d  CO LO  r~r-LO  1828  d  rr d d  rs LO  0.41  is  CO LO  d>  10000  0.25  d  0.41  CO CO LO  CO  1864  0.97  d  co  1639  LO  584.7  CM  CM  583.8  583.8  d  LL  583.1  CM  1392  584.7 527.33  CO  CM LO  0.72  CD  LO  SO  Hole  CO CM  so  DDH91 DDH91 DDH91 DDH91 DDH91 DDH91 DDH91 DDH91 DDH91 DDH91 DDH91 DDH91 DDH91 DDH91 DDH91 DDH91 DDH91  DDH91 DDH93 DDH93  209  CJ)  LO  co  LO  LO  LO CD  rv  rv  CD  LO  o  ID  LO  rf f-  rv  rf  rf  ID  LO  00 LO  CO  OJ  oj oj  LO  rf  o  ID  o  rf  CO rf  O  CO  LO  rf O)  co  LO rf  CO  LO  CO  LO  O) rf  LO  rv  CO  CO  LO  LO  CO LO  co  CJ)  oj oj  CO  rf  LO  LO  iv rf  CN  CO  oj oj  CD  cn  LO  LO  rf rf  CJ)  rf rf  ID  CO  CJ  LO CO  LO  CD CD  OJ  OJ OJ  r-  CJ  oi oj  1517  ID  2857  rf  3998  rf  5855  CJ  6885  OJ  OJ  LO  CO  CJ)  CD rf  co  rv  LO  O)  oj  rf  rv  LO  CO LO  00  rf  OJ  CO  o  00  rv  rf  CO rf  LO  *-  LO CO  o  CJ)  o  CD  LO  oo  CD  co CO co  LO  CN rf  OJ  o co  LO  rf  rf  co  oj  ppm  ppm  ppm  ppm  ppm  Q-  rf  oj co  LO  OJ  z  rf  oj  oj  OJ  CJ  oj  OJ  CO CM  o  CD rf  PPm  CO  CO  oj co  PPm  co  co  CJ)  co  rf  PPm  CO rf  LO  rv  PPm  0.41  2018  0.01 0.04  rv  0.03  cn  00 OJ  2  oj oj  r~-  0.03  0.34  CO  oj  Pb  OJ  oj  oj  Sb  2.97  1.42  0.86  0.63  0.59  1296  rf  2181  oj oj  0.21  *"  ppm  0)  Sr  co  Sc  1215  O)  CJ  1405  0.57  OJ  Mo  o>  I  LO  rf  oi oi  2586  <>  LO  0.13  CJ  CO  PPm  00  LO  LO  Mn  0.11 0.11 0.13 0.12 0.19 0.18 0.15 0.13 0.17 0.24 0.33  0.16  1.13 1.38 1.73  1256  1.19 4.44  4062 4373  3.99 3.21  00  2808  OJ  2630  2.33  O  2941  2.51  OJ  1471  3.52  O  0.14  PPm Cu  o>  ppm  CO  rf  0.04  o  CJ)  co  0.04  LO  rf  ppm  cq  cn  CO  0.17  CO  CO  CO  d  0.17  rf  co  CO  CD  ppm  rf  rf  rv  La  CJ) rf  OJ  CD  5.35  OJ  CD  CO  14489  rf  |v  7.78  CJ)  co  00 d  rf  12503  LO  rf  6.82  CO  3.35  rv  1.39  rf CO  o  LO  11520  rf  |v OJ  2.09  rf  LO  rf  2.87  o co  OJ  1115  rf  OJ  2.67  CO  rv  CJ)  1622  rv  co  CO rf  4.08  LO  OJ  7638  rf  OJ  Cr  00  O)  Co  ppm Ba  OJ  rf  OJ  19.7  CO  rf  o  rf  1.49  OJ  OJ  CO  19.9  o  CO  1.51  rv  OJ  3.94  CO  CO CO  OJ  rf  od  LO  CO  6872  OJ  rv OJ  00  ppb  3.65  2.86 CO  o  o  rf  d  7.82  6.95  rv  rf  |v  rf  O  co  rf  CJ  Fe  9.43 2.55 3.78 2.11  cj  0.68  0.36  1070  0.29 0.34 0.41  1.03  |v  rf  CO  < <  d  oJ  rf rf  OJ  E E  To  co  oo co  CO  LO  1549  0.37 0.48 d  LO  OJ  2051  0.51 rf  d  LO  ppm  cq  OJ  As  d  co  ppm  d  3912  0.56 0.49 0.45  1.51  LO  d  0.09  1.17  LO  ppm  o co  d  Au  O)  LO  2.38  531.87  CO LO  530.9  o  rv  OJ  0.72  LO  d  o  CO  d  O  OJ  PPm  OJ  d  rv  OJ  rf  Cd  1.38 0.96  d  ci  rf LO  co  Ca  537.22  d  533.05  536.45  iv  OJ  535.6  d  534.52  O  CO  CJ  d  3840  rf  CO LO  LO CO rf  0.55  00  rv d  CO  CO  6655  d  LO CO  0.67  LO  rf rf  oo  rf  4278  co  536.45  537.22  CO CO LO  0.95  CO LO  o  2125  1.08  d  cn  d  LO  6128  0.16 1.63  d  d  CO  o  10000  543.35  541.02  542.15  co oi  540.05  ci  o  10000  OJ  rf CO  2374  0.08  ci  540.05 541.02  0.03  542.15 543.35  ci  535.6  DDH93 DDH93 DDH93  rf LO  LO  534.52  DDH93 DDH93  LO rf LO  533.05  DDH93  rv  531.87  DDH93  CJ) rf LO  LO rf LO  530.9  DDH93 DDH93 DDH93 Hole  rf LO  From  DDH93 DDH93 DDH93 DDH93 DDH93 DDH93 DDH93  h-  0.01  210  I- > N  C  rf  0.058 I  CM OO CM CO O) CM  0.089  DDH96  0.513 d  is  CM  ppm  ppm  o  Cu  5  rr 1.32  LO CM  |s CM CO  CO CM s  00  o 00  CO  LO  LO  LO  co  LO LO O) |s  rr CM LO  o  co cn co O)  CM  rr  o  o  CM  o  o d  LO  d  CD  d  rs  d  rr CO  CO  CM  CO  CO  o  CO  CM  LO  o  CM LO CO  is  00 CM CM  O O CM  0.16  rr CM  ppb  CM  3.56  LO  4380  LO  1.46  CM  ^5  c>  2  Ul  ppm  ppm  ppm  ppm  