The Marcona - Mina Justa District, South-Central

THE MARCONA - MINA JUSTA DISTRICT, SOUTH-CENTRAL

PERÚ: IMPLICATIONS FOR THE GENESIS AND DEFINITION OF

THE IRON OXIDE-COPPER (-GOLD) ORE DEPOSIT CLAN

Huayong Chen

A thesis submitted to the Department of Geological Sciences and Geological Engineering

In conformity with the requirements for

the degree of Doctor of Philosophy

Queen’s University

Kingston, Ontario, Canada

May, 2008

“I am looking for IOCG deposits”

Frontispiece: The mysterious Nazca Lines (a hummingbird) in the Cañete basin, 70 km north of Marcona

ii ABSTRACT

The Marcona district of littoral south-central Perú represents the largest concentration of iron oxide-copper-gold deposits in the Central Andes. Hydrothermal activity occurred episodically from 177 to 95 Ma and was controlled by NE-striking faults.

At Marcona, emplacement of massive magnetite orebodies with subordinate, overprinted magnetite-sulphide assemblages coincided with a 156-162 Ma episode of eruption of andesitic magma in the Jurassic arc, but mineralization is hosted largely by underlying, Lower Paleozoic metaclastic rocks. The magnetite orebodies exhibit smoothly curving, abrupt contacts, dike-like to tubular apophyses and intricate, amoeboid interfingering with dacite porphyry intrusions, interpreted as evidence for the commingling of hydrous Fe oxidic and silicic melts. An evolution from magnetite - biotite - calcic amphibole ± phlogopite assemblages, which are inferred to have crystallized from an Fe-oxide melt, to magnetite - phlogopite - calcic amphibole - sulphide assemblages coincided with quenching from above 700°C to below 450°C and with the exsolution of aqueous fluids with magmatic stable isotopic compositions. Subsequent, subeconomic chalcopyrite - pyrite - calcite ± pyrrhotite ± sphalerite assemblages were deposited from cooler fluids with similar δ34S, δ18O and δ13C values, but higher δD, which may record the involvement of both seawater and meteoric water.

The much younger (95-110 Ma), entirely hydrothermal, Mina Justa Cu (-Ag) deposit is hosted by Middle Jurassic andesites intruded, on a district scale, by small dioritic stocks at the faulted SW margin of an Aptian-Albian shallow-marine volcano-sedimentary basin. Intense albite-actinolite alteration (ca. 157 Ma) and K-Fe metasomatism (ca. 142 Ma) long preceded the deposition of magnetite-pyrite assemblages from 500-600°C fluids with a magmatic isotopic signature. In contrast, ensuing chalcopyrite - bornite - digenite - chalcocite - hematite - calcite mineralization was entirely the product of non - magmatic, probably evaporite-sourced, brines.

iii Marcona and Mina Justa therefore represent contrasted ore deposit types and may bear minimal genetic relationships. The former shares similarities with other Kiruna-type magnetite

(-apatite) deposits. In contrast, the latter is a hydrothermal system recording the incursion of fluids plausibly expelled from the adjacent Cañete basin. Non-magmatic fluids are inferred to be a prerequisite for economic Cu mineralization in the Cu-rich IOCG deposits in the Central Andes and elsewhere.

iv ACKNOWLEDGEMENTS

This thesis research would not have been completed without the support of many people. First I present my esteem and appreciation to my co-supervisor, Dr. Alan Clark, who originally gave me the chance to come to Canada to continue my graduate study and his fatherly caring for my studies during these years in Kingston, and more important, for his extremely strict but definitely reasonable guidance in my Ph.D. project. Finally, although not completely, he opens my “hard granite brain” using his sharp axe, to make me more like an economic geologist, who should know everything. I also greatly thank my co-supervisor, Dr. Kurt Kyser, for his great patience in allowing an “over proud” but mindless Ph.D. student to grow up slowly. I will remember his endless understanding of my difficult situation and more important, his guidance in my research, concise but extremely important. I also thank my supervisors for their funding for both my research and living, dominantly from their grants funded by Natural Sciences and Engineering

Research Council of Canada (NSERC), and partly from the scholarships offered by Queen’s

University.

Shougang Hierro Perú SA, Chariot Resources and Rio Tinto are thanked for their cooperation in my field work, particularly for comfortable lodging in San Juan. Mr. Yuming

Chen, former chief geologist of Marcona Mine, is especially thanked for his great help at the very beginning of this project. I thank Mr. Nicholas Hawkes and Timothy Moody for their understanding and permission to sample Mina Justa drill cores, and also the protection afforded by their local mine team, who ensured my safe return to Canada, without falling into ancient mining workings.

At Queen’s, I benefited greatly from the advice and help of Kerry Klassen for isotope analysis, April Vuletich for Laser-TOF-ICP-MS analysis, and Alan Grant for X-ray studies. Dr.

Gema Olivo is thanked for her help and permission to use the fluid inclusion equipment and digital-camera microscope at Queen’s. Roger Innes and Jerzy Advent prepared thin and polished v sections. I also thank Thomas Ullrich and Peter Johns for their help and advice in Ar-Ar and microprobe analysis at UBC and Carleton University, respectively. I especially give my thanks to

Joan Charbonneau, who was the bridge between Alan and me during my thesis correction, and other Geology staff members, Dianne Hyde, Linda Brown and Ellen Mulder for their kind help.

Thanks to many grad students and postdocs for their useful discussions and help. Special thanks are due to Al Montgomery, Greg Lester, Jorge Benavides, Chan Quang, Amelia Rainbow, Mike

Cooley, Dave Love, Farhad Bouzari, Rui Zhang, Jingyang Zhao and Luis Cerpa.

Last, but not least, I would like thank my dear wife here in Kingston for her support, encouragement and sacrifices, especially when she has to work hard on her own Ph.D. project.

Also I thank all my Chinese friends in Kingston, especially those from the Kingston Chinese

Alliance Church: my life will be difficult without their help.

vi STATEMENT OF ORIGINALITY

I hereby certify that all of the work described within this thesis is the original work of the author.

Any published (or unpublished) ideas and/or techniques from the work of others are fully acknowledged in accordance with the standard referencing practices.

Huayong Chen

May, 2008

vii TABLE OF CONTENTS Abstract……………………………………………………………………………………………iii Acknowledgments………………………………………………………………………………...v Statement of Originality…………………………………………………………………………vii Table of Contents………………………………………………………………………………...viii List of Figures…………………………………………………………………………….....…….x List of Tables…………………………………………………………………………………….xiv

Chapter 1. Introduction The “Iron Oxide-Copper-Gold” Ore Deposit Clan: Problematic Definition and Genesis………………………………………………………………………………………1 The Scientific Contributions of This Study…………………………………...... ……18 Thesis Organization………………………………………………………………………...21

Chapter 2. The Longlived, Marcona-Mina Justa Iron-Copper District, Perú: Implications for the Origin of Cu-poor and Cu-rich IOCG Mineralization in the Central Andes 2.1 Abstract…………………………………………………………………………………22 2.2 Introduction……………………………………………………………………………..24 2.3 Regional and District Geologic Setting…………………………………………………31 2.4 The Marcona Magnetite Deposit………………………………………………………..40 2.5 Paragenetic Relationships of the Marcona Orebodies…………………………….…….53 2.6 The Mina Justa Cu (-Ag-Au) Deposit……………………………………...... ……69 2.7 Stable Isotope Geothermometry………………………………………………………...84 2.8 40Ar/39Ar Geochronology………………………………………………………….……88 2.9 Discussion 2.9.1 An Oxide Melt Origin for the Main Marcona Magnetite Orebodies?...... ……97 2.9.2 Evolution of the Marcona-Mina Justa District…………………………………..103 2.10 Conclusions…………………………………………………………………………..115

Chapter 3. Contrasted Fluids and Reservoirs in the Contiguous Marcona and Mina Justa Iron-Oxide Cu (-Au-Ag) Deposits, South-Central Perú 3.1 Abstract………………………………………………………………………………..119 3.2 Introduction……………………………………………………………………………120

viii 3.3 Ore Deposit Geology………………………………………………………...... ……123 3.4 Alteration and Mineralization 3.4.1 Marcona Magnetite Deposit……………………………………………………..126 3.4.2 Mina Justa Cu Deposit…………………………………………………………..129 3.5 Sampling and Analytic Techniques……………………………………………………130 3.6 Results 3.6.1 Fluid Inclusions…………………………………………………………….……133 3.6.2 Stable Isotope Geochemistry…………………………………………………….150 3.7 Discussion 3.7.1 Fluid Evolution in the Marcona Deposit………………………………………...160 3.7.2 Fluid Evolution in the Mina Justa Deposit………………………………………165 3.7.3 Implications for Cu-mineralizing Fluids in IOCG Deposits……………….……169 3.8 Conclusions……………………………………………………………………………173

Chapter 4. Conclusions 4.1 Marcona – A Unique “Kiruna-type”, Magnetite Deposit in the Middle Jurassic Metallogenetic Sub-province of the Central Andes…………………………………...174 4.2 Mina Justa Cu (-Ag) Deposit – A Major Cu-rich IOCG Deposit in the Cretaceous Central Andes…………………………………………………………………….……………177 4.3 The Protracted History of Alteration and Mineralization in the Marcona-Mina Justa District and Other IOCG Centres……………………………………………………...183 4.4 Implications for the Genesis of IOCG Deposits: A Redefinition and Reclassification of the IOCG Clan…………………………………………………………………………185

References………………………………………………………………………………………194

Appendix A: Analytical Techniques………………………………………………….....……230 Appendix B: Summerized 40Ar/39Ar Analytical Data for Hydrothermal Minerals from Marcona and Mina Justa…………………………………………………………………...234 Appendix C: Fluid Inclusion Database………………………………………………….……243 Appendix D: LA-TOFICP-MS Database……………………………………………………..263

ix LIST OF FIGURES

Chapter 1 Figure 1-1. Global Distribution of IOCG and districts and important deposits……………………7 Figure 1-2. Grade-tonnage data for Cu-rich IOCG deposits and Cu-poor Fe oxide deposits…………………………………………………………………………………….9 Figure 1-3. The relationships between Cu-equivalent ore metal resource and deposit size for Cu-rich IOCG, porphyry Cu-Mo and porphyry Cu-Au deposits…………………………10 Figure 1-4. The relationships between Cu-equivalent ore metal resource and deposit size for Cu-rich IOCG globally, porphyry Cu (-Mo-Au) deposits in Chile-Perú, SW USA-Sonora and British Columbia, Canada……………………………………………………………11 Figure 1-5. Alteration and mineralization zonation in Cu-rich IOCG deposits…………….…….13 Figure 1-6. Alteration and mineralization zonation in Cu-poor iron oxide deposits……….…….16 Figure 1-7. Schematic cross section illustrating the model for alteration zoning in IOCG deposits…………………………………………………………………………………...17 Figure 1-8. Locations of the Cu-rich IOCG deposits, principal iron deposits and manto-type deposits in Perú and Chile………………………………………………………………...19

Chapter 2 Figure 2-1. Locations of Cu-rich IOCG deposits, principal iron deposits and manto-type deposits in Perú and Chile………………………………………………………………………….25 Figure 2-2. Geology of the Marcona-Mina Justa district…………………………………………30 Figure 2-3. Summarized stratigraphic column for the Marcona-Mina Justa district……….…….32 Figure 2-4. Major host-rocks of the Marcona and Mina Justa deposits…………………….…….34 Figure 2-5. Schematic stratigraphic columns of the Río Grande Formation in the Cañón Río Grande, Marcona and Pampa de Pongo areas…………………………………………….37 Figure 2-6. Geology of the area surrounding the Marcona deposit and Mina Justa prospect…………………………………………………………………………………...41 Figure 2-7. Schematic cross-section of Marcona mine area……………………………………...42 Figure 2-8. Cross-section of the Mina 1 and Mina 4 orebodies, Marcona………………….……43 Figure 2-9. Three-dimensional imaging of the areal relationships of the E-Grid Marcona magnetite orebodies………………………………………………………………………44 Figure 2-10. NE-striking, NW-dipping Mina 7 orebody…………………………………………46 x Figure 2-11. Dike-like apophyses of magnetite extending from the hanging wall of the Mina 11 orebody and cutting strongly foliated and folded Marcona Formation………….……….46 Figure 2-12. Megascopic features of the contacts between magnetite orebodies and Marcona Formation host rocks in the Marcona mine…………………………...... ……47 Figure 2-13. Large-scale relationships between magnetite orebodies and dacite porphyry in the Marcona deposit…………………………………………………………………………..48 Figure 2-14. Panorama of part of Mina 3 open pit showing crudely planar or convoluted contacts between massive magnetite bodies and their host rocks………………………………….49 Figure 2-15. Crudely spheroidal bodies of dacite porphyry enclosed by massive magnetite at the contact between the Mina 3 magnetite orebody and dacite porphyry…………………….50 Figure 2-16. Angular blocks of massive magnetite enclosed in a matrix consisting of ovoid bodies of fine-grained magnetite…………………………………………………………………51 Figure 2-17. Vuggy texture of magnetite and euhedral octahedral magnetite crystals……..…….51 Figure 2-18. Dislocation of magnetite orebody by late Mina Justa and Huaca normal faults……52 Figure 2-19. Alteration and mineralization paragenesis of the Marcona deposit………………...54 Figure 2-20. Stage I precursor alteration at Marcona…………………………………………….56 Figure 2-21. Marcona Stage M-II Na metasomatism…………………………………………….58 Figure 2-22. Mineralogical and textural relationships of Marcona main magnetite Stage (M-III) and magnetite-sulphide stage (M-IV)…………………………………………………….60 Figure 2-23. Mineralogical and textural relationships of Marcona polymetallic sulphide stage (Stage M-V)………………………………………………………………………………64 Figure 2-24. Marcona Stage M-VI chlorite-talc-serpentine alteration……………………..…….65 Figure 2-25. Marcona late veins (Stage M-VII)………………………………………………….66 Figure 2-26. Hypogene alteration and mineralization paragenesis of the Mina 11 orebody.…….67 Figure 2-27. Paragenetic relationships of Mina 11 orebody………………………………..…….68 Figure 2-28. Geological map of Mina Justa Cu deposit, hosted by the upper Río Grande Formation…………………………………………………………………………………70 Figure 2-29. Cross-sections through major Mina Justa orebodies………………………………..72 Figure 2-30. Mineralogical and structural zonation of the Mina Justa orebodies, based on logging of selected drill cores……………………………………………………….…………….73 Figure 2-31. Alteration and mineralization paragenesis of the Mina Justa deposit………………75 Figure 2-32. Albitization and actinolite alteration (Stage J-I) and K-Fe metasomatism (Stage J-II) at Mina Justa……………………………………………………………………………...77

xi Figure 2-33. Mineralogical and textural relationships of Mina Justa actinolite alteration (Stage J-III)………………………………………………………………………………………78 Figure 2-34. Platy Stage J-V magnetite (after Stage J-IV hematite) intergrown with calcite, quartz and chalcopyrite……………...... …………...79 Figure 2-35. Mineralogical and textural relationships of Mina Justa magnetite-pyrite alteration (Stage J-V)………………………………………………………………………………..81 Figure 2-36. Mineralogical and textural relationships of Mina Justa Cu mineralization (Stage J-VI)………………………………………………………………...... ……83 Figure 2-37. Laser-induced 40Ar/39Ar age spectra, with Ca/K and Cl/K ratios for each heating step, and inverse isochron plots for samples from Marcona alteration and mineralization stages……………………………………………………………………………………..90 Figure 2-38. Laser-induced 40Ar/39Ar age spectra, with Ca/K and/or Cl/K ratios for each heating step, and inverse isochron plots for samples from Mina Justa alteration and mineralization stages……………………………………………………………………………………...95 Figure 2-39. Potential compositional trajectories of volcanic rock suites from different locations…………………………………………………………………………………102 Figure 2-40. Laser-induced 40Ar/39Ar plateau ages for Marcona and Mina Justa alteration and mineralization stages……………………………………………………………….……105 Figure 2-41. Cartoon of the evolution of the Marcona deposit…………………………….……106 Figure 2-42. Cartoon of the evolution of the Mina Justa deposit……………………………….113

Chapter 3 Figure 3-1. Geology of the area surrounding the Marcona deposit and Mina Justa prospect…………………………………………………………………………….……123 Figure 3-2. Geological map of the Mina Justa Cu deposit……………………………………...125 Figure 3-3. Alteration and mineral paragenesis of the Marcona deposits………………………127 Figure 3-4. Mineralogical and textural relationships of fluid inclusion-hosting assemblages from the Marcona……………………………………………………………………………..128 Figure 3-5. Alteration and mineral paragenesis of the Marcona deposits………………....……129 Figure 3-6. Mineralogical and textural relationships of fluid inclusion-hosting assemblages from the Mina Justa deposits………………………………………………………………….130 Figure 3-7. Fluid inclusion types in the Marcona and Mina Justa deposits……………………..135

xii Figure 3-8. Histograms of homogenization temperatures (Th), eutectic temperatures (Te) and

ice-melting temperatures (Tm(ice)) for fluid inclusions of various paragenetic stages from Marcona and Mina Justa………………………………………………………………...140

Figure 3-9. Relationships among Te, Tm(ice) and Th for fluid inclusions of the Marcona major sulphide stage (M-V)……………………………………………………………………143

Figure 3-10. Relationships among Te, Tm(ice) and Th for fluid inclusions of the Mina Justa Cu mineralization stage (J-VI)……………………………………………………………...146 Figure 3-11. LA-TOF-ICP-MS data for cation concentrations in fluid inclusions of the Mina Justa Cu ore-forming fluid……………………………………………………………………149 Figure 3-12. Interelement correlations based on TOF data……………………………………..151 Figure 3-13. δ34S values of sulphides and calculated ore-forming fluids from Marcona and Mina Justa mineralization and alteration stages……………………………………………….156 Figure 3-14. Calculated δD and δ18O values of fluids at Marcona and Mina Justa………..……161 18 Figure 3-15. δ Ofluid/melt-temperature relationships for main mineralization and alteration stages at Marcona and Mina Justa and other IOCG deposits……………………………………..162 Figure 3-16. δ34S values of ore fluids at Marcona and Mina Justa and other major IOCG deposits…………………………………………………………………………….……162 Figure 3-17. Calculated δ13C and δ18O values of hydrothermal fluids, Marcona and Mina Justa mineralization and alteration stages………………………………………………..……164 Figure 3-18. Cartoon showing the evolution of the main magnetite mineralization and major sulphide stage at Marcona………………………………………………………….……165 Figure 3-19. Genetic model for the magnetite-pyrite alteration and Cu mineralization at Mina Justa……………………………………………………………………………………...168

Chapter 4 Figure 4-1. Model for the Middle-to-Late Jurassic mineralization in the Central Andes……….177 Figure 4-2. Model for the Early Cretaceous mineralization in the Central Andes……………...181 Figure 4-3. Cartoon illustrating the settings of IOCG deposits………………………….……192

xiii LIST OF TABLES

Chapter 1 Table 1-1. Tonnages and Grades of Selected IOCG and Allied Deposits…………………………3

Chapter 2 Table 2-1. Tonnages and Grades of Selected IOCG and Allied Deposits………………………..26 Table 2-2. Selected Tonnage/Grade Data for Marcona Orebodies………………………….……42 Table 2-3. Representative Electron Microprobe Data for Alteration Minerals, Marcona Iron Deposit………………………………………………………………………………55 Table 2-4. Representative Electron Microprobe Data for Hydrothermal Silicates and Sulphides from Mina Justa………………………………………………………………………76 Table 2-5. Oxygen and Sulphur Isotope Geothermometry, Marcona and Mina Justa…………...86 Table 2-6. Summary of 40Ar/39Ar Ages from Marcona and Mina Justa…………………….……89

Chapter 3 Table 3-1. Locations and Types of Fluid Inclusions, Marcona and Mina Justa………………...138 Table 3-2. Summary of Fluid Inclusion Petrography and Microthermometric Data, Marcona and Mina Justa…………………………………………………………………….………...139 Table 3-3. Average Laser Ablation Time of Flight ICP-MS Analyses of Single Quartz -hosted Fluid Inclusions, Mina Justa……………………………………………………….……148 Table 3-4. Stable Isotopic Compositions of Minerals from Marcona…………………………..152 Table 3-5. Stable Isotopic Compositions of Minerals from Mina Justa………………………...154 Table 3-6. Summary of Sulphur Isotopic Composition of Minerals and Ore-forming Fluids……………………………………………………………………………………155 Table 3-7. δ18O, δD and δ13C Values of Minerals and Fluids from Various Stages at Marcona and Mina Justa………………………………………………………………………………159 Table 3-8. Copper Ore-forming Fluids Involved in IOCG Mineralization and Other Deposit Types…………………………………………………………………………………….168

Chapter 4 Table 4-1. Characteristics of Major Cu-poor “Kiruna-type” Iron Deposits……………….……176 Table 4-2. Salient Features of the Major Central Andean Cu-rich IOCG deposits……………..179

xiv Table 4-3. Proposed Revised Classification of IOCG Deposits………………………………...188

xv Chapter 1

INTRODUCTION

THE “IRON OXIDE-COPPER-GOLD” ORE DEPOSIT CLAN:

PROBLEMATIC DEFINITION AND GENESIS

The past half-century has seen major progress in the clarification of the genesis of the majority of the long-established classes of hydrothermal ore deposits, including porphyry copper

(-molybdenum and/or gold), epithermal base and precious metal, skarn/carbonate replacement,

Mississippi Valley-type, and both volcanic- and sediment-hosted massive sulphide systems. At the same time, newly recognized forms of mineralization, such as Carlin-type gold-silver and unconformity-controlled uranium deposits, have been extensively documented, even if they attract continued genetic argument. Against this background, the “iron oxide-copper-gold”

(IOCG) clan, first promulgated by Hitzman et al. (1992), stands out as a focus of fundamental controversy, extending even to doubts as to its coherence and specificity. IOCG deposits do not necessarily represent strictly new discoveries: several such deposits in southern and northern

Africa have been sources of copper for over 2000 years, while, among sulphide-poor deposits widely assigned to the clan, Malmberget and Kiirunavaara, Norbotten, Sweden, have been major sources of iron since, respectively, 1888 and 1901 (Grip, 1978), and the Grängesberg district of central Sweden was a major producer from the 18th Century to 1980. Moreover, vein systems rich in hematite and/or magnetite, such as have long been mined for copper and gold in the Central

Andes, have been included in the clan by Sillitoe (2003) and others.

It could, however, be argued that an abundance of such extremely common minerals as magnetite or hematite constitutes a vulnerable basis for the definition of a copper sulphide ore deposit class. Thus, why should IOCG deposits not be subsumed under the magnetite-rich

1 copper-gold sub-class of the porphyry clan (e.g., Ulrich et al., 2002), and what are their relationships to magnetite- and hematite-rich gold (-copper) porphyry mineralization (Vila and

Sillitoe, 1991)? Similarly, the widespread development of calc-silicate alteration, largely amphibolitic but locally rich in diopside and garnet, and the calcareous host-rocks of some IOCG deposits, suggest affinities with skarn mineralization. As an extreme example, is the small

Amolanas deposit of northern Chile (3 Mt@ ~ 2.7 percent Cu) an IOCG system? There, nodular clusters of chalcocite, bornite and chalcopyrite intergrown with and rimmed by hematite were emplaced along flow bands and ductile-to-brittle fractures during vesiculation of a subaerial rhyolite lava flow erupting at the margin of a Paleocene intra-arc graben hosting thick gypsiferous evaporites (Lortie and Clark, 1976). Although differing in many respects from those in which most IOCG deposits occur, this environment retains several key features, including the dominance of Cu-rich sulphides, scarcity of pyrite and association with hematite, which are characteristic of the clan.

Indeed, it was only the 1975 discovery of the Olympic Dam hematitic breccia complex in

South Australia (Woodall, 1993; Haynes, 2006), subsequently a world-class source of Cu and Au as well as the largest single uranium producer (Hitzman and Valenta, 2005), that prompted the establishment of the IOCG clan as a distinct entity. The direct association (Haynes et al., 1995) of major copper sulphide mineralization with an anorogenic alkali feldspar granite stock, rather than with the inherently copper-rich intermediate (granodioritic - quartz dioritic), orogenic granitoid suite which globally hosts the superficially similar porphyry copper deposits, was seen as evidence for a fundamental ore genetic distinction.

Over the past several decades, intensive exploration for deposits broadly comparable to

Olympic Dam has met with only modest success, with the greenfield discovery of only a handful of large copper-rich examples in both Precambrian and Phanerozoic terranes of both orogenic and anorogenic origin (Table 1-1 and Fig. 1-1). Among these, significant copper has been produced at

Table 1-1. Tonnages and Grades of Selected IOCG and Potentially Allied Iron and Copper Deposits and Prospects

Deposit * Age Tonnage Fe Cu Au Ag Other Data (Ma) (Mt) (%) (%) (g/t) (g/t) metals etc. (%) source Gawler Craton - Curnamona district, Australia

Olympic Dam ~1590 7738 ne (Hem) 0.87 0.3 1.6 U3O8 1 (0.29kg/t) Acropolis-Oak Dam ~1600 560 50 (Mt) ? ? 2, 3 Prominent Hill ~1600 102 ne (Hem) 1.5 0.5 3.5 3, 4 Carrapateena ~1600 ? 905 m1) ne (Hem) 2.1 1.0 5 Kalkaroo ~1605 30 ne (Mt) 0.28 0.14 6 Cloncurry district, Australia Ernest Henry 1504-1530 167 ne (Mt) 1.1 0.54 Co (0.05)7 Osborne ~1540 15.2 ne (Mt) 3.0 1.05 6 Starra ~1503 6.9 ne (Mt) 1.7 4.8 6 Mount Elliott Mesoproterozoic 3.3 ne (Mt) 3.6 1.8 6 Eloise ~1530 3.1 ne (Mt) 5.5 1.4 8 Tennant Creek district, Australia Warrego ~1830 5 ne (Mt-Hem) 2.6 7.0 Bi (0.3) 6 Gecko ~1830 4.7 ne (Mt-Hem) 3.8 0.7 14.0 Bi (0.2) 6 Peko ~1830 3.5 ne (Mt-Hem) 4.0 3.5 6 Peruvian Coastal Belt Raúl-Condestable ~115 >32 ne (Mt>Hem) 1.7 0.9 6.0 9 Monterrosas ~115 1.9 ne (Mt) 1.1 6 20 10, 11 Marcona 156-162 1940 55.4 (Mt) 0.12 trace Zn 12, 13 Mina Justa 95-104 347 ne (Mt-Hem)0.71 0.03 3.83 13, 14 Pampa de Pongo <109 953 44.7 (Mt) trace trace 15, 16 Chilean Coastal Belt Mantos Blancos 2) ~140 500 ne (Mem) 1.0 present 17 Tocopilla ~165 2.4 ne (Mt) 3.1 trace 11, 18 Chilean iron belt 110-130 2000 60 (Mt) trace 19 Teresa de Colmo 70 ne (Mt) 0.8 trace 11, 20 Cerro Negro 249 ne (Mt) 0.4 0.15 11 Mantoverde 117-128 400 ne (Mt-Hem) 0.52 0.11 21, 23 Santo Domingo Sur 172 ne (Mt) 0.57 0.08 23 La Candelaria 110-116 470 ne (Mt>Hem) 0.95 0.22 3.1 24 Punta del Cobre 110-117 >120 ne (Mt>Hem) 1.5 0.2-0.6 Zn 24 Productora ~130 30-70 1) ne (Mt) 0.3-0.6 trace U, Co 25 * significant Cu (>1 Mt) and/or Au (> 1 Moz) producers and prospects underlined. 3 Deposit/Location Age (Ma) Tonnage Fe Cu Au Ag Other Data (Mt) (%) (%) (g/t) (g/t) metals (%) source Panulcillo ~115 ~15 ne (Mt) ~ 1.45 ≤ 0.1 11, 20 El Espino ~108 30 ne (Mt) 1.2 0.15 11, 26 El Soldado 2) ~108 >200 ne (Hem) 1.4 27 Andacollo 3) 98-104 300 ne (Hem) 0.7 >0.25 28 Andean Cenozoic arcs Antauta 23 ? ne (Hem) trace trace trace Mo, W, REE 29 Amolanas 62 >3 ne (Hem) 2.5-3.0 trace trace 30 El Laco ~2.3 500 >60 (Mt) trace 31 Carajás district, Brazil Salobo ~2580 789 ne (Mt) 0.96 0.52 55 32, 33 Igarapé Bahia/Alemão ~2575 219 ne (Mt) 1.4 0.86 34 Sossego 2200-2300 245 ne (Mt) 1.1 0.28 35 Cristalino ~2719 500 ne (Mt) 1.0 0.3 35, 36 Mexico San Fernando ~100 31 m 1) ne (Mt) 0.96 37 Boleo 2) < 8 445 ne (Hem) 0.71 Co (0.06) 38 Zn (0.69) Cerro de Mercado Oligocene >100 62 (Mt) trace 39, 2 Southwest U.S.A Salton Sea Pleistocene ne (Hem) present Zn (present) 40 Copperstone Mid-Tertiary 2.1 ne (Mt) present 16.2 2 Iron Springs Mid-Tertiary 450 47 (Mt) trace 2 Pumpkin Hollow ~170 312 12.3 (Mt) 0.44 41 Southeast Missouri, U.S.A

Pea Ridge ~1470 200 55 (Mt) trace trace REE 42, 2 Boss-Bixby Mesoproterozoic 70 20 (Mt) 0.7 42, 2 Pilot Knob Mesoproterozoic 22 45 (Mt) trace 42, 2 Adirondacks - U.S.A Benson Neoproterozoic 200 25 (Mt>Hem) trace 43, 2 Lyon Mt. Neoproterozoic 25 25 (Mt) trace 43, 2 Mineville Neoproterozoic >10 42 (Mt) trace REE 43, 2 Dover Mesoproterozoic 26 49 (Mt) trace 43, 2

Sanford Lake Mesoproterozoic 30 Mt-Ilm 36 (TiO2), 44 V and P Mid-Atlantic U.S.A Cornwall Triassic >100 45 (Mt) 0.2 trace Co 45, 2 Grace Triassic 100 44 (Mt) 0.06 Co (0.2) 45, 2 Michigan Native Cu Mesoproterozoic 5 Mt Cu4) ne (Mt-Hem) 0.6-2.6 trace As 46

Deposit/Location Age (Ma) Tonnage Fe Cu Au Ag Other Data (Mt) (%) (%) (g/t) (g/t) metals (%) source Canada Nico (Great Bear) 1850-1880 21.8 1.08 Co- 0.13 47 Bi- 0.16 Sue Diane (Great Bear) 1850-1880 17 ne (Mt) 1.72 2.7 48 Wernecke ~1600 >20 Mt-Hem 0.35 0.17 49 Minto Early Jurassic 16.7 ne (Mt) 1.54 0.56 5.95 50,51 Iron Range ~1470 ~1 m 1) ne (Mt) 1.81 1.0 18 Pb, Zn 52 Coppercorp Mesoproterozoic 1.1 ne (Hem) 1.46 trace >200 53

Lac Tio Mesoproterozoic 120 hemo-ilmenite 32 (TiO2), V 54 Kwyjibo ~972 61 m 1) ne (Mt-Hem) 0.26 1.6 (max) 29 (max) REE, U 55, 49 Mont-de-l’Aigle ~400 11 m 1) ne (Mt-Hem) 1.0 2.2 (max) 56 Fennoscandinavia Kiirunavaara-Luossavaara ~1880 2600 62 (Mt) trace (P - ~1.0) 2 Malmberget Paleoproterozoic 840 55 (Mt-Hem) trace (P - 0.7) 2 Rakkurijärvi 1853-1862 ? ne (Mt) trace trace 57 Pahtohavare ~1880 1.7 ne (Mt) 1.9 0.9 58, 2 Tjårrojåkka ~1770 53 (3.2) 52 (Mt) trace (0.9) 59 Aitik 3) 1730-1890 606 ne (Mt) 0.38 0.21 60, 61 Raajärvi, Misi region 2017-2123 6.6 47 (Mt) trace trace V 62 Grängesberg Paleoproterozoic 400 55 (Mt) trace trace W 63, 2 Bidjovagge, Norway Paleoproterozoic? ? ne (Mt) ? ? 64 Russia-Central Asia Magnitogorsk, S. Ural Devonian 500 45 (Mt) 2 Peschnask, S. Ural Devonian 173 46 (Mt) 0.61 2 Teyskoe, Altai-Sayan Devonian 373 32 (Mt) trace REE, U ? 2 Kachar, Kazakhstan Carboniferous 2000 45 (Mt) 2 Sokolovsk, Kazakhstan Carboniferous 967 41 (Mt) 2 Korshunovsk, Siberia Permo-Triassic 609 34 (Mt) trace 2 Other regions in Eurasia Bayan Obo, China 420-555 1500 35 (Mt) REE (0.6) 65, 66 Nb (0.13) Hankou (Daye), China Jurassic >330 55 (Mt) 0.2 2 Luohe, China Jurassic >100 ne (Mt) 0.41 67 Sin Quyen, Vietnam Mesoproterozoic 52.8 ne (Mt) 0.91 0.44 REE (0.7) 68, 2 Khetri, India Neoproterozoic 140 ne (Mt-Hem) >1.1 0.5 69, 2 Madhan, India Neoproterozoic 66 ne (Mt) >1.1 >0.3 >2 69, 2 Bafq district, Iran Cambrian 1500 20-60 (Mt) trace P 70

5 Deposit/Location Age (Ma) Tonnage Fe Cu Au Ag Other Data (Mt) (%) (%) (g/t) (g/t) metals (%) source Avnik, Turkey Ordovician >104 14-58 (Mt) trace 2 Ossa Morena, Spain 330-350 Ma >180 25-66 (Mt) 0.11-0.4 ≤ 15 71 Africa Guelb Moghrein 5), 550-720 23.6 ne (Mt) 1.9 1.4 Co (143g/t) 72 Mauritania Phalaborwa, S. Africa ~2060 850 ne (Mt) 0.5 REE, Ni 73,74 O’Okiep, S. Africa2) ~1100 >25 ne (Mt) 1.7 Co, Zn 75 Chimiwungo, Zambia2) Pan-African 766 ne (Hem) 0.66 0.1 Co (0.01) 76, 2 Malundae, Zambia2) Pan-African 161 ne (Hem) 0.89 0.03 Co (0.014) 77, 2 Kasempa, Zambia Pan-African 229 66 (Mt) 77, 2 Kalengwa, Zambia Pan-African 1.6 ne (Hem-Mt) 6.5 77, 2 Mhangura, Zimbabwe2) Mesoproterozoic 76 ne (Mt) 1.2 15 78, 79

Hem hematite, Mt magnetite, Ilm ilmenite. 1) mineralized drill-core intersection. 2) commonly classified as manto-type or sediment-hosted Cu deposits. 3) predominant mineralization classified as porphyry Cu-Au type. 4) total production of copper to 1968. 5) Akjoujt district. ne - not economic. References: 1- BHP Billiton, 2007; 2- Williams et al., 2005; 3- Skirrow et al., 2002; 4- Oxiana Resources, 2006; 5- Teck Cominco Limited, 2007; 6- Williams and Pollard, 2003; 7- Mark et al., 2006; 8- Baker et al., 2001; 9- de Haller et al., 2006; 10- Injoque, 2002; 11- Sillitoe, 2003; 12- Shougang Hierro Perú SA., 2003; 13- this study; 14- Chariot Resources, 2006; 15- Cardero Resource Corp., 2005; 16- Hawkes et al., 2002; 17- Ramírez et al., 2006; 18- Ruiz and Peebles, 1988; 19- Oyarzun et al., 2003; 20- Hopper and Correa, 2000; 21- Benavides et al., 2007; 22- Vila et al., 1996; 23- Far West Mining Ltd. 2007; 24- Marschik and Fontboté, 2001; 25- Ray and Dick, 2002; 26- Correa, 2003; 27- Boric et al., 2002; 28- Reyes, 1991; 29- Clark and Kontak, 2004; 30- Lortie and Clark, 1976; 31- Rhodes et al., 1999; 32- Souza and Vieira, 2000; 33- Requia et al., 2003; 34- Tallarico et al., 2005; 35- Monteiro et al., 2008; 36- Huhn et al., 2000; 37- Cruise et al., 2007; 38- Conly et al., 2001; 39- Lyons, 1988; 40- McKibben and Elders, 1985; 41- Gander et al., 2007; 42- Day et al., 2001; 43- Friehauf et al., 2002; 44- Force, 1991; 45- Rose et al., 1985; 46- White, 1968; 47- Fortune Minerals, 2007; 48- Corriveau, 2005; 49- Hunt et al., 2005; 50- Sherwood Copper Corp., 2007; 51- Tafti and Mortensen, 2003; 52- Eagle Plains Resources Ltd., 2005; 53- Richards and Spooner, 1989; 54- Gross et al., 1997; 55- Gauthier et al., 2004; 56- Simard et al., 2006; 57- Smith et al., 2007; 58- Lindblom et al., 1996; 59- Edfelt et al., 2005; 60- Wanhainen et al., 2003; 61- Wanhainen et al., 2005; 1999; 62- Niiranen et al., 2005; 63- Ripa, 1999; 64- Ettner et al., 1994; 65- Smith and Henderson, 2000; 66- Smith and Wu, 2000; 67- Pan and Dong, 1999; 68- Mclean, 2002; 69- Knight et al., 2002; 70- Torab and Lehmann, 2006; 71- Tornos et al., 2005; 72- Kolb et al., 2006; 73- Vielreicher et al., 2000; 74- Groves and Vielreicher, 2001; 75- Stumpfl et al., 1976; 76- Equinox Minerals Ltd., 2007; 77- Nisbet et al., 2000; 78- Maiden et al., 1984; 79- Maiden and Master, 1986.