ppm  ppm  ppm  ppm  ppm  z  Q_  Pb  Sb  Sc  CO l-  >  Zn  d  ppm  0.09 17481  2.31  3.93  15945 CM  17348  0.09 7806  15508  20000 20000 LO  LO  0.32  0.47  0.14  0.17  iq cn d  20000  0.76  3.36 4516  iq d  LO  0.22  0.56  5.31 6931  co  20000  0.22  0.11  0.09  0.18  4350  0.19  2238  CM  Mo  O CM  1.39  rf  20000  Oi CM  0.47  rf CM  16588  O CM r--  ""  0.43  CM  3236  o  CM rf  rr LO CM  5) LO CM  0.55  s  0.59  3.39 1895  rf CM  co LO oo CD  co iq iq  rs CJ) CO 1  CM  1014  6.02 3751  CO  d  CM  rs  iq iq rs  rs CM  iq iq co  cn 1  1153  0.18  5097  LO  50  Ui  O)  50  X  0.23  d  0.28  co 0.12  LO O CO  0.12  cn CO .CM  0.13  4744 0.16  CO LO  0.24  rr 4.76  CO  co  5.31  CM CD LO  6048  CO rs  5.43  LO CO is LO  5766  CM CM  in  4.96  CO  O LO  o  iq d CD  1477  CM CM  1.43  CO  co  1.63  CO CM  CM |s  1.22  CM  co  2.35  0.56  2000  00  LO CO  2.54  ~  CO  CO  cn  ppm  o  2.56  CO  2.47  CO CO CO  CM  cn  LO  Mn  CO LO T  1.22  LO  CO  CO CO  ppm  00 CO  0.86  1162  0.22 CM CO CO  co cd  CM  La  rr CO CM CM  4.18  rs CO r--  2.88  1181  0.25 o  cn  Fe  ppm  Co  As  o  0.42  o r^  ZL  ppm  < <  cr-  Cd  •-S  co  Ca  1008  rr d  1243  cn rr  CO  ppm  d d  ppm  0.32 CO 00  d  Ba  0.37  0.58  0.55  0.64  0.28  0.27 CO  0.23  CM  CM O CO  0.39  cn CM  ppm  CM  rr  0.081  en 00  0.064  O) CM  1286  rr  0.54  5.011  4.526 d  0.38  11.492  297.35  d 0.42  0.693  2.276  6.95  2.516  1.092  1.847  4.347  10.935  298.05  d 0.43  305.6 306.75  305.6  304.2  302.02  301.85  300.75  299.95  298.94  d  296.5  DDH96  d  295.35  DDH96 304.2  DDH96 302.02  DDH96 301.85  DDH96 300.75  DDH96 299.95  DDH96 298.94  DDH96  d  298.05  DDH96  d  297.35  DDH96  d  296.5  DDH96  d  295.35  DDH96  d  ppm  290.88  E  To  DDH96 CO CO CM  290.88  DDH 96  d  Au  290.05  DDH96  E  From  Hole  21  d  d  d LO  1  ~  CO  rr  co  CM  CM  CM CM  rr  o> LO  d  CD CO CM  cn is  co  LO  co  rs CM  iq CJ) CM  CO CO  rf CM  co co CO CM LO CM rf rf  o o rr  LO rf CM  CO CM CD  r--  CO  rf  CJ> CO  10  00 CO  LO  LO CM  CM  CM  CM  CM  iq  CM  iq  iq  iq  CM  CM  cn cn co  CM  CM  iq iq  CO CJ)  rr  CM  CO CO  LO CM CO  oo co LO CM  00 CO |s  rr O O  CO CD  o CM CM  CM CO CM LO CM  cn CO  rr  r» OO CO CO CO  rr  co  CM  CM  rr  CM  00  LO  co  CD  CM  CO  rr  -i—  CM  o co  CM  co  LO CM  1344 CM  CD LO  co  CO  r r  1108  CM  1485  CM  O)  CM  LO  LO  CM  CM  rf  LO  LO  CM  CM  LO  LO  CM  CM  LO  CO  rf  Is  LO  *~  rf  1293  rs  CO  1451  LO  CO  1394  r r  LO  CM  CO  CO  o  r r  2167 2809  CM  2396  co  CM  Is  co  1  co  |s  co o  d  CD  co  |s  r r  co  oo  rf  o  LO  CM  LO  CO  CO CO  IS  CO  rs  o  CM  o  r r  LO  LO  CM  CM  LO  LO  CM  CM  LO  IO  CM  CM  LO  Lp  CM  CM  LO  LO  CM  CM  CD  Lp  CM  r r CM  CM  CO IS  CT)C O  |s |s  CO  CO CD  O  rf  rf  r r  r r  CO  r r  10  CM  00  CM  o  LD  LO  CT)  CM  CM  rf CM  LO  o  CM  r r  r r  rs  LO  o  CD  CD  r r  -8  ppm  ppm  ppm  ppm  ppm  z  Q.  H  >  Zn  CM  Sr  CM  CO  ppm  LO  ~  CD  CM  CM  00  2032  1.73  1.53  CM  6614  0.63 CM  2520  0.65 0.73 1.21 CO  rs  CO  1479  1.04  15277  0.52  4398  0.74  rs  0.71  0.13 0.15  rs  CM  LO CM  Sc  Cu  O)  X  00  LO  ppm  ppm  ppm  Cr  ^8  r r  0)  rf  Sb  ppm Co  rs  r r  CM  |s  ppm  ppm Cd  LO O)  co  Pb  ppm  -8  Ca  CM  Ba  o co  CO  CM  ppm  CM  O)  o  CM  5031  CO  rf  o  LO  CM  2830  o  00  CM  d  0.09  cd  co  LO  ppm  r r |s  O)  CM  ppm  CM  LO  CO  rs  Mo  |s  CD O LO  CO  CM  LO  Mn  rs  O  o  CD  r r  0.33  rs  rf  d  ID CM  0.16  CT)  CM  LO CM  0.11  CM  00  CD  Mg  rf  CM  rs  La  r r  CM  CO  O CM  ppm  r r  1.66  CO CO  LO CO  00  0.13  rs  r r  0.92  9.31  o  CT) CM  CD  9.45  LO CD  CM  |s  CT>o r r  CO CO  LO  1.08  o  |s  