Ernest Henry, Queensland (167 Mt @ 1.1% Cu, 0.54g/t Au); La Candelaria (470 Mt @ 0.95% Cu,

0.22g/t Au) and Mantoverde (400 Mt @ 0.52% Cu, 0.11g/t Au), Chile; Sossego (245 Mt @ 1.1%

Figure 1-1. Global distribution of major IOCG districts and important deposits, including potentially allied iron and copper deposits (modified from Corriveau, 2005). Detailed data are listed in Table 1-1.

7 Cu, 0.28g/t Au), Brazil; the Khetri district, India (>200 Mt @ 1.1% Cu, 0.5g/t Au); and (perhaps a dubious candidate: Wanhainen, 2005) Aitik, Sweden (606 Mt @ 0.38%Cu, 0.21g/t Au), while the Salobo, Carajás, Brazil (789 Mt @ 0.96%Cu, 0.52g/t Au) and Prominent Hill, Australia (102

Mt @ 1.5%Cu, 0.5g/t Au) deposits will begin production in 2008. Elsewhere, the major

Phalaborwa (Palabora) Cu (-magnetite-apatite-vermiculite-pentlandite-baddeleyite) deposit,

South Africa, although associated with a Paleoproterozoic carbonatite centre, has been persuasively assigned to the IOCG clan (Groves and Vielreicher, 2001). Further, the iron oxide-rich Mantos Blancos (Chile), Chimiwungo (Zambia) and Mhangura (Zimbabwe) copper deposits, commonly described as of “manto” or “sediment-hosted” type, share some similarities with copper-rich IOCG deposits (Maiden and Master, 1986; Maksaev and Zentilli, 2002;

Williams et al., 2005).

It is, however, apparent that the vast majority of deposits which have been assigned to the clan, including Kiirunavaara itself, are magnetite-rich but sulphide-poor, containing, if any, only subeconomic copper and gold. Williams et al. (2005) have therefore recommended that the

IOCG designation be restricted to deposits with economic copper and/or gold (Fig. 1-2), although maintaining a genetic connexion between Cu-rich and Cu-poor mineralization. The IOCG clan thus suffers from an inherent dichotomy: most such systems failed to generate significant sulphide mineralization, no matter how intense the “precursor” magnetite deposition. Further, with the exception of Olympic Dam, these deposits are dwarfed by supergiant and behemothian

(sensu Clark, 1993) porphyry copper systems, whether Mo- or Au-rich (Fig. 1-3A). Nonetheless, in terms of the relationships between copper-equivalent ore metal resource and deposit size (Fig.

1-3), IOCG’s compare closely to both Cu-Mo and Cu-Au porphyry systems. Further, even omitting from consideration the uranium in the Olympic Dam deposit, the metal resource vs. ore tonnage correlation coefficient (Figs. 1-3 and 1-4) implies that the aggregate processes of base and precious metal concentration in Cu- and Au-rich IOCG deposits may have been more

Figure 1-2. Tonnage-grade data for Cu-rich and Cu-poor deposits which have been assigned to the iron oxide-copper-gold clan (modified from Williams et al., 2005). Most Fe deposits contain traces of Cu. Those commonly described as manto-type or sediment-hosted Cu deposits are bracketed. The Aitik and Andacollo deposits are underlined because their major mineralization is of porphyry Cu-Au type (see text).

Figure 1-3. A. The relationships between Cu-equivalent ore metal resource and deposit size for Cu-rich IOCG, porphyry Cu-Mo (PCMD) and porphyry Cu-Au (PCGD) deposits (cf. Clark, 1993). B. an enlarged part of A to show the relationships where tonnage < 4500 Mt. IOCG, PCMD and PCGD all show good correlations (R2 > 0.8) between Cu-equivalent metal and tonnage. The slopes are: IOCG (0.011) > PCMD (0.0083) > PCGD (0.0074). If uranium is included for Olympic Dam (OD) the slope for IOCG mineralization would be much higher (A). IOCG data are from references listed in Table 1-1. Data for porphyry deposits are from USGS (2005). Metal prices (Cu, Mo, Au, Ag and U) at 2007 levels.

Figure 1-4. A. The relationships between Cu-equivalent ore metal resource and deposit size for Cu-rich IOCG deposits globally, and porphyry Cu (-Mo-Au) deposits (PCD) in Chile-Perú, SW USA-Sonora and British Columbia. B is an enlarged part of A to show the relationships where tonnage < 4500 Mt. IOCG and PCD all show good correlations (R2 > 0.85) between Cu-equivalent metal and tonnage. The slopes are: IOCG (0.011) >= Chile-Perú PCD (0.01) > SW USA-Sonora PCD (0.0056) > BC (0.0044). IOCG data are from references listed in Table 1-1. Porphyry deposits are referenced from Clark (1993) and USGS (2005). OD-Olympic Dam.

11 efficacious than in porphyries, including those of the Central Andes. In part, this must reflect the sulphur-, and hence pyrite-deficient, nature of IOCG’s but, as suggested by Clark (1993), may be evidence for unusually high ore metal contents in the hydrothermal fluids.

Nonetheless, although aggressive exploration for Cu- and Au-rich IOCG mineralization continues, it is improbable that many deposits as large and rich as Olympic Dam remain to be discovered. Similarly, Kiirunavaara is likely to remain the largest magnetite deposit in the expanded clan. It should be emphasized that, in the broadly analogous porphyry copper field, many of the largest deposits, e.g., Bingham, Morenci, El Teniente and Chuquicamata, were similarly among the first to be discovered and delimited. However, a large number of major porphyry copper deposits have been identified and developed in the past half-century, and many of these are much larger than all known IOCG’s except for Olympic Dam (Fig. 1-3A). Large

IOCG’s, and especially examples rich in copper and gold, must therefore be considered, in global terms, to be both scarce and of significantly lower inherent potential than porphyry copper systems. These deficiencies may directly reflect an origin dependant on an improbable conjunction of geological processes.

Ore Genetic Uncertainties

Proximity to granitoid stocks, intense, commonly high-temperature hydrothermal alteration, and extensive hydrothermal brecciation have been interpreted by many (e.g., Sillitoe, 2003; Pollard,

2006) as supporting a direct genetic relationship between IOCG deposit development and hydrous fluid exsolution from crystallizing silicate melts, the abundance of Fe, Cu, Au and, locally, Co being ascribed to a mafic parental magma. Such basically magmatic-hydrothermal models prompt analogies with, particularly, molybdenite-poor porphyry copper-gold deposits, the vast majority of which are magnetite-rich (e.g.,Ulrich et al., 2002; Pollard and Taylor, 2002).

13 However, several characteristic, if not ubiquitous, features of Cu-rich IOCG’s, such as intense albitization, Ca metasomatism (commonly amphibolization), hematite- and calcite-rich sulphide mineralization (Fig. 1-5) and the scarcity of pyrite, as well as the overall sequence of alteration-mineralization events (e.g., Ullrich and Clark, 1999), are difficult to reconcile with thermally retrograde melt-aqueous fluid equilibria (Candela, 1989a and b). In addition, many major IOCG provinces of both Proterozoic and Phanerozoic age exhibit a close correlation with mid-latitude sedimentary/ volcanic basins which either incorporate evaporite sequences or preserve their metamorphic relics, i.e., regional scapolite-albite (-tourmaline) assemblages

(Frietsch et al., 1997). Such relationships underlie the proposal of Barton and Johnston (1996;

2004) that the brines responsible for Cu (-Au) sulphide mineralization in IOCG deposits were derived wholly or in part through the dewatering of intra-orogenic or anorogenic rift basins.

Stable isotope evidence for such a non-magmatic origin, at least for sulphur and oxygen, was first presented by Ullrich and Clark (1999) and Ullrich et al. (2001) for La Candelaria, the most important Cu-rich IOCG deposit in the Central Andes, but the incursion of basinal fluids had earlier been documented at Olympic Dam by Haynes et al. (1995). A preliminary classification of hydrothermal IOCG deposits by Hunt et al. (2007) ascribes all major IOCG deposits to hybrid magmatic – non-magmatic fluids. Although the ore metals, i.e., Cu and Au, in these deposits may be, at least in part, magma-derived (Ullrich and Clark, 1999), such inherently complex genetic models have direct implications for the occurrence of, and hence exploration for, Cu-rich IOCG’s, as is exemplified by the studies of Benavides et al. (2006, 2007) in the Mantoverde district of northern Chile.

A requirement for “exotic” sulphur in the development of Cu-rich IOCG’s would highlight not only the problematic interrelationships of “precursor magnetite” and sulphide mineralization, but also the processes of magnetite formation per se. Whereas bodies of massive magnetite antedating Cu sulphide deposition in numerous IOCG deposits clearly resulted from intense

14 hydrothermal Fe metasomatism (e.g., La Candelaria; Ullrich and Clark, 1997; Rakkurijärvi:

Smith et al., 2007; Mantoverde: Benavides et al., 2007), magnetite ± apatite ± actinolite mineralization elsewhere has been interpreted as the product of phosphatic and silicic Fe oxide melts. Such an origin, first advocated at Kiirunavaara by Geijer (1910), was convincingly documented by Lundberg and Smellie (1979) for the nearby Mertainen deposit. The involvement of oxide melts has since been proposed for numerous orebodies in the Cretaceous Chilean iron belt (Nyström and Henríquez, 1994; Henríquez et al., 2003), but is best exemplified by the

Pliocene magnetite volcanic flows in the El Laco district of northern Chile (Park, 1972; Naslund et al., 2002). A growing body of experimental data has confirmed the development of stable immiscibility between silicate and Fe oxide liquids in the system Na2O + K2O + Al2O3 + MgO –

FeO + MnO + TiO2 + CaO + P2O5 – SiO2 (Philpotts, 1982), and the geological occurrence of oxidic melts with up to 60 weight percent FeO, 37 weight percent TiO2 and 22 weight percent

P2O5 has been demonstrated (Clark and Kontak, 2004). Apparently magmatic magnetite - apatite bodies are areally juxtaposed with unambiguously hydrothermal hematitic Cu-Au veins and breccias in numerous districts, e.g., the Carmen area of northern Chile (Gelcich et al., 2005), and commonly exhibit similarities in their alteration facies. This implies that Fe oxide melts may be integral to IOCG-type sulphide mineralization, perhaps through their vesiculation through either aqueous fluid (Matthews et al., 1995) or chloride (Broman et al., 1999) saturation.

However, the importance of oxidic melts and, therefore, the recognition of the magmatic/hydrothermal “interface”, remain controversial (Rhodes et al., 1999; cf. Sillitoe and

Burrows, 2002). Similarly problematic is the relationship between “Kiruna-type” magnetite-apatite deposits and ilmenite- or rutile-rich “nelsonite” bodies, which are only rarely associated with copper sulphide mineralization (e.g., McLelland et al., 1994; Clark and Kontak,

2004).

15 In the light of the controversy regarding the contribution of Fe oxide melts to IOCG mineralization, it is paradoxical that the extensive hydrothermal alteration enveloping sulphide-poor iron deposits such as Kiirunavaara, the Chilean iron belt and El Laco (Fig. 1-6) is

Figure 1-6. Alteration and mineralization zonation in Cu-poor iron oxide deposits. A - Generalized cross-sectional reconstruction of the Kiruna district, Sweden (modified from Hitzman et al., 1992). B - Schematic diagram of intrusive magnetite mineralization and alteration envelope that formed in andesitic host rock, El Laco, Chile (Rhodes et al., 1999). (Ab-albite, Act-actinolite, Apt-apatite, Bar-barite, Cal-calcite, Cpx-clinopyroxene, Fl-fluorite, Hem-hematite, Mt-magnetite, Qtz-quartz, Ser-sericite)

16 remarkably similar to that associated with Cu-rich deposits incorporating metasomatic and, hence, hydrothermal magnetite, including Ernest Henry and La-Candelaria – Punta del Cobre (Fig. 1-5), thereby permitting integrated ore deposit models such as that in Figure 1-7.

Figure 1-7. Schematic cross-section illustrating the model for alteration zoning in IOCG deposits proposed by Hitzman et al. (1992), with the addition of representative Cu-poor systems. The three major alteration zones, viz. Na-Ca, K (Kspar/biotite) and hydrolitic, show vertical and, locally, lateral zonation. NB. “Na-Ca alteration” is commonly represented by albitization, marialitization and Ca-amphibolitization (and, less widely, garnet and diopside formation in Ca-poor protoliths), and hence records either Na- or Ca- metasomatism. Hematite and breccias are dominant in the upper levels, magnetite and veins (breccias) at depth. Cu-rich IOCG and Cu-poor iron oxide mineralization is focused at various levels, with different alteration patterns and mineral assemblages.

The Present Study

The problematic roles of both Fe-oxide melts and “exotic” hydrothermal fluids have hindered the development of a uniform model for the genesis of IOCG mineralization, as well as the establishment of exploration protocols for Cu-Au – rich members of this clan. The research documented in this thesis addresses these two salient problems in the context of a major 17 concentration of unmetamorphosed Mesozoic IOCG deposits in the Central Andean orogen.

Hosted in part by Middle Jurassic andesitic arc volcanics, the Marcona magnetite deposit, south-central Perú (Fig. 1-8), is much the largest-known locus of iron oxide mineralization along this well-studied convergent plate margin, and probably represents the second-largest

“Kiruna-type” deposit globally (Table 1-1). With only subeconomic chalcopyrite, Marcona is directly analogous to the iron deposits of the Cretaceous Chilean iron belt. In contrast, the contiguous Mina Justa Cu(-Ag) prospect (Fig. 1-8), the only demonstrably economic mineralization of this type discovered over the past three decades in the Central Andes and hosted entirely by Middle Jurassic andesites, is representative of Cu-rich IOCG deposits globally.

Together with the nearby Pampa de Pongo magnetite deposit (Hawkes et al., 2002), these contrasted deposits occur in well-defined stratigraphic and tectonic contexts and provide a basis for clarification of the processes responsible for magnetite and copper concentration in the IOCG clan.

In a wider context, this research has implications for the delimitation of this controversial family of ore deposits. Given the extreme variations in salient features of deposits assigned to the

IOCG clan, particularly with regard to geological setting and the types and sequence of alteration, it must be asked whether such inconsistencies go beyond those exhibited by porphyry copper deposits, for which Gustafson (1978) coined the term “variations on a theme”.

The Scientific Contributions of This Study

This research project in the Marcona-Mina Justa area has resulted in several distinct, but interrelated, contributions. These are:

Figure 1-8. Locations of Cu-rich IOCG deposits, principal iron deposits and manto-type deposits in Perú and Chile (from Clark et al., 1990; Hawkes et al., 2002; Maksaev and Zentilli, 2002; Sillitoe, 2003; Oyarzun et al., 2003 and Benavides et al., 2007). NB. The twelve major orebodies occurring in an area of 25 km2 at Marcona incorporate a magnetite tonnage similar to that in the five largest Cretaceous Chilean iron mines. 19

1- the characterization and redefinition of paragenetic relationships in the Marcona magnetite deposit, and clarification of the relationships between magnetite and multiple-stage, but non-economic, sulphide precipitation. Megascopic relationships strongly suggest an immiscible melt origin for the Marcona orebodies, a model supported by fluid inclusion and stable isotope analysis, although the incursion of modified seawater is evident in the subsequent polymetallic sulphide stage;

2- the first documentation of the paragenetic relationships of the Mina Justa deposit, where Cu mineralization significantly postdated magmatic-hydrothermal magnetite-pyrite alteration. The various reservoirs of the fluids responsible for the magnetite and the later Cu mineralization stages are defined on the basis of fluid inclusion and stable isotope analysis, as well as other geological information. Low-temperature basinal brines are inferred to be responsible for the formation of the Mina Justa Cu orebodies;

3- the recognition of the protracted and multistage evolution (from 177 to 95 Ma) of the district, incorporating the Jurassic (Callovian) Marcona Fe deposit and the Cretaceous (Albian) Mina

Justa Cu deposit, on the basis of paragenetically-constrained 40Ar/39Ar geochronology. Marcona and Mina Justa, although spatially associated, represent two entirely independent mineralization events, with the suggestion that the intra- or retroarc environment may be favorable for IOCG mineralization at widely separated intervals; and

4- a proposed reclassification of the broad and problematically-defined IOCG clan.

Thesis Organization

This dissertation is constructed in a manuscript format which fulfills the requirements of the

Queen’s University School of Graduate Studies and Research. Two major manuscripts are prepared for publication and form the main body of this thesis:

Chapter 2:

Chen, H., Clark, A.H., and Kyser, T.K., The longlived, Marcona-Mina Justa iron-copper district,

Perú: implications for the origin of Cu-poor and Cu-rich IOCG mineralization in the Central

Andes: to be submitted to Economic Geology

Chapter 3:

Chen, H., Kyser, T.K., and Clark, A.H., Contrasted fluids and reservoirs in the contiguous

Marcona and Mina Justa iron-oxide Cu (-Au-Ag) deposits, south-central Perú: to be submitted to

Mineralium Deposita

General conclusions are drawn in Chapter 4, addressing the significance of the contrasted

Marcona and Mina Justa IOCG mineralization, summarizing the Mesozoic metallogenesis of the

Central Andes, and concluding with a new classification of IOCG deposits.

Chapter 2

THE LONGLIVED MARCONA-MINA JUSTA IRON-COPPER DISTRICT,

PERÚ: IMPLICATIONS FOR THE ORIGIN OF CU-POOR AND CU-RICH

IOCG MINERALIZATION IN THE CENTRAL ANDES

2.1 Abstract

The IOCG subprovince of littoral south-central Perú incorporates Marcona, the preeminent

Andean magnetite deposit (1.9 Gt @ 55.4% Fe and 0.12% Cu), and, 3-4 km distant, Mina Justa, one of the few major Andean IOCG deposits with economic copper (346.6 Mt @ 0.71% Cu, 3.8 g/t Ag and 0.03 g/t Au). Fe oxide and Cu sulphide mineralization was controlled by NE-striking reverse (Marcona) and normal (Mina Justa) faults transecting a Middle Jurassic

(Aalenian-to-Oxfordian) andesitic shallow-marine arc and a succession of contiguous, plate boundary-parallel, Late Jurassic to mid-Cretaceous, volcano-sedimentary basins. Detailed documentation of alteration and mineralization relationships, supported by stable isotope geothermometry and 40Ar/39Ar geochronology, reveals an episodic history of Mesozoic magmatic and hydrothermal processes extending for at least 80 m.y., from ca. 177 to 95 Ma.

At Marcona, initial hydrothermal activity occurred in the Aalenian (177 Ma) and Bajocian

(171 Ma), when high-temperature Mg-Fe metasomatism of Paleozoic siliciclastics underlying the nascent Río Grande Formation arc, generated, respectively, cummingtonite and phlogopite-magnetite assemblages. Subsequent, widespread albite-marialite alteration (Na-Cl metasomatism) largely predated, but overlapped with, the emplacement of an en echelon swarm of massive magnetite orebodies with subordinate, overprinted magnetite-sulphide assemblages.

Magnetite and uneconomic Cu and Zn sulphide mineralization coincided with a 156-162 Ma

22 episode of andesitic eruption which terminated the growth of the arc, but was hosted largely by siliceous metaclastic rocks approximately 300 m below the ocean-floor. The magnetite orebodies exhibit abrupt, smoothly curving contacts, dike-like to tubular apophyses and intricate, amoeboid interfingering with dacite porphyry intrusions, and there is no convincing megascopic or microscopic evidence for large-scale Fe metasomatism. The largest, 400 Mt Minas 2-3-4 orebody is interpreted as a bimodal magnetitite-dacite intrusion comprising immiscible melts generated through the dissolution of metasedimentary quartz in parental andesitic magma.

From 162 to 159 Ma, an evolution from magnetite-biotite-calcic amphibole ± phlogopite ± fluorapatite assemblages, which are inferred to have crystallized from a hydrous Fe-oxide melt containing < 30 weight percent combined CaO, MgO, K2O and SiO2, to magnetite - phlogopite - calcic amphibole - pyrrhotite - pyrite assemblages coincided with quenching from above 800°C to below 450°C and the concomitant exsolution of dilute aqueous brines. Subsequently, at

156-159 Ma, chalcopyrite - pyrite - calcite ± pyrrhotite ± sphalerite ± galena assemblages were deposited from lower-temperature (≤ 360oC) brines that may have incorporated modified seawater.

Hydrothermal activity was thereafter focused in the Mina Justa area, where Middle Jurassic andesites experienced intense albite-actinolite alteration at ca. 157 Ma and K-Fe metasomatism at ca. 142 Ma. The development of the Mina Justa deposit proper began much later, adumbrated by ca. 109 Ma actinolitization and the 101-104 Ma deposition of bodies of massive, metasomatic, magnetite and pyrite from 500-600oC hydrothermal fluids. Mid-Cretaceous hydrothermal activity was areally associated with the intrusion of small diorite stocks along the faulted southwestern margin of the Aptian-Albian, arc-parallel, Cañete basin. Finally, at 95-99 Ma, hypogene argentian chalcopyrite-bornite-digenite-chalcocite mineralization, with abundant calcite and hematite, was emplaced as two ~ 400 m long, ~ 200 m thick, gently-dipping, tabular arrays of breccia and stockwork, cored by earlier magnetite-pyrite lenses.

23 Marcona and Mina Justa, although contiguous, represent contrasted ore deposit types. The former, like El Laco, the second-largest Andean magnetite deposit, is interpreted as the product of

Fe oxide melts coexisting with dacitic magmas within a failing andesitic arc, albeit in a subaqueous rather than subaerial environment. The weak associated magmatogene Cu-Zn sulphide mineralization, generated through melt vesiculation, contrasts with the considerably higher-grade Cu- and Ag-rich orebodies at Mina Justa, which were the product of cool, oxidized, hydrothermal fluids plausibly expelled from the Cañete basin during its tectonic inversion. This major IOCG district demonstrates that the largest, orthomagmatic, Central Andean magnetite deposits may bear no direct genetic relationship to copper-rich centres, while the 142 Ma K-Fe alteration at Mina Justa is evidence that IOCG-type hydrothermal activity may itself terminate before generating significant Cu sulphide mineralization, despite occurring in a favourable tectonic and stratigraphic environment.

2.2 Introduction

Iron oxide-copper-gold (“IOCG”) mineralization, first formally defined by Hitzman et al. (1992), has been a major exploration target since the discovery of the enormous Olympic Dam Cu-U-Au

(-REE) deposit in 1975. However, few large IOCG centres with economic gold or, even, copper grades have been identified, and the clan remains incompletely defined or delimited. Because most early-identified IOCG systems, e.g., these of the Gawler Block of South Australia, the

Eastern Mount Isa Block of Queensland and the northern Fennoscandinavian Shield, are of

Proterozoic age, it was initially proposed that this type of mineralization was restricted to that period (Hitzman et al., 1992). However, the Central Andean orogen, and especially the volcano-plutonic arcs of Jurassic and Cretaceous age exposed in the Cordillera de la Costa of northern Chile and central and southern Perú, are now recognized as hosting major IOCG mineralization (Fig. 2-1). Nonetheless, the great majority of the Andean deposits assigned to the

24 IOCG clan by Sillitoe (2003) and others are exploited only for magnetite, having negligible sulphide contents. Only the La Candelaria – Punta del Cobre (Ryan et al., 1995) and Mantoverde

(Vila et al., 1996; Zamora and Castillo, 2001) districts in Chile and the much smaller Raúl –

Condestable deposit cluster in Perú (de Haller et al., 2006) and Panulcillo deposit, Chile (Hopper and Correa, 2000), as well as numerous vein systems (Ruiz, 1965), have produced significant copper, and only La Candelaria – Punta del Cobre and Raúl – Condestable have been important sources of gold (Table 2-1). In the Andean context, however, Mantos Blancos and other so-called

Figure 2-1. A - Locations of Cu-rich IOCG deposits, principal iron deposits and manto-type deposits in Perú and Chile (from Clark et al., 1990; Hawkes et al., 2002; Maksaev and Zentilli, 2002; Sillitoe, 2003; Oyarzun et al., 2003 and Benavides et al., 2007). B - Simplified geologic map of the IOCG mineralization belt of south-central Perú (modified from Vidal et al., 1990), illustrating the extent of the mid-Cretaceous Cañete intra-arc extensional basin (Atherton and Aguirre, 1992).

Table 2-1. Tonnages and Grades of Selected IOCG and Allied Deposits in the Central Andes

Deposit Tonnage Fe (%) Cu (%) Au (g/t) Ag (g/t) Data source (Mt) Raúl-Condestable >32 ne 1.7 0.3 6.0 de Haller et al., 2006 Marcona 1 ~1940 2 55.4 0.12 trace Shougang Hierro Perú SA., 2003 3 Mina Justa 347 ne 0.71 0.03 3.83 Chariot Resources, 2006 3 Pampa de Pongo 953 44.7 trace Cardero Resource Corp., 2005 3 Mantos Blancos 4 500 ne 1.0 Ramírez et al., 2006 El Laco 5 500 >60 trace Rhodes et al., 1999 Chilean iron belt 6,7 2000 60 trace Oyarzun et al., 2003 Teresa de Colmo 70 ne 0.8 trace Hopper and Correa, 2000 Cerro Negro 249 ne 0.4 0.15 Sillitoe, 2003 Mantoverde 8 400 ne 0.52 0.11 Benavides et al., 2007 Santo Domingo Sur 140 ne 0.59 Far West Mining Ltd. 2007 3 Candelaria 470 ne 0.95 0.22 3.1 Marschik and Fontboté, 2001 Punta del Cobre 9 >120 ne 1.5 0.2-0.6 Marschik and Fontboté, 2001 Panulcillo ~15 ne ~ 1.45 ≤ 0.1 Hopper and Correa, 2000 El Espino 30 ne 1.2 0.15 Sillitoe, 2003 1) Annual production ~5.5 Mt concentrates (2003-2005); 2) including 389 Mt ore production (1952-2002); 3) unpublished reports; 4) classified as a hematite-rich hydrothermal breccia deposit or manto-type; 5) sulphide-barren Pliocene iron deposit; 6) including five large deposits (200-400 Mt), viz. Boquerón Chañar, Los Colorados, Algarrobo, Cristales and El Romeral; and many small deposits with 20-100 Mt; 7) Annual production ~7.3 Mt concentrates (2003-2004; Compania Minera del Pacifico annual report, 2004); 8) Hypogene protore; 9) comprising several discrete deposits; ne: not economic

manto, stratabound or breccia-type Cu (-Ag) deposits (Fig. 2-1A), although incorporating only minor hematite (Ramírez et al., 2006), may be genetically related to Fe oxide – rich IOCG systems (Vivallo and Henríquez, 1997), with which they are locally juxtaposed (Orrego et al.,

2006).

IOCG deposits are defined primarily by an abundance of magnetite and/or hematite.

Whereas the latter may be unambiguously associated with Cu mineralization, as at Olympic Dam and Mantoverde, the genetic link between magnetite and chalcopyrite deposition is tenuous in

26 many centres, and it is not clear why most deposits in the Andes and elsewhere did not evolve from a magnetite-dominated (“IO”) to an economically significant chalcopyrite±gold - rich

(“CG”) stage. This uncertainty in the ore-genetic modeling directly reflects controversy regarding the nature and, particularly, sources of the hydrothermal fluids which are clearly implicated in the

Cu sulphide facies of numerous deposits. Thus, Pollard (2000, 2001, 2006), Perring et al. (2000),

Marschik and Fontboté (2001) and Sillitoe (2003) interpret IOCG’s on the basis of magmatic- hydrothermal models, and hence as broadly analogous to the magnetite-rich porphyry Cu-Au clan, the evolution of which is controlled primarily by silicate melt-aqueous fluid equilibria. Oliver et al. (2004) and Marshall and Oliver (2006) concur in the importance of magma-derived fluids, although concluding that their extensive interaction with country rocks is a prerequisite for the development of the ore-forming brines. In contrast, Barton and Johnson (1996, 2000, 2004) argue that the global geological and paleogeographic setting and evolution of IOCG systems imply that the incursion of “exotic”, in part evaporite-sourced, brines is essential to economic Cu (-Au) mineralization. The involvement of diverse fluids has indeed been recognized in several sulphide-rich IOCG deposits, including Olympic Dam (Haynes et al., 1995; Johnson and

McCulloch, 1995), Ernest Henry (Mark et al., 2006; Kendrick et al., 2007), Tennant Creek

(Skirrow and Walshe, 2002), Tjårrojåkka (Edfelt et al., 2005) and Sossego (Monteiro et al., 2008).

Moreover, light stable isotope and fluid inclusion data specifically supporting the involvement of evaporite- or seawater- derived brines during Cu (-Au) mineralization have been reported by

Ullrich and Clark (1999) and Ullrich et al. (2001) for La Candelaria (but cf. Marschik and

Fontboté, 2001; Pollard, 2006), by Ripley and Ohmoto (1977) and de Haller et al. (2002) for

Raúl-Condestable, by Wanhainen et al. (2003) for Aitik, by Hunt et al. (2005, 2007) for the

Wernecke breccias, and by Benavides et al. (2006, 2007) for Mantoverde and its satellites.

Predictably, therefore, a commodious classification which incorporates magmatic and

27 non-magmatic hydrothermal fluid origins has been recently advocated for IOCG deposits by

Williams et al. (2005).

A radically different perspective on the genesis of IOCG mineralization is provided by the proposal that the majority of magnetite-dominated, so-called “Kiruna-type” (Geijer, 1931), deposits are the product of silica-poor, iron oxide - rich melts (e.g., Nyström and Henríquez, 1994;

Naslund et al., 2002; Henríquez et al., 2003), the existence of which has considerable field, petrographic and experimental support (e.g., Park, 1961; Philpotts, 1967; Lundberg and Smellie,

1979; Lester, 2002; Clark and Kontak, 2004; Lledó, 2005). Whereas the contribution of oxide melts to IOCG genesis remains controversial (cf. Rhodes et al., 1999; Sillitoe and Burrows, 2002,

2003), Naslund et al. (2002) argue that the extensive hydrothermal alteration associated with bodies of inferred magmatic magnetite directly reflects a high volatile content which permits dense oxide melts to rise to the shallow crust. Moreover, although their involvement has not been confirmed in IOCG deposits with economic Cu-Au mineralization, it is possible that oxide melts may have been the source of either hypersaline melts generated through Cl saturation

(Broman et al., 1999) or brines released on second boiling (Naslund et al., 2002).

Our purpose herein is to contribute to this argument through documentation of the

Marcona-Mina Justa district, south-central Perú, an Andean IOCG centre with both major iron oxide and significant Cu sulphide mineralization. Representing the largest-known concentration of high-grade magnetite ore in the Central Andes (Table 1-1), the Marcona deposit is centred in

Ica Department at Latitude 15o 12' S, Longitude 75o 7' W (Figs. 2-1 and 2-2), 10-15 km from the

Pacific coast and at elevations of below 800 m a.s.l. Hosted by Paleozoic metasedimentary and

Jurassic andesitic and sedimentary strata, it has present reserves of 1551 Mt grading 55.4 percent

Fe and 0.12 percent Cu, and had produced 358 Mt of ore by 2002 (Shougang Hierro Perú SA.,

2003). Annual magnetite production in recent years has exceeded 2 Mt of magnetite concentrates, attaining 4.5-5.0 Mt in 2005 and 2006. Cu, Co, Ni, Zn, Pb, Ag and Au are enriched in parts of the 28 deposit but, with the exception of artisanal copper mining, have not been recovered. The sulphur and copper contents of magnetite pellets are restricted to, respectively, 0.010 and 0.015 weight percent through flotation.

3-4 km NE of the Marcona mine (Fig. 2-2), the Mina Justa Cu-(Ag) prospect has an indicated resource of 346.6 Mt at an average grade of 0.71 percent Cu (soluble and insoluble), ~

3.8 g/t Ag and ~ 0.03 g/t Au at a cut-off grade of 0.3 percent Cu, and an inferred resource of

127.9 Mt at 0.6 percent Cu (Mining Journal, 2006). Located at Latitude 15o 10' S, Longitude 75o

5' W, the deposit was discovered by Rio Tinto Mining and Exploration and is under development by Chariot Resources. The wider district includes (Fig. 2-1) a second giant magnetite deposit,

Pampa de Pongo, hosted largely by Jurassic andesites and intercalated sedimentary strata.

Located 30 km southeast of Marcona-Mina Justa (Fig. 2-2; Hawkes et al., 2002), this incorporates an inferred resource of 953 Mt grading 44 percent Fe and with erratically distributed chalcopyrite and gold (Cardero Resource Corporation, 2005). In addition to these exocontact deposits, numerous magnetite and/or hematite-rich veins, some rich in Cu and Au, cut dioritic - to - monzogranitic plutons of the mid-Cretaceous Coastal Batholith in the Acarí - Cobrepampa district

(Fig. 2-2; Caldas, 1976; Injoque, 1985), which includes the formerly productive La Argentina mine.

Although Hudson (1974) interpreted the Marcona magnetite deposit as a partially remobilized,

Paleozoic sedimentary ironstone, it has been generally considered to represent an epigenetic replacement-type or skarn system (e.g., Atchley, 1956; Atkin et al., 1985; Injoque, 1985; Vidal et al., 1990). Injoque (2002) and Hawkes et al. (2002) assigned it to the IOCG clan but, despite extensive study, key aspects of the mineralization have been neglected, in particular the forms and contact relationships of the numerous orebodies which constitute the magnetite reserve.

Moreover, the heterogeneity of the metasedimentary and volcanic host-rocks has hindered the recognition of alteration/metasomatism directly related to mineralization. We herein document in 29

Figure 2-2. Geology of the Marcona-Mina Justa district (modified from Caldas, 1978; Hawkes et al., 2002)

30 detail the magnetite bodies and their host-rocks, in the context of a revised paragenetic model,

40Ar/39Ar geochronology and light stable isotopic constraints on the temperatures of ore formation.

We specifically address the possible contribution of Fe oxide melts to the formation of this largest

Andean IOCG centre, and attempt to identify the timing of the magmatic-hydrothermal transition.

Complementary studies of fluid inclusion microthermometry and chemistry, and an assessment of the hydrothermal reservoirs, will be reported elsewhere. Detailed documentation of the paragenetic and age relationships of the Mina Justa Cu (-Ag, Au) deposit incorporates the observations of Moody et al. (2003) and Baxter et al. (2005).

2.3 Regional and District Geological Setting

The wider Marcona-Mina Justa district (Figs. 2-1 and 2-2) exposes erosional remnants of a succession of volcano-plutonic arcs which regionally range in age from latest-Triassic to

Holocene, evidence for a protracted but episodic history of supra-subduction zone magmatism along the convergent margin of the South American Plate. However, the local magmatic record is dominated by Middle Jurassic volcano-sedimentary and hypabyssal units and by mid-Cretaceous granitoid plutons, both associated with IOCG-type mineralization. Stratigraphic relationships in the wider Marcona area, incorporating data from Caldas (1978), Vidal et al.

(1990), Hawkes et al. (2002) and this study, are summarized in Figure 2-3.

Pre-Mesozoic units

The IOCG belt of southern Perú is underlain by high-grade metamorphic rocks of the allochthonous Paleoproterozoic-to-Mesoproterozoic (Wasteneys et al., 1995; Loewy et al., 2004)

Arequipa Massif, the northern part of the Arequipa-Antofalla basement domain, which was accreted onto the Amazonian Craton during the ca. 1.0-1.3 Ga Grenville-Sunsas orogeny (Loewy et al., 2004; Chew et al., 2007). Comprising schists, gneisses, granites and migmatites cut by

31 basic and pegmatitic dikes, the basement complex is unconformably overlain by Neoproterozoic and Paleozoic sedimentary strata and, more extensively, volcanic and sedimentary rocks of

Mesozoic age (Fig. 2-2; Caldas, 1978; Hawkes et al., 2002). The Neoproterozoic strata comprise

Figure 2-3. Summarised stratigraphic column for the Marcona-Mina Justa district. (modified after Caldas, 1978, Injoque, 1985; Hawkes et al., 2002, and Loewy et al., 2004). 32 the ca. 700Ma glacial diamictites of the Chiquerío Formation and the San Juan Formation, its dolomitic cap (Chew et al., 2007). The overlying, unfossiliferous but probably Ordovician,

Marcona Formation (Caldas, 1978) hosts the majority of the economic magnetite orebodies at

Marcona. This ~ 1,500 m-thick metasedimentary package was described by Atchley (1956) as dominated by phyllites, “hornfelsic phyllites” and hornfelses, but clastic tectures are widely preserved and spaced cleavage and decussate microscopic textures are more extensively developed than penetrative foliation. Recognizable protoliths include siltstones, sandstones and minor quartz arenites, all dominated by quartz with lesser interstitial plagioclase, biotite, sericite and chlorite (Fig. 2-4A). Dolomitic marble and crystalline limestone occur in the north and southeast parts of the deposit, but are not major ore-hosts. This unit was considered by Atchley

(1956) to be overlain by a ~ 850 m-thick succession of metasedimentary and metatuffaceous strata assigned to a Paleozoic Cerritos Formation, a term still employed on mine geological maps

(Marcona Mining Company, 1968). However, the great majority of the rocks thus described are exceptionally rich in magnetite, actinolite and alkali feldspar, and are herein considered to represent either metamorphosed and metasomatized Marcona Formation lithologies or, more widely, strongly metasomatized Jurassic andesites and intercalated volcaniclastic and clastic sediments. The Marcona Formation is intruded by monzogranites, granodiorites and gabbro-diorites of the post-kinematic, 425 ± 4 Ma (Mukasa and Henry, 1990; Vidal et al., 1990),

San Nicolás batholith (Fig. 2-2). In the mine area metaclastic and metacarbonate members widely exhibit cordierite + biotite ± muscovite (Fig. 2-4B) and tremolite ± quartz (Fig. 2-4C) assemblages, respectively, but fine-grained diopside- and forsterite felses also occur in apparently unmetasomatized sections. Nonetheless, neither Atchley (1956) nor Injoque (1985) ascribed these apparent hornblende hornfels and pyroxene hornfels assemblages to thermal metamorphism by the San Nicolás intrusions.

Figure 2-4. Major host-rocks of the Marcona and Mina Justa deposits. A - Weakly-altered Marcona Formation arkosic siltstone: fine-grained original quartz clasts (>75%) and finer-grained interstitial plagioclase, with minor secondary chlorite and sericite (MA3-4-A; Mina 3 open pit, south wall, 580 m; transmitted light, crossed nicols). B - Cordierite + biotite ± muscovite hornfels, Marcona Formation (MA3-4-B; Mina 3 open pit, south wall, 580 m; plane-polarized transmitted light). Cordierite occurs as ovoid poikiloblasts with muscovite + quartz envelopes. Biotite is distributed in matrix. These assemblages commonly form mafic laminae in arkosic siltstone. C - Tremolite ± quartz in metacarbonate horizon of Marcona Formation. Late magnetite + actinolite assemblage replaced metamorphic tremolite and residual calcite (DDM4-6-1; drill hole DDM4-6; 124 m; transmitted light, crossed nicols). D - Amygdular Rio Grande Formation andesite hosting the Mina Justa orebodies, with vesicles filled by chlorite and carbonates. Plagioclase (An30) and hornblende occur as phenocrysts (MJ-4; surface sample; trench in the upper zone; scanned polished-thin section). E - Ocöite dike cutting Mina Justa orebodies. Plagioclase (An~50; ≤ 1.5 cm) and hornblende occur as phenocrysts, with intense sericite-chlorite alteration. The matrix is composed by fine-grained plagioclase, quartz, augite, hornblende and minor magnetite (MJ-2; surface sample; trench in the upper zone; plane-polarized transmitted light). F - Dacite porphyry at Marcona. Plagioclase (An30-40) commonly occurs as phenocrysts. Hornblende, quartz and fine-grained plagioclase constitute the groundmass (MA7-10; Mina 7 open pit, NW wall, 680 m; transmitted light, crossed nicols)

Mesozoic Stratigraphy

The Marcona iron deposit is, at least in part, of mid-Jurassic age (K-Ar data of Injoque, 1985) and is partially hosted by Middle Jurassic shallow-marine sedimentary and volcanic strata. The relative ages of these units and the mineralization are therefore critical to an understanding of the environment of ore formation. Caldas (1978) comprehensively documented the Mesozoic stratigraphic relationships in the San Juan 10,000 quadrangle, in which both Marcona and Mina

Justa lie, and in the contiguous Acarí and Yauca quadrangles. The Jurassic and Cretaceous strata of the wider Marcona area are subdivided (Figs. 2-2 and 2-5), with decreasing age, into the

Río Grande, Jahuay, Yauca and Copara Formations (Caldas, 1978). The ages of the three older formations are well established on faunal grounds, but those of the Copara Formation and the dominantly hypabyssal andesitic-dacitic Bella Unión complex which intrudes it, as well as the post-Yauca Formation hypabyssal Tunga Andesite, are poorly defined (Caldas, 1978).

The Río Grande Formation hosts the Mina Justa deposit and several orebodies of the

Marcona mine (Injoque, 1985; Hawkes et al., 2002; Moody et al., 2003). The type-section of this ~ 3000 to 4000 m-thick, generally NE-striking and NW-dipping, succession is exposed in the

Monte Grande area in the Cañón Río Grande, NW of Marcona (Fig.2-2; Rüegg, 1956, 1961). It incorporates (Fig. 2-5) a 500 m lower member comprising a polymictic basal conglomerate overlain successively by mudstones, sandstones, limestones, rhyolitic to andesitic breccias and rhyolitic to andesitic flows (Romeuf et al., 1993). This association, probably corresponding to

Atchley’s (1956) “Cerritos Formation” at Marcona (Fig. 2-5), is overlain by at least 2000 m of gently folded red sandstones, shales, limestones and brecciated andesitic flows with high-K calc-alkaline – to – shoshonitic compositions (Fig. 2-4D; Aguirre, 1988). The formation is quasi-pervasively affected by non-deformational zeolite facies metamorphism (Aguirre and

Offler, 1985; Aguirre, 1988).

Figure 2-5. Schematic stratigraphic columns of the Río Grande Formation in the Cañón Río Grande, Marcona and Pampa de Pongo areas (Atchley, 1956; Rüegg, 1956; Caldas, 1978; Injoque, 1985; Aguirre, 1988).

Faunal assemblages (Cox, 1956; Rüegg, 1956) define a Middle Jurassic (Dogger) age for the lower, largely sedimentary, member of the formation, and the occurrence of planammatoceras and hammatoceras (W.J. Arkell, in Rüegg, 1956) and the Bredya manflasensis faunal assemblage

(Roperch and Carlier, 1992) indicates that sedimentation was underway in the Aalenian, i.e., prior

+1.0 +1.0 to 174.0 -7.9 Ma and after 178.0 -1.5 Ma (Pálfy et al., 2000). Río Grande Formation volcanism

+3.1 persisted into the Oxfordian (Cox, 1956; Rüegg, 1956; Caldas, 1978) i.e., ca. 156.5 -5.1 to 154.7

37 +3.8 -3.3 Ma (Pálfy et al., 2000), but sedimentation and volcanism in the upper part of the formation were apparently interrupted between ca. 166 and 164 Ma (Fig. 2-5). The very low-grade metamorphism affects andesitic units above this unconformity, and therefore either took place, or resumed, after the Bathonian.

In the Pampa de las Treinta Libras area NE of Marcona, the overlying ~ 1,000-m - thick,

NW-striking, Jahuay Formation (Rüegg, 1961; Figs. 2-2 and 2-3) incorporates a basal, largely andesitic succession, overlain by a sequence of limestones and sandstones, the latter exhibiting lateral transitions to conglomerates, intercalated with dacitic porphyry flows and cut by sills and plugs of andesite. The upper members of the formation are dominantly shales and limestones.

+2.5 +3.4 Faunal assemblages extend from the lower to the upper Tithonian, i.e., 141.8 -1.8 to 150.5 -2.0 Ma

(Pálfy et al., 2000), but sedimentation may have begun earlier in the late Kimmeridgian (Caldas,

1978). Variable discordance with the underlying Río Grande Formation is interpreted by Caldas

(1978) as the result of the development of a tectonic dome, centred on the Marcona mine area and recorded by thick conglomerates in the Jahuay Formation and the local absence of Tithonian strata. The Jahuay Formation accumulated in the NW-trending pre-Andean depression (Caldas,

1978).

The succeeding Neocomian Yauca Formation crops out (Fig. 2-2) to the east of the Marcona deposit and comprises 1,500 m of strongly-faulted shales, mudstones and sandstones (Caldas,

1978; Injoque, 1985; Hawkes et al., 2002). The ~ 1000 m thick, probably Aptian to lower Albian,

Copara Formation, unconformably overlying the Yauca Formation, is composed of conglomerates with volcanic clasts, feldspathic sandstones, graywackes and shales (Caldas, 1978;

Injoque, 1985), underlying a dominantly intermediate pyroclastic sequence, all intruded by the andesitic Bella Unión complex. The Copara Formation accumulated in the apparent southern extremity of the Cañete Basin (Atherton and Aguirre, 1992). This mid-Cretaceous, plate

38 boundary-parallel, intra-arc extensional feature is, to the northwest, infilled by calc-alkaline, high-alumina and K-rich andesites and dacites.

The dike swarms, sills and small plugs assigned to the Tunga Andesite intrude the Yauca

Formation and older units (Caldas, 1978; Fig. 2-3). The most characteristic lithology is a coarsely

– porphyritic rock with large (≤ 1.5 cm) glomerocrysts of labradorite and sparse augite phenocrysts, termed “ocöites” by Rio Tinto geologists (Hawkes et al., 2002) by analogy with the broadly contemporaneous, strikingly porphyritic andesites of the Ocoa Formation in the Copiapó area of northern Chile (Thomas, 1958). Essentially identical textures are, however, shown by several Río Grande Formation andesitic flows in the Mina Justa area, a potential source of stratigraphic confusion.

Coastal Batholith

Granitoid plutons of the Coastal Batholith (Pitcher et al., 1985) intrude Neocomian and older strata in the Acarí-Cobrepampa area (Fig. 2-2; Dunin-Borkowski, 1970; Caldas, 1978). U-Pb zircon age data are lacking for this part of the Arequipa segment of the batholith, but K-Ar

(Cobbing, 1998) and Rb-Sr (Sánchez, 1982) dates for, respectively, the Acarí diorite and

Cobrepampa monzonite-monzogranite suggest that large-scale intrusion locally began at ca. 109

± 4 Ma, shortly after emplacement of the Bella Unión complex. Small, undated, dioritic stocks,

7-8 km ESE and SE of the Mina Justa prospect (Caldas, 1978), may be correlative with the larger intrusions to the east.

Cenozoic stratigraphy and landforms

The Mesozoic and older strata and intrusions are overlain discordantly by a Tertiary sequence locally assigned to the Pisco, Millo and Sencca Formations (Fig. 2-2). The 450 m-thick, Miocene to Pliocene, Pisco and Millo Formations, composed of shallow-water marine sediments and

39 subordinate fine-grained volcaniclastics (Caldas, 1978; Hawkes et al., 2002), extend from the

Pisco basin in the north (Devries, 1998) to the south of Lomas (Fig. 2-2), where they contain abundant whale fossils (Brand et al., 2004). The Lower Miocene volcaniclastics of the Pisco

Formation record an early Neogene pulse of volcanic activity recognized throughout the Central

Andes (Noble et al., 1990). A vestigially preserved, ≤ 50m, rhyodacitic ash-flow tuff, assigned arbitrarily to the Sencca Formation (Caldas, 1978; Injoque, 1985), gives a 40Ar/39Ar biotite plateau age of 9.13 ± 0.25 (2 σ) Ma (Quang et al., 2001). This unit overlies the Loma de Marcona

Surface, the youngest of four regionally developed erosional pediments preserved in the district.

Subsequently, the area was depressed below sea-level, blanketed by uppermost Pliocene beach deposits (Ortlieb and Macharé, 1990), and again uplifted, a process recorded by a remarkable series of tectonically-controlled marine terraces (Broggi, 1946), the younger of which formed in response to the local subduction of the aseismic Nazca Ridge.

2.4 The Marcona Magnetite Deposit

The Marcona mine now comprises eight open pits in a ~ 25 km2 area elongated from WNW to

ESE (Fig. 2-2). Traditionally, a crudely en echelon array of 12 major magnetite orebodies

(“minas”) and 55 smaller “cuerpos” are recognized (Fig. 2-6). However, the three zones exploited by the largest, 3 km-long pit, viz. Mina 2, Mina 3 and Mina 4, represent interconnected segments of a single orebody (Table 2-2). Approximately 60 per cent of the reserve, comprising the so-called “E-grid” orebodies, is hosted by the Marcona Formation (Fig. 2-4B), and the remainder, the N-13 type orebodies, by the lower members of the Río Grande Formation (i.e.,

Atchley’s (1956) Cerritos Formation; Fig. 2-7). The hypogene grades (Table 2-2) of the larger orebodies hosted by the Paleozoic metasediments average 57-58 per cent Fe, significantly exceeding those of 41-48 per cent for the orebodies in Jurassic strata. Whereas the total sulphur content of the mineralization is consistent at ~ 3 wt. percent, the copper content is more variable,

Figure 2-6. Geology of the area surrounding the Marcona deposit and Mina Justa prospect. Line A-A’ illustrates the cross-section (see Figure 2-7) through the Marcona mine (Modified from Rio Tinto, Marcona JV exploration report, June 2003). Insert shows area of Figure 2-28 41

Figure 2-7. Schematic cross-section of Marcona mine area (A-A’ in Figure 2-6). Ornaments as in Figure 2-6. The magnetite orebodies are extensively dislocated by faults (modified from Hawkes et al., 2002).

Table 2-2. Selected Tonnage/Grade Data for Marcona Orebodies 1

orebody Minas 2-3-4 Mina 5 Mina 7 Minas 9-10 Mina 14 Mina 11 Mancha N-13 Host rock Marcona Marcona Marcona Marcona Marcona Marcona Río Grande Fm. Fm. Fm. Fm. Fm. Fm. Fm. Reserve (Mt)2 399 190 18 110 110 35 224 Fe grade (%)3 58.5 60.2 57.3 58.1 57.0 54.4 41.9 Cu grade (%)3 0.17 0.06 0.06 0.11 0.08 0.45 0.04 S content (%)3 3.55 2.57 3.10 2.51 2.97 3.51 2.86 Zn grade (%)3 ------0.5 ------1) from “The Resource Estimate of the Marcona Iron Mine”, Shougang Hierro Perú, unpub. report, 2003 (in Chinese); 2) in 2003; 3) Fe, Cu, Zn and S grades of hypogene ore, Pb grade is not available.

averaging 0.06-0.18 percent, but attaining 0.4 percent in Mina 1 and 0.9 percent in the upper part of the easternmost, Mina 11, orebody (Fig. 2-6). Pyrrhotite occurs mainly in the lower, and chalcopyrite in the upper levels of the orebodies. As exemplified by the schematic cross-section of the Mina 4 orebody (Fig. 2-8A), most orebodies at Marcona exhibit higher Cu grades as well as elevated total sulphide contents in their upper parts, although sulphides are locally enriched in the lower parts of some orebodies. Sphalerite and galena, normally subordinate to chalcopyrite, are abundant in the Mina 14 orebody.

The large-scale relationships of the major magnetite orebodies are here illustrated by photographs taken in the open pits, the locations of most of which are shown in Figures 2-9A and

B. The orebodies are dominated by essentially massive magnetite, and most original contacts

Figure 2-8. A - Cross-section of the Mina 4 orebody, Marcona. Copper grade distribution on the right is for the >50% Fe orebody. Porphyritic andesite and andesitic-basaltic dikes are common. The main and subsidiary orebodies are controlled by NE-striking and NW-dipping Repetición faults, and displaced by later Mina Justa and Huaca system faults. B - Cross-section of the Mina 1 orebody, Marcona. Two sets of Repetición faults are recognized, one post-mineralization and displacing the orebody, in turn cut by Mina Justa system faults, the other controlling the emplacement of the orebody and dacite intrusions. Porphyritic andesite dikes are displaced by late Repetición and Huaca faults. Locations of sections are shown in Figure 2-6 (modified after Shougang cross-sections of Mina 4 and Mina 1, 2004)

with both Paleozoic and Jurassic host rocks are abrupt, only locally complicated by disseminated mineralization, stockwork veining or hydrothermal breccias. Despite extensive segmentation by

Figure 2-9. Three-dimensional imaging of the areal relationships of the E-Grid Marcona magnetite orebodies. A - Panorama of the main open pits (image from Google Earth 2006; looking north). Locations and directions of observation of subsequent views are shown. All orebodies strike NE and dip NW, parallel to the Repetición faults. B - Looking SW along the SW half of the 3 km-long Minas 2-3-4 pit (image from Google Earth 2006). 44 post-mineral faults (Figs. 2-7 and 2-8), the original forms of the orebodies, i.e., lensoid to tabular in both plan and vertical section, are evident (Fig. 2-10). The 40-60o strikes and ca. 45° NW dips of the orebodies reflect control by the dominantly reverse Repetición faults, which strike ~ 45o and dip 30-60o to the north (Fig. 2-6), but those with steeper dips are widely subparallel to stratigraphy. The Repetición faulting persisted after both magnetite and sulphide mineralization, displacing the orebodies by up to 70 m (Fig. 2-8B). Where the hanging walls of the orebodies are exposed, as in the Mina 11 orebody, dike-like bodies of massive magnetite (± actinolite) transect the foliations and folds in the Marcona Formation and extend for over a hundred metres (Fig.

2-11). On a smaller scale, particularly in the Marcona Formation, the boundaries of the main magnetite bodies have smoothly curving contacts (Figs. 2-12A, B), from which extend crudely tubular apophyses (Fig. 2-12C). These, in turn, are flanked by spheroidal, apparently isolated magnetite bodies (Fig. 2-12D).

The mineralized area is intruded by a swarm of hypabyssal bodies (Fig. 2-8). These range from apparently syn- to clearly post-mineralization and, in composition, from silicic to, rarely, ultramafic (hornblende pyroxenite: Atchley, 1956), but magmatic chemistry and mineralogy are almost everywhere disguised by alteration. Thus, some units described as porphyritic andesites which cut the orebodies contain relict labradoritic plagioclase and probably represent basalts or basaltic andesites. However, other late-stage, 20-50 m wide, dikes, assigned to the Tunga suite, are unambiguously andesitic (Fig. 2-4E). The most widespread intrusive units are dacite or rhyodacite porphyries. Injoque (1985) referred to these as trachyandesitic on the basis of whole-rock geochemistry, but his descriptions of analyzed samples record extensive biotitization.

The original mineral assemblage (Fig. 2-4F) comprised phenocrystic plagioclase (An30-40), hornblende and quartz set in a granophyric groundmass. Injoque (1985) interpreted immobile trace element suites as evidence for a “transitional volcanic arc” environment. He further inferred

Figure 2-10. NE-striking, NW-dipping Mina 7 orebody, controlled by Repetición faults and both cutting and paralleling the foliation in the Marcona Formation host-rocks; looking SW.

Figure 2-11. Dike-like apophyses of magnetite extending from the hanging wall of the Mina 11 orebody and cutting strongly foliated and folded Marcona Formation (Photograph of SE end of open pit taken in 2001).

Figure 2-12. Megascopic features of the contacts between magnetite orebodies and Marcona Formation host rocks in the Marcona mine. A, B - Smoothly curving contacts (Mina 3 open pit, 650 m bench; arrow points to a backpack in A). C - Cross-sections of crudely tubular magnetite bodies (Mina 3 open pit, 620 m bench; arrow points to a hammer). D - Apparently spheroidal, 8m diameter, magnetite body (Mina 2 open pit, 680 m bench).

that the dacitic bodies were feeder dikes for the basal felsic flows of the Río Grande Formation, but their intimate association with the main magnetite orebodies is evidence for a younger age.

Rare in the Río Grande Formation, the porphyry bodies form a dense network of elliptical to amoeboid bodies in the underlying Marcona Formation (Fig. 2-13A). Although some porphyry bodies cut unmineralized metasediments, several of the main orebodies exhibit an intimate interfingering of massive magnetite and porphyry. This is examplified by Minas 2-3-4, where steeply N to NW - dipping bodies of dacite are interlaminated with massive magnetite (Fig.

2-13B). Here, the dacite bodies strike ~ NE, and dip concordantly with the tabular magnetite orebodies. Many do not extend beyond the orebodies to the WNW or ESE, the magnetite and dacite bodies therefore constituting a composite mass surrounded by, and intrusive into, the 47

Figure 2-13. Large-scale relationships between magnetite orebodies and dacite porphyry in the Marcona deposit. A - Simplified map showing the main orebodies and dacite intrusions. B - A close-up of Minas 2-3-4 showing the intimate interfingering, amoeboid, relationship between the orebodies and steeply N-dipping bodies of dacite porphyries. Vertical projection of Figure 2-14B is shown (modified from Geology and Drill Hole Locations, Marcona JV Rio Tinto-Shougang report, Rio Tinto Mining and Exploration Limited, 2003)

Marcona Formation. The contacts between porphyry and massive magnetite are abrupt and range from crudely planar (Fig. 2-14A), through smoothly curving to convoluted (Fig. 2-14B). In several areas, coherent rounded bodies of dacite are enclosed in magnetite; elsewhere, rounded proturbances of magnetite extend into dacite (Fig. 2-15). Whereas locally such contact

Figure 2-14. A - A panorama showing crudely planar or curviplanar contacts between magnetite ore zones, dacite porphyry and the Marcona Formation (Mina 3 open pit, south wall, 2001; 629 m bench, looking SE). B - Part of the Mina 3 open pit south wall (September 2004) showing the convoluted contacts between massive magnetite bodies (Mt) and their host rocks, including dacite (Da) and metasediments of the Marcona Formation. A late diabase (Dia) dike cuts a magnetite orebody and dacites. The magnetite body on the right is truncated by a Mina Justa system fault. 49 relationships reflect late tectonic modification of the margins of magnetite bodies, in others there is no evidence of such post-mineral disturbance. In the Marcona Formation, dike-like bodies comprise angular blocks of massive magnetite and ellipsoidal bodies of finer-grained magnetite in a matrix of friable magnetite (Fig. 2-16). Much of the magnetite mineralization is made up of intergrown subhedral crystals, but locally a vuggy texture is developed, with euhedral octahedra projecting into voids (Figs. 2-17A, B).

Figure 2-15. Crudely spheroidal bodies of dacite porphyry are enclosed by massive magnetite at the contact between the Mina 3 magnetite orebody and dacite porphyry (arrow points to a 30 cm long hammer). The contacts are generally abrupt and smoothly curving (Mina 3 open pit, 620 m, south wall, see Figure 2-9A).

Fault systems

Three principal fault systems were documented in the Marcona mine by Atchley (1956) and

Hawkes et al. (2002), but new observations show that at least four are represented. The oldest, the

Pista normal faults, including the regionally extensive Lechuza Fault (Fig. 2-2), strike ~295° and dip ~60° to the north, have been ascribed to detachment along the boundaries of uplifted blocks 50

Figure 2-16. Angular blocks of massive magnetite enclosed in a matrix consisting of ovoid bodies of fine-grained magnetite (Mina 1 open pit, centre part, 700 m bench)

Figure 2-17. A - Vuggy texture of magnetite (MA3-22, Mina 3 open pit, south wall; 650 m). B – SEM image showing that euhedral octahedral magnetite crystals occur as aggregates in voids. Striations are locally developed on the crystal surfaces.

of pre-Mesozoic basement (Pope, 2003). Together with the coeval or younger Repetición faults, they are inferred to record ESE-WNW contraction during the Jurassic, perhaps linked to sinistral shear along the regionally important, NW-trending, Treinta Libras fault zone northeast of

Marcona (Figs. 2-2 and 2-6). The faults controlling the main magnetite bodies and Cu

51 mineralization at Mina Justa, herein termed Mina Justa Faults, have strike directions similar to those of the Repetición faults at Marcona, but they dip shallowly southeast and show normal displacement. They have been recognized herein in the Marcona mine where they segment the orebodies, and also locally show reverse displacement (Figs. 2-8A and 2-18). These faults may record a change to dextral transtension on the Treinta Libras fault zone (Pope, 2003). The youngest, Huaca, normal faults strike ~335° and dip ~60° to the east. They are post - mineralization at both Marcona and Mina Justa, but are commonly followed by porphyritic andesitic dikes (ocöite).

Figure 2-18. Dislocation of magnetite orebody (Mt) by late Mina Justa and Huaca system normal faults (Mina 4 open pit, 600 m bench, looking NW, see figure 2-9A) (arrow points to a backpack).

Supergene effects

Most magnetite orebodies in the Marcona mine were overlain by 10-40 m-thick supergene oxidation profiles comprising a lower 4-6m thick horizon of sulphate-rich “transitional” ore (Fig.

2-8A) in which martitized magnetite is intergrown with jarosite, botyrogen, amarantite and

52 parabutlerite, and a surficial leached, martite-dominated zone. Geomorphological relationships show that oxidation largely preceded erosion of the Loma de Marcona pediment and hence occurred before ca. 9 Ma.

2.5 Paragenetic Relationships of the Marcona Orebodies

Numerous stages of hydrothermal alteration and mineralization are recognized herein, largely on the basis of megascopic and microscopic textural relationships and mineral assemblages (Fig.

2-19). However, as in other IOCG deposits, paragenetic relationships are locally ambiguous, in part owing to the repeated development of hydrothermal amphibole, trioctahedral mica and albite.

Figure 2-19 therefore incorporates evidence for the ages of alteration assemblages provided by

40Ar/39Ar geochronology, as discussed below. Representative electron microprobe analyses of alteration minerals are recorded in Table 2-3, complementing the data of Injoque (1985). To avoid confusion, paragenetic stages at Marcona are prefixed by “M” and those from Mina Justa by “J”. In addition, those from the Marcona Mina 11 orebody (Fig. 2-6) are distinguished as

“M11”. Alteration facies are designated according to the mineral assemblages generated, without implication for metasomatic exchange relationships.

Stage M-I: Early Mg-silicate alteration. Felted aggregates of fine-grained cummingtonite

(Table 2-3; Figs. 2-20A-C) are assigned herein to paragenetic Stage M-IA (Fig. 2-19). Compared to cummingtonite of igneous origin (Deer et al., 1997), that at Marcona has higher SiO2 (57.6 vs.

52.3 weight percent) and MgO (21.0 vs. 14.8 weight percent), lower CaO (0.13-0.69 vs. 2.27 weight percent), Na2O (0.13 vs. 0.49 weight percent) and Al2O3 (0.80 vs. 3.02 weight percent), but similar K2O contents (0.09 vs. 0.07 weight percent). Metamorphic cummingtonite (Deer et al.,

1997) exhibits wide compositional ranges, overlapping with that at Marcona in CaO, Al2O3, MgO,

K2O, N2O and F contents, but with lower SiO2 (53.0-55.6 vs. 57.6 weight percent) and higher

FeO contents (18.9-25.0 vs.13.0-14.0 weight percent). The unambiguously hydrothermal

Figure 2-19. Alteration and mineralization paragenesis of the Marcona deposit.