CM  CO  0.35  co  CM  CM CM  CM  r r  0.32  cp  d  0.07  2624  rf  CM  0.09  co  o  o  0.15  LO  0.24  rf  0.14  CO CM  CM CM  0.13  rf  CM CM  0.12  CO  00  is  CM CM  0.15  co o  00  o  0.09  LO OO  CM  co  CO  0.08  2.75  1377  1.86 1.43 1.81  CO  2.56  CO CM  2.61  CO  LO  d  d  2.73  ppm As  O) CM  O CM  o  O)  co  CO  1575  ppm O)  CO  rs  2.68  ppm Au  < <  O CM  LO  o  CM  rs  CD CM  ppb  E  co  LO  00  r r  4.02  0.395  E  CO CM  co  o  is  CM  r -  LO  r r  CO  rs  r co co o  r r  CO  CO CM  CO  r~rs  CM  CO CM  rs  LO  LO  o  LO  CM  r r  2.03  rf  00  rf  1.61  9.62  LO  co  Fe  0.58 0.43 0.48 0.34  1588  0.45 0.36 0.27 0.21  r r  307.8  DDH96  CO  rf  d  rf  d  CO  To  d  0.21  1.598  d  0.22  0.673  309.3  CM  d  rf  O) CD  O)  CT)  0.26  0.589  CO  r r  1142  0.247  311.45 o  CO  0.28  425.65 424.53 423.12 317.83  425.65  DDH96 DDH96  424.53 423.12  DDH96 DDH96  422.34 311.45  DDH96  309.3  DDH96  d  307.8  DDH96 DDH96  LO  d  CO  co  o  r r  d  co  CO  d  d  CO CO  306.75  DDH96  o co  d  o  From  Hole  co  d  r r  0.68  0.244  LO  d  r r  is rf  |s  o  CO  rf  CM  is  o  d  CD  CO  CM CM  0.51  co L O co co  rs  d  d  CD  co  0.39  0.062  cp  co co  r r  o  2455  0.08  CO  co  rs  CM  d  LO  2753  0.097 0.044 0.014 0.019 0.017  cp  428.24  LO  DDH96  rf  r r  00  CO  LZ  cp  CM  LO  CO  oo  09  cp  DDH96  d  CM  r r  2163  1.358  d  DDH96  rs  CD  CO  1515  431.35  0.614 0.038 0.258  430.25  481.5 429.3  DDH96 DDH96  430.25  479.2  DDH96  d  428.24  DDH96  d  429.3  212  1031  O  ^9 cr-  ppm  ppm  -8  < <  As  Ba  Ca LO  o |v CO  o LO  co CO OO  CO  rf  CO  rf rf  CO CD  CN  CO rv CD CO  5 CO  CM  •<f  rf  CN  iv  rf  "5  CD CD  CO  cr-  CO CO  CD LO  rCD  oo  O CO  CD CN  d  X  CO  d  cf-  CO  OO  rv  co  cr-  2  D) CO  d  d LO CJ CO  LO CO LO  LO  CD  LO CJ  CO LO LO CO LO CM LO  LO CD  co  LO  CO iv  oo iv CN LO CN |v LO CD CO  LO CD  rf  rv LO  CD CO 00 CO  LO  co  co  CD  LO CO  co  CO  CD  ppm  ppm  >  Zn  r  LO |v  LO 00  co CN  d  d  d  d  ""  CN [v  rf LO  cj oo  rv OJ  T—  co O) LO  co |v OJ  LO 00 CO LO  rf  LO  o CD CD  CO  CO LO CM CO  CO 00 LO  LO  CO  ~  oi  CO 00 LO  CO |v CO LO  y  OJ  CO  ,—  LO  LO  oj  oj  oi oj  ip CO oi rf CD  co co IV CD  LO oo oi oi rf rv CO  LO  OJ CO  co  LO CO  oj oj  oi  ppm  oo  CM  Sr  0.85  0.04  co d  LO  Sc CD  ppm  iz Q. LO  ppm  d oj rf  ppm  00 LO CO  ppm  o 00  Pb  CO rv  Sb  20000  1.63  0.14 20000  20000  20000 14468 rv  8082  LO  20000  1.59  0.12 0.12  LO rv  ppm  1.18  0.22 20000  0.04  CO  LO  20000  0.15  0.11  0.05  0.14  0.18  2951 CN  0.13  0.86  3.85  3.87 3.84  2.82 1.93  2.87  4.02  d  LO  5399  0.16  0.23  2.98  CO  3956  0.52  |v  0.49  d  1553  CD  0.56  rf  oo d  0.43  rf CD OO  20000  rf  ~  T  0.56  rco  LO  11495  d CD CN  oo d  ppm  CD CO CD  •<f CD CD CD rv  LO  d  ppm  o LO rv  CO  rv  LO  Mo  CO LO  co  CO CD Ol  CM  0.39  oo  o  LO  10321  CO  rf  00  0.12  CD  0.07  rv LO  oo  0.17  CO  rv  11735  oo  3810  ID  0.65  CO rv  1.64  oo  CD  ppm  CO  d  CD  3.24  rv CO  oo co o  1.39  4.83 rv  rv  1.53  11.6  1.44 CO CO  CO  5.97  5.81  1.02  1.19 CO  •<t  4.28  3.48  1.32  oo  co  oo  Mn  0.39 rf  1.28  CO rv  CO  rf  ppm  CD  CO CD LO  La  CN CO 00 CN  LO  0.11  LO CN  6429  CO CD LO  rf  ppb  LO  4.95  CD CN CD  ppm  LO rv  Fe  rv LO CD  Cu  CO CO  co co  ppm  iv rv  Cr  r*  rv  5-  ppm  CD LO  LO rv CO  Co  00 CN CN  o> CO  7.57  5 o  10.6  CO  co  2.46  O  r*  2.37  4204 2700  0.39 1124  0.28 CD  ppm  1350  0.22  o o oo  Cd  3851  0.37  0.43  0.32  co d  0.85  0.24  0.05  0.699  1.393 CN  d  CD LO  2.04  0.32  0.014 d rv  0.71  ppm  E E CD  rco  0.47  5.027  d  ppm d  •>t  Au  497.5 2.673  10.043  8.569  3.866  0.328  0.809  