54 Table 2-3. Representative Electron Microprobe Data for Alteration Minerals, Marcona Iron Deposit Mineral Cum Cum Gre Bt Bt Trm Trm Act Phl Phl Phl Phl Phl Phl Phl Act Trem Trem Trem Trem Trem Chl Chl Tlc Stage M-I M-I ? M-III M-III M-III M-III M-III M-III M-III M-IV M-IV M-IV M-IV M-IV M-V M-V M-V M-V M-V M-V M-VI M-VI M-VI Sample MA5-9 MA5-9 MA5-9 MA5-9 MA5-9 MA3-18 MA7-23 DDM3- MA3-18 MA3-18 MA3-19 MA3-19 MA3-19 MA3-19 MA3-19 DDM5 DDM5 DDM3 DDM3 DDM5 MA91 MA3-19 MA3-19 MA3-18 number II-2 I-1 II-3 II-1 I-2 II-2 II-2 3-1-II II-1 I-1 I-1 I-2 II-2 II-3 II-1 -4-2-2 -4-2-1 -3-8 -3-1-I -4-3 -1 I-4 I-4 I-2

SiO2 57.68 57.64 35.72 39.3 40.19 58.54 58.7 55.39 44.73 43.51 45.41 44.89 44.51 43.35 41.91 55.61 58.63 57.8 59.77 58.19 57.98 36.74 37.59 62.11

TiO2 0.01 n.d. 0.01 1.23 1.44 n.d. n.d. 0.06 0.03 0.04 0.07 0.05 0.32 0.69 1.62 0.02 0.01 0.02 0.05 0.05 n.d. n.d. 0.03 n.d.

Al2O3 0.83 0.8 0.09 12.46 12.32 0.56 0.48 2.43 10.99 12.62 11.2 10.88 11.43 12.73 13.09 0.22 0.33 1.17 0.29 0.73 0.37 10.23 8.83 0.37 FeO* 13.04 14.01 49.98 17.03 15.71 1.44 2.25 7.26 2.08 2.24 2.45 1.91 2.45 2.29 6.58 14.69 3.83 3.6 1.8 2.83 2.55 5.17 5.72 1.11 MnO n.d. 0.05 0.56 0.02 0.06 0.01 0.08 0.04 n.d. n.d. n.d. 0.04 0.01 0.02 0.07 0.11 0.06 0.08 0.08 0.05 0.05 0.08 0.06 0.02 MgO 21.12 20.91 2.84 15.11 16.07 23.51 23.01 19.08 27.25 26.42 27.17 27.44 26.81 26.15 22.03 14.42 22.15 22.33 23.49 22.92 21.59 33.63 33.52 29.68 CaO 0.69 0.13 0.03 n.d. n.d. 13.66 13.34 13.44 0.01 n.d. n.d. n.d. n.d. n.d. n.d. 12.81 13.27 13.9 13.73 12.68 13.45 0.11 0.1 0.04

Na2O 0.13 0.13 0.01 0.07 0.04 0.2 0.12 0.32 0.1 0.18 0.11 0.1 0.05 0.06 0.08 0.1 0.12 0.25 0.09 0.22 0.15 0.05 0.13 0.13

K2O 0.09 0.09 0.01 9.45 9.43 0.12 0.13 0.16 9.65 10.24 9.79 9.79 10.11 9.88 9.74 0.03 0.09 0.17 0.06 0.1 0.11 0.14 0.05 0.1 Cl 0.03 0.05 0.92 1.06 0.98 n.d. 0.02 0.05 0.14 0.18 0.14 0.16 0.12 0.12 0.3 0.01 0.03 0.04 0.01 0.02 0.02 0.05 0.03 0.01 F 0.35 0.28 0.01 0.53 0.73 0.47 0.4 0.37 1.97 1.55 1.77 1.8 1.64 1.59 1.14 0.31 0.35 0.54 0.53 0.58 0.58 0.66 0.78 0.71 Total 93.95 94.09 90.19 96.27 96.97 98.5 98.52 98.6 96.94 96.97 98.1 97.05 97.44 96.87 95.67 98.32 98.86 99.89 99.89 98.38 96.85 86.85 86.82 94.28

Si 8.28 8.29 4.20 5.91 5.96 7.96 7.99 7.74 6.24 6.08 6.25 6.25 6.19 6.05 5.97 8.03 8.00 7.85 8.01 7.96 8.05 7.03 7.21 8.02 Al 1) 0.00 0.00 2.09 2.04 0.04 0.01 0.26 1.76 1.92 1.75 1.75 1.81 1.95 2.03 0.00 0.00 0.15 0.00 0.04 0.00 0.97 0.79 0.00 Al 2) 0.14 0.14 0.01 0.12 0.11 0.05 0.07 0.14 0.05 0.16 0.07 0.03 0.06 0.15 0.17 0.09 0.06 0.04 0.07 0.08 0.16 1.33 1.21 0.10 Ti 0.00 0.00 0.00 0.14 0.16 0.00 0.00 0.01 0.00 0.00 0.01 0.01 0.03 0.07 0.17 0.00 0.00 0.00 0.01 0.01 0.00 0.00 0.00 0.00 Fe 1.57 1.68 4.92 2.14 1.95 0.16 0.26 0.85 0.24 0.26 0.28 0.22 0.29 0.27 0.78 1.77 0.44 0.41 0.20 0.32 0.30 0.83 0.92 0.12 Mn 0.00 0.01 0.06 0.00 0.01 0.00 0.01 0.01 0.00 0.00 0.00 0.01 0.00 0.00 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.00 Mg 4.52 4.48 0.50 3.39 3.55 4.76 4.67 3.98 5.67 5.50 5.58 5.69 5.55 5.44 4.68 3.10 4.51 4.52 4.69 4.67 4.47 9.59 9.59 5.72 Ca 0.11 0.02 0.00 0.00 0.00 1.99 1.95 2.01 0.00 0.00 0.00 0.00 0.00 0.00 0.00 1.98 1.94 2.02 1.97 1.86 2.00 0.02 0.02 0.01 Na 0.04 0.04 0.00 0.02 0.01 0.05 0.03 0.09 0.03 0.05 0.03 0.03 0.01 0.02 0.02 0.03 0.03 0.07 0.02 0.06 0.04 0.02 0.05 0.03 K 0.02 0.02 0.00 1.81 1.78 0.02 0.02 0.03 1.72 1.82 1.72 1.74 1.79 1.76 1.77 0.01 0.02 0.03 0.01 0.02 0.02 0.03 0.01 0.02 Cl 0.01 0.01 0.18 0.27 0.25 0.00 0.00 0.01 0.03 0.04 0.03 0.04 0.03 0.03 0.07 0.00 0.01 0.01 0.00 0.00 0.00 0.02 0.01 0.00 F 0.16 0.13 0.00 0.25 0.34 0.20 0.17 0.16 0.87 0.68 0.77 0.79 0.72 0.70 0.51 0.14 0.15 0.23 0.22 0.25 0.25 0.40 0.47 0.29 R** 0.74 0.73 1.6 1.8 0.97 0.95 0.82 23.6 21.2 19.9 25.9 19.1 20.1 6.0 0.64 0.91 0.92 0.96 0.94 0.94 0.1 0.1 Values are reported in weight percentage (%). * Total iron. Number of ions calculated on the basis of F, Cl, and 23 O for cummingtonite (Cum), tremolite (Trem) and actinolite (Act); 22 O for biotite (Bt) and phlogopite (Phl); 28 O for chlorite (Chl), 22 O for talc (tlc) and 14 O for greenalite. R**: Mg/Mg+Fe ratio for amphiboles; Mg/Fe ratio for micas, Fe/Fe+Mg for chlorites. 1) Al-tetrahedral; 2) Al-octahedral. n.d. – undetected

Figure 2-20. Stage I precursor alteration at Marcona. A - Cummingtonite (Stage M-IA), partially altered to greenalite, occurs interstitially to coarse-grained biotite and magnetite (Stage M-III) (# MA5-9, Mina 5 open pit, 670 m, main orebody; plane-polarized transmitted light). B - Fine-grained (< 10μm), felted aggregates of cummingtonite (Stage M-IA) occur with magnetite (# MA5-9, Mina 5 open pit, 670 m, main orebody; transmitted light, crossed nicols). C - Electron backscatter image illustrates the replacement of cummingtonite (darker) by fine-grained Stage M-III biotite (paler). D - Stage M-IB phlogopite and magnetite (Mt-1) replaced by late-stage pyrite and magnetite (Mt-2). Pyrite was emplaced along the cleavage of early phlogopite (#MA5-2, Mina 5 open pit, 670 m, main orebody; plane-polarized reflected light).

cummingtonite at Marcona has an Mg/Mg+Fe ratio of 0.74, exceeding those of most metamorphic (0.55 to 0.69) and all igneous (0.52) occurrences. The low anhydrous total, 94.1 weight percent (Table 2-3) is probably the result of alteration to greenalite (Fig. 2-20B), which would also account for the low SiO2 and MgO and high FeO and Cl contents (Table 2-3). The cummingtonite is also locally replaced by biotite and magnetite (Fig. 2-20C). Cummingtonite alteration is megascopically indistinguishable from the more widespread actinolitic facies, but is apparently restricted to the envelopes of orebodies in the upper Marcona Formation. 56 Coarse-grained phlogopite, intergrown with magnetite, but also replaced by magnetite and pyrite (Fig. 2-20D), talc and chlorite, also developed at an early stage in the alteration envelopes of Mina 5 and other orebodies hosted by the Marcona Formation. This magnesian mica alteration is assigned to Stage M-IB (Fig. 2-19).

Mass balance calculations have not been carried out, but the silicic nature of the protoliths indicates that alteration stages M-IA and M-IB record Mg-Fe metasomatism.

Stage M-II: albite-scapolite alteration. At Marcona, Na-Cl metasomatism widely generated albite and subordinate Na-rich scapolite, particularly in Marcona Formation siliciclastics and lower Río Grande Formation sedimentary units and andesites. Patches of coarse, white albite with clusters of bladed white scapolite are widely developed along the foliation of metaclastic host rocks, in places adjacent to bodies of massive magnetite (Fig. 2-21A). In such zones, scapolite is restricted to within 1-1.5 m of the magnetite bodies. Albite and scapolite do not occur within the latter, however, and much of this alteration is inferred to have predated magnetite mineralization and is assigned to alteration Stage M-II. Nonetheless, pink albite occurs around plagioclase phenocrysts (Fig. 2-21B) both in andesites and in the dacite porphyries which (see below) are interpreted as contemporaneous with Stage M-III magnetite mineralization (Fig. 2-19). Rock staining and X-ray study are commonly required to distinguish this pink albite alteration from the widespread K-feldspathization. Elsewhere, complete replacement of dacite by coarse-grained

(0.5-1 mm) albite generated white rocks. Albitization everywhere predated K-feldspar development which was, in turn, overprinted by actinolite-sulphide alteration (Fig. 2-21C).

Na-rich scapolite locally replaced original feldspars in dacite in contact with the orebodies, and, with a composition of Meionite29-38 and 2.8 to 3.3 weight percent Cl, has been identified in the lower Río Grande Formation north of the Marcona mine (Injoque, 1985), where it was subsequently replaced by amphibole and magnetite.

Figure 2-21. Marcona Stage M-II Na metasomatism. A - Albitization of Marcona Formation metasediments. Coarse white albite (locally with pockets of bladed scapolite) is concentrated along the foliation (Mina 2 open pit, 700 m level, south wall; see Fig. 2-14A). B - Albitization of fine-grained andesite. Pink-red albite occurs around white-grey relict plagioclase (pl) phenocrysts (# MA2-13, Mina 2 open pit, 700 m, ~ 50m from the main orebody; scanned polished hand-specimen). C - Stage M-II albitized dacite (white) are cut by Stage M-III K-feldspar (microcline; pink-red) veins which are reopened by Stage M-V actinolite (+ sulphides, dark-green). The major sulphide is pyrite (# MA3-24, Mina 3 open pit, 600m. ~ 30m from the main magnetite orebody).

Stages M-III and M-IV: main magnetite and magnetite-sulphide mineralization. Magnetite in the massive orebodies and local stockwork-breccia mineralization is associated with variable

58 proportions of calcic amphibole, phlogopite, biotite, K-feldspar, apatite, calcite, diopside and sulphides. Shougang Hierro Perú do not report minor or trace element chemistry of magnetite pellets, but Injoque et al. (1988) record up to 30 ppm Ti, 110 ppm V and 120 ppm Cr on the basis of electron microprobe analysis of representative magnetite grains. It is not, however, clear which paragenetic stages are reprsented by these data.

The major mineral associations in the main magnetite orebodies are magnetite - actinolite (or tremolite) ± phlogopite and magnetite-biotite (± actinolite), both assigned to sulphide-free Stage

M-III, and magnetite-actinolite (or tremolite)-sulphides (± apatite ± calcite), and magnetite - phlogopite - sulphides (± actinolite or tremolite ± apatite ± calcite), assigned to Stage M-IV. In both stages, magnetite occurs as euhedral to subhedral, 0.3-5 mm grains and as massive aggregates, commonly intergrown (Fig. 2-22A) with fine-grained (0.1-0.5mm), light-green tremolite (Mg/Mg+Fe = 0.95 to 0.97: Table 2-3) or dark-green actinolite (Mg/Mg+Fe = 0.82:

Table 2-3) in the classification of Leake et al. (1997). However, these amphiboles rarely coexist in the same hand-specimen. The major Stage M-IV sulphides are pyrite, chalcopyrite and pyrrhotite, both hexagonal and monoclinic, occurring largely as subhedral to anhedral crystals interstitial to magnetite, calcic amphibole and phlogopite (Fig. 2-22B), but locally showing microscopic replacement textures (Fig. 2-22C). Although no unambiguous replacement of pyrrhotite by pyrite and chalcopyrite was observed, the common association of pyrrhotite and magnetite without other sulphides, especially in the lower parts of the orebodies, and the absence of pyrrhotite in late sulphide veins, suggest that it largely formed prior to pyrite and chalcopyrite.

Lit par lit replacement by magnetite-sulphide-amphibole is displayed at the contacts of Stage

M-IV assemblages with Marcona Formation, but has not been observed for M-III, sulphide-free assemblages. Red-brown Stage M-III biotite (Ann55-63), with Mg/Fe between 1.6 and 1.8 and high

Cl (1.0-1.1 weight percent: Table 2-3), commonly occurs as coarse flakes in the main orebodies

Figure 2-22. Mineralogical and textural relationships of Marcona main magnetite stage (M-III) and magnetite-sulphide stage (M-IV). A - Stage M-III tremolite (with actinolite) occurs interstitially to subhedral magnetite. Chloritization of amphibole is common. (# MA3-22, Mina 3 open pit, 600 m, combined reflected and transmitted light). B - Stage M-IV magnetite, pyrite, tremolite and phlogopite. The smooth contacts suggest contemporaneous formation (# MA2-9, Mina 2 open pit, 600 m, orebody, combined reflected and transmitted light). C - Pyrrhotite and pyrite coexist with magnetite and actinolite, but locally pyrrhotite occurs as veinlets cutting magnetite crystals (# DDM4-6-4, drill core DDM4-6, 201 m, main orebody of Mina 4; plane-polarized reflected light). D - Coarse - grained Stage M-III phlogopite with interstitial magnetite and apatite. Chlorite and minor talc replace phlogopite along cleavages (# MA3-18, Mina 3 open pit, 600 m, orebody, transmitted light, crossed nicols). E - Stage M-III alteration in dacitic host rocks. Biotite-pyrite-quartz veins cut massive fine-grained biotite and actinolite-diopside-pyrite. Biotite in matrix and veins is red-brown but the latter is commonly subhedral to euhedral with good cleavage. Biotite and amphibole are partly sericitized and chloritized (# DDM5-4-9, drill core DDM5-4, 537 m, close to the contact between Mina 5 orebody and dacitic host rock; combined reflected and transmitted light). F - Electron backscatter image shows the local replacement of Stage M-III biotite (paler, massive aggregates and veins) by (darker) Stage M-IV phlogopite (# MA3-19, Mina 3 open pit, 600 m). G - Alteration at the contact between magnetite (Mt) orebody and dacite, zoned outward from biotitization to K-feldspathization and albitization (all with or without minor magnetite) (Mina 3 open pit, east end of south wall, 620 m). H - K-feldspar (-magnetite) alteration in dacite. K-feldspar commonly occurs as envelopes around original plagioclase (# MA14-5, Mina 14 open pit; 740 m, plane-polarized transmitted light).

61 and their envelopes. Phlogopite, locally occurrring in Stage M-III, but more abundant in Stage

M-IV assemblages, has a composition of Ann4-5 with Mg/Fe = 19.1-25.9 (Table 2-3), and high F

(1.6-2.0 weight percent; Table 2-3) in both the orebodies and their alteration envelopes. Generally coarse-grained, with dark green–light brown pleochroism, it is widely replaced by chlorite and talc (Fig. 2-22D). The contrasted F and Cl contents in the micas are characteristic of hydroxyl-bearing ferromagnesian silicates, in which members with high Mg/Fe ratios tend to incorporate more F than ferroan members, whereas Cl exhibits the inverse relationship, i.e., “Fe-F and Mg-Cl avoidance” (Munoz, 1984). Calculated log [fHCl / fHF] ratios (Munoz, 1984) for the crystallization of biotite and phlogopite have ranges of 3.4 to 3.5 and 1.8-3.1, respectively. Stage

M-III fluids were therefore fluorine-poor, but HF fugacity may have increased significantly as magnetite-sulphide deposition progressed and phlogopite replaced biotite in Stage M-IV.

Accessory minerals include fine-grained Stage M-III fluorapatite (Fig. 2-22D) and Stage M-IV calcite, both widely coexisting with magnetite and/or sulphides.

Biotitization and, more locally, phlogopitization were the major alteration processes in hanging wall and footwall Marcona Formation metasediments. Aggregates of dark-brown, massive, fine-grained (< 0.1mm) biotite are commonly cut by narrow brown biotite veins, some with pyrite and quartz. Locally, green actinolite-diopside alteration intervened temporally and spatially between these two biotite-bearing associations (Fig. 2-22E). Sericitization of biotite and chloritization of actinolite are common. Replacement of both massive and vein biotite by phlogopite is observed in electron backscatter images (Fig. 2-22F).

Stage M-III biotitic and K-feldspar alteration, with or without magnetite, is dominant in dacite and fine-grained andesite, biotitization commonly giving way outwards to the other alteration facies (Fig. 2-22G). Fine-grained tremolite or actinolite, with minor pyrite and magnetite, locally developed in dacite in contact with magnetite orebodies. Secondary K-feldspar in dacite generally has the orthoclase structure and is locally associated with fine-grained

62 magnetite (Fig. 2-22H), whereas that occurring as haloes to actinolite-sulphide veins which cut albitized dacite (Fig. 2-21C) is microcline-dominant and not associated with magnetite.

K-feldspathization also commonly affected porphyritic andesitic dikes that postdate the main magnetite mineralization. Phlogopite - magnetite - actinolite alteration occurs locally in Marcona

Formation marble.

Stage M-V: Polymetallic sulphide mineralization. The major sulphides in Stage M-V are again pyrite, chalcopyrite and pyrrhotite. Sulphide-rich veins, commonly with calcic amphiboles

(actinolite and tremolite: Table 2-3), occur in the upper parts of the orebodies and cut massive

Stage M-III and M-IV magnetite-amphibole associations (Fig. 2-23A). However, the relationship to Stage M-IV sulphides is rarely clearcut. Stage V sulphides and coexisting minerals commonly occur as aggregates replacing Stage M-III magnetite-amphibole. The common assemblages include chalcopyrite - pyrite - calcic amphibole (± pyrrhotite) and, less widely, chalcopyrite - pyrite - calcite (Fig. 2-23B). The chalcopyrite-pyrite-calcic amphibole-calcite assemblage occurs locally. Stage M-V sulphides are generally euhedral to subhedral and coarse-grained, and commonly have planar contacts with amphibole and calcite, which may indicate broadly coeval precipitation. Microscopically, however, the textural relationships between Stage M-V sulphides and Stage M-III or M-IV magnetite, or between Stage M-V amphiboles and calcite and Stage

M-III or M-IV magnetite, are ambiguous, even in the case of veins cutting Stage M-III and M-IV associations. Pyrrhotite mainly occurs as aggregates replacing Stage M-III or M-IV magnetite-amphibole and is subordinate to chalcopyrite and pyrite in sulphide veins. Accessory

Stage M-V sulphides include sphalerite, which is abundant in the Mina 14 orebody and is commonly associated with pyrite and chalcopyrite. Coarse-grained pyrite - prehnite - calcite - actinolite veins locally cut metasediments, and a coarse-grained, euhedral, fluorapatite-pyrite association is superimposed on Stage M-III magnetite. Stage M-V carbonates include ankerite and intergrown calcite.

Figure 2-23. Mineralogical and textural relationships of Marcona polymetallic sulphide stage (Stage M-V). A - Actinolite-tremolite-sulphide veins cut Stage M-III massive magnetite-calcic amphibole aggregates. The major sulphides are chalcopyrite and pyrite. Stage M-V actinolite-tremolite is commonly coarse-grained. Almost no magnetite occurs in Stage M-V veins (# DDM5-4-2, drillcore DDM5-4, 210 m, main orebody of Mina 5). B - Stage M-V Pyrite, chalcopyrite and calcite occur as aggregates superimposed on Stage M-III magnetite (# MA5-3, Mina 5 open pit, 670 m, orebody). C - Late- Stage M-V sulphide-barren tremolite veins cut magnetite-amphibole-sulphide aggregates (Mina 14 open pit, an isolated block of the orebody on the bench of 720 m).

Calcic amphibole formed extensively in Stage M-V as tremolite (Mg/Mg+Fe = 0.91 to 0.96:

Table 2-3) and actinolite (Mg/Mg+Fe = 0.64), both coexisting with sulphides. Tremolite, without associated metallic minerals, also developed late in Stage M-V, forming veins cutting massive magnetite orebodies (Fig. 2-23C). Hydrothermal breccias, in which coarse-grained, late-Stage

M-V tremolite cements magnetite-sulphide clasts, are widespread in the Cu-poor Mina 5 and

64 Mina 7 orebodies, but are only locally developed elsewhere. Tremolite which formed late in

Stage M-V has a lower iron content than that associated with sulphides (Table 2-3).

Stage M-VI: Chlorite-talc-serpentine alteration. Talc commonly occurs as fine-grained aggregates and replaces or cuts calcite, locally also replacing Stage M-IV phlogopite (Fig. 2-24A).

Lizardite and, locally, chrysotile also replace Stage M-V actinolite and tremolite (Fig. 2-24B), and talc and serpentine replace coarse-grained Stage M-V apatite. Serpentine veins commonly cut magnetite and sulphides in the cores of the orebodies. Whereas the chlorite (commonly talc-chlorite, in the classification of Hey (1954))-talc-serpentine assemblage records the retrograde alteration of phlogopite, actinolite and tremolite, the replacement of calcite and calcic amphiboles by talc and serpentine is evident for Mg metasomatism following the main Stage

M-V sulphide precipitation.

Figure 2-24. Stage M-VI chlorite-talc-serpentine alteration, Marcona. A - Fine-grained talc replaces Stage M-IV phlogopite (# MA3-11, Mina 3 open pit, 580 m, close to a magnetite orebody; transmitted light, crossed nicols). B - Serpentine (gray-white interference colour) replaces Stage M-V actinolite (brown-red interference colour) (# MA3-8, Mina 3 open pit, 580 m; transmitted light, crossed nicols).

Stage M-VII: Late veins. Late-stage hydrothermal veins are abundant at Marcona, but their mutual age relationships are ambiguous. Fine-grained subhedral magnetite and sulphides form narrow veins cutting both late-Stage M-V tremolite and Stage M-III magnetite (Fig. 2-25A). Rare late-stage chalcopyrite veins lacking gangue minerals cut late magnetite veins. Late magnetite is

65 widely weathered to powdery hematite, but some hematite (± gypsum) veins which cut the main magnetite orebodies and late quartz veins (Fig. 2-25B) are interpreted as hypogene. The major sulphides in hematite veins are pyrite and chalcopyrite. Locally, chalcopyrite - pyrite - amphibole veinlets with albite-rich envelopes cut porphyritic andesite dikes which transect the magnetite orebodies (Fig. 2-25C). Although the main mineralization stages at Marcona are almost free of quartz, barren quartz ± calcite veins cut magnetite orebodies and host-rock alteration zones (Fig.

Figure 2-25. Late veins (Stage M-VII) at Marcona. A - Late magnetite veins (Mt-2) cut massive magnetite (Mt-1) and late tremolite. Magnetite in veins is commonly fine-grained and locally associated with calcite (# DDM3-3-3, drill core DDM3-3, 343 m, Mina 3 orebody). B - Late quartz vein (locally associated with calcite and Mn-oxides) cuts massive magnetite-amphibole-sulphide assemblages. A hematite vein cuts both (# MA3-35, Mina 3 open pit, 580 m, south wall). C - Late amphibole-chalcopyrite-pyrite veins cut porphyric andesite (ocöite) dikes which postdated magnetite mineralization. Pink albite occurs as envelops of these veins or as patches in ocöite (Mina 1, open pit, 780 m, southeast side).

66 2-25B). Rhodochrosite locally occurs in these veins. Calcite veins up to 5 cm thick cut the magnetite orebodies and reopen serpentine veins. Tourmaline-quartz-pyrite veins locally cut

Marcona Formation metasediments.

Paragenetic sequence in the Mina 11 orebody.

Mina 11, the highest in altitude and stratigraphic level, and most Cu-rich orebody at Marcona, differs paragenetically from the other orebodies. Five main stages are identified (Fig. 2-26). Of these, Stage M11-I magnetite ± calcic amphibole occurs as dike-like bodies and veins cutting the hanging wall metasediments (Fig. 2-11), but is less abundant than in the other Marcona orebodies.

Stage M11-I veins commonly display envelopes of intense albitization. Stage M11-II, comprising

Figure 2-26. Hypogene alteration and mineralization paragenesis of the Mina 11 orebody

magnetite + calcic-amphibole (± phlogopite) + chalcopyrite + pyrite, dominates the massive orebody, locally forming breccias and stockworks along its contacts. Magnetite + chalcopyrite + pyrite veins assigned to Stage M11-III cut Stage M11-I and Stage M11-II assemblages (Fig.

2-27A). Stage M11-IV calcic amphibole (±magnetite) veins cut massive Stages M11-I and

M11-II bodies, but their temporal relationships with Stage M11-III are uncertain. Stage M11-V anhydrite veins, commonly replaced by gypsum and bassanite, cut all previous stages. Some lack sulphides (Fig. 2-27A), but coarse-grained anhydrite – pyrite - chalcopyrite and anhydrite-pyrite

67 stockworks are widely developed (Fig. 2-27B). Locally, anhydrite with abundant pyrite and minor chalcopyrite forms the matrix of hydrothermal breccias. Tourmaline-quartz-sulphide veins locally reopen earlier K-feldspar (± biotite) veins. The tourmaline is usually a coarse-grained, euhedral to subhedral, dark-green dravite, occurring interstitially to quartz, pyrite, chalcopyrite and molybdenite, with accessory fluorite, rutile and titanite. Such veins were in turn reopened by late anhydrite veins, but their relationships to other stages are unclear.

Figure 2-27. Paragenetic relationships of Mina 11 orebody. A - Late anhydrite (Stage M11-V, replaced by supergene gypsum) veins cut fine-grained late magnetite veins (Mt-2, Stage M11-III) which, in turn, cut the main magnetite (Mt-1, Stage M11-II)-amphibole-sulphide orebody. The main sulphide is chalcopyrite (# MA91-2, drillcore MA91, 92 m, Mina 11). B - Gypsum (after anhydrite) and coarse-grained pyrite (Stage M11-V) cement brecciated Stage M11-II magnetite- calcic-amphibole-sulphide assemblages (# MA91-3, drill core MA91, 95 m, Mina 11).

Compared with the general paragenetic sequence for Marcona (Fig. 2-19), the Mina 11 mineralization is dominated by Stages M11-II and M11-III (Fig. 2-26) which correspond,

68 respectively, to Stage M-IV and Stage M-V assemblages in the other orebodies, but they incorporate more sulphides, especially chalcopyrite. Pyrrhotite is rarely observed at Mina 11, and anhydrite, with or without sulphides, is much more abundant than elsewhere.

Comparison with earlier paragenetic models

Paragenetic sequences for alteration and mineralization in the Marcona deposit have been proposed by Atchley (1956), Injoque (1985) and Hawkes et al. (2002). The major differences between the previous models and that presented herein are: (1) the precursor cummingtonite, phlogopite and magnetite alteration systems are newly identified; (2) early Na metasomatism, characterized by albitization and scapolite formation (without magnetite), and preceding both potassic alteration and the main magnetite and sulphide mineralization, was not recognized by

Atchley (1956) and Injoque (1985), while Hawkes et al. (2002) assigned albite development to the main magnetite-Ca-amphibole mineralization stage; (3) concomitantly, the association of the main magnetite mineralization (Stage M-III) with biotite-phlogopite and K-feldspar alteration was largely unrecognized in previous studies; and (4) calcic amphiboles, dominantly actinolite and tremolite, are shown to be intimately related, not only to magnetite deposition, but also to sulphide (pyrrhotite, pyrite and chalcopyrite) emplacement, a relationship underestimated in previous models.

2.6 The Mina Justa Cu (-Ag) Deposit

The Mina Justa orebodies are hosted by the upper Río Grande Formation (Fig. 2-28), dominated by porphyritic andesite flows and medium to fine-grained andesitic volcaniclastics with minor horizons of sandstone, siltstone and limestone with a Callovian to Oxfordian fauna (Caldas, 1978;

Hawkes et al., 2002; Baxter et al., 2005). The volcaniclastic rocks locally incorporate rounded plagioclase phenoclasts in a much finer-grained matrix. Subordinate host rocks include

Figure 2-28. Geological map of Mina Justa Cu deposit, hosted by the upper Río Grande Formation. B-B’ and C-C’ show locations of the Figure 2-29 cross-sections (Modified from Rio Tinto 1: 10,000 mapping of Mina Justa Prospect, February 2003, unpublished report). Ab-albite; Kfs- K-feldspar; Act-actinolite

plagioclase- and hornblende-phyric andesite with vesicles filled by chlorite and carbonates (Fig.

2-4D). Lens-like marble bodies occur mainly in the southeast part of the area, but host no economic mineralization. The host rocks at Mina Justa are pervasively altered, and although

Injoque (1985) described the majority of the volcanic rocks in the Marcona area as shoshonites or

70 latitic andesites, but the absence of petrographic evidence for magmatic sanidine and, particularly, the abundant secondary hydrothermal biotite and sericite argue against such a marginally alkaline nature (Hawkes et al., 2002). However, the andesites at the base of the upper section of the Río

Grande Formation in Cañón Río Grande (Fig. 2-5) are K-rich and have high Cu contents (average

400 ppm: Aguirre, 1988).

Two principal arrays of orebodies, the main and upper, constitute the Mina Justa deposit

(Figs. 2-28 and 2-29A). The mineralized bodies incorporate massive magnetite-sulphide cores enclosed by hydrothermal breccias, comprising strongly altered host rock clasts in a magnetite+sulphide matrix, in turn surrounded by extensive stockwork (Fig. 2-30). They are controlled by subparallel, northeast-trending and shallowly southeast-dipping faults and range from 10 m to 200 m in vertical extent (Baxter et al., 2005). The main mineralized body crops out as a 400 m long, discontinuous belt of Cu oxides and albite - K-feldspar - actinolite alteration (Fig.

2-28), which dips 10° to 30° to the southeast. It has been intersected to a maximum depth of 500 m, where it remains open (Fig. 2-29A). The upper mineralized body, cropping out subparallel to and approximately 400 m southeast of the main zone (Fig. 2-28), has a similar elongate to oval shape in section, and a similar dip of 10° to 30° to the southeast. On surface, this zone has been identified over a distance of at least 400 m. and it has been intersected to a maximum depth of

300 m (Fig. 2-29A). The northeast-trending and southwest-dipping magnetite lenses are also exposed on surface (Fig. 2-28). They commonly contain minor Cu oxides and are locally cut by the southeast-dipping Mina Justa normal faults (Fig. 2-29A). Copper oxides dominate the upper

200 m of the deposit, giving way gradually to sulphides with depth (Figs. 2-29A and B). In individual orebodies, the major sulphides are zoned upwards, and locally laterally (cf. Moody et al., 2003), from pyrite-chalcopyrite to bornite-chalcocite (± digenite), with an increase in Cu grade (Figs. 2-29A and B). Around the magnetite-sulphide orebodies, the alteration is zoned outwards from potassic (K-feldspar - dominant), through calcic (actinolite) to sodic (albite).

71 K-feldspathization and calcite development are spatially associated, respectively, with iron oxide and Cu sulphide mineralization. Hematite commonly occurs in the upper parts of the zones of Cu mineralization.

Figure 2-29. Cross-sections through the major Mina Justa orebodies. A - NW-SE section through the Main and Upper orebodies (from Baxter et al., 2005). B - SW-NE section through the Main orebody (sulphide zones modified after Moody et al., 2003).

Figure 2-30. Mineralogical and structural zonation of the Mina Justa orebodies, based on logging of selected drill cores. The locations of holes MA-64, MA-17, MA-35 and MA-27 are shown in Figure 2-29A. MA-45 and MA-89 are collared 600 to 800 m southeast of the upper zone and out of the map area in Figure 2-28. * Magnetite either occurs erratically as haloes around coarse-grained pyrite or is absent in this zone. Mt-magnetite, Bn-bornite, Cp-chalcopyrite, Cc-chalcocite, Py-pyrite

73 Paragenetic relationships

Seven stages of hydrothermal alteration and mineralization are recognized at Mina Justa (Fig.

2-31).

Stage J-I: Albite-actinolite alteration. The earliest hydrothermal event at Mina Justa is albite-actinolite alteration. Light pink albite and green, fine-grained, actinolite (Mg/Mg+Fe = 0.70:

Table 2-4) replace both plagioclase phenocrysts and the matrix of andesites (Fig. 2-32A;

Reynolds, 2002a), recording Na-metasomatism.

Stage J-II: K-feldspar – magnetite alteration. Rocks affected by this alteration generally appear massive in hand-specimen, and range from pink to black. K-feldspar commonly occurs as extremely small grains (< 0.05 mm) replacing both fresh and previously albitized plagioclase (Fig.

2-32A), and the associated magnetite is mainly fine- to medium- grained (0.05-0.1 mm), locally forming aggregates interstitial to the feldspar (Fig. 2-32B). Stage J-II alteration, unambiguously the result of K-Fe metasomatism, was probably contemporaneous with the development of lenses of sulphide-free magnetite which strike NE and dip NW, locally crosscut by massive magnetite - pyrite bodies (Fig. 2-28). Overprinting by Stage J-III actinolite and Stage J-V coarse-grained

K-feldspar ± magnetite is common (Fig. 2-32C).

Stage J-III: Actinolite (± magnetite ± diopside) alteration. Green actinolite (Mg/Mg+Fe =

0.74: Table 2-4), associated with minor magnetite, occurs throughout the deposit, commonly as massive aggregates along the contacts of Stage J-V magnetite bodies or as coarse, acicular crystals in veins cutting Stage J-II K-feldspar-magnetite alteration. More locally, it forms the matrix of hydrothermal breccias (Fig. 2-33A) incorporating clasts of K-feldspar-magnetite - altered host rocks, termed “red-green breccias” by Hawkes et al. (2003). Along the contacts of the

Stage J-V magnetite bodies with their actinolitic alteration haloes, actinolite relics occur as irregular clasts in a magnetite-sulphide matrix. A temporal evolution is evident from K-feldspar – magnetite, through actinolite, to magnetite-pyrite alteration (Fig. 2-33B). Actinolite is strongly

Figure 2-31. Alteration and mineralization paragenesis of the Mina Justa deposit (Note: supergene minerals are omitted).

Table 2-4. Representative Electron Microprobe Data for Hydrothermal Silicates and Sulphides from Mina Justa

Mineral Act Act Chl Dg Dg Bn Cc Stage J-I J-III J-V J-VI J-VI J-VI J-VI Sample MA64 MA89 MA89 MA64 MA64-4 MA64 MA64 number -3 -4-1 -4-2 -4-II-1 -II-1-A -4-II-2 -4-I-1

SiO2 52.88 56.14 32.62 As n.d. 0.12 n.d. 0.10

TiO2 0.20 0.02 0.02 S 20.99 21.38 25.18 20.13 Al2O3 3.25 0.95 14.41 Fe 0.14 0.08 11.17 0.04 FeO* 11.76 10.73 20.41 Ni n.d. n.d. n.d. n.d.

MnO 0.10 0.10 1.06 Zn n.d. n.d. n.d. n.d. MgO 15.71 16.88 17.71 Ag 0.24 0.26 0.12 0.14 CaO 12.72 13.16 1.56 Cu 78.40 78.25 63.41 80.46 Na2O 0.44 0.15 0.12 Co n.d. n.d. n.d. n.d.

K2O 0.23 0.06 0.03 Cl 0.10 0.04 0.04 F 0.33 0.17 0.32 Total 97.70 98.41 88.31 Total 99.773 100.088 99.880 100.865

Si 7.63 7.95 6.67 1) Al 0.37 0.05 1.33 2) Al 0.18 0.10 2.14

Ti 0.02 0.00 0.00

Fe 1.42 1.27 3.49 Mn 0.01 0.01 0.18 Mg 3.38 3.56 5.40

Ca 1.97 2.00 0.34

Na 0.12 0.04 0.05 K 0.04 0.01 0.01 Cl 0.02 0.01 0.01

F 0.15 0.08 0.21 R** 0.70 0.74 0.39

*Total iron. Number of ions calculated on the basis of F, Cl, and 23 O for actinolite (Act); 28 O for chlorite (Chl), Dg-digenite, Bn-bornite, Cc-chalcocite. R**: Mg/Mg+Fe ratio for amphiboles; Fe/Fe+Mg for chlorite. 1) Al-tetrahedral; 2) Al-Octahedral. Detection limits for sulphides (in weight percent): Fe - 0.02; Ag - 0.04; As - 0.07; Cu - 0.02; S - 0.02; Ni, Co and Zn - 0.06. n.d. – undetected.

Figure 2-32. Albitization and actinolite alteration (Stage J-I) and K-Fe metasomatism (Stage J-II) at Mina Justa. A - Light-pink albite (not stained) and fine-grained actinolite extensively replaces original phenocrystic and groundmass plagioclase (stained pink to red). Stage J-II red microcline (stained yellow) replaces albite. Stage J-III actinolite is superimposed on albite and microcline (# MA64-7, drill core MA64, 394.4 m, 80 m from main orebody). Staining method is documented in Appendix I. B - Fine-grained microcline coexists with magnetite in a clast cemented by Stage J-III actinolite (see Figure 2-34A). Subhedral to euhedral actinolite crystals locally replace microcline (# MA64-3, drill core MA64, 220.1 m, plane-polarized transmitted light). C - Magnetite-sulphide-calcite veins with K-feldspar haloes (red) cut Stage J-III actinolite and Stage J-II fine-grained K-feldspar-magnetite (grey to pink) alteration. Actinolite is extensively chloritized. (# MA17-7, drill core MA17, 364 m).

chloritized and carbonatized, and locally replaced by quartz. It clearly replaced both albite and

K-feldspar, evidence for Ca metasomatism. Diopside is spatially associated with actinolite in the albitized and K-Fe metasomatised host rocks, but is also locally replaced by it. 77

Figure 2-33. Mineralogical and textural relationships of Mina Justa actinolite alteration (Stage J-III). A - “Red-green breccia” in which stage J-III actinolite (green) matrix cements clasts of stage J-II fine-grained K-feldspar - magnetite (Mt-1) (pinkish red to dark gray). Coarse-grained Stage J-V magnetite (Mt-2) occurs with actinolite and locally as veins (# MA64-3, drill core MA64, 220.1 m). B - Stage J-V magnetite (Mt-2)-bornite-chalcocite assemblage occurs as a matrix to pinkish-red Stage J-II K-feldspar - magnetite (Mt-1) altered clasts cut by Stage J-III actinolite (green) veins. The magnetite-sulphide matrix was reopened and partially replaced by late specularite (# MA64-6, drill core MA64, 276 m).

Stage J-IV: Early hematite-calcite alteration: “mushketovite”, i.e., magnetite unambiguously pseudomorphous after specular hematite, occurs commonly in the main magnetite bodies, evidence for a now covert hematite-dominant stage which temporally separated the actinolite alteration and the main magnetite alteration in andesite. The hematite originally formed fractured plates (Fig. 2-34). Anhedral to subhedral, and medium to coarse-grained calcite is intergrown with the pseudomorphs, and is locally replaced by quartz, magnetite and chalcopyrite. Rarely, coarse-grained, subhedral to euhedral allanite (Stage J-V) occurs in Stage J-IV calcite in contact with Stage J-V magnetite and pyrite.

Figure 2-34. Platy Stage J-V magnetite (after Stage J-IV hematite) occurs with calcite, quartz and chalcopyrite. Chalcopyrite extensively replaces pyrite and locally occurs along fractures in magnetite. Chloritized Stage J-III actinolite relics occur between the magnetite crystals. Stage J-IV calcite grains have planar contacts with platy magnetite, but are locally replaced by Stage J-V granular magnetite and quartz veins. (# MA17-6 from drill core MA17, 355.1 m, combined reflected and transmitted light).

Stage J-V: Magnetite-pyrite alteration: The massive, lensoid and brecciated magnetite-pyrite bodies which host the highest-grade copper sulphide mineralization at Mina Justa (Fig. 2-30) were controlled by the NE-striking, SE-dipping, Mina Justa system faults, but are dislocated by the NW-striking, NE-dipping Huaca faults and associated ocöite dikes. Magnetite-pyrite veins, varying from 0.1 to 5 cm in width, cut alteration assemblages of Stages J-II and J-III adjacent to the massive magnetite bodies. Hydrothermal breccias commonly occur in altered host rocks in contact with the magnetite bodies, and comprise a magnetite-pyrite – dominant matrix and angular clasts of andesite altered to microcline (Stage J-II) or actinolite (Stage J-III) (Fig. 2-35A).

However, the textural relationships between Stage J-V magnetite-pyrite and Stage J-III actinolite are complex. Locally, unambiguous crosscutting or replacement textures are shown (Fig. 2-35B), but magnetite and pyrite more commonly occur as elongated, locally ellipsoidal, aggregates in massive actinolite in contact with the main magnetite bodies, and magnetite-pyrite intergrowths replace Stage J-III actinolite (Fig. 2-35B). The aggregates comprise pyrite cores enclosed by 79 magnetite haloes. Actinolite veins peripheral to the main magnetite bodies are in places reopened by magnetite and pyrite, which clearly replaced actinolite. Rarely, Stage J-V magnetite-rich alteration occurs as spots in altered host rocks peripheral to the main magnetite bodies (Fig.

2-35C).

Magnetite and pyrite of Stage J-V are medium to coarse-grained (0.5 to 10 mm; locally over 1 cm in the case of pyrite) and subhedral to euhedral. Magnetite commonly occurs interstitially to pyrite and has planar contacts in general (Fig. 2-35D). The major mineral associated with magnetite-pyrite alteration in the main magnetite bodies is quartz, which occurs as 0.1–1 mm, subhedral to euhedral, crystals interstitial to magnetite and pyrite and commonly with actinolite inclusions (Fig. 2-35D). Accessory calcite is generally anhedral to subhedral and medium-grained, coexisting with magnetite, pyrite and quartz. Pink to red K-feldspar, predominantly microcline, is a common alteration mineral in rocks associated with Stage J-V magnetite-pyrite mineralization, forming haloes to magnetite-pyrite veins or patches incorporating medium to fine-grained magnetite crystals and superimposed on early alteration (Fig. 2-32C). Chlorite, largely clinochlore and diabantite (Reynolds, 2002a; classification of Hey, 1954; Table 2-4), extensively replaces actinolite or diopside (Fig. 2-35E), and locally occurs in veins with magnetite, pyrite and quartz. In the massive actinolite aggregates in contact with magnetite orebodies, chloritization of actinolite is commonly focused around the ellipsoidal magnetite-pyrite aggregates. Although characteristically related to magnetite-pyrite mineralization, chlorite veins which cut actinolite are in turn cut by magnetite-pyrite-quartz veins. Locally, chlorite-albite-quartz-calcite-hematite veins related to late Cu mineralization (Stage J-VI) transect microcline grains associated with magnetite-pyrite mineralization. Titanite commonly forms medium to coarse-grained subhedral crystals or aggregates in chlorite, and euhedral grains in magnetite - pyrite - quartz - chlorite veins. Fluorapatite locally occurs in Stage J-V mineralized veins, but more commonly forms

Figure 2-35. Mineralogical and textural relationships of Mina Justa magnetite-pyrite alteration (Stage J-V). A - Hydrothermal breccia at the margin of Main Orebody. Magnetite (Mt-2)-sulphide occurs as a matrix around angular Stage J-II microcline-magnetite (Mt-1) clasts. Actinolite relics occur in matrix (# MA35-0, drill core MA35, 484.3 m). B - Replacement of Stage J-III actinolite by Stage J-V magnetite-pyrite. Strong chloritization of actinolite is locally evident (lower-right) (# MA27-2, drill core MA27, 366.9 m; transmitted light, crossed nicols). C - Spotty magnetite-chalcopyrite-quartz mineralization in earlier actinolite and microcline-magnetite (gray to pink) - altered host rocks. Chalcopyrite coexists with magnetite and quartz (# MJ-38, drill core MA54, 341.8 m). D - Magnetite-pyrite-quartz alteration. Quartz is coarse-grained and euhedral. Stage J-III actinolite crystals occur as relics in quartz grain (# MA17-6, drill core MA17, 355.1 m, combined reflected and transmitted light). E - Magnetite alteration and related chloritization. Magnetite coexists with pyrite, quartz and chlorite. Chlorite (locally with quartz) extensively replaces Stage J-III actinolite (# MA89-4, drillcore MA89, 360.2 m, plane-polarized transmitted light).

81 coarse-grained subhedral to euhedral crystals in Stage J-III actinolite in contact with magnetite orebodies. The textural relationships of apatite and actinolite are ambiguous.

Stage J-VI: Copper sulphide mineralization: Stage J-V magnetite alteration, although rich in pyrite, lacks inherent Cu sulphides. Cu sulphide–bearing veins, assigned to Stage J-VI, locally cut altered host rocks and Stage J-V magnetite-pyrite-quartz (Fig. 2-36A) but, more commonly, Cu sulphides and associated assemblages occur in massive magnetite-pyrite bodies or veins and exhibit unambiguous microscopic replacement textures with the latter (Fig. 2-36B). Locally,

Stage J-V magnetite-pyrite aggregates in Stage J-III actinolite veins have been almost completely replaced by chalcopyrite or bornite, giving rise to the common actinolite-Cu sulphide association.

The main hypogene Cu sulphides at Mina Justa are, in decreasing abundance, chalcopyrite, bornite, chalcocite and digenite. Except for chalcopyrite, these are concentrated above, or in the upper parts of, the main magnetite bodies (Fig. 2-30), commonly as veins that cut the host rocks and earlier alteration assemblages. Supergene covellite occurs mainly in the oxide zone, replacing bornite or chalcopyrite (Fig. 2-36C). Chalcocite, digenite and bornite typically form large patches with vermicular, eutectic-like intergrowths (Fig. 2-36D), such as are inferred to form through noncoherent exsolution at low temperature (< 250oC) and under protracted cooling (Brett, 1964).

Similar relationships are documented at Olympic Dam by Roberts et al. (1983). Cu sulphides exhibiting vermicular textures are all rich in Ag (Table 2-4), and represent the major host of Ag in the ores. Chalcopyrite occurs as veinlets in the upper parts of the main magnetite bodies, but beneath the bornite-chalcocite zones. Chalcocite, bornite and chalcopyrite locally occur together with no unambiguous mutual replacement relationships. Rarely, chalcopyrite veins cut bornite bodies, evidence for repeated Cu mineralization in the upper levels. Unambiguous replacement textures between chalcopyrite and earlier magnetite and, particularly, pyrite are common (Fig.

2-36B). Such textures also locally occur between bornite-chalcocite aggregates and magnetite.

Figure 2-36. Mineralogical and textural relationships of Mina Justa Cu mineralization (Stage J-VI). A - Chalcopyrite-calcite veins cut altered host rocks. Microcline occurs as haloes around calcite veins and locally cuts calcite (# MA45-6, drillcore MA45, 404.2 m). B - Chalcopyrite replaces Stage J-V pyrite and magnetite (# MA17-9, drillcore MA35, 507.9 m, plane-polarized reflected light). C - Supergene covellite replaces chalcopyrite (# MA14-3, drillcore MA14, 394.7 m, plane-polarized reflected light). D - Fine-grained bladed hematite coexists with bornite, digenite and chalcocite with vermicular, eutectic-like textures, occurring as patches in a magnetite vein which cuts host rocks (# MA64-4, drill core MA64, 248.3 m, plane-polarized reflected light).

Accessory Stage J-VI sulphides include sphalerite, galena, molybdenite and rare fine-grained (<

25 μm), carrollite (Reynolds, 2002b). These generally coexist with chalcopyrite and locally replace pyrite, but some sphalerite surrounds chalcopyrite with an ambiguous replacement texture.

The iron oxide associated with both chalcopyrite and bornite-chalcocite mineralization is dominantly fine-grained platy hematite, commonly occurring as aggregates around Cu sulphides

83 (Fig. 2-36D). Locally, Stage J-VI hematite formed with chalcopyrite along the boundaries of earlier magnetite grains or Stage J-IV coarse-grained hematite (“mushketovite”). Calcite is the dominant gangue mineral associated with Cu mineralization, generally occurring in veins which cut the host rocks and magnetite mineralization (Fig. 2-36A). Calcite-Cu sulphide assemblages dominate these veins but give way upwards to hematite-bearing assemblages. Albite (± microcline) locally occurs in chalcopyrite-calcite veins cutting altered andesite host rock, or replaces magnetite - pyrite - quartz and chalcopyrite bodies. Sparse epidote veins with chalcopyrite or bornite cut Stage J-V magnetite-pyrite mineralization and associated chloritic alteration zones. Locally, the epidote-calcite-prehnite assemblage occurs in altered host rocks, replacing Stage J-III actinolite. Red microcline and subordinate albite occur as narrow haloes around calcite-chalcopyrite veins, and locally cut calcite, evidence that they partially postdated

Cu mineralization (Fig. 2-31). Epidote and clinozoisite commonly occur in calcite in contact with

K-feldspar. Barite locally occurs in these late K-feldspar veins, but the temporal relationships between K-feldspar veins and specular hematite (Stage J-VII) are ambiguous.

Stage J-VII: specular hematite: In the upper parts of the orebodies, a Cu sulphide-barren hematite stage locally developed. Medium to coarse-grained (0.1-0.5 mm), specular hematite forms veins cutting Stage J-III actinolite alteration zones and Stage J-V magnetite mineralization.

Locally, Stage J-VII hematite replaces Stage J-V (magnetite) and Stage J-VI (Cu) mineralization in hydrothermal breccias. The replacement and crosscutting relationships between Stage J-VII hematite and Cu sulphides confirm the distinction of Stage J-VII.

2.7 Stable Isotope Geothermometry

Determination of the P/T conditions of mineralization at Marcona is rendered difficult by a lack of transparent minerals amenable to fluid inclusion microthermometry, particularly in the earlier,

84 magnetite-dominant paragenetic stages. Hydrothermal quartz is rare and apatite, diopside and actinolite-tremolite do not host inclusions of adequate size. Estimation of the temperatures of mineralization therefore relies largely on light stable isotope (O and S) fractionation between coexisting minerals (Table 2-5). Isotopic analyses were also conducted on Mina Justa samples to complement the microthermometric data. Most magnetite-H2O fractionation factors (Zheng, 1991;

Cole et al., 2004), when combined with silicate-H2O factors, give temperatures exceeding the

o upper thermal stability limits of both the Stage M-III biotite (Ann55-63; 800-850 C between 1.0 and 2.1 kbar at the QFM buffer, and 750-800oC and 650-670oC at the Ni-NiO and magnetite-hematite buffers, respectively: Wones and Eugster, 1965) and the actinolite

o (Mg/Mg+Fe = 0.83; 780-830 C at 1.0 to 2.0 kbar: Lledó, 2005). In contrast, the magnetite-H2O fractionation factor determined by Bottinga and Javoy (1973) gave acceptable temperatures for the Marcona magnetite stage (Table 2-5). The analytical errors for O and S isotopic data are ± 0.2 and ± 0.3‰, respectively, corresponding to uncertainities of ~ ± 30oC and ± 50oC in the temperature estimates.

Marcona deposit

At Marcona, five oxygen isotope pairs were used to constrain the temperature of the main magnetite stage (Stage M-III, Fig. 2-19; Table 2-5). Samples MA5-9 (magnetite-biotite), MA3-22

(magnetite-apatite and magnetite-tremolite), MA3-7 (magnetite-phlogopite) and DDM3-3-1

(magnetite-actinolite), which are either sulphide-free or contain only minor pyrrhotite, give temperatures ranging from 700oC to 800oC, all but the MA3-22 magnetite-apatite pair exceeding

750oC. These data are inferred to represent broadly the conditions under which magnetite crystallized in equilibrium with amphibole, biotite, phlogopite and apatite. It should be emphasized, however, that isotopic reequilibration during cooling is a common feature for

85 igneous rocks of all types and the calculated temperatures are inferred to be minimum values.

Stage M-III may therefore have been initiated significantly above 800oC.

Table 2-5. Oxygen and Sulphur Isotope Geothermometry, Marcona and Mina Justa

Stage Sample δ18O Mineral Pair δ34S Mineral Pair Temperature Marcona Stage M-III: DDM3-3-1 Act (8.6‰) - Mt (4.4‰) 800 oC Main magnetite MA5-9 Bt (8.7‰) - Mt (5.2‰) 800 oC MA3-7 Phl (8.0‰) - Mt (4.2‰) 770 oC MA3-22 Apt (7.2‰) - Mt (3.7‰) 760 oC MA3-22 Trm (8.9‰) - Mt (3.7‰) 700 oC Stage M-IV: MA3-19 Phl (9.3‰) - Mt (2.9‰) 600 oC magnetite-sulphide MA1-9 Py (4.0 ‰) - Cp (3.4 ‰) 590 oC DDM4-7-4 Cal (13.2) – Mt (4.8) 570 oC MA91-4 Py (2.5 ‰) - Cp (1.6 ‰) 430 oC DDM4-6-5 Po (2.4 ‰) - Cp (2.1 ‰) 430 oC Stage M-V: MA2-12 Py (4.7 ‰) - Cp (3.6 ‰) 360 oC Polymetallic sulphide DDM4-7-9 Po (3.5 ‰) - Cp (2.7 ‰) 160 oC Mina Justa Stage J-V: MA45-3 Qtz (12.6 ‰) - Mt (4.7 ‰) 600 oC magnetite-pyrite MA89-2 Apt (8.0 ‰) - Mt (2.9 ‰) 580 oC -quartz MA45-2 Apt (9.0 ‰) - Mt (3.3 ‰) 540 oC MA89-1 Qtz (13.5 ‰) - Mt (2.7 ‰) 460 oC Stage J-VI: MA89-5 Ab (12.6 ‰) - Cal (13.3 ‰) 530 oC Cu mineralization MA45-6 Ab (13.2 ‰) - Cal (13.1 ‰) negative

Oxygen isotope fractionation factors: Mt-Qtz (Clayton and Keiffer, 1991); Mt-Apt (Valley, 2003); Mt-H2O (Bottinga and Javoy, 1973); Qtz-H2O (Clayton et al., 1972); Ab-Cal (Zheng, 1993a); Act-H2O (Zheng, 1993b); Trm-H2O (Zheng,

1993b); Cum-H2O (Zheng, 1993b); Bt-H2O (Bottinga and Javoy, 1973); Phl-H2O (Zheng, 1993b). Sulphur isotope fractionation factors: Py-Cp (Kajiwara and Krouse, 1971); Po-Cp (Kajiwara and Krouse, 1971). Mineral abbreviations: Act-actinolite; Mt-magnetite; Bt-biotite; Phl-phlogopite; Apt-apatite; Trm-tremolite; Py-pyrite; Cp-chalcopyrite; Po-pyrrhotite; Qtz-quartz; Ab-albite; Cal-calcite.

One phlogopite-magnetite oxygen isotope pair, from sample MA3-19, gives a temperature of

600oC for Stage M-IV (Table 2-5), close to a sulphur isotope temperature (590oC) estimated for a

86 pyrite-chalcopyrite pair from sample MA1-9, and a oxygen isotope temperature (570oC) for a magnetite-calcite pair from sample DDM4-7-4 from the same stage. However, another pyrite-chalcopyrite pair and one pyrrhotite-chalcopyrite pair from Stage M-IV magnetite-sulphide assemblages give identical temperatures of 430oC. In spite of the marked difference (100 - to -

370oC) in the temperature estimates for Stages M-III and M-IV, no apparent mesoscopic mineral assemblage zonation is observed in the magnetite orebodies and no obvious replacement textures are identified between minerals assigned to these Stages. This implies that the hydrothermal system was cooling rapidly during the initial deposition of sulphides.

Two magnetite-free Stage M-V sulphur isotope pairs give temperatures of 360oC and 160oC, significantly lower than those for Stages M-III and M-IV oxides and sulphides. The difference of over 200oC for the formation of similar sulphide assemblages indicates that Stage M-V sulphides may have precipitated in an environment in which the temperatures rapidly decreased, in conformity with the mixing of two fluids, one cooler than 160oC.

Mina Justa deposit

Four oxygen isotope pairs were analyzed from the main magnetite stage (J-V) of the Mina Justa deposit, giving temperatures of from 460oC to 600oC (Table 2-5). In sample MA89-1, which gives the lowest temperature, chalcopyrite is abundant, replacing Stage J-V pyrite and magnetite.

Minor fine-grained hematite is associated with chalcopyrite and replaces magnetite. Such replacement textures indicate that the analyzed magnetite may be partially altered to hematite, which is depleted in 18O compared to magnetite (Zheng and Simon, 1991), thus giving an unreliable temperature. Although apparently in textural equilibrium, two albite-calcite pairs from the Mina Justa Cu mineralization stage (J-VI) yielded, respectively, a temperature of 530oC and a negative value, and we infer that the albite was not in equilibrium with the calcite. Similarly, several pyrite-chalcopyrite pairs from Mina Justa yielded apparent temperatures exceeding

87 1000oC, far too high for the Cu mineralization, supporting the textural evidence that Stage VI chalcopyrite replaced pyrite, and implying that sulphur from different sources was involved in the formation of these sulphides.

2.8 40Ar/39Ar Geochronology

Sensibly concordant conventional K-Ar ages of 154 ± 4 (2σ) Ma and 160 ± 4 Ma for, respectively, sericite and phlogopite from Marcona were reported by Injoque (1985) but, although supporting a mid-Jurassic age for the main period of magnetite mineralization, these data provide no evidence of the duration of hydrothermal activity in this multi-stage centre. Moreover, the age relationships of the Marcona and Mina Justa deposits have remained undefined. Laser-induced incremental-heating 40Ar/39Ar techniques were herein applied to high-quality mineral separates from samples well documented in terms of field relationships and mineralogy. They include biotite, phlogopite, cummingtonite, actinolite, tremolite and microcline from Marcona, and actinolite and microcline from Mina Justa. The locations and descriptions of the dated Marcona and Mina Justa samples are documented in Table 2-6. The age spectra and corresponding Ca/K and Cl/K ratio and inverse isochron plots for Marcona and Mina Justa are illustrated in Figures

2-37 and 2-38, and analytical details are provided in Appendix I. All dates are quoted with an uncertainty of ± 2σ (95 percent confidence level). An age plateau is defined (Dalrymple and

Lanphere, 1974) as at least three separate outgassing steps with ages that are concordant at 2σ errors and that account for at least 50 percent of the 39Ar released.

40Ar/39Ar age spectra and age relationships

Marcona: Two acceptable and concordant plateau ages of 177.0 ± 1.5 Ma and 175.2 ± 2.3 Ma, accounting for 61 and 74 percent of the 39Ar released, respectively, are given by Stage M-IA

Table 2-6. Summary of 40Ar/39Ar Ages from Marcona and Mina Justa Sample Location Sample Description Mineral Stage Plateau Age dated ± 2σ (Ma) Marcona MA5-9A* Marcona, Mina 5 open pit Fine-grained Mt interstitial to Cum and Cum Stage M-I-A 177.0 ± 1.5 (670 m a.s.l, NE corner) Bt 175.2 ± 2.3 MA5-2 Marcona, Mina 5 open pit Fine-grained Mt with sulphide, Cum Phl Stage M-I-B 171.5 ± 1.11) (620 m a.s.l, centre) and Phl MA5-9B Marcona, Mina 5 open pit Fine-grained Mt interstitial to Cum and Bt Stage M-III 161.4 ± 0.9 (670 m a.s.l, NE corner) Bt MA3-30 Marcona, Mina 3 open pit Fine-grained Mc alteration in dacite Mc Stage M-III 109.2 ± 0.6 (580 m a.s.l, south wall) MA3-24 Marcona, Mina 3 open pit Mc veins cut albitized dacite and Mc Stage M-III 101.0 ± 0.6 (580 m a.s.l, SW corner) refilled by act-sulphide veins MA3-19 Marcona, Mina 3 open pit Coarse grained Mt-sulphides-Phl Phl Stage M-IV 159.7 ± 0.8 (580 m a.s.l, SW corner) DDM3-3-8 Marcona, Mina 3 Trm-sulphide aggregates replacing Trm Stage M-V 158.5 ± 1.9 Drill core DDM3-3; 418 m massive Mt MA91-2* Marcona, Mina 11, Fine-grained Mt interstitial to Trm and Trm Stage M11-II 157.3 ± 3.2 Drill core MA91, 92 m sulphides 156.8 ± 2.9 DDM5-4-2 Marcona, Mina 5 Act-sulphide veins cutting Mt Act Stage M-V 156.6 ± 4.2 Drill core DDM5-4; 210 m DDM3-3-1 Marcona, Mina 3 Trm cementing Mt-sulphide-Act clasts Trm Stage M-V 156.2 ± 2.4 Drill core DDM3-3; 340 m Mina Justa MA45-2* Mina Justa, Act alteration in andesite Act Stage J-I 157.3 ± 3.5 Drill core MA45, 262 m 154.3 ± 5.5 MJ-6 Mina Justa, Hem-Cp-Ep-Ab veins cutting Stage II Mc Stage J-II 142.4 ± 6.7 Drill core MA64 Mc-altered andesite (+ Ab) MA64-3* Mina Justa, Coarse-grained Act (± Mt) brecciating Act Stage J-III 110.9 ± 0.7 Drill core MA64, 220.1m K-Fe metasomatised andesite 109.9 ± 1.0 1) MA17-7 Mina Justa, Py-Mt-Mc veins Mc Stage J-V 103.7 ± 0.6 1) Drill core MA17; 364 m MA14-3 Mina Justa, Mt-Py-Mc aggregates in Act-altered Mc Stage J-V 101.5 ± 0.7 Drill core MA14; 394.7 m andesite MA45-6 Mina Justa, Cal-Cp veins (with Ab-Mc envelopes) Mc Stage J-VI 99.1 ± 0.9 Drill core MA45; 404.2 m (+ Ab) MA17-9 Mina Justa, Cal-Cp veins and late-refilling Mc Mc Late-Stage 95.0 ± 0.6 Drill core MA17; 408.7 m veins cut Act veins J-VI * Samples with double runs. 1) Plateau represents less than 50% of 39Ar released. Mineral abbreviations: Cum-cummingtonite; Ep-epidote; others as in Table 2-5.

Figure 2-37. Laser-induced 40Ar/39Ar age spectra, with Ca/K and Cl/K ratios for each heating step, and inverse isochron plots for samples from Marcona alteration and mineralization stages.

90 cummingtonite (Mg/Mg+Fe = 0.74; sample MA5-9, Mina 5; Table 2-3). Alteration to greenalite and chlorite was probably responsible for the “staircase” lower-temperature heating steps (Figs.

2-37A, B). The plateau ages and inverse isochron ages are concordant within 2σ error, with initial

40Ar/36Ar ratios of 292 and 294, close to the atmospheric value. The higher error in Figure 2-37B reflects the smaller number of heating steps. The cummingtonite is locally replaced by Stage

M-III biotite (Ann57; Figs. 2-20B and C), which gave a significantly younger plateau age of

161.42 ± 0.89 Ma (Fig. 2-37E; Table 2-6). The 175-177 Ma age for the cummingtonite hosted by the upper Marcona Formation demonstrates that it represents a precursor alteration event during the initial deposition of the overlying Río Grande Formation in the Aalenian, i.e., prior to 174.0

+1.0 +1.0 -7.9 Ma and after 178.0 -1.5 Ma (Pálfy et al., 2000). Hydrothermal activity, albeit not demonstrably associated with either magnetite or sulphide deposition, was therefore underway in a very near-surface, probably littoral environment, plausibly within a rift developed on eroded Marcona

Formation metasediments.

Stage M-IB phlogopite (sample MA5-2) gives two very similar “staircase” spectra, in which all low-temperature steps represent mixtures of gases derived from sites that have lost Ar and/or

K (Fig. 2-37C). A quasi-plateau at 171.5 ± 1.1 Ma represents only 35.5 percent of the 39Ar released, but the relatively good inverse isochron age (171.0 ± 1.9 Ma) obtained for a duplicate sample (Fig. 2-37D) indicates that the “plateau age” is probably acceptable. In the latter sample, the age spectrum exhibits a configuration suggestive of reactor-induced 39Ar recoil, and therefore the apparent age of the highest-temperature step, i.e., 172.6 ± 1.8 Ma, may represent a maximum for the age of crystallization. The relatively high Ca/K ratios of all steps (Figs. 2-37C and D) suggest that the phlogopite is partially altered to talc and chlorite, which would release K.

Coarse-grained Stage M-III biotite (Ann57, sample MA5-9B; Fig. 2-37E) gives a good plateau age (80 percent of the 39Ar released) of 161.42 ± 0.89 Ma (Table 2-6), which is concordant with the inverse isochron age within error. The minor “staircases” in the initial

91 heating steps in the spectrum may record later hydrothermal disturbance. The two highest-temperature steps, which were not included in plateau age calculation, gave slightly younger ages of 160.08 ± 1.19 and 159.19 ± 2.48 Ma respectively. The larger errors of ages in these steps probably reflect the small amounts of 39Ar they represent (Fig. 2-37E). Two Stage

M-III K-feldspars (maximum microcline, MA3-30 and MA3-24) from potassic alteration zones in dacite give acceptable plateau ages (75.9 and 61.2 percent of the 39Ar released, respectively) of

109.18 ± 0.64 Ma and 101.04 ± 0.56 Ma, concordant with the inverse isochron ages within error

(Figs. 2-37F and G). These ages are markedly younger than those for M-III biotite. Microcline has a low closure temperature of 130oC-160 oC (Harrison and McDougall, 1982), and the two

Marcona Stage M-III apparent age plateaus could therefore record resetting by late hydrothermal events at ca. 101-109 Ma, evidence that all of the Marcona deposit area was reheated to over

130oC in the mid-Cretaceous. The earliest step in MA3-30 may record a second, ca. 90 Ma resetting (Fig. 2-37F). The slightly “hump-shaped” age spectra of the microclines (Figs. 2-37F and G) could be related to microtextural features, e.g., extensive turbidity or compositional variation, which may disrupt existing diffusion domains (McLaren et al., 2007).