6.442  1.571  1.143  d  CD CO CO  0.41  0.03  4.444  4.858  432.13  rf  To  496.35 496.35  495.35  486.46  494.35  488.6  487.52  d  485.54  493.77  DDH96  492.5  DDH96  491.5  DDH96  d  CD  484.45  DDH96  495.35  DDH96 494.35  DDH96 493.77  DDH96  d  488.6  492.5  DDH96 o  487.52  491.5  DDH96  d  486.46  DDH96 o  485.54  DDH96  d  484.45  DDH96  d  0.49  483.1  434.27  433.03  483.1  DDH96  d  0.29  481.5  DDH96  433.03  DDH96  432.13  DDH96  d  431.35  DDH96  d  From  d  d  CN LO LO  SO  Hole  d  CD  oo LO CO  ip o oi CN 00 rv  213  rf LO CN LO CO LO  co CO  OJ  oi oi rf  LO  CO  CD  oj o  LO  LO  6.653  13.471  5.028  0.024  |v  E  E  ppm  ppm  Au  < < 9573  1594  d  LO  -s  ppm |v LO LO CO  CO  OJ OJ  OJ  CD CO  co  cr-  o  co OJ O  o  rT 00 CO  CD rv |v  CJ CO  o  CJ CO  cr-  I  co  cr-  LO CO OJ  LO  OJ  --9 cr-  PPm  ppm  ppm  ppm  ppm  ppm  ppm  PPm  O)  C  z  0_  Sb  Sc  c/5  f-  >  Zn  CJ CO  Pb  CO CJ rT  co  3454  d  ppm  LO  1677  CD  ppm  d  ppm  LO  0.005  |v  0.12 d LO  2 2  2.42 14143  1.97 11182  0.09  d LO  rf IV rf LO  co LO  OJ  LO  OJ  CO  d  CD  LO  3337  0.005 13835  7680  0.08  9800  co |v  |v LO  1  d  OJ  d 00 CO  rf  co o  LO CD  LO OJ  rr  OJ  oi  rf OJ CO  CD LO LO  00  rr  00 fv CD CJ  |v  •<* LO co oi |v CD  LO  oj oi  LO  00 rv  CO  CO  rf LO  LO  oj  LO |v rf LO  25000  |v  00  1389 CO  oj co  7809  6121 |v  rf  o  co  CD rT  |v CD  o CD LO OJ  cn LO  o  CO  OJ CD  o  OJ  co  co oi o rT  LO CD rv 00 00 00  LO rT  co LO OJ  LO CD OJ CD  o OJ co  LO |v CD  6636  14559  1.79  1712  O CO rf LO OJ CD rf  co  rv CJ CO CO LO  |v  •<r |v LO  cj  LO OJ LO  LO  rT rT rf  cj  LO  LO CD LO  oi  cn OJ  cj o  o  cj co  oi cn OJ  1  o  23000!  LO  1596  d  3328  d  LO  1956  5.86 3298  0.005  2050  co d  LO  LO  LO rf  5  CD  6535  o  d  CD  45000  3.69 5320  cn LO CJ cn d  LO  d  LO  3541  2280  O) LO Ol  5386  12338  0.005 CD rT  0.005  LO  0.005  0.03  O LO  d  0.02  oj  0.01  0.06 0.13  0.03  0.08  0.12 0.06  0.06  0.04  00  Mo  0.04  •>r d  0.06  0.07  LO rf  d  0.92  0.14  OJ  ppm  |v rf CD  La  rr  0.12 LO  oi  0.21  LO  5.48  rf  6.18  co co  4.75  LO  Ol rf  0.05  76.6  0.37  2.92  CO  -<r  2560  LO CO CO CJ  5880  co  co  3.54  OJ  cn  11720  CO  co  o  1.96  OJ  7780  LO LO  4.43  co  CO  CO  1320  OJ  LO CO  2.66  d  co  ->t  CD CO  4320  rf  d  CO 00  4.05  OJ  co  rT  23980  O)  co  CO CO  6.66  r-  co  7360  co  CO CD  0.92  0.65  4.77  OJ  LO  ppb  OJ OJ  ppm  0.52  cq  ppm  o 45.2  13.3  52.6  00  ppm  6.51 rv  0.14  6.18  OJ  0.03  oo  0.58  LO  LO  0.22  00 rf  ppm  |v LO  Co  cn CO  Cd  cr-  co cn  0.38  co co  38.1  rf  0.68  IO  0.09  CO OJ rf  Ca  rf CO CO  |v  0.54  rf  ppm  CO  CO |v  Ba  0.17 1133  0.18 CO  0.17  0.31 2774 4993  7959  0.57 6476 co  0.12  0.38  3.943  282.65  CO  0.11  0.45  2.651  1.479  3.618  1.493  282.15  DDH97  o  CO  OJ  00009  CO  r-  CD  Fe  OJ  CD  Cu  CO  1070  rf LO  cd  0.34  282.15  280.45  d  0.15  DDH97  279.65  280.45  DDH97  279.65  DDH97  278.95  DDH97  d  0.19  2.266  278.95  278.45  DDH97  4.755  278.45  277.8  DDH97  7.135  277.8  277.3  DDH97  rf  9.693  277.3  276.9  DDH97  cn  1.627  276.9  276.1  DDH97  rf CO  2.02  276.1  274.54  DDH97  d  0.36  1.407  274.54  274.15  DDH97 00 OJ  0.32  IO  As  cd LO  0.51  •* d  0.34  274.15  273.2  272.7  O)  To  273.2  DDH97  272.7  DDH97  272.15  DDH97  LO  497.5  DDH96  CO  From  CO Cd  iv CO  SO  Hole  214  o LO  00  CO  rv  CO  3611 24000  cn LO oj co CO  05 co  LO CO cj rv CJ  CD co co  LO o oj co  o LO CM rr tf  cn Ol  rv  LO 00 CO oi co  PPm  PPm  0.  