A Stage M-IV phlogopite (Ann4-5, sample MA3-19; Fig. 2-37H) gives a good plateau age

(97.8 percent of the 39Ar released) of 159.69 ± 0.84 Ma (Table 2-6), which is concordant with the inverse isochron age within error. This slightly younger age relative to that for Stage M-III biotite is in conformity with the textural evidence for replacement of biotite by phlogopite (Fig. 2-22F).

The unacceptable inverse isochron plot for MA3-19 phlogopite (Fig. 2-37H) reflects the very low

36Ar content.

Stage M-V tremolite (Mg/Mg+Fe = 0.92; DDM3-3-8) and actinolite (Mg/Mg+Fe = 0.64;

DDM5-4-2), both associated with chalcopyrite and pyrite, give acceptable and similar plateau ages (98.7 and 80.5 percent of the 39Ar released, respectively) of 158.5 ± 1.9 and 156.6 ± 4.2 Ma

(Figs. 2-37I and J). A second Stage M-V tremolite (Mg/Mg+Fe = 0.96; DDM3-3-1), unassociat

92 with sulphides, gives a plateau age of 156.2 ± 2.4 Ma (100 percent of the 39Ar released; Fig.

2-37K), which is concordant with its inverse isochron age within 2σ error.

An actinolite from the Mina 11 orebody (MA91-2), assigned to Stage M11-II (Table 2-6), which probably corresponds to Stages M-IV or M-V at Marcona, gave a good plateau age of

157.3 ± 3.2 Ma (99.88% percent of the 39Ar released; Fig. 2-37L). An identical plateau age of

156.8 ± 2.9 Ma (99.89% percent of the 39Ar released; Fig. 2-37M) was obtained for a duplicate sample. The inverse isochron ages, with large errors, are either younger or older than the plateau ages (Figs. 2-37L and M), probably reflecting excess atmospheric argon released in the earlier steps (McDougall and Harrison, 1999). The Mina 11 actinolite ages are concordant with those for

Stages M-IV and M-V actinolite from Marcona, demonstrating that, although mineralogically distinctive, the Mina 11 orebody is an integral part of the Marcona deposit.

Despite the possibility of resetting of biotite due to its relatively low closure temperature (ca.

310°C: Harrison et al., 1985), the ages for alteration minerals of Stages M-III, M-IV and M-V at

Marcona are in conformity with paragenetic relationships (Fig. 19), and imply that the main episode of base-metal mineralization shortly followed the emplacement of the economic magnetite bodies. Both the magnetite and sulphide stages at Marcona are inferred to have taken place between 156 and 162 Ma, in agreement with the K-Ar age of phlogopite from the magnetite mineralization stage reported by Injoque (1985). A ca. 160 Ma mineralization age for the magnetite orebodies, hosted largely by the Marcona Formation, but in part by the Lower Río

Grande Formation (Fig. 2-5), may be compared to the Callovian-Oxfordian (151.4 - 161.5 Ma:

Pálfy et al., 2000) fossil ages of the upper section of the Río Grande Formation (Caldas, 1978;

Fig. 2-5) and the K-Ar whole-rock age (164 ± 4 Ma Ma) of an andesite flow at the base of the upper member of the Río Grande Formation in the Cañón Río Grande area (Aguirre, 1988).

Major magmatic and hydrothermal activity at Marcona is therefore inferred to have been

93 contemporaneous with the latest volcanism of the Río Grande Formation, i.e., following the

Bathonian stratigraphic hiatus (Fig. 2-5).

Mina Justa: Two actinolite and five alkali feldspar separates were dated from Mina Justa

(Table 2-6). X-ray powder diffraction study (Wright, 1968) shows that all of the feldspars are strongly ordered, with the microcline structure (Fig. 2-38A).

Two fine-grained Stage J-I actinolites (MA45-2) give good plateau ages of 157.3 ± 3.5 and

154.3 ± 5.5 Ma (95.2% and 99.82% percent of the 39Ar released, respectively; Figs. 2-38B and C).

However, the inverse isochron ages are older than the plateau ages, with large errors and low initial 40Ar/36Ar ratios of 284 and 288 (Figs. 2-38B and C), which may be related to an excess atmospheric component in the lower-temperature steps. The Stage J-I actinolite ages are, within error, concordant with Marcona magnetite mineralization and polymetallic sulphide precipitation ages (Table 2-6).

An alkali feldspar separate (MJ-6, Fig. 2-38D) from the envelope of a Stage J-VI hematite - chalcopyrite - epidote vein cutting altered Río Grande Formation andesite gives a plateau age

(89.4 percent of the 39Ar released) of 142.4 ± 6.7 Ma (Table 2-6), concordant within 2σ error with the inverse isochron age, which has an initial 40Ar/36Ar ratio of 297.4, close to the atmospheric value. The low K content results in a large error (Fig. 2-38D). Feldspar staining and microscopic textures reveal fine-grained microcline intergrown with albite, and the relatively high Ca/K ratio may be record albite or minor epidote in the separates. The Berriasian age is therefore interpreted as that of Stage J-II (Fig. 2-31), implying that the hematite, chalcopyrite and epidote in the sample were emplaced during Stage J-VI through reopening of an early fracture with previously feldspathized margins.

A Stage J-III coarse-grained actinolite gives a disturbed apparent age spectrum (Fig. 2-38E), with a configuration suggestive of Cl-derived excess argon (McDougall and Harrison, 1999) in the low temperature steps, as is supported by the high Cl/K ratios. The minor calcite in actinolite

Figure 2-38. A - X-ray powder diffraction data for hydrothermal K-feldspars from Mina Justa (Data fall in the microcline fields in the Al/Si ordering diagram of Wright, 1968). B-J - Laser-induced 40Ar/39Ar age spectra, with Ca/K and/or Cl/K ratios for each heating step, and inverse isochron plots for each sample.

95 separates may be responsible for concomitant high Ca/K and Cl/K ratios (Fig. 2-38E), the Cl perhaps being contributed by the abundant fluid inclusions (Harrison et al., 1993) in calcite. The age of the highest-temperature step, 110.89 ± 0.72 Ma, may therefore approach the true age. Four low temperature steps gave an inverse isochron age of 99.4 ± 2.7 Ma (Fig. 2-38E), much younger than the age calculated from the high temperature steps. A duplicate sample similarly failed to yield an acceptable age plateau: the disturbed spectrum incorporates a trough at 109.90 ± 0.97 Ma

(28.2 percent of the 39Ar released) (Fig. 2-38F), which is concordant within 2σ error with the IIA

(107.2 ± 9.1 Ma) given by the three steps constituting a plateau (Fig. 2-38F).

Microcline associated with the main Mina Justa magnetite stage (J-V; MA17-7, Table 2-6) has a quasi-plateau age of 103.73 ± 0.64 Ma (33.9 percent of the 39Ar released: Fig. 2-38G). The spectrum is disturbed and 39Ar recoil may be recorded by the low temperature steps for this very fine-grained (< 50 μm) sample (McDougall and Harrison, 1999). In contrast, a second Stage J-V microcline (MA14-3) gave a good plateau age of 101.49 ± 0.67 Ma (83.3 percent of the 39Ar released: Fig. 2-38H). Microcline (+ albite) from the Stage J-VI Cu mineralization (MA45-6) associated with calcite and chalcopyrite but not with magnetite, similarly gives an acceptable plateau age (84.9 percent of the 39Ar released) of 99.05 ± 0.9 Ma (Fig. 2-38I), concordant with the inverse isochron age, which defines an initial 40Ar/36Ar ratio of 296, close to the atmospheric value. The plateau age is slightly younger than those of the two dated Stage J-V microclines. A microcline vein (MA17-9) which reopened a calcite-chalcopyrite vein gives an age of 95.04 ± 0.6

(with 72.6 percent of the 39Ar released, Fig 2-38J), concordant within error with the inverse isochron age, which, however, defines an initial 40Ar/36Ar ratio of 371, higher than the atmospheric value.

Microcline has a low closure temperature and the five apparent age plateaus determined for this mineral at Mina Justa could be interpreted as recording resetting by post-mineralization thermal events, ascribable to the numerous post-mineralization ocöitic andesite dikes assigned to

96 the Upper Cretaceous Tunga andesite (Caldas, 1978; Fig. 2-6). However, the considerably older plateau age yielded by the MJ-6 microcline indicates that Ar resetting of triclinic feldspar was not pervasive within the deposit. In addition, the plateau age of 109.90 ± 0.97 Ma for Stage J-III actinolite defines the lower age limit for Stages J-V and VI alteration and mineralization, and the younger feldspar dates are therefore interpreted as representing crystallization ages, particularly because they conform to the defined paragenesis (Table 2-6). We conclude that the main hydrothermal events (Stages J-IV through J-VI) at Mina Justa, i.e., those responsible for the formation of hematite and magnetite as well as the Cu mineralization, occurred at ca. 100 Ma, in the Albian. Stage J-I (albitization) and J-II (K-feldspathization), however, were much older and were not directly related to development of the orebodies.

2.9 Discussion

2.9.1 An Oxide Melt Origin for the Main Marcona Magnetite Orebodies?

As noted in the introduction, numerous authors have proposed that “Kiruna-type” magnetite deposits are products of the crystallization of Fe oxide-dominated melts or magmas (Naslund et al., 2002, and references therein). This possibility must be considered in the case of the

Marcona deposit, despite the extensive hydrothermal alteration associated with the magnetite orebodies hosted by the Marcona Formation siliciclastics and overlying Río Grande Formation andesites. Whereas the elevated temperatures estimated herein, i.e., ≥ 700°C, or even > 800oC, for Stage M-III, do not preclude a hydrothermal origin, and are indeed shared by numerous magnetite-rich porphyry Cu-Au deposits with unambiguous hydrothermal affinities (e.g., Bajo de la Alumbrera: Ulrich et al., 2002), the megascopic forms and contact relationships of the massive magnetite bodies are, we argue, best explained as recording the intrusion of dense, albeit volatile-rich, melts (e.g., Fig. 2-12). The dike-like apophyses of the Mina 11 orebody (Fig. 2-11)

97 could also be interpreted as products of melt injection. More diagnostic are the intimate intercalations of smoothly-bounded magnetite and dacitic/rhyodacitic porphyry bodies (Figs. 2-13

– 2-15), best shown by the Minas 2-3-4 orebody (Fig. 2-13). This largest mineralization centre, in particular, appears to have been emplaced coherently as a bimodal system into the Marcona

Formation, and the amoeboid contact relationships (Fig. 2-13) closely resemble those of commingled mafic and felsic silicate melts in polyphase intrusions worldwide. Critically, neither the magnetite or dacite bodies comprising Minas 2-3-4 extend individually into the Marcona

Formation host-rocks. Moreover, comparable magnetite/dacite relationships are not observed in porphyry Cu-Au deposits, in which the magnetite is demonstrably metasomatic (e.g., Arancibia and Clark, 1996; Ullrich et al., 2001). However, a broadly similar environment has been documented by Lundberg and Smellie (1979) for the Painirova and, particularly, Mertainen magnetite-dominated deposits of the Kiruna district, Sweden. There, magnetite occurs as massive lenses, breccia matrices and disseminations restricted to a single, probably subaerial, sequence of lavas and tuffs with a Fe- and Na-rich trachytic composition. Critically, massive magnetite lenses at Mertainen are surrounded by low-grade sections in which magnetite occurs largely as small (~

1-10 mm diameter) subspherical globules, many surrounded by magnetite-depleted trachyte.

Immicibility between P-rich Fe oxide and alkali- and Fe-rich trachytic liquids is the most plausible interpretation. The Minas 2-4 orebody (Fig. 2-13) would, on this basis, be interpreted as analogous, despite the apparent absence of small-scale globular magnetite textures.

The magnetite breccia body in Mina 1, in which the ellipsoidal bodies and angular blocks of massive magnetite are enclosed by a matrix of powdery magnetite (Fig. 2-16), unambiguously intruded Marcona Formation metasediments and may be analogous to the tuffisitic ores documented by Nyström (1985) and others in the Kiruna district. Such association are tentatively interpreted as the product of the crystallization and explosive vesiculation of a hydrous

Fe oxide melt at shallow depth. The ellipsoided bodies may represent either pebbles produced by

98 abrasion or hypogene exfoliation or, more probably, pillows of originally molten magnetite forming in a water-rich environment, possibly resulting from the downward penetration of seawater. A vuggy texture, identified as a characteristic of magmatic magnetite bodies (Naslund et al., 2002), is also locally developed in the Marcona orebodies (Fig. 2-17A). The euhedral octahedra are interpreted as having developed in vesicles, and exhibit striations (Fig. 2-17B) which may represent columnar forms, elsewhere considered as evidence of a magmatic origin

(Henríquez and Martin, 1978). It should, however, be emphasized that many smaller-scale textural features suggestive of a melt origin (e.g., Nyström and Henriquez, 1994; Naslund et al.,

2002) are not shown by the massive magnetite-rich assemblages at Marcona.

Despite these uncertainties, we propose that the greater part of the high-temperature magnetite mineralization at Marcona, i.e., Stage M-III, probably records the emplacement of oxidic melts which were commingling with silicic melts. The ubiquitous fine intergrowths of magnetite and calcic amphibole (actinolite and/or tremolite), a characteristic of many of the

Cretaceous iron deposits of Chile (e.g., Bookstrom, 1977), could represent the crystallization of relatively Si-, Ca- and Mg-rich oxide melts, such as were synthesized at 650-800°C at a pressure of 1 kbar by Lledó (2005). Although a feature not documented by that author, the locally abundant biotite and phlogopite in the main Marcona magnetite orebodies imply that the oxide melts also contained significant K2O. The intense proximal biotitization and weak distal actinolite-tremolite alteration directly associated with Stage M-III orebodies would, in this model, be ascribed to a substantial water content in the melts. In this aspect, Marcona differs from the majority of Kiruna-type magnetite bodies for which a melt origin has been proposed, e.g., Sierra

La Bandera (Lledó, 2005), Kiruna (Frietsch, 1978) and El Laco (Naslund et al., 2002), in which

K-silicate alteration is much less widespread than amphibole- or clinopyroxene-rich facies, and where the orebodies contain only minor, erratically distributed trioctahedral mica.

99 In this framework, the albite-scapolite alteration (i.e., Na-Cl metasomatism) at Marcona, especially that developed in the dacite porphyry hosting the magnetite orebodies, is inferred to have been generated by brines released from the iron oxide melt. The poorly-developed actinolite

- tremolite alteration associated with Stage M-III magnetite mineralization would, conversely, be in conformity with Ca retention in the melt phase due to extensive amphibole precipitation (Lledó,

2005). A feature common to Marcona, Kiruna, El Laco and Sierra La Bandera is that demonstrably metasomatic magnetite replacement is volumetrically minor, whereas in lower-temperature, entirely hydrothermal, IOCG deposits, e.g., La Candelaria (Ullrich and Clark,

1999), Mantoverde (Benavides et al., 2007), the Tennant Creek district (Skirrow and Walsh, 2002) and Starra (Williams et al., 2001), intense Fe metasomatism was responsible for the development of massive bodies of so-called “ironstone” (Williams et al., 2001). The high-T (700->800°C), sulphide-free, Marcona Stage M-III magnetite-amphibole and magnetite-biotite (± amphibole) assemblages indicate that the inferred oxide melt crystallized in a high fO2 (Log fO2: –9 to -15)

o and low fS2 (Log fS2: < – 4) environment, as defined in the Fe-O-S-Si system at 800 C and 2 kbar by Shi (1992), whereas the markedly lower-T (~ 430-600°C), sulphide-rich Stage M-IV assemblages could record quenching of hypersaline fluids released through vesiculation of the commingling oxide (e.g., Broman et al., 1999) and silica-rich melts. Magnetite coexists with

o Fe-sulphides over a wide fO2 - fS2 field at temperatures of ≤ 600 C and at 2 kbar (Shi, 1992).

Origin of oxidic melt: Interpretation of the processes responsible for formation of the proposed Fe oxide melts at Marcona must accommodate the larger-scale petrological environment, and in particular the evidence for a diversity of silicate magmas broadly contemporaneous with mineralization. However, although modestly potassic (Aguirre, 1988), the dominant magmas at this stage, i.e., the andesites of the upper Río Grande Formation, exhibit compositional features typical of Central Andean arcs. Acceptance of an origin through melt immiscibility for the magnetite bodies implies that they crystallized from a silica-poor melt

100 coexisting with a silica-rich melt, plausibly represented by the associated dacitic/rhyodacitic porphyry. Roedder (1951) was the first to document a germane stable liquid immiscibility field in the system fayalite - leucite - silica. Visser and Koster van Groos (1979) subsequently demonstrated that even modest concentrations of P2O5 markedly increase the extent of this field at elevated temperatures, and Philpotts (1982) delimited an even wider two-liquid field on the basis of natural igneous suites and in terms of the expanded system Na2O + K2O + Al2O3 + MgO

– FeO + MnO + TiO2 + CaO + P2O5 – SiO2. Andean Mesozoic and Cenozoic calc-alkaline – to – shoshonitic volcanic and plutonic (granitoid) suites are dominated by intermediate members, i.e., andesites and quartz diorites, and reflect crystal fractionation and magma mixing relationships, as well as varying degrees of assimilation of upper crustal lithologies. Their compositional trajectories would not normally intersect the stable immiscibility field outlined by Philpotts (1982;

Fig. 2-39). However, Lledó (2005) has shown experimentally that, owing to the strong melt polymerization characteristic of the final stages of crystallization, even P2O5-poor andesitic magmas readily generate immiscible Fe-P-oxide melts. Alternatively, mixing of andesitic and more silicic melts may have prompted immiscibility. Thus, Matthews et al. (1995) observed the development of globules of Fe oxide melt in the Quaternary Lascar flow, Chile, resulting from the mixing of basaltic andesite and strongly oxidized, anhydrite-phyric, dacitic melts. Although the compositional trajectories of the Lascar flow would not intersect the stable immiscibility field outlined by Philpotts (1982; Fig. 2-39), an enlarged immiscibility field plausibly generated by high fO2 (pyrrhotite/anhydrite buffer: Matthews et al., 1995) may have existed (Fig. 2-39). Such enlargement of the stable immiscibility field at high fO2 is supported by experiments (Naslund,

1983). In contrast, Clark and Kontak (2004) documented the formation at Antauta, Perú, of coexisting Fe oxidic and rhyolitic melt droplets through the dissolution of quartz phenocrysts in an exceptionally reduced rhyodacitic magma undergoing commingling and mixing with a more oxidized absarokite melt. Such specific dissolution of quartz would have an effect on parental

101 melt composition differing radically from that generated through crustal assimilation caused by crystal factionation. Eleswhere, and also on a microscopic scale, Naslund et al. (2004) record the trapping of inclusions of rhyolitic glass with subordinate magnetite globules and apatite and pyroxene daughter crystals in orthopyroxene phenocrysts in the El Laco magnetite-pyroxene flows.

Figure 2-39. Potential compositional trajectories of volcanic rock suites from different locations. Two immiscibility fields are shown, as defined by Roedder (1951) and Philpotts (1982). Part of an immiscible boundary at high fO2 is inferred from Matthews et al. (1995, 1999). The Andean trend, which represents processes combining assimilation-fractional crystallization (AFC) and magma mixing, shown for comparison, appears not to intersect an immiscibility field. However, an andesite melt (upper Río Grande Formation) may experience immiscibility by either quartz dissolution (cf. Clark and Kontak, 2004) or through enlargement of the immiscible field at high fO2 (cf. Matthews et al., 1995). 102

At Marcona, therefore, volumes of upper Río Grande Formation andesitic magma may have generated immiscible Ca, Mg and Si-bearing Fe oxide and felsic silicic melts as a result of their extensive crystallization and polymerization, possibly abetted by enrichment of volatiles such as

P and, given the scarcity of apatite in the orebodies, more plausibly S (Lester, 2002). An open system is envisaged, wherein proportionately subordinate volumes of magnetitite and dacite/rhyodacite melt accumulated within Repetición faults which channelled the parental mafic magmas which fed the 2000-3000 m-thick upper Río Grande Formation andesitic lave pile.

Further, quartz dissolution around bodies of andesitic magma ponded in the Marcona Formation siliciclastics (Clark and Kontak, 2004: Fig. 2-39) may have contributed to immiscibility, perhaps explaining the restriction of the largest magnetite orebodies to the Paleozoic basement beneath the arc. The high proportion of magnetitite to dacite/rhyodacite porphyry in the most important,

Minas 2-3-4, orebody (cf. Clark and Kontak, 2004) may reflect the denser, less buoyant nature of the former melt.

2.9.2 Evolution of the Marcona-Mina Justa District

The Mesozoic magnetite-rich mineralization of the Marcona – Mina Justa district, extending over an area of ~ 75 km2 and encompassing numerous dispersed orebodies, bears a superficial resemblance to the two Central Andean IOCG centres, La Candelaria – Punta del Cobre and

Mantoverde, which have produced significant copper over the past two decades. Of these, at least the former was the site of Fe-Cu-Au mineralization over a protracted interval, i.e., ≥ 10 m.y.

(Pop et al., 2000). However, each of these districts may be considered to comprise a cluster of interrelated deposits, by analogy with the Río Blanco – Los Bronces (Deckart et al., 2005) and

Los Pelambres – El Pachón (Bertens et al., 2006) porphyry Cu-Mo districts. In contrast, in the

Marcona – Mina Justa district differs in that hydrothermal alteration appropriate for IOCG systems occurred episodically over a period of ca. 80 m.y., extending from 176 to 95 Ma.

103 Moreover, the Mina Justa Cu mineralization is shown to have been emplaced at least 60 m.y. after the Marcona Fe (-Cu) deposit (Fig. 2-40) and cannot, therefore, be genetically related to it.

In the absence of U-Pb or Re-Os age data, clarification of the history of alteration and ore formation in the Marcona – Mina Justa district relies almost entirely on 40Ar/39Ar geochronology.

Whereas even perfect 40Ar/39Ar apparent age plateaus for hydrothermal micas in giant, longlived porphyry centres (e.g., Río Blanco: Deckart et al., 2005) have been shown to record resetting by late-stage, high-temperature, hydrothermal or magmatic activity, the ages recorded herein for hydrothermal calcic amphibole, trioctahedral micas and even alkali feldspars conform satisfactorily to observed paragenetic relationships. Further, stable isotope data and fluid inclusion microthermometry (Chapter 3) demonstrate that much of the late subeconomic Cu sulphide mineralization at Marcona and all economic Cu sulphide mineralization at Mina Justa took place at or below ~ 250°C, a temperature at which complete Ar-outgassing of the dated

K-bearing minerals would be unlikely.

Figures 2-41 and 2-42 illustrate the major events in the district recorded by the newly defined paragenetic and geochronologic relationships. It should be emphasized that the geochronological coverage of this complex mineralized area remains sparse and, given the extended history of alteration and mineralization, it is very probable that the number and duration of hydrothermal episodes are incompletely defined. We therefore subdivide the evolution of the district entirely on the basis of the available 40Ar/39Ar data, with the possibility of either over- or under-discrimination between events. It should be emphasized that the plateau ages determined for alteration minerals with high Ar-retention temperatures (i.e., cummingtonite, phlogopite and actinolite-tremolite) permit reliable assignment of alteration events at Marcona to specific stages in the eruption of the Río Grande volcanic arc.

104

Figure 2-40. Laser-induced 40Ar/39Ar plateau ages for Marcona and Mina Justa alteration and mineralization stages.

105

Figure 2-41. Cartoon of the evolution of the Marcona deposit.

106 (a) Precursor alteration, ca. 177 Ma (Fig. 2-41A): Cummingtonite-dominated alteration

(Stage M-IA, Fig. 2-19) in the upper metasedimentary units of the Marcona Formation (~ 300 m vertically below the paleosurface) was apparently unassociated with either magnetite or Cu sulphide mineralization. Development of an alkali-poor Fe-Mg amphibole implies either that the metaclastic host rocks had earlier suffered feldspar destruction without the formation of muscovite (cf. Fig. 2-4A), possibly as a result of chloritization around hot springs at the sediment-ocean interface, or, more probably, may directly record Fe-Mg metasomatism. Similar early development of Mg-Fe clinoamphibole has been documented from several other IOCG deposits, including: Vähäjoki, Finland (Liipo and Laajoki, 1991); Osborne, Cloncurry (Adshead,

1995); La Candelaria, Chile (Ullrich and Clark, 1997, 1999); Salobo, Carajás (Requia and

Fontboté, 1999; Requia et al., 2003); and Guelb Moghrein, Mauritania (Kolb et al., 2006). An

Mg-Fe metasomatic origin is accepted in the majority of these examples. Moreover, cummingtonite develops through post-metamorphic, high-temperature (> 500oC), replacement of earlier carbonate-bearing alteration zones in some giant hydrothermal hematite deposits such as those of the Krivoy Rog district, Ukraine (Belevtsev, 1973; Dalstra and Guedes, 2004).

Despite Cenozoic partial resetting (Figs. 2-37 A, B), the two Ar/Ar age spectra for cummingtonite confirm that this precursor alteration took place in the Aalenian, contemporaneous with the initial sediment-dominated, shallow-marine, accumulation of the Río Grande Formation.

Hydrothermal activity is therefore inferred to have taken place at very low pressure and at temperatures considerably below 760°C (Evans and Ghiorso, 1995). A plausible environment would be the SW margin of an arc-parallel rift delimited by segments of the NE-dipping Pista fault system (Fig. 2-41A), recording large-scale NE-SW extension within the West Peruvian trough at the outset of the Middle Jurassic (Benavides - Cáceres, 1999).

The lowermost units of the Río Grande Formation, both at Marcona and in the Cañón Río

Grande (Fig. 2-5), lack major volcanic or pyroclastic units, but this earliest alteration may have

107 been promoted by intermediate or basic magmas which did not breach the ocean floor. Although seawater incursion probably occurred, it would be unlikely to have resulted directly in the formation of cummingtonite.

(b) Precursor phlogopite-magnetite mineralization, 171 Ma (Fig. 2-41B): The disturbed

40Ar/39Ar age spectra and inverse isochron ages for hydrothermal phlogopite coexisting with magnetite in the upper Marcona Formation are interpreted as evidence either for a discrete alteration-mineralization event (M-IB) in the Bajocian or, less probably, for a continuation of the hydrothermal activity recorded by the 177 Ma cummingtonite. The mineral assemblage is indicative of intense local K-Fe metasomatism. At this stage, intermediate – to – silicic pyroclastic activity was well underway and a significant thickness of the lower Río Grande

Formation had locally accumulated, in permissive agreement with a magmatic contribution to this apparently minor phase of magnetite mineralization. Localization by continued displacement on the Pista Fault system is favoured (Fig. 2-41B).

(c) Albite-scapolite alteration (Stage M-II and early-Stage M-III; Fig.2- 41C): No age data are available for the Na-Cl metasomatism which generated the widespread Stage M-II albite ± marialitic scapolite alteration predating Stage M-III: this can only be bracketed between ca. 171 and ca. 162 Ma. Albitization at Marcona is markedly less intense than that in other IOCG centres, e.g., the La Candelaria deposit (Ullrich and Clark, 1999), but is similarly unassociated with either magnetite or sulphide deposition. As at La Candelaria, moreover, albitization at

Marcona involved the destruction of Ca- and K-bearing minerals and, although incomplete, cannot be designated “Na-Ca alteration”. Albite (-scapolite) alteration is also developed in both dacite porphyry and Marcona Formation rocks in immediate contact with the Minas 2-4 magnetite orebody, and therefore persisted into Stage M-III (Fig. 2-41C).

(d) Main magnetite and magnetite-sulphide mineralization (Stages M-III and M-IV), 159-162

Ma (Figs. 2-41C-E): The excellent 40Ar/39Ar age plateaus determined for Stage M-III biotite

108 (161.4 Ma) and Stage M-IV phlogopite (159.7 Ma) directly associated with, respectively, the major Mina 5 and Mina 3 magnetite orebodies, and coinciding with the 160 Ma K-Ar phlogopite date reported by Injoque (1985), indicate that the most important magnetite mineralization at

Marcona took place close to the Bathonian – Callovian boundary (Pálfy et al., 2000). Although the two Ar/Ar dates are almost coincident at 2σ error, they are in agreement with the development of biotite largely prior to, and at higher temperatures than, the phlogopite (Fig. 2-19; Table 2-5).

These data demonstrate that the initial, minor, deposition of chalcopyrite at Marcona occurred at ca. 160 Ma. Both the extremely high temperatures (≥ 700oC) of the major magnetite mineralization (Stage M-III) and the moderate – to – high temperatures for the magnetite-sulphide assemblages (430-600oC, Stage M-IV) predicate the involvement of magma, either as a reservior of iron oxide melt or as a source of magmatic-hydrothermal fluids. The lack of Ar-resetting of the precursor cummingtonite and phlogopite alteration assemblages in the envelope of the Mina 5 magnetite orebody may reflect either the high Ar-retention temperatures of these minerals (i.e., >550 and >500oC, respectively: Harrison, 1981; Kelly and Wartho, 2000;

Jenkin et al., 2001), or the rapid emplacement of the oxidic melts.

Iron oxide and sulphide mineralization at Marcona therefore shortly followed the ca. 164-166

Ma hiatus in andesitic volcanism documented in the upper Río Grande Formation (Figs. 2-5 and

2-41C). This erosional event records a regional, terminal-Bajocian (ca. 166 Ma), uplift event which interrupted the longterm subsidence of the Western Peruvian Trough (Hosmer, 1959).

Continued or renewed uplift between 163 and 160 Ma is documented by Quang (2003) for the

Cocachacra area (Lat. 17°S), where aluminum – in – hornblende geobarometry records exhumation of granitoid plutons prior to the Callovian reestablishment of the marine Guaneros

Formation arc, correlative with the upper Río Grande Formation (Fig. 2-5). In the Marcona area, uplift was accommodated by the predominantly reverse, ENE-trending Repetición Faults, which developed along the main linear array of the earlier Pista Faults as a result of ESE-WNW

109 contraction, probably as a response to major sinistral transcurrent displacement on the Treinte

Libras fault system to the northeast (Figs. 2-2 and 2-41C). Following the uplift and erosion, the late Bathonian-Oxfordian (ca.155-164 Ma) submarine upper-Río Grande Formation volcanic and volcanosedimentary units (Fig. 2-5; Aguirre, 1988), broadly coeval with the marine Chala and

Guaneros Formation arcs to the south (Romeuf et al., 1993), accumulated within a newly-developed rift close to the continental margin. Magnetite and the subsequent sulphide mineralization at Marcona were emplaced along the Repetición Faults during accumulation of the upper Río Grande Formation from ca. 162 Ma (Fig. 2-41D).

The intimate association of bodies of massive magnetite and dacitic porphyry, particularly in the largest, Minas 2-3-4, orebody, implies a direct genetic relationship between Stage M-III magnetite mineralization and intermediate magmatism. Melt fractionation under high fO2 conditions, possibly accompanied by the dissolution of quartz from Marcona Formation siliciclastic host-rocks (cf. Clark and Kontak 2004), generated coexisting silicic (dacitic) and oxidic (Fe3O4 ≥ 70%) immiscible melts, the commingling of which constructed the major Fe orebodies (Stage M-III: Fig. 2-41D). Crystallization of the oxide melts at 700-800oC to form magnetite-actinolite (±biotite) aggregates released the brines responsible for the intense

K-metasomatic haloes around the orebodies. Subsequent quenching of the magmatogene fluids through 400oC at ca. 159-161 Ma (Stage M-IV: Fig. 2-41E) precipitated further magnetite and considerable pyrrhotite and pyrite, but only minor chalcopyrite, with the continued formation of calcic amphibole, the termination of K-feldspathization and the supplanting of biotite by phlogopite. These potassic and calcic alteration assemblages were superimposed on moderately to strongly albitized rocks, and record closely linked K- and Ca - metasomatism.

(e) Polymetallic sulphide mineralization (Stage M-V), 156-159 Ma (Fig. 41F): the main base-metal sulphide stage in the Marcona deposit saw a decreased deposition of pyrrhotite and a marked increase in pyrite and chalcopyrite, particularly in the Mina 11 orebody (Fig. 2-6; Table

110 2-2). Sphalerite is abundant (Table 2-2) in the Mina 14 orebody (Fig. 2-6), where it is associated with minor galena. Polymetallic sulphide mineralization accompanied continued crystallization of calcic amphiboles (Fig. 2-19), both as pervasive alteration zones and veins, but also saw a major increase in calcite deposition, commonly in cavities in vuggy magnetite.

Chalcopyrite-pyrite enrichment was therefore associated with calcic alteration. Although much of the chalcopyrite in Mina 11 formed with magnetite and amphiboles (Fig. 2-26), a significant proportion was deposited with late-stage anhydrite, more abundant here than in more westerly orebodies. In comparison to the high-temperature and strongly oxidized conditions prevailing during earlier stages, the magnetite-free and sulphide-dominated Stage M-V assemblages formed

o in a relatively reduced (pyrite-pyrrhotite fO2 buffer), and much cooler (ca. 160 – 360 C) environment, which may record the incursion of non-magmatic fluids.

Three Stage M-V calcic amphiboles from the Mina 3 and Mina 5 orebodies yield acceptable

Ar/Ar age plateaus in the range 156.2 to 158.5 Ma (Table 2-6). Although two of the dates overlap within error with those for Stage M-III micas, suggesting that sulphide deposition shortly followed the main magnetite mineralization, the youngest, 156.2 ± 2.4 Ma, implies that amphibole formation continued into the Oxfordian. Therefore, intense hydrothermal activity at

Marcona, initiated following the ca. 164-166 Ma uplift event at the Bajocian-Bathonian boundary, persisted until the termination of andesitic volcanism in the extensional marginal rift (Fig. 2-41F).

The ages of the ensuing Stages M-VI and M-VII are undefined. Because zeolite facies assemblages occur in the andesite flows of the upper Río Grande Formation (Aguirre, 1988), but not in the overlying Jahuay Formation, this nondeformational metamorphism was, at least in part, broadly contemporaneous with the Marcona Fe and sulphide mineralization, recording either seawater/rock interaction (Aguirre, 1988) or diastathermal processes (Alt, 1999), i.e., intense heat flow during the initial development of a basin in which the Jahuay Formation sediments and andesites were deposited (Caldas, 1978). This close temporal association of Fe- and Cu

111 sulphide mineralization with very low-grade regional metamorphism contrasts with the sequence of events documented by Benavides et al. (2007) in the Mantoverde district, Chile, where IOCG mineralization postdated metamorphism in the mid-Cretaceous back-arc basin.

(f) K metasomatism in the Mina Justa area, 142 Ma (Fig. 2-42A): this probably weak

Tithonian-Berriasian alteration event is recorded by a single Ar/Ar plateau for microcline from an andesite of the upper Río Grande Formation. The structural context of the potassic alteration is uncertain, but it is likely that it was focused by faulting related to renewed movement on the

Treinte Libras fault system to the NE (Figs. 2-2 and 2-42A). By ca. 142 Ma, Jahuay Formation volcanic activity had apparently terminated (Caldas, 1978) and no potentially parental intrusive bodies have been recognized in the immediate area. Injoque (1985, p. 251), however, reports whole-rock K-Ar dates of 137.4 ± 3 and 136.4 ± 3 Ma for two “basic dykes” in the Marcona mine area, evidence for an otherwise undocumented Neocomian magmatic event. K metasomatism may also have been a response to fluid expulsion resulting from continued or renewed detachment faulting (cf. Chapin and Lindley, 1986; Roddy et al., 1988) along the SW margin of the evolving Jahuay basin, or to latest-Tithonian inversion of the basin.

(g) Magnetite and Cu sulphide mineralization at Mina Justa, 95 - 104 Ma (Figs. 2-42B and

C): One Ar/Ar date for actinolite and four for microcline record the major events which generated the Mina Justa orebodies, defining a protracted Albian-to-Cenomanian history of hydrothermal activity (Figs. 2-42B and C). Thus, magnetite-pyrite-quartz – dominated assemblages assigned to Stage J-V give ages of 103.7 and 101.5 Ma, significantly postdating the 109 Ma Stage J-III actinolite alteration. These thermal events apparently also affected the Marcona deposit, in which

Stage M-III microclines were reset at 101-109 Ma. Feldspathic alteration envelopes to chalcopyrite-calcite veins representing the main, Stage J-VI, Cu mineralization formed at 99.1

Ma, and terminal, late-Stage VI microcline veins associated with minor chalcopyrite-calcite veins

112

Figure 2-42. Cartoon of the evolution of the Mina Justa deposit.

113 record the apparent termination of sulphide deposition at 95.0 Ma (Table 2-6). As noted previously, the Mina Justa microclines have Ar/Ar ages conforming with paragenesis, precluding wholesale resetting. The age of the younger, barren, hematite veins is uncertain. By the late

Albian, accumulation of the dominantly clastic Copara Formation had probably ceased in the

Cañete Basin (Caldas, 1978); the closest outcrop of this unit is ~ 12 km east of Mina Justa (Fig.

2-42B). A short-lived uplift and erosional event preceeding deposition of the Copara Formation, recorded by an unconformity with the underlying Yauca Formation (Fig. 2-3), exhumed the upper

Río Grande Formation in the Mina Justa area (Fig. 2-42B). Further, the hypabyssal intrusion of the andesitic Bella Unión Complex (Caldas, 1978), extensively exposed north and east of Mina

Justa, probably significantly antedated mineralization. In contrast, the age range for hydrothermal microclines coincides well with the early stages of emplacement of the Coastal

Batholith in the Acarí area to the east (Fig. 2-2), i.e., ≤ 109 ± 4 Ma.

Numerous authors (e.g., Jaillard et al., 2000; Oyarzun et al., 2003) have emphasized the geodynamic impact of the reorganization of plate interactions in late-Neocomian time, probably triggered by the development of a Mid-Pacific superplume (Larson, 1991). Coupling of the converging plates, with a northeast vector, ended a long period of orthogonal extension and sinistral transtension along the central South American littoral, which gave rise to dextral transtension and the formation of mid-Cretaceous basins (Polliand et al., 2005). The subsequent increasing coupling rate resulted in uplift and basin inversion, accompanied by emplacement of the earliest, ca. 109 Ma, plutons of the Arequipa Segment of the Coastal Batholith (Vidal et al.,

1990; Atherton, 1990). In the Marcona area, dextral transtension on the Treinte Libras fault system (Pope, 2003) generated the ENE-striking Mina Justa normal faults which controlled the emplacement of both magnetite lenses and Cu orebodies (Figs.2- 42B and C). However, Cu mineralization at Mina Justa (Stage J-VI; 95.0 - 99.1 Ma) significantly postdated magnetite-pyrite alteration (Stage J-V; 101.5 - 103.7 Ma).

114 It is unlikely that exotic, basin-derived fluids contributed to the early high-temperature

(~600oC) magnetite-pyrite alteration stage at Mina Justa, despite the proximity of the Cañete

Basin to a well-developed fault system in the Río Grande Formation (Fig. 2-42B) which could have served as a pathway for fluid circulation. The high-T, Cu-barren, magnetite-pyrite assemblage may be evidence for either a Cu-poor magmatic fluid or the suppression of Cu sulphide deposition by high temperatures (~600oC: Hezarkhani et al., 1999). During further compression and basin inversion, exotic, possibly basinal fluids, are inferred to have invaded the

Jurassic andesite succession along the deep detachment fault depicted in Figures 2-42B and C.

However, these exotic fluids were plausibly driven by heat from deep-seated intrusions, represented by the dioritic stocks in the Mina Justa area, and the possibility of mixing with magmatic fluids cannot be excluded. Ore-forming fluids rose along the Mina Justa fault system and replaced the Stage J-V magnetite-pyrite bodies, generating the Cu mineralization (Fig.

2-42C). The coexistence of hematite with bornite-chalcocite-chalcopyrite assemblages indicates a relatively low T and P (i.e., <500oC and < 1kbar: Hemley et al., 1992) and a medium -to- high pH

(calcite stable) environment for Cu mineralization. This mineralization event is unrecorded in the

Marcona deposit, where the minor Stage M-VII hematite veins are sulphide-free. The Stage J-VI hydrothermal system was therefore restricted arealy to the margin of the Cañete basin.

2.10 Conclusions

Alteration and mineralization persisted episodically for at least 80 m.y. in the Marcona-Mina

Justa district, i.e., from 177 to 95 Ma. The earliest hydrothermal alteration at Marcona, represented by cummingtonite and magnetite-phlogopite assemblages (Stage M-I), occurred in the Aalenian, whereas the massive magnetite-Ca amphibole - trioctahedral mica bodies did not form until the late Bathonian. The 159-162 Ma magnetite orebodies are interpreted as recording an evolution from high-temperature (700->800°C), hydrous, vesiculating, Fe oxide-dominated

115 melts (Stage M-III), to rapidly cooling (400-600°C) endogenous hydrothermal fluids (Stage M-IV) which precipitated Fe sulphides with magnetite. Intense K metasomatism of host rocks, represented by biotite and K-feldspar alteration, was intimately associated with magnetite mineralization. Subsequently, at 156-159 Ma, magnetite-free sulphide - calcic amphibole and calcite assemblages, accompanied by a modest enrichment in Cu, Zn and Pb, formed at markedly lower temperatures (160-360oC) in a relatively reduced environment (pyrite-pyrrhotite buffer), possibly triggered by the incursion of seawater. Simultaneously, at 156-157 Ma, hydrothermal activity in the Mina Justa area generated albite-actinolite alteration (Stage J-I), followed at ca.

142 Ma by K-Fe metasomatism (Stage J-II). However, the major high-T (540-600oC) magnetite-pyrite alteration (Fe metasomatism) and low-T Cu mineralization at Mina Justa were not initiated until ca. 104 Ma, following precursor actinolite alteration (Ca metasomatism) at ca.

109 Ma. The 95-99 Ma Mina Justa Cu (-Ag) orebodies, dominated by Cu sulphide - calcite - hematite assemblages, are inferred to be the product of high fO2 and low-temperature (≤ 250°C) fluids which partially replaced the ca. 2.5 m.y. - older magnetite-pyrite bodies.

The Bathonian-Callovian Marcona Fe deposit formed in and beneath a failing Middle

Jurassic andesitic arc, following a brief eruptive hiatus at 164-166 Ma. The emplacement of massive magnetite bodies and subsequent polymetallic sulphide precipitation were contemporaneous with eruption of the upper Río Grande Formation in a shallow-marine environment. Parental andesitic or basaltic andesitic magmas are envisaged as experiencing mixing with more silicic melts and silicification through the dissolution of quartz from the

Marcona Formation siliciclastics, resulting in the development of coexisting, immiscible, Fe oxide and dacitic melts. Similar processes may have occurred in the Pliocene magnetite flows in a subaerial andesitic arc at El Laco, northern Chile and, in a shallow-marine to subaerial environment, in the Cu-poor iron deposits of the Cretaceous Chilean iron belt (CIB). Marcona, like El Laco and the CIB deposits, is therefore interpreted as a largely endogenous product of arc

116 magmatism, with minimal intervention by non-magmatic fluids. However, these Andean,

“Kiruna-type”, iron deposits formed in different tectonomagmatic settings, viz.: Marcona during basin formation in an extensional andesitic-dacitic arc; the Cretaceous Chilean iron deposits during the inversion of basins contiguous to an extensional arc; and El Laco during the eruption of a subaerial andesitic-dacitic arc in an extensional environment (Kay and Kay, 1993; Marrett et al., 1994).

In contrast, the upper Albian Mina Justa Cu deposit, and other major Cu-rich central Andean

IOCG deposits, viz. La Candelaria-Punta del Cobre, Mantoverde, and Raúl-Condestable, were entirely products of hydrothermal activity adjacent to mid-Cretaceous basins undergoing tectonic contraction and inversion, stimulated by marginal granitoid plutons. The latter may have contributed modest volumes of magmatic-hydrothermal fluid, but the incursion of exotic fluids with evaporite-sourced sulfur was probably a prerequisite for economic Cu mineralization, as proposed by Ullrich and Clark (1999) and Benavides et al. (2007).

The Marcona magnetite and Mina Justa iron oxide-rich Cu (-Ag) deposits, therefore, although areally associated, represent two temporally separated mineralization centres and contrasted ore deposit types, and bear no direct genetic relationships. The only modest associated magmatogene Cu-Zn sulphide mineralization at Marcona may imply a limited partitioning of Cu and Zn from parental arc magma into immiscible Fe oxide melts. A requirement for the involvement of external fluids for Cu mineralization in Mina Justa and other major Andean IOCG deposits underlies the lack of Cu enrichment in the majority of hydrothermal Cu-poor IOCG centres in the Central Andes and elsewhere, in which a high-T magmatic fluid predominated in the ore-forming system.

The protracted history of alteration and mineralization documented herein for the Marcona and Mina Justa deposits, i.e., ~ 20 m.y. at Marcona and ~ 60 m.y., respectively, recording the

117 multiple stages of the ore-forming systems, has been recognized in other major IOCG centres in the Andes and elsewhere. Early-stage, IOCG-like hydrothermal alteration with or without minor mineralization, significantly predating the main economic mineralization, is common in the major

IOCG centres.

118 Chapter 3

CONTRASTED FLUIDS AND RESERVOIRS IN THE CONTIGUOUS

MARCONA FE AND MINA JUSTA IRON OXIDE-CU (-AU-AG) DEPOSITS,

SOUTH-CENTRAL PERÚ

3.1 Abstract

The Marcona-Mina Justa deposit cluster, Ica Department, hosted by Lower Paleozoic metaclastics and Middle Jurassic shallow-marine andesites, incorporates the most important known magnetite mineralization in the Andes at Marcona (1.9 Gt @ 55.4% Fe and 0.12% Cu) and one of the few major Andean IOCG deposits with economic Cu grades (346.6 Mt @ 0.71%

Cu, 3.8 g/t Ag and 0.03 g/t Au) at Mina Justa. The Middle Jurassic Marcona deposit is located 3-4 km southwest of the mid-Cretaceous Mina Justa Cu-(Ag, Au) prospect.

At Marcona, magnetite-biotite-calcic amphibole assemblages, assigned to paragenetic Stage

M-III, are inferred to have crystallized at 700->800oC from a Fe-oxide melt with δ18O of +5.2 to

+7.7‰. Overprinted Stage M-IV magnetite-phlogopite-calcic amphibole-sulphide assemblages are associated with the exsolution of 430-600oC aqueous fluids with dominantly magmatic isotopic compositions (δ34S=+0.8 to +5.9‰; δ18O = +9.6 to +12.2‰; δD = -73 to -43‰ and δ13C

= -3.3‰). Stages M-III and M-IV account for over 95% of the magnetite mineralization at

Marcona. Subsequent, non-economic, lower-temperature, sulphide-calcite-amphibole assemblages (Stage M-V) were deposited from fluids with similar δ34S (+1.8 to +5.0‰), δ18O

(+10.1 to +12.5‰) and δ13C values (-3.4‰), but higher δD values (average -8‰). Several groups of low-temperature (< 200oC, with a mode at 120oC) and high-temperature (> 200oC) fluids are recognized in the main polymetallic sulphide Stage M-V and may record the involvement of modified seawater.

119 At Mina Justa, early, non-economic, magnetite-pyrite assemblages precipitated from a magmatic fluid (δ34S= +0.8 to +3.9‰; δ18O= +9.5 to +11.5‰) at 540-600oC. In contrast, ensuing chalcopyrite-bornite-digenite-chalcocite-hematite-calcite mineralization was the product of non-magmatic, probably evaporite-sourced, brines with δ34S ≥ +29‰, δ18O = 0.1‰ and δ13C =

-8.3‰. Two groups of fluids are identified at the Cu mineralization stage, viz. Ca-rich with a low temperature (ca. 140oC) and high salinity (basinal brine), and Na (-K) – dominant, with a similar temperature but low salinity (meteoric water). Laser Ablation-Time of Flight-ICP-MS analyses show that fluids at the magnetite-pyrite stage were probably Cu-barren, but that those associated with later stages were enriched in Cu, Ag and Zn.

3.2 Introduction

Iron oxide-copper-gold (“IOCG”) deposits have become a major exploration target in the last 30 years, resulting in significant clarification of their salient geological features. However, controversy regarding the nature, sources and timing of the ore-forming fluids have impeded definitive ore-genetic modeling, resulting in a broad, ill-defined clan of deposits and few widely applicable exploration tools. Characterization of the nature of IOCG ore fluids has included conventional microthermometric analysis of fluid inclusions in several major deposits, including

Olympic Dam (Oreskes and Einaudi, 1992) and Ernest Henry (Mark et al., 2000) in Australia,

Aitik in Sweden (Wanhainen et al., 2003) and La Candelaria in Chile (Ullrich and Clark 1999), but more sophisticated analytical techniques have been applied only to small representatives of the clan. Thus, proton-induced X-ray emission (PIXE) analyses of individual fluid inclusions have been documented for Starra (Williams et al., 2001) and Lightning Creek (Perring et al.,

2000) in Australia, and for the Wernecke centres in Canada (Gillen et al., 2004), and Laser

Raman Spectroscopic methods have been applied at Pahtohavare, Sweden (Lindblom et al., 1996),

Bidjovagge, Norway (Ettner et al., 1994) and Starra (Williams et al., 2001). On the basis of this 120 limited database, some researchers interpret IOCG deposits as magmatic-hydrothermal systems

o involving high-temperature (~300-450 C), high-salinity and CO2-rich ore-forming fluids (e.g., Xu,

1999; Pollard, 2000, 2001, 2006; Fu et al., 2003). Such fluids have been identified at La

Candelaria (Ullrich and Clark, 1999), Ernest Henry (Mark et al., 2000) and Eloise (Baker, 1998), in all cases directly associated with chalcopyrite-gold mineralization. High-temperature fluid inclusions were also trapped in the pre-mineralization “ironstone” stage at Starra (Williams et al.,

2001) and Lightning Creek (Perring et al., 2000). Nevertheless, many IOCG deposits, particularly at their Cu-Au mineralization stages, are characterized by medium to low-temperature (< 300oC), high-salinity, Ca-rich fluids with variable CO2 contents, ascribed to non-magmatic reservoirs, as at Olympic Dam (Oreskes and Einaudi, 1992), Pahtohavare (Lindblom et al., 1996), Aitik

(Wanhainen et al., 2003), and Wernecke (Gillen et al., 2004; Hunt et al., 2005, 2007).

Multi-source mixing models, involving fluids with variable temperatures, salinities and chemical compositions, have been proposed by Haynes et al. (1995), Ullrich and Clark (1999), Williams et al. (2001, 2005), Barton and Johnson (2004) and Hunt et al. (2005, 2007) for the Cu-Au stage of

IOCG deposits.

Stable isotope geochemistry has been used to identify fluid reservoirs in several IOCG deposits, including Olympic Dam (Oreskes and Einaudi, 1992), Raúl-Condestable (Ripley and

Ohmoto, 1979; de Haller et al., 2002, 2006), Bidjovagge (Ettner et al., 1994), Starra (Rotherham et al., 1998; Williams et al., 2001), La Candelaria (Ullrich and Clark, 1999; Ullrich et al., 2001;

Marschik and Fontboté, 2001), Lightning Creek (Perring et al., 2000), Ernest Henry (Mark et al.,

2000, 2006), Eloise (Baker et al., 2001), and Mantoverde (Benavides et al., 2007). However, the major sources of sulphur and metals remain enigmatic in many examples (Williams et al., 2005) and, as with fluid inclusion studies, interpretations of the source of the mineralizing fluids advocate two contrasted models: a dominance of magmatic-hydrothermal fluids (Pollard, 2000,

2001, 2006); and external non-magmatic fluids as a prerequisite for mineralization (Barton and

121 Johnson, 1996, 2000, 2004). Thus, at Ernest Henry (Mark et al., 2000, 2006) and Eloise (Baker et

34 al., 2001), the majority of calculated δ Sfluid values are magmatic (~ 0‰), whereas at

Raúl-Condestable (de Haller et al., 2006), La Candelaria (Ullrich and Clark, 1999; Ullrich et al.,

34 2001) and Mantoverde (Benavides et al., 2007), the calculated δ Sfluid values for Cu-Au mineralization attain much higher values predicating external fluid involvement.

Consequently, the major controversy regarding the genesis of IOCG deposits focuses on the nature and source of fluids responsible for Cu-Au mineralization and their relationship to those involved at the precursor magnetite stage. The Marcona-Mina Justa deposit cluster in south-central Perú, incorporating, respectively, the largest concentration of magnetite and one of the most important IOCG-type Cu deposits in the Central Andes, provides an excellent basis for comparison of the ore-forming fluids in IOCG systems with and without economic Cu mineralization (Fig. 3-1). The Marcona iron mine has present reserves of 1551 Mt grading 55.4 percent Fe and 0.12 percent Cu, and annual production in recent years has averaged 4-5 Mt of magnetite concentrates. Cu, Co, Ni, Zn, Pb and Au are enriched in parts of the deposit, but not recovered. The Mina Justa Cu-(Ag-Au) prospect, under development by Chariot Resources and with a 346.6 Mt resource at an average grade of 0.71 percent Cu, 3.83 g/t Ag and 0.03 g/t Au, at a cut-off grade of 0.2 per cent Cu (Mining Journal, 2006), is located 3-4 km northeast of the

Marcona mine. We herein present microthermometric data for fluid inclusions in hydrothermal calcite from Marcona and in quartz and calcite from Mina Justa, as well as compositions for quartz-hosted inclusions obtained through Laser Ablation Time-of-Flight (TOF) ICP-MS analysis.

These data are integrated with stable isotope analyses of sulphur, oxygen, hydrogen and carbon to define the relative contributions of various reservoirs to the magnetite- and sulphide-rich assemblages of the two deposits.

122

Figure 3-1. Geology of the area surrounding the Marcona deposit and Mina Justa prospect (modified from Rio Tinto, Marcona JV exploration report, June 2003). Insert A shows the location of study area and B the local major fault systems. Area of Figure 3-2 is shown.

3.3 Ore Deposit Geology

The Marcona-Mina Justa district is underlain by metamorphic rocks of the Proterozoic Arequipa

Massif, covered by Neoproterozoic and Paleozoic sedimentary strata and, more extensively, volcanic and sedimentary rocks of Mesozoic age which accumulated in a succession of intra- and back-arc rift basins (Fig. 3-1; Hawkes et al., 2002; Atherton and Webb 1989). Granitoid rocks of the Lower Palaeozoic (425±4 Ma: Mukasa and Henry, 1990; Vidal et al., 1990) San Nicolás

Batholith intruded the Lower Palaeozoic metasedimentary Marcona Formation that hosts much of the Marcona magnetite mineralization. The overlying Mesozoic volcano-sedimentary sequence

123 was intruded by the Coastal Batholith in the late Aptian (de Haller et al., 2006). The Marcona

Formation consists dominantly of thermally-metamorphosed siltstone, sandstone, minor quartz arenite and limestone. In contrast, all Cu-(Ag-Au) mineralization at Mina Justa is hosted by the

Middle Jurassic Río Grande Formation (Figs. 3-1 and 3-2), which is dominated by plagioclase-phyric andesite flows and andesitic volcaniclastic units with minor sandstone, siltstone and limestone lenses (Caldas, 1978; Hawkes et al., 2002; Baxter et al., 2005).

Massive magnetite orebodies at Marcona are lensoid to tabular, generally strike northeasterly and dip to the northwest, and are segmented by faults (Fig. 3-1). Most orebodies exhibit higher Cu grades as well as elevated total sulphide contents at shallow levels. The megascopic forms of the major magnetite bodies strongly suggest that they formed from intrusive oxide melts rather than through hydrothermal replacement.

Two principal arrays of orebodies, the main and upper, controlled by subparallel, northeast-trending and shallowly southeast-dipping faults, have been identified at Mina Justa (Fig.

3-2), 100 to 200 m apart, and both are 10 to 200 m in vertical extent (Baxter et al., 2005). The

Mina Justa orebodies comprise massive magnetite-pyrite cores surrounded by hydrothermal breccias, with strongly-altered host rock clasts in a magnetite-sulphide matrix, in turn surrounded by extensive magnetite-sulphide veining. Hypogene sulphides in individual orebodies are zoned upwards, and locally laterally, from pyrite-chalcopyrite to bornite-chalcocite±digenite, with a concomitant increase in Cu grade.

The mineralized area at Marcona is intruded by a swarm of hypabyssal bodies. These range from apparently syn- to clearly post-mineralization and, in composition, from silicic to, rarely, ultramafic (Atchley, 1956; Injoque, 1985). The most widespread intrusive units are dacite or rhyodacite porphyries, which display intimate interdigitation with the magnetite bodies, suggestive of melt commingling. Late-stage, 20-50m wide andesitic dykes, assigned to the Tunga

Andesite Suite, cut the orebodies at Mina Justa.

124

Figure 3-2. Geological map of the Mina Justa Cu deposit, hosted by the Middle Jurassic upper Río Grande Formation (modified from Rio Tinto 1: 10,000 mapping of Mina Justa Prospect, February 2003, unpublished report). Ab-albite; Kfs-K-feldspar; Act-actinolite

Four principal fault systems are identified in the Marcona-Mina Justa area. The oldest, the Pista normal faults, strike 295o and dip 60o northeast. The coeval or younger Repetición fault system, which controls the distribution of magnetite orebodies, comprises a series of reverse-slip faults which strike 045o and dip 30-60o northwesterly (Fig. 3-1). The main orebodies at Mina

Justa are controlled by the Mina Justa faults, with strike directions similar to those of the

125 Repetición Fault system but with shallow southeast dips and normal displacements. The youngest,

Huaca, normal faults strike 335o and dip 60o to the east and are commonly filled by andesite dykes.

3.4 Alteration and Mineralization

3.4.1 Marcona Magnetite Deposit

Seven stages of alteration and mineralization are identified at Marcona (Fig. 3-3). The earliest hydrothermal alteration (Stage M-I) comprises characterized by aggregates of 175-177 Ma old cummingtonite and 171 Ma phlogopite (-magnetite), which locally occur in the upper section of the Marcona Formation. Stage M-II Na metasomatism occurred throughout the mine area and is represented by albite and minor Na-rich scapolite. Stage M-III magnetite and Stage M-IV magnetite-sulphide mineralization, which together form the massive iron orebodies accounting for over 95% of the magnetite mineralization at Marcona, comprise a variety of opaque and gangue minerals including magnetite, calcic amphiboles, biotite, phlogopite, K-feldspar, calcite, apatite, diopside and sulphides. Major mineral assemblages assigned herein to Stage M-III are magnetite-actinolite (or tremolite±phlogopite±apatite) and magnetite-biotite (±actinolite). The main magnetite orebodies exhibit smoothly curving, abrupt contacts, dyke-like to tubular apophyses and intricate, amoeboid interfingering with dacite porphyry intrusions, and local vesicular texture, interpreted as evidence for the commingling of hydrous Fe oxidic and silicic melts. The main magnetite orebodies are haloed by biotite and K-feldspar alteration in metasedimentary rocks and dacites. Sulphide deposition was initiated at Stage M-IV in association with magnetite-actinolite (or tremolite±calcite, Fig. 3-4A) and magnetite-phlogopite

(±actinolite or tremolite±calcite) assemblages. The major sulphides of Stage M-IV are pyrrhotite and pyrite, occurring commonly as subhedral to anhedral crystals intergrown with magnetite, silicates and calcite.

126

Figure 3-3. Alteration and mineral paragenesis of the Marcona deposit. Supergene minerals are omitted

In Stage M-V, sulphide veins, commonly with calcic amphibole, formed in the upper levels of the orebodies and cut massive Stage M-III magnetite-calcic-amphibole assemblages. More commonly, Stage M-V sulphides and coexisting minerals occur as aggregates in cavities in Stage

M-III and Stage M-IV magnetite-amphibole assemblages (Fig. 3-4B). Common Stage M-V assemblages include chalcopyrite-pyrite± pyrrhotite-amphibole and chalcopyrite-pyrite-calcite.

Late-stage, M-VII, hydrothermal veins developed in multiple episodes with ambiguous age relationships, and include tourmaline-quartz-sulphide, magnetite and sulphide, quartz (±calcite), calcite and hematite facies (Fig. 3-4C). Both Stage M-V and M-VII represent post-magnetite mineralization hydrothermal activites and only contain minor non-economic Cu mineralization.

127

Figure 3-4. Mineralogical and textural relationships of fluid inclusion-hosting assemblages from the Marcona deposits. A - Characteristically fine-grained amphibole (actinolite+tremolite), sulphide and calcite occur interstitially with magnetite in Stage M-IV mineralization (# DDM4-7-4, drill core DDM4-7 at 415.1m, Mina 4 orebody). B - Stage M-V pyrite, chalcopyrite and calcite occur as aggregates superimposed on Stage M-III magnetite (# MA5-3, Mina 5 open pit, 670 m, orebody). C - Late quartz vein, with erratic calcite and Mn-oxides, cuts massive magnetite-amphibole-sulphide assemblages, and is in turn cut by a hematite vein (# MA3-35, Mina 3 open pit, 580 m, south wall). Amph-amphibole, Act-actinolite, Cal-calcite, Cp-chalcopyrite, Hem-hematite, Kf-K-feldspar, Mt-magnetite, Py-pyrite, Qtz-quartz

128 3.4.2 Mina Justa Cu Deposit

Seven stages of hydrothermal alteration and mineralization are also recognized at Mina Justa (Fig.

3-5). Stage J-I albite-actinolite alteration is represented by commonly light-pink albite and green fine-grained actinolite replacing the plagioclase phenocrysts in andesites. Rocks affected by Stage

J-II K-feldspar-magnetite are megascopically massive, the K-feldspar commonly occurring as extremely fine-grained crystals replacing plagioclase. Stage J-II magnetite is usually fine- to medium-grained and locally forms aggregates interstitial to K-feldspar. Green to dark-green Stage

J-III actinolite generally forms coarse, elongated crystals which commonly occur as veins cutting

Stage J-II K-feldspar-magnetite alteration zones or, locally, as the matrix of breccias with clasts of K-feldspar-magnetite-altered host rocks. The deposition of Stage J-IV specular hematite is revealed entirely by pseudomorphs of Stage J-V magnetite, i.e. “mushketovite”. Stage J-V magnetite-pyrite alteration generated the massive, lensoid, magnetite bodies which nucleated the

Figure 3-5. Alteration and mineral paragenesis of the Mina Justa deposit. Supergene minerals are omitted

129 subsequent Cu mineralization, and are haloed by magnetite-pyrite veins and hydrothermal breccias. Quartz occurs interstitially to magnetite and pyrite in the main magnetite bodies (Fig.

3-6A). Other constituents include K-feldspar, chlorite, titanite and fluorine-rich apatite. Stage

J-VI Cu sulphide-bearing veins locally cut altered host rocks (Fig. 3-6B) and Stage J-V magnetite-pyrite-quartz but more commonly, Cu sulphides and associated assemblages occur as replacive bodies in the magnetite-pyrite lenses. The main hypogene Cu sulphides are, in order of decreasing abundance, chalcopyrite, bornite, chalcocite and digenite. Calcite and hematite, locally intergrown with epidote and albite, are directly associated with Cu mineralization (Fig. 3-5).

Figure 3-6. Mineralogical and textural relationships of fluid inclusion-hosting assemblages from the Mina Justa deposits. A - Stage J-V Magnetite-pyrite (partly replaced by chalcopyrite)-quartz veins cut actinolite - altered andesite at Mina Justa. Actinolite is chloritized. (# MA45-3, drill core MA45, 267.5 m). B - Stage J-VI Chalcopyrite-calcite veins cut actinolite (-K-feldspar-magnetite) - altered andesite. Microcline occurs as haloes around calcite veins and locally cuts calcite (# MA45-6, drill core MA45, 404.2 m).

3.5 Sampling and Analytical Techniques

Samples for fluid inclusion microthermometry and chemical analysis were chosen based on paragenetic relationships (Figs. 3-3 and 3-5). At Marcona, emphasis was placed on the post-melt evolution of the hydrothermal system from Stages M-IV through M-VII. Although several minerals were examined, only calcite from Stages M-IV, M-V and M-VII, and quartz from Stage

130 M-VII were found to host fluid inclusions amenable to study (Fig. 3-4). At Mina Justa, data were obtained from quartz of Stage J-V and calcite of the Cu mineralization stage (J-VI) (Fig. 3-6).

Microthermometry was conducted using a Linkam THMS-G-600 heating-freezing stage

(-180oC – 600oC), calibrated using Synflinc synthetic fluid inclusion standards. The errors associated with temperature measurement below 30ºC and above 100ºC were ±0.2 and ±2ºC, respectively.

The LA-TOF-ICP-MS technique was chosen for inclusion chemical analysis because it has the ability to analyze numerous elements in the transient signal that results from the opening of a single fluid inclusion by laser ablation. Data were collected using a Renaissance axial geometry

LA - TOF - ICP-MS (LECO Corporation) with 266-nm (UV) laser system, under standard operating conditions (0.9-1.8mJ pulse energy, 20 Hz repetition rate, 10–150 μm spot size). NIST

612 SRM glass was used as an external standard, and concentrations were calculated from the element ratios via an internal standard element, Na, the concentration of which was estimated prior to ablation by microthermometric measurement. The typical detection limit and error associated with this technique are on the order of 100 ppb and less than 30 percent, respectively

(Mahoney et al., 1996; Leach and Hieftje, 2001; Balcerzak, 2003). The high error is generated by calculation of Na concentration from the measured ice melting temperatures. Spatial resolution of up to 10 μm is attained using the laser beam (Günther et al., 2000, Olivo et al., 2006). Samples were selected for analysis based on the size of their fluid inclusions and the potential to isolate each inclusion type. At Mina Justa, Stage J-V magnetite-pyrite-quartz veins contain the most suitable fluid inclusions for LA-TOF-ICP-MS analysis because of their large size, typically 8-25

μm. Fluid inclusions in calcite could not be analyzed because of the possibility of contamination from host ions such as Ca and Mg during the opening of the inclusions, the tendency of calcite to cleave during analysis, and, most importantly, the lack of suitable standards for calcite hosts.

131 The selection of samples for stable isotope analysis (Tables 3-4 and 3-5) was similarly based on detailed petrographic relationships. Pyrite, chalcopyrite, pyrrhotite and molybdenite were chosen for sulphur isotope analysis from the Marcona magnetite-sulphide stage (Stage M-IV), the main polymetallic sulphide stage (Stage M-V) and the late vein stage (Stage M-VII). In addition, sulphides and gypsum (replacing anhydrite) samples were analyzed from Mina 11, which differs paragenetically from the other orebodies. At Mina Justa, pyrite samples from the magnetite-pyrite stage (J-V), and bornite and chalcopyrite from the Cu mineralization stage (J-VI) were selected.

Magnetite, amphibole, biotite, apatite and phlogopite, from the Marcona magnetite-rich stages (M-III and M-IV) were selected for oxygen isotope analysis, as well as amphiboles from the major sulphide stage (M-V). Magnetites and amphiboles from Mina 11 were studied for comparison. At Mina Justa, magnetite, quartz and apatite from the magnetite-pyrite stage (J-V) were selected for oxygen isotope analysis.