Sc  CM  6413  c/5 r-  LO  ppm  ppm  rf  >  Zn  15700  rf  97000  LO CO CO oi OJ  37000  CM  2514  CO LO CO CM CD CD  11300  OJ  ppm  ppm  LO LO LO CM oo CM CJ  ppm  ppm Mo  z  cn  Sb  ppm  es 2  Mn  rt  -8 ey-  CD LO CO M •>* CM C CO  PPm  CO LO o LO OJ d  OJ  26000  0.005  LO LO CD d CO  CO cn CO CO CD CM co CD  2307  LO cn d  LO CO  CM CD  4964  ,—  CM LO CO co d •<t  CD rv cn  16700  rt  cn cn OJ  2000  rf  o  CD LO LO LO CM CM o  53900  20000 rf  r- CO  ppm  0.66 0.005  00 LO OJ d CD  LO LO d  LO LO CD CM OJ CO  O  CD  1183  6205  0.32 0.005 0.005  rv  rv  rv  LO iv d  OJ LO CO cn d  LO LO CO CO oj oi  1443  CJ LO  0.005  1.33  rt  0.005  14107  OJ  0.01  X  CO  0.005  c>  rf  cn  Pb  2.96  20000  05 CD  0.64  12648  LO LO LO co oi CM LO  1.95  rf  1.15  Oi  2047  0.14 0.21  1834  rv  0.15  LO o co  PPm  PPm  Fe  CD LO  cn CD LO LO CM  La  ppm  Cu  rf  o  0.12 0.08 0.03 0.02 0.09 0.09 0.07 0.06 0.09 0.09 0.14  ppm Co  6.37  ppm Cd  rf  0.06  6.97 7.84  9480  OJ  6.21  0.64 0.38  10.1  PPm  0.43  4.68  6  13.698  rf  CO  -8 cr-  cr-  en  CD  co CM  •-8  rv  OJ CO  cn  ppm  rf  CO  O  Ca  CJ CO CO  CO o CO  CM O CD  CO  Ba  rf  co CO co CO  05 CM cn  CO  9740  26.1 rf  LO co oo cn LO  rf  CO  CJ CO rv o 00  CO CO  LO  0.09  5.99  4486  1.72 1.38 1.34 1.71 2.12  CM CD LO O  rv  3147  2.59  CO LO  CD  2.68  470.5  rf  151.7  1.43 0.33 0.95 0.75  OJ  -8  cr-  CD  CO  cn  As  co  rf  LO  rf rf  11540  3.75  r-  CO  •t  0.68  rf  CO OJ CJ CJ CD  0.78  LO  0.75  1420  0.25 0.21 0.22 0.28  1034 1864  0.28 0.28  2915 1216  0.39 0.37  LO  CO  cn OJ 00 cn  ppm  284.05  CO LO OJ CO  LO  CO  < <  283.3  rt  CJ  00 CO  E E  282.65  00 OJ co  oj  CD  50000  CO OJ  CO  49540  OJ OJ co CD  d  rv  32680  4.11  rv  rf  CO CO  9540  co CO CO cn  CO  7560  CD OJ CO  rf  rf  LO  ppb  LO d  CO CO  05  8.84  O) r f o O CO rf  rv  3.18  0.39  0.54 0.63  1411  0.51 0.36  CO OJ CM CO CO d  0.25  CM o o CO CO CO  Au  18.9  rv  CO CJ o co d  2501  cn  cn cn CO  00 (D o cn co rv  1480  35.9  14.1  10.5  CO in  CO  0.18  1.422  d  rf  1162  1.046 0.714  d  LO  ppm  414.2  3.812 0.029 0.079 0.219  d  12.818  289.5 285.4  284.6 284.05 283.3  co d  d  rf  OJ CO CD o cj r-  From  284.6  285.4  r00 OJ  286.15  r~ CO OJ  287.45  287.7  CO CO OJ  286.15  287.45  287.7  CO cn CO CO OJ OJ  d  rv  18.718  cn CO OJ  rv  31.44  289.5  o cn OJ  To  DDH97 Hole  o cn cn OJ OJ  rv  CM O cn  SO  DDH97  DDH97  DDH97  DDH97  DDH97  DDH97  DDH97  DDH97  DDH97  DDH97  DDH97  DDH97  DDH97  Oi  CD cn OJ OJ  0.013  d  11.806  413.3 rt  9.055  DDH97  CO  d  413.3  DDH97  215  CO | s  CD  cn  CO  IS  OJ  QS  CO  X  OJ  ppm  ppm  ppm  ppm  z  0-  Sb  Sc  cn  ppm  |s  4157  4027  0.06  co 0.04 2169  rf  0.12 11706  0.18 rs  0.06  0.28  0.02  2501  0.02  LO LO CM CO  5 CM  cn CO  ^~ LO CO rf  d *—  CO | s rf LO  LO CM CM  d Ul  LO  d  d  CO  CM  cn  LO  CM  LO CO  00  rs CO  OJ  00  CT> OJ CD CT)  o  o  cn  d  CM  w H  rf  co rs CM  00  CO CO  o  rf CM CO  LO OO  LO  LO OJ  LO  CO LO  rf  co oi  rf  o CT)  co  CM  rf OJ  r-  CT)L O CM OJ oi CT)  CO  LO CO LO  CO CM rf  is O OJ CD CO LO  rs Ol is  5 LO CM LO  LO OJ rf OJ  LO  LO OJ rf CO  CD  CO  >  OJ is OJ LO  CD LO  LO LO |s rs  LO rf rs  co CD  LO | s CO rf  CT)  LO CO LO CO  LO | s | s lO IO  LO rf CO CM  Ul  LO CM CD OJ  LO LO |S  CT)  CD rs  CO  CT) cn  CD  31000  co oi  oi  oi  oi  oi  oi  oi  oi  oi  oj  oi  oi  00  oi u>  10  o  LO CM  oi co  12000  CO CM  LO CO  1662  00  OJ  ppm  CM CM  ppm  OJ CO  d  Zn  rs  9915  LO  17200  0.005 0.005 cn d  0.005  0.005  0.005 CO CO  0.005  OJ  rf  5138  0.005  co  o  29000  0.005  0.07  OJ  00  1417  0.005 Ol CD CD  0.005  4.31 1900 0.05  2040  oo  ppm  2.41 1840 0.02  1600 0.04  3.32 1220 0.12  8400  0.04  1055  0.11  4104  0.07  7.95  0.14  0.03  0.02 0.04  2.94  |s  Pb  O  d  ppm  rf OJ LO  3238  O LO  d d  ppm  co co LO  d  ppm  CO LO  d  Mo  IO rf ID  d  Mn  CO  in d  0.09  Is  co CO  0.29  LO LO  1064  co o  CM  3.18  CO  rf  2080  co co o  0.29  6.21  CO  4220  LO  2.97  OJ OJ  0.31  rf  cn  CM  15660  rf  oi  ppm  OJ is rf  rf Ol LO  to CO OJ  3.14  CO CO  CO  CO  6880  oi LO  F^  4.52  LO  oj CO  LO  2423  OJ  d  OJ  12880  cn d cn to  CT)  3.96  0.87  CO  d  OJ  2820  1.05 is  OJ  3.72  0.23 | s  0.23  OJ  |s  cd  uo d  Mg  rs  to d  LO  LO  r--  La  cn  d  OJ  ppb  OJ  to d  CD  7.95  OJ  Fe  co  ppm  O CO LO  ppm  CO  r-  Cu  rf CD  0.37  LO  0.25  s.  co co o d  d  152.2  CO rf  o  to  27.4  LO  0.35  co CO to  OJ OJ  10.6  0.17  0.14 0.07 | s  0.61  1856  0.663  425.1  424.12  DDH97  LO  LO  0.88  5912  0.07 co o to  0.11  1.312  424.12  423.15  DDH97  LO OJ  to  1.19  5426  5149 rf  3971  1783  8.612  423.15  422.91  DDH97  CO  ppm  <  Ca  9936  0.24  11.94  422.91  421.84  DDH97  cn  r~  Cr  ppm  O) <  Ba  LO  | s LO  Co  5411  ppm  QS  As  CO  rf CO  0.66  oj 0.47  rf  0.15  8.409  421.84  rf  0.69  7.901  2.81 14.4  420.92  420.45  420.92  DDH97  LO  0.38  32.2  420.45  DDH97  420.31  DDH97  to to  0.18  45.7  6.656  420.31  419.8  DDH97  IS  0.27  20.5  2.919  419.8  418.79  DDH97  rf  0.41  71.3  3.959  418.79  417.88  DDH97  CO  0.69  2.299  2.36  2.504  417.88  417.1  416.7  417.1  DDH97  416.7  DDH97  CO rs  416.2  DDH97  LO  ppm  0.71  2.558  416.2  415.74  DDH97  rf  5.326  415.74  414.9  DDH97  oj  Cd  24.9  ppm  --8 t>  Au  414.9  E  22.794  414.2  DDH 97  E  ppm  To  From  Hole  216  cn  in  co  00  CO  LO  LO CM  LO  LO CM  CM  LO  oi  CO  LO  LO  LO  CO  o  1  rr  CD  LO  LO  CO CD  CO  0)  CO  rr  LO  LO 04  s.  CO  oi  CO CO  10  rs  CO  LO  oo  CD  00 LO  rs  |s LO  co  CO OJ  CM CM  oo  CO  38000  LO  rs  CO  CO  LO  CM CM CO  o  co  LO  CO CD OJ  CO  CM  LO  rs CD  LO  rs  oi  1753 1335  Ol  oi oi r r  oi oi  oi  LO  CO  CO  CD  o  co r-  LO  |s  ID  o  rs  LO  o  CD CO  ID  CD  r-  LO  O  OJ rs  LO  ppm  ppm  CO  ppm  oj  LO  z  CL  Sc  d  co  co H  Ul  CM  LO  o  CD  CO CM  CO  LO  oi  CD CO  co oi rr rs  CD  oi  oi  ppm  d  9689  d  3628  d  rT  4655  co d  3781  d  CD  LO  ppm  0.93 1.02 0.27 2.89 2.69  o  Sb  Ul  I  LO  CM rs  ppm  CO  LO  Pb  rr rr  oi oi  ppm  Ol  LO  d  OJ  ppm  d  o  LO  LO  3944  CO  LO  LO  LO CO CO  3183  rf  o  CO  oi oo  11447  Ol  LO  IS  7061  OJ  o  ppm  OJ  CO  CO CM  ppm  CO rs  co  Mo  co  o  LO CM  oj  Mn  Ol  LO  CO  2.58  OJ  00  CN  ip d  1.25  5.36  CO  OJ CM OJ  rf  OJ  2.86  rT  LO  0.39  CO is  LO  CO  oi oi  0.01  OJ  CD  0.005  Ol  rT  OJ LO OJ  0.005  CO  LO  CO  LO  oi oi  CD  OJ  co oi  rT  LO  0.005  LO ID  r- co  rf  8840  162.6  co  rr  4320  cp  oi  LO  6360  OJ  Ol  3.65  7.85  LO CO  Ca  CD  0.47  CO  o  0.35  co  rT  0.29  LO |s  o  0.14  CO  CD Ol  La  1.34 3.05 3.53  CO  1197  co  2.38  ID  1.21  o co  1.27  CO  o  1.73  CO Ol  co  1.22  o> ri  CO  3.56  co  LO  rf  ppm  ppm  Fe  CD OJ  oi  oi o co  CD  >  Zn  ppb  Cu  7.63  CN  1845  OJ  rs  1057  ppm  Co  6  o  r-  1360  0.05  ppm  ejs  CO  CO  d  2149  0.07  1000  ppm  QS  CO CO  1  LO  7284  6360  3.23  d  CO  d  CO  CO  OJ  6716  6.04  LO  CO  d  OJ  rf CO  3480  0.09  CO  CO  d  CO  0.83  0.24 0.15 0.17 0.26 0.24 0.19 0.15 0.11 0.08 0.13 0.08 0.08 0.11 0.11  o  rr  3281  CN  00  o  CO  3264 o  CD  ip d  2338 1081  ppm  1564 ppm  Ba  1200  0.39  As  13.7  0.34  13.1 ppm  0.19  0.811 ppm Au  CD CO  0.12  1.457  To  DDH97  426.85 425.8  CD < <  425.8  E  DDH97  425.1  1  E  From  CD  CO  LO  rs  CO CD  00  LO  CO LO  CD |s  rr  Hole  oo  rT  rT  rr  rr  CO  CD  rr  rr  co o  CD  10  |s OJ  LO  CO LO  2517  1.267  d  0.37  co d  26.4  0.369 0.44  d  co  OJ  "-  1.32  CO  CN  |s  rT  7.24  0.49  o  rs LO  LO OJ  d  7.44  OO LO  rr  co cp OJ oi  ppm  0.57 0.62  o  0.24  427.85  CD  O OJ  rf  LL  428.6  429.25  rr  1.026  CO  00 CM  rr  d  0.875  o  o  3592  o co co r r rr  rf  CD  CD  Cd  0.61 0.37 0.44  rr d  CN  rr  d  o  |s LO CO  0.32  0.253  co co  rr  432.6  LO  0.142  d  1.632  CN CO  CO  |s OO  oo r r  rs  0.26  2.829 1.448 0.804  0.056  437.08  0.152 0.537  434.25  6.444  433.5 432.9  d  427.45  DDH97 DDH97  434.25 433.5  DDH97  d  432.9  DDH97  d  432.6  DDH97  rs CO  d  428.6 429.25  DDH97  d  rr co rr  427.85  DDH97  rr  427.45  DDH97 DDH97 DDH97 DDH97  LO  rr  426.85  DDH97  co co  CO rf  O CO  d  co  CO OJ  OJ  90  DDH97  LO  o  CO  LL  DDH97  CO CO  rr  CD O OO  d  90  CO CO rf  rr  90  DDH97  rs  CO CO  £  E  ppm  ppm  Au  ra < <  0.44  co o rf  CM  ^9  ppm  ppm  Ba  Cd  CM  cn rf  O CO  0.12  6  ^9  I  6s  2  ppm  ppm  ppm  ppm  ppm  z  0.  Pb  Sb  Sc  LO  d co co  co CO  rf co  LO  co cvi  =  LO CM  LO CM LO CM CO  LO  LO eg  cn  LO CM LO  LO CM  co co  o  co H  ppm  0.74  LO  ppm  2.58 2393  0.11  f~- d CO  >  Zn  3592  1.06  CM  ppm  ppm  a>  ppm  1—  Mo  cn 3.32  CO  3.01  CM  1916  0.09 rf  ppm  rf CO  Mn  CO  0.06  2.24 CO CM  ppm  LO  o rf  La  10 CM  rf  2.85  CM CO  ppb  CM CM  2.49  rf rf  ppm  o CM  Fe  CO CM  ppm  co rf  Cu  o  ppm  LO CO  oo d  Co  co  Ca  o  3942  0.29 CD CO  3687  0.35  d  1511  d  ppm  CO  As  oo  0.33  0.043  0.249  0.262  0.088  d  To  438.25  438.25  437.9  DDH97  d  437.6  DDH97 437.9  DDH97 437.6  DDH97  OJ  From  00  437.08  Hole  218  CM CD CO  CO O 00  o  CM  o  o  o  219  Appendix D: Manganese sulphide mineral electron microprobe data (see section 3.5.3 for operating conditions)  220  I co  Gi T <? ^ CO CM co co  rt 00 co d ai LO rT CO  ^  o o  T-  LO co TCM oj CO  rt  i - CO iv 00 r r O CD CM CO CD  CO CM d LO  o> OJ  cb CO  rJ  CO  d oi LO rt  ri  CM 00 T - CD LO rt CD CM CO CD  rr 00 ^ CO CM CO CD  CO  T-  rv ai rt  1 0  CD  rv  °° o r?Ci  ^ O o  CD N  ^  O CT) LO rt  ri  c  .21 rt  CO  I  CM  co o d  CD CM CO CO  o  O  cc i co JO JD co cd oco co •  JO co  o  CO CM  I  JO co o co  JO co  o  co co  i  JO CO  o  CO CD cr, O? °> O CD T Q co co  LO o ai  rv CM rt rt o a> LO rt  rv O  O CD CD CO  LO O LO rv CM o d ai LO rt  o o co CM C O o — rv' ^ d co co  CM r- rv 00 CM o CD CM d CO CD  CD  LO CO rv CD rv o CD CM CO CD  LO rf ai Gi  -t  o co ai CD  00 CM rt CM CD d ai o LO rt d  q  CD CD  rt CM LO 00 CO o CD CM CO CD d  a> o d  g  CO CM LO 00 CD d o i r i UO rt  OJ  ai rt  LO o  CD  CD O d  CM CO CM CO  rv co CO CD o O)'. LO rt  O Too uo cb CM co co  CM 00 o CM rv o d ai d LO r r  CD rv cb co  O CT) LO rt  O co c\i co  I  JO co o CO  00 rt CD CD rt o CO CM CO CD  CO  C  2  oo  -  *° 9-  LO  rr§  o ro ,-; LO rt  CD =3 C JO LL O N Q. sfl  vO  sO  o~- o  o~- o  -  NO  +-  o  o  °  C/D c  o  J)  vP  3  c  u  rv  rv  iv Gi  oo CD CM CO CO  O d  rv  O d  5 L L O N Q . vj  s«  sP  rv LO LO LO rt CD CM o CO CD  co  C  fl)  rv  o  o o d  •  C  J J  i?  sS  s?  -2  o  CM  LO o d  T-  ai rt  d LO  rv o  a> oo rt O Gi rt  ,  2  ai a>  2LLOND.vo  o  o  CO CM  o  o  0 0n  C  CD o d  <D 3 c LL O N  H CL  221  co co  CD  TCD r t LO iv T - : d r f LO  q  o  °> o cn LO co g CO LO Q LO °  rt  oo o  CD  O  CD  TCD CO  CO CM CD O)  co n ri co co  co rv  rt O CD O)  [ V CO  CO LO LO r t  O CD rv CM CD  co  rv rv CD CD O J O CO L O d LO r t  O CO  CD  rv CM , - ; CD CO  rt O LO CT) O J CO L O LO r t  O CM  CO CO  oj  LO CM CD CD  ^  rt  . °° o  CD  oS CD  rv CM CD CO  ri  oo oo  —i.  rv CM CD CO  ri  CO c  o  CO  O  I  rv CD O T-  LO rt  o  CD O rt LO  rt c\i  i  JO  co o  rt 00  rt CO LO rt  99.44  o  co o o o  Tr f oo  o  CO L O d LO r t  d  LO CO  CD d CD  ^ ^ o  c CD O i  in in - a? 8 0  CD Q. rv CM CD CO O  ri  g  M D in CD CD  CM  o  d  CM  rt  CM _ j .  co co  CO  CD JO  oj CD CJ) CM  CD CM  CO  cd L O LO r t  CD  co o  oS  co •  JO CO  o  rv CM Cri CD  S r ? 0  rv  co CM co  co  ,-;  co o  rv r t CM CD rv! c\i CD CO  o  cd rb LO r t  CM CM CD CD  co rv CD d 00 CD CO  00 0 0 CM CO LO LO  CD O CD CD  CD r t CD rv CM CD CO  CO CD rv r f  CD  T-  rt  CO  co o co  o l  co L O oo m r t cd uo LO r t  o  CM OO CO LO  co o  LO CO LO rt  CO S? ?  5?  CM CD  O CM  CD CD rv CM  CD CD  N  CO  S?  S5  rt co O  CM rt CD CD  "cC  Tol  JO  o  %%^% %  CO  o  CO  C  CD  co  CM  o  CD  i/-.  co O oo CD CM v oo O d) CD cb cb CD co co  g  co CD  CO CD r t CD CD  SS c n  o  CO LO lO r t  cz sz N CO  £ CO 2  (1) Li-  ^  S« 5^  S?  co  a. o  CO LO CO LO ri LO r t  CO CO CD CO d  co co  T  .  cz CD  Z  c c N CO S5  S5  "ca  o  1-  CD rv d CD  CO v?  rt •x—  co o  o  CO co d  c v ov°  d  CD = cz tz L L Z N W V?  > °  222  Appendix E: Pb-isotope data for the Potrerillos district  The following unpublished lead isotope analyses from the Potrerillos district were provided R . M . Tosdal (written communication, 1998).  Material sampled Pre-Jurassic igneous rocks  Jurassic andesites  Pre-Jurassic sedimentary rocks  Jurassic sedimentary rocks  m  Pb/ Pb 18.902 18.582 19.338 19.346 19.056 19.345 19.674 18.744 18.876 19.947 18.79 19.11 18.891 18.809 18.556 18.814 18.581 18.619 18.709 18.274 18.687 18.921 19.568 19.34 18.469 18.45 18.65 18.8 18.871 18.858 18.813 18.77 18.822 18.897 18.825 19.534 2 U 4  2 U /  Pb/ Pb 15.655 15.651 15.671 15.674 15.655 15.675 15.698 15.666 15.661 15.724 15.642 15.661 15.662 15.642 15.611 15.667 15.621 15.607 15.617 15.58 15.647 15.715 15.732 15.709 15.661 15.645 15.668 15.655 15.652 15.688 15.648 15.641 15.63 15.644 15.637 15.691 / U 4  20«  p b /  204  39.138 38.666 39.415 39.608 39.083 39.546 39.677 38.708 38.849 39.915 38.828 39.224 38.955 38.928 38.574 38.894 38.588 38.499 38.599 38.419 39.06 39.155 39.785 39.319 38.649 38.512 38.756 38.87 39.003 39.019 38.917 38.754 38.66 38.78 38.713 39.313  p b  223 18.874 19.559 19.651 18.994 Eocene to Oligocene 18.551 porphyries 18.53 18.537 18.534 18.508 Cobre porphyry 18.561 sulphide and 18.527 sulphosalt minerals* 18.6 18.588 18.643 18.595 18.602 El Hueso pyrite 18.602 18.603 18.603 18.603 * Data reported in Marsh (1997)  15.632 15.672 15.689 15.646 15.603 15.591 15.594 15.592 15.582 15.606 15.598 15.63 15.618 15.64 15.63 15.604 15.606 15.618 15.608 15.608  38.636 38.802 39.096 38.686 38.506 38.456 38.47 38.459 38.415 38.516 38.49 38.614 38.576 38.538 38.569 38.543 38.552 38.591 38.559 38.559  

Cite

Citation Scheme:

        

Citations by CSL (citeproc-js)

Usage Statistics

Share

Embed

Customize your widget with the following options, then copy and paste the code below into the HTML of your page to embed this item in your website.
                        
                            <div id="ubcOpenCollectionsWidgetDisplay">
                            <script id="ubcOpenCollectionsWidget"
                            src="{[{embed.src}]}"
                            data-item="{[{embed.item}]}"
                            data-collection="{[{embed.collection}]}"
                            data-metadata="{[{embed.showMetadata}]}"
                            data-width="{[{embed.width}]}"
                            async >
                            </script>
                            </div>
                        
                    
IIIF logo Our image viewer uses the IIIF 2.0 standard. To load this item in other compatible viewers, use this url:
http://iiif.library.ubc.ca/presentation/dsp.831.1-0089378/manifest

Comment

Related Items