Hydrogen isotope analyses were conducted on Stage M-III biotite, amphibole and phlogopite,

Stage M-IV phlogopite and amphibole, Stage M-V amphibole and Stage M-11-I amphibole from

Mina 11. Fluids trapped in inclusions in three quartz samples and released by thermal decrepitation were analyzed for their hydrogen isotopic compositions. Calcites from the Marcona magnetite-sulphide and major sulphide Stages and the Mina Justa Cu mineralization Stage IV were analyzed for carbon and oxygen isotopes.

Stable isotope analyses were performed at the Queen’s Facility for Isotope Research (QFIR).

Minerals were extracted from a crushed and washed fraction of the sample or by drilling. Sulphur was extracted online with continuous-flow technology, wherein 0.2-0.3 mg of sulphide samples

34 was converted to SO2 in a Carlo Erba Element Analyzer NCS 2500, with CuO as an oxidant. δ S values are reported relative to Canyon Diablo Troilite (CDT) standard. Oxygen isotope analysis of silicates and iron oxides were carried out using the BrF5 method of Clayton and Mayeda

(1963). Hydrogen isotope compositions of quartz-hosted fluid inclusions and silicates were

132 determined using the methods of Kyser and O’Neil (1984). δ18O and δD values are reported relative to Vienna Standard Mean Ocean Water (V-SMOW).

18 13 δ O and δ C values for carbonate were measured on CO2 released from 5 to 10 mg powdered carbonate samples reacted with 100% phosphoric acid, using the acid fractionation factor of Sharma and Clayton (1965). The δ13C values are reported relative to the V-PDB standard, and δ18O values to V-SMOW. All isotopes were measured using the QFIR Finnigan

MAT 252 isotope-ratio mass spectrometer and reported in per mil units (‰). δ34S, δ18O, δ13C and

δD analyses are reproducible to ± 0.3, ± 0.2, ± 0.1 and ± 3‰, respectively.

Isotope fractionation factors used in this study are as follows: for sulphur: pyrite-chalcopyrite: Kajiwara and Krouse (1971), pyrrhotite-chalcopyrite: Kajiwara and Krouse

(1971); for oxygen: magnetite-quartz: Clayton and Keiffer (1991), magnetite-apatite: Valley

(2003), magnetite-H2O: Bottinga and Javoy (1973), quartz-H2O: Clayton et al. (1972), calcite-H2O: O’Neil et al. (1969), amphibole-H2O: Zheng (1993), biotite-H2O: Bottinga and

Javoy (1973); phlogopite-H2O: Zheng (1993); for hydrogen: amphibole-H2O: Suzuoki and

Epstein (1976), phlogopite-H2O: Suzuoki and Epstein (1976), biotite- H2O: Suzuoki and Epstein

(1976); and for carbon: calcite-CO2: Ohmoto and Rye (1979).

3.6 Results

3.6.1 Fluid inclusions

There is no clear evidence of inclusions being trapped along growth zones in either quartz or calcite crystals from Marcona or Mina Justa, and therefore primary fluid inclusions cannot be identified using the criteria of Roedder (1984) and Goldstein (2003). Many fluid inclusions are distributed in three dimensions in host crystals without clear crosscutting relationships (Fig. 3-7A) and Fluid Inclusion Assemblages (FIA: Goldstein and Reynolds, 1994; Goldstein, 2003) cannot be used to distinguish groups of inclusions trapped coevally. These fluid inclulsions are herein

133 interpreted as uncertain in origin (“U.O.”). It should be emphasized, however, that similarly randomly distributed inclusions are assumed to be primary elsewhere (e.g., Hedenquist et al.,

1998; Wanhainen et al., 2003; Seedorff and Einaudi, 2004; Bouzari and Clark, 2006), even though their true origin is ambiguous (Bodnar, 2003; Goldstein, 2003). Locally, unambiguously secondary fluid inclusions are trapped along microfractures cutting crystals (Fig. 3-7B).

Five fluid inclusion types have been defined on the basis of phases observable at room temperature (Figs. 3-7C-E). Type A liquid-vapour two-phase inclusions, which account for over

80% of the total, contain an aqueous liquid with a vapour bubble. The volumetric liquid:vapour ratio is greater than unity in most such inclusions (Fig. 3-7C). Type A inclusions are predominantly spherical or irregularly shaped and less than 10 μm in size. All such inclusions homogenize into the liquid phase. Type B-1 liquid - vapour - halite three-phase inclusions contain liquid, a vapour bubble and a halite crystal identified on the basis of its cubic shape, isotropism and refractive index (Fig. 3-7D). Such inclusions commonly have irregular shapes and are typically less than 15 μm in diameter. They generally homogenize into the liquid. Type B-2 three-phase inclusions host a daughter mineral other than halite, i.e., neither cubic nor transparent.

Some solid phases have medium- to high- relief and high birefringence and may be an iron chloride or carbonate (Fig. 3-7E: Smith and Henderson, 2000; Bouzari and Clark, 2006). Other

Type B-2 inclusions have irregular solid phases and inconsistent solid/fluid ratios, and probably record accidental trapping (Goldstein, 2003). Rare multiple-phase (n>3) Type C inclusions consist of liquid, a vapour bubble and more than one daughter or accidental mineral, including halite, other salts and solids. Some contain three or four solid phases (Fig. 3-7F). Dark or transparent single-phase inclusions (Type D) have rounded and elliptical shapes (Fig. 3-7G), and show no response to heating or freezing. They are thought to be vapour-dominant (dark) or liquid - dominant (transparent). These, especially the vapour-rich, are generally larger than other types.

134

135

Figure 3-7. Fluid inclusion types in the Marcona and Mina Justa deposits. A - Fluid inclusions with uncertain origin are randomly distributed in three dimensions (quartz from #MA89-1, drill core MA89, 322.1 m, Mina Justa). B - Secondary fluid inclusions occur along fractures, some crossing crystal boundaries (calcite from MA7-23, Mina 7 open pit, southwest corner, 680 m, Marcona). C - Type A: Liquid-vapour two-phase inclusion (quartz from #MA89-1, drill core MA89, 322.1 m, Mina Justa). D - Type B-1: Liquid-vapour-halite three-phase inclusions (quartz from #MA89-1, drill core MA89, 322.1 m, Mina Justa). E - Type B-2: Liquid - vapor - non-halite solid three-phase inclusions. The rounded-rhombic solid with high relief and birefringence may be either accidental or a Fe chloride or carbonate daughter mineral (quartz from #MA89-1, drill core MA89, 322.1 m, Mina Justa). F - Type C: Multiple-phase (n>3) inclusions. At least four solids occur in this inclusion. The cubic, low-relief and low-birefringence daughter mineral is halite. The rhombic, high-relief and high-birefringence daughter may be a Fe chloride or carbonate. Halite melted at 375oC, but the other solids had not homogenized at the 600oC thermal limit of the heating stage (quartz from #MA89-1, drill core MA89, 322.1 m, Mina Justa). G - Type D: vapour (dark) fluid inclusion (calcite from MA2-12, Mina 2 open pit, south side, 650 m, Marcona).

136 There is no evidence for liquid CO2 or clathrates in any inclusions. Inclusions suite showing evidence of boiling were not identified.

In Marcona Stage M-IV, fine-grained calcite interstitial to magnetite and sulphides (Fig.

3-4A) hosts many inclusions along fractures and crossing grain boundaries, evidence for a secondary origin, although others are randomly distributed and have an uncertain origin. The secondary fluid inclusions are dominantly Type A, but the randomly distributed inclusions are mainly Type A and Type B (B-1 dominant), with a few Types C and D examples (Table 3-1).

Stage M-V coarse-grained calcite (Fig. 3-4B), however, has no extensive fractures and hosts fluid inclusions which are mainly of Type A and are randomly distributed in three dimensions. Stage

M-VII quartz and calcite crystals have no extensive fractures and the Type A fluid inclusions are commonly randomly distributed or isolated, with an uncertain origin.

At Mina Justa, Stage J-V quartz crystals, which are pervasively fractured and occur interstitially to magnetite and pyrite (Fig. 3-6A), host mainly Type A inclusions, with some Type

B and a few Type C. Fluid inclusions in calcite-sulphide veins from the J-VI Cu mineralization

(Fig. 3-6B) are randomly distributed in three dimensions, with an uncertain origin. Most of these fluid inclusions are Type A, while the remainder are Type B-2 (Table 3-1).

Microthermometric data representative of most alteration-mineralization stages at Marcona and Mina Justa are summarized in Table 3-2. At Marcona, the majority of Type A fluid inclusions

o in Stage M-IV calcite exhibit homogenization temperatures (Th) of between 73 and 193 C (Table

3-2), most falling between 100 and 140oC (Fig. 3-8). Secondary Type A fluid inclusions from

Stage M-IV have a similar Th range and distribution, as do inclusions in calcite from Stage M-V and in quartz and calcite from Stage M-VII veins (Fig. 3-8). The similarity in the Th of inclusions hosted by mineral assemblages of different mineralization stages implies that all or most of the

Type A fluid inclusions of uncertain origin in Stage M-IV calcite represent fluids trapped at later

137

Table 3-1. Locations and Types of Fluid Inclusions, Marcona and Mina Justa (Paragenetic stages shown in Figures 3-3 and 3-5). Sample no. Location Paragenetic stage Fluid inclusion Inclusion types hosted minerals MA7-23 Marcona, Mina 7 open pit; SW corner; 680 m Stage M-IV Calcite U.O. and S*; type A (90%) + B (10%) + C DDM4-7-4 Marcona, drill core DDM4-7; 300 m Stage M-IV Calcite U.O and S; type A MA2-12 Marcona, Mina 2 open pit; South side; 650 m Stage M-V Calcite U.O.; type A + D MA2-14 Marcona; Mina 2 open pit; South side; 650 m Stage M-VII Calcite U.O.; type A MA3-35 Marcona; Mina 3 open pit; South side; 580 m Stage M-VII Quartz U.O.; type A MA45-3 Mina Justa; drill core MA45; 267.5 m Stage J-V Quartz U.O.; type A MA89-1 Mina Justa; drill core MA89; 322.1 m Stage J-V Quartz U.O.; type A (90%) + B (10%) + C MA89-5 Mina Justa; drill core MA89; 366.5 m Stage J-VI Calcite U.O.; type A (95%) + B-2 (5%) MA27-5 Mina Justa; drill core MA27; 521.8 m Stage J-VI Calcite U.O.; type A MA17-2 Mina Justa; drill core MA17; 329.1 m Stage J-VI Calcite U.O.; type A (95%) + B-2 (5%) * U.O. – uncertain origin; S - secondary

138

Table 3-2. Summary of Fluid Inclusion Petrography and Microthermometric Data, Marcona and Mina Justa o o o o Stage Mineral Types Size Filling Thomogenization ( C) Tfinal ice melting ( C) Teutectic ( C) Tfinal hydrohalite melting ( C) Marcona Min Max n Min Max n Min Max n Min Max n Stage M-IV Cal Type A (U.O.) 3-20 μm 55-75% 73 193 123 -52.2 -3.1 101 -77.3 -28.6 83 -49.1 -45.0 3 Type A (S.) 3-9 μm 55-75% 117 185 29 -41.1 -6.0 28 -64.8 -37.9 26 Type B-1 (U.O) 7-20 μm 60-75% 117 220 17 -37.6 -24.3 9 -65.3 -43.1 9 Type B-2 7-11 μm 20-75% 102 143 6 -41.2 -41.2 1 (U.O.) Stage M-V Cal Type A (U.O.) 4-25 μm 50-80% 75 321 126 -31.0 -0.8 88 -68.1 -15.0 80 -48.0 -35.0 15 Stage M-VII Qtz-Cal Type A (U.O.) 4-12 μm 60-75% 70 170 41 -32.9 -3.9 40 -57.5 -8.6 38 Mina Justa Stage J-V Qtz Type A (U.O.) 3-28 μm 50-85% 57 374 176 -41.8 -0.1 163 -87.4 -16.7 128 -49.7 -23.7 17 TypeB-1 (U.O.) 7-25 μm 50-75% 123 400 12 -38.1 -9.0 10 -72.1 -53.9 9 TypeB-2 (U.O.) 7-25 μm 50-75% 96 146 6 -30.0 -21.6 5 -66.6 -50.0 4 Stage J-VI Cal Type A (U.O.) 4-30 μm 60-80% 87 198 107 -50.4 -0.4 78 -75.1 -11.3 69 -45.6 -37.6 8 Type B-2 4-25 μm 60-70% 88 219 6 -49.7 -22.1 3 -75.6 -52.1 3 (U.O.) Cal-calcite, Qtz-quartz; U.O. uncertain origin, S. secondary

139

Figure 3-8. Histograms of homogenization temperatures (Th), eutectic temperatures (Te) and ice-melting temperatures (Tm(ice)) for fluid inclusions of various paragenetic stages from Marcona and Mina Justa

stages. Trapping of late fluids may also be recorded by inclusions in Stage M-V minerals, although no unambiguously secondary inclusions were identified in calcite from this stage.

140 o However, the rare higher-temperature (Th = 220-320 C) Type A inclusions occur only in Stage

M-V calcite (Fig. 3-8).

Type A U.O. fluid inclusions in Stage M-IV samples have eutectic temperatures (Te) of between -77.3 and -28.6oC (Table 3-2), with modes at -65oC, -50oC and -30oC (Fig. 3-8). Their

o final ice melting (Tm(ice)) temperatures range from -52.2 to -3.1 C, with at least three modes, i.e.,

o o o -10 C, -25 C and -40 C (Fig. 3-8). Secondary Type A inclusions in Stage M-IV have similar Te and Tm(ice) (Fig. 3-8), further supporting the secondary origin of the U.O. inclusions. The major Te and Tm(ice) modes of inclusions in Stage M-V all appear in Stage M-IV inclusions (Fig. 3-8), but

o o extremely low Te (-65 C) and Tm(ice) (-40 C) values occur only in Stage M-IV hosts. Fluid inclusions in the late veins of Stage M-VII, dominantly Type A with an uncertain origin, have eutectic and ice melting temperatures of between -57.5 and -8.6oC and -32.9 and -3.9oC, respectively (Table 3-2). Both Te and Tm(ice) are bimodal and differ from those for Stage M-V fluid

o o inclusions which show a Te peak of -35 C and a Tm(ice) peak at -20 C (Fig. 3-8). However, the possibility that Stage M-V calcite trapped low-temperature fluids from Stage M-VII cannot be excluded.

Type B fluid inclusions occur only in Stage M-IV samples. Of these, Type B-1 inclusions have homogenization temperatures of between 117 and 220oC (Table 3-2), with apparent peaks at

120oC and 200oC. In contrast, Type B-2 fluid inclusions, as with both U.O. and secondary Type A inclusions in Stage M-IV (Table 3-2), have a single peak at 120oC. Consequently, Type B-2 and low-temperature Type B-1 inclusions are interpreted as secondary, but they are not identified in the subsequent Stages M-V and M-VII. The high-temperature (200oC) peak in Type B-1 populations may be evidence for another distinct fluid, either a primary Stage M-IV fluid or high-temperature fluids trapped during Stage M-V (Fig. 3-8). Ice first melted at -43.1 to -65.3oC in Type B-1 fluid inclusions (Table 3-2), indicating the presence of Ca or Mg+Fe.

141 On the basis of lithological relationships described by Hawkes et al. (2002), a lithostatic pressure of 1.5-2 kbar, estimated from the overlying 4-5 km-thick Río Grande Formation and part of the underlying Marcona Formation, can be applied to Marcona magnetite-rich stages (M-III and M-IV). Such pressures entail a correction of approximately 100oC in the estimation of entrapment temperature for the fluid inclusions (Bodnar, 2003). Therefore, for Marcona Stage

M-IV, a maximum entrapment temperature of ~ 300oC is estimated from the plausibly primary

o fluid inclusions (Type B-1, Th ≤ 220 C) hosted by calcite. Such temperatures are, however, much lower than those of 430-600oC indicated by sulphur and oxygen isotope geothermometers for this stage (see below), which suggests that Type B-1 inclusions may not represent the Stage M-IV primary fluid. However, the temperatures indicated by sulphur isotope geothermometry for Stage

M-V fall in a range from 160oC to 360oC, broadly in agreement with the homogenization temperatures of fluid inclusions, which cluster at 120oC with a maximum of 320oC (Table 3-2 and

Fig. 3-8), evidence that the pressure correction for entrapment temperature does not exceed 50oC for this stage.

Th, Te and Tm(ice) data for Stage M-V inclusions are plotted in Figure 3-9. About 90% of the

o inclusions have low homogenization temperatures (< 200 C) but a wide Te range, evidence for complex fluid compositions (Goldstein and Reynolds, 1994; Fig. 3-9). On the basis of known stable and metastable eutectic temperatures and observed Te ranges for various aqueous brine systems (Crawford, 1981; Davis et al., 1990; Goldstein and Reynolds, 1994; Baldassaro, 2000), the low-temperature fluids may be assigned to Na-rich, Mg-Fe - rich and Ca-rich populations, with Mg-Fe - rich compositions predominating (Fig. 3-9). The major cations identified from the

Te data are in permissive agreement with the main Stage M-V mineral assemblage, i.e., Ca-Mg - rich amphibole + calcite + Fe-Cu sulphides (Fig. 3-3). This complex fluid system also exhibits a wide Tm(ice) range (Fig. 3-9), which implies mixing of fluids with varied salinities. Na-dominant fluids, poor in Mg, Fe and Ca, have generally low salinities, whereas the Ca-rich fluids are

142

Figure 3-9. Relationships among Te, Tm(ice) and Th for fluid inclusions of the Marcona major sulphide stage (M-V). The eutectic temperatures of systems with different components are indicated. The observed Te ranges of the systems NaCl-CaCl2-H2O and NaCl-MgCl2-H2O, and the observed metastable Te ranges of the systems NaCl-H2O and NaCl-MgCl2-H2O are shown in Th vs. o Te space. Lower- and higher-temperature populations are divided at 200 C on the basis of homogenization temperatures. Low-, medium- and high-salinity inclusions are differentiated on o o o the basis of final ice melting temperatures (Tm(ice)) (0- -10 C, -10 - -20 C and < -20 C, respectively). The histograms of the Te populations, with inferred dominant cations (see text), are shown on the right.

143 commonly highly saline, and Mg-Fe - rich fluids exhibit both low- and high-salinity populations

(Fig. 3-9). Fluids with moderate salinities are inferred to be products of the mixing of such end-member fluids. Consequently, at least four groups of low-temperature fluids can be recognized, viz., Na-dominant with low salinity, Mg-Fe - rich with low salinity, Mg-Fe - rich with high salinity, and Ca-rich with high salinity. Fluids with relatively high homogenization temperatures, i.e. > 200oC, observed only in Stage M-V calcite, also exhibit wide ranges of composition and salinity (Fig. 3-9). The limited data, however, preclude the identification of distinct fluid populations.

At Mina Justa, type A fluid inclusions of uncertain origin in quartz from the magnetite-pyrite stage (J-V) at Mina Justa have homogenization temperatures ranging widely from 57 to 374oC

(Table 3-2), with a mode at 140oC (Fig. 3-8). Fluid inclusions in calcite of the subsequent Stage

o o J-VI have Th values of between 87 and 198 C (Table 3-2), but again with a 140 C mode (Fig. 3-8).

Eutectic and final ice melting temperatures for both Stages J-V and J-VI are characterized by a comparably wide range (Fig. 3-8). These similarities, as well as the extensively-developed fractures in Stage J-V quartz, strongly suggest that most Type A fluid inclusions in Stage J-V quartz are secondary, trapped during the subsequent Cu mineralization. Consequently, the low-temperature inclusions hosted by both Stage J-V quartz and Stage J-VI calcite may represent the fluids responsible for J-VI Cu mineralization. In contrast, the sparse Type A inclusions in

o Stage J-V quartz with high Th (200–380 C; Fig. 3-8) are unlikely to represent Stage J-VI fluids and may have been trapped at the Stage J-V magnetite-pyrite alteration stage. Type B-1 and Type

C fluid inclusions, which occur only in Stage J-V quartz, have homogenization temperatures of between 123 and 400oC (Table 3-2), the majority exceeding 200oC, indicating that most of the

Type B-1 (and C) inclusions could be related to the high-Th Type A population, representing primary fluids trapped during Stage J-V. Type B-2 inclusions, occurring in both Stage J-V quartz

144 o and Stage J-VI calcite, have low Th values, with a mode at 140 C, similar to those for Type A inclusions.

The entrapment pressure for the ore-forming fluids at Mina Justa is uncertain because no boiling fluid inclusions are identified and detailed paleostratigraphic information is unavailable.

However, on the basis of the pervasive hydrothermal brecciation in the orebodies, the coarse-grained quartz and calcite crystals and the large fluid inclusions (up to 30 μm), a near-surface and hence low-pressure environment is favoured, and the pressure correction for Th is probably minor. The maximum Th determined for the magnetite-pyrite stage (Stage J-V) is ~

400oC, evidence for an entrapment temperature of at least that value. The temperatures inferred from oxygen isotope geothermometry for this stage range from 540oC to 600oC (see below), requiring a pressure correction equivalent to ≤ 1 kbar (Bodnar, 2003), consistent with the inferred shallow ore-forming environment. The maximum Th for the later Cu-mineralization stage (Stage

o o J-VI) is less than 200 C and most Th measurements cluster around 140 C, accepted as approximating the entrapment temperature.

The Cu mineralization stage (J-VI) fluid inclusions also exhibit eutectic and ice melting temperatures clustering around -45 to -50oC and -20 to -25oC, respectively (Fig. 3-8), evidence for a significant Ca content and a high salinity. “Browning” or “orange peel” development

(Goldstein, 1994) were commonly observed at the eutectic temperature. Such features were identified in the Olympic Dam by Oreskes and Einaudi (1992) and at Eloise by Baker (1998), where they were also accepted as evidence for Ca-rich ore-forming fluids. However, the

o extremely low Te values (< -50 C) recorded herein for many fluid inclusions may represent metastable phenomena, which are prevalent in the system H2O-NaCl-CaCl2 (Davis et al., 1990).

Apparently, Ca-rich fluids dominated the Cu mineralization stage at Mina Justa (Fig. 3-10). At least two groups of Stage J-VI fluids can be identified, viz., Ca-rich with a low temperature and high salinity (Stage J-VI-1 in Table 3-3) and Na (-K) - dominant with a low temperature and low

145

Figure 3-10. Relationships among Te, Tm(ice) and Th for fluid inclusions of the Mina Justa Cu mineralization stage (J-VI). The eutectic temperatures of systems with different components are indicated. The observed Te ranges of the systems NaCl-CaCl2-H2O and NaCl-MgCl2-H2O, and the observed metastable Te ranges of the systems NaCl-H2O and NaCl-MgCl2-H2O are shown in Th vs. o Te space. Lower- and higher-temperature populations are divided at 200 C on the basis of homogenization temperatures. Low-, medium- and high-salinity inclusions are differentiated on o o o the basis of final ice melting temperatures (Tm(ice)) (0- -10 C, -10 - -20 C and < -20 C, respectively). The histograms of the Te populations, with inferred dominant cations (see text), are shown on the right.

146 salinity (Stage J-VI-2 in Table 3-3). Other inclusions may record the mixing of these two groups

(Fig. 3-7).

LA-TOF-ICP-MS Results

LA-TOF-ICP-MS data were obtained for 31 fluid inclusions hosted by Mina Justa Stage J-V quartz (Table 3-3). Four Type B-1 fluid inclusions with relatively high homogenization temperatures (313-400oC), such as occur only in Stage J-V quartz crystals and may represent the primary fluids of the Mina Justa magnetite stage, have high average concentrations of 11.3% for

Na, 1.2% for Ca, 0.4% for Fe and 0.1% for K, in addition to significant B (0.1%), Ni (169 ppm),

Zn (126 ppm) and Ba (107 ppm). In contrast, the ore metals, Cu and Ag, are below detection and

7 ppm, respectively (Table 3-3). In comparison, Stage J-VI fluid inclusions, interpreted as the low-temperature, high-salinity and Ca-rich fluids associated with Cu mineralization (Stage J-VI-1 in Table 3-3 and Figure 3-11A), have lower Na (7.8%) and higher Ca (2.5%) and Fe (0.6%) concentrations. They are also strongly enriched in Cu (average of 100 ppm), Ag (55 ppm), Zn

(483 ppm) and Au (69 ppm) (Table 3-3 and Figure 3-11A). One low-temperature, low-salinity, and Na-K - dominant fluid inclusion, interpreted as meteoric water from Stage J-VI (Table 3-3 and Figure 3-11B), has much lower contents of Cu, Sb, Cr, Al, K and Fe. These data indicate that whereas the Stage J-V fluids and the meteoric fluid involved in Stage J-VI were Cu-barren, Cu as well as the precious metals are markedly concentrated in the Stage J-VI low-temperature, high-salinity and Ca-rich fluid.

The LA-TOF-ICP-MS data (Table 3-3) suggest that, although the Na/Ca ratios in Stage J-VI

Cu mineralization fluid inclusions vary markedly from 1.2 to 24.1 (n=26), the majority are below

5.0. In contrast, the Na/Ca ratios estimated from the melting temperature of hydrohalite (Thh,

Table 3-2) and ice-melting temperature in the system H2O-NaCl-CaCl2 (Oakes et al., 1990) are less than unity, with most below 0.5. Such low estimates may result from the imprecision of Thh

147 Table 3-3. Average Laser Ablation Time of Flight ICP-MS Analyses (ppm) of Single Quartz-hosted Fluid Inclusions, Mina Justa Stage Types No.1) Fluid features Li B Na Mg Al P K Ca Sc Ti V Cr Mn Fe Co Ni J-V B-1 4 High T, high S, Na-rich 95 1040 112700 97 426 852 1253 12461 116 349 46 55 234 3532 26 168 J-VI A 26 Low T, high S, Ca-rich 359 811 77799 106 2971 2793 1480 25330 115 3453 88 506 307 5814 57 337 J-VI A 1 Low T, low S, Na-rich 397 n.d. 38479 14 n.d. 384 52 9525 85 6373 3 n.d. 158 485 39 233 Stage J-VI-12) / Stage J-V 3.8 0.8 0.7 1.1 7.0 3.3 1.2 2.0 1.0 9.9 1.9 9.3 1.3 1.7 2.2 2.0 Stage J-VI-23) / Stage J-VI-1 1.1 0.0 0.5 0.1 0.0 0.1 0.0 0.4 0.7 1.9 0.0 0.0 0.5 0.1 0.7 0.7 Stages Types No.1) Fluid features Cu Zn As Rb Sr Y Zr Nb Mo Ag Cd Sn Sb Ba Ce Nd J-V B-1 4 High T, high S, Na-rich n.d. 126 70 30 73 20 44 12 64 7 26 22 27 107 19 5 J-VI A 26 Low T, high S, Ca-rich 100 483 205 51 77 23 57 29 129 55 235 94 73 181 17 41 J-VI A 1 Low T, low S, Na-rich n.d. 336 n.d. n.d. 52 2 50 30 26 20 244 22 n.d. 37 7 7 Stage J-VI-12) / Stage J-V >10.0 3.8 2.9 1.7 1.1 1.1 1.3 2.4 2.0 7.4 9.0 4.3 2.7 1.7 0.9 8.3 Stage J-VI-23) / Stage J-VI-1 0.0 0.7 0.0 0.0 0.7 0.1 0.9 1.1 0.2 0.4 1.0 0.2 0.0 0.2 0.4 0.2 Stages Types No.1) Fluid features Sm Eu Gd Tb Dy Ho Er Tm Yb Lu W Au Pb Th U Br 4) J-V B-1 4 High T, high S, Na-rich 20 2 14 9 9 8 30 2 47 11 97 34 39 14 21 55 J-VI A 26 Low T, high S, Ca-rich 37 19 76 12 50 15 54 11 48 16 97 69 105 20 17 307 J-VI A 1 Low T, low S, Na-rich 11 20 72 3 32 11 11 19 35 10 15 35 5 6 14 29 Stage J-VI-12) / Stage J-V 1.8 8.1 5.6 1.4 5.4 1.8 1.8 4.8 1.0 1.5 1.0 2.0 2.7 1.4 0.8 6 Stage J-VI-23) / Stage J-VI-1 0.3 1.1 0.9 0.3 0.7 0.8 0.2 1.7 0.7 0.6 0.2 0.5 0.1 0.3 0.8 n.d. 1) Number of analyses; 2) Stage J-VI-1 – low-T, high-Salinity and Ca-rich fluid in stage J-VI; 3) Stage J-VI-2 – Low-T and low-Salinity fluid, without Ca; 4) using average counts/second as units because no standard is available. n.d. – undetected.

148

Figure 3-11. A - LA-TOF-ICP-MS data for cation concentrations in fluid inclusions of the Mina Justa Cu ore-forming fluid (Stage J-VI-1: low-temperature, high-salinity and Ca-rich) and the Stage J-V primary fluid. R value indicates Stage J-VI-1/stage J-V ratio. B - Comparison of elemental concentrations in the Mina Justa Cu ore-forming fluid (Stage J-VI-1) and in inferred meteoric water (Stage J-VI-2: low-temperature, low-salinity and Na-rich). R value indicates Stage J-VI-1/stage J-VI-2 ratio.

149 measurement, but other cations (e.g., Fe, Zn, Mn, etc.) in the fluids may invalidate the estimation of Na/Ca ratios based on the system H2O-NaCl-CaCl2.

Figure 3-12 illustrates the correlations of Na, Ca and Cu with other cations revealed by the

TOF data for the Cu mineralization Stage J-VI inclusions. Na has a negative or no correlation with most other cations, especially the ore metals, Cu, Ag and Au (Fig. 3-12A). Ca is negatively correlated with Na (Fig. 3-12A), but has a positive correlation with Cu, Ag, Au and Zn (Fig.

3-12B). Cu also has a positive correlation with Zn, Co, Ag and Au (Fig. 3-12C), in agreement with the occurrence of sphalerite and carrollite in the Cu-rich assemblages, as well as the high Ag contents in the Cu sulphides. However, the concentration of each element in contiguous Stage

J-VI fluid inclusions is variable.

3.6.2 Stable Isotope Geochemistry

Sulphur, oxygen, hydrogen and carbon isotopic compositions are shown in Tables 3-4 and 3-5, incorporating temperatures estimated based on S and O isotope geothermometry (Table. 2-5).

Sulphur isotopic compositions of minerals and hydrothermal fluids: The δ34S values of

Marcona Stage M-IV sulphides, intergrown with magnetite, amphibole and trioctahedral micas, range from +0.8 to +5.9‰ (Table 3-6), with the mode at ~ +2‰ (Fig. 3-13A). The δ34S values of

Marcona Stage M-V sulphides, commonly coarse-grained and associated with calcite and amphibole, are more restricted, varying from +1.8 to +5.0‰ and clustering near +5‰ (Table 3-6 and Fig. 3-13A). One late vein stage (M-VII) pyrite has a δ34S value of +7.4‰ (Table 3-6 and Fig.

3-13A). δ34S values of sulphides from Mina 11 fall into the ranges shown by the other Marcona orebodies (Table 3-4). Two gypsum samples recording supergene alteration of the Mina 11 anhydrite-sulphide stage (M11-V) have δ34S values of +14.5 and +15.8‰.

At Mina Justa, pyrite from the magnetite stage (J-V) has δ34S values ranging from +0.8 to

+3.9‰ (Table 3-6). The δ34S values of chalcopyrite and bornite from the ensuing Cu

150

Figure 3-12. Inter-element correlations based on TOF data (A) Na versus other cations (B) Ca versus other cations (C) Cu versus other cations. r represents correlation coefficient.

151

Table 3-4. Stable Isotopic Compositions of Minerals from Marcona Sample Stage T (oC) δ 34S δ 18O δ D δ 13C MA5-9 M-IA 8.7 (Cum) -62 (Cum) MA5-9 M-IB 9.4 (Phl) -59 (Phl) DDM 3-3-1 M-III 800 4.4 (Mt); 8.6 (Amph) -67 (Amph) DDM 3-3-1 M-III 8.9 (Amph) -63 (Amph) MA3-7 M-III 770 4.2 (Mt); 8.0 (Phl) -67 (Phl) MA3-22 M-III 700-760 3.7 (Mt); 7.2 (Apt); -60 (Amph) 8.9 (Amph) DDM4-7-1 M-III 3.9 (Mt) DDM5-4-7 M-III 4.6 (Mt) MA5-9 M-III 800 5.2 (Mt); 8.7 (Bt) -71 (Bt) MA10-4 M-III 8.3 (Amph) -64 (Amph) MA3-19 M-IV 600 2.9 (Mt); 9.3 (Phl) -62 (Phl) DDM 3-3-1 M-IV 3.8 (Py) DDM 3-3-1 M-IV 3.7 (Py) MA3-8 M-IV 5.9 (Py) MA3-18 M-IV 3.9 (Py) DDM 4-6-5 M-IV 430 2.4 (Po); 2.1 (Cp) DDM 4-7-4 M-IV 570 1.8 (Py) 13.2 (Cal); 4.8 (Mt) -6.0 (Cal) DDM 4-7-10 M-IV 5.7 (Mt) DDM 4-7-11 M-IV 2.4 (Po) DDM 5-4-3 M-IV 1.2 (Py) DDM 5-4-4 M-IV 3.5 (Mt) DDM 5-4-5 M-IV 2.6 (Py) DDM 5-4-9 M-IV 2.7 (Py) MA5-2 M-IV 2.3 (Cp) 4.8 (Mt) MA5-9 M-IV 14.5 (cal) -7.2 (Cal) MA5-10 M-IV 1.9 (Py) 12.9 (Cal) -8.2 (Cal) MA1-9 M-IV 590 4.0 (Py); 3.4 (Cp) MA7-13 M-IV 9.0 (Amph) -64 (Amph) MA7-23 M-IV 2.8 (Py) 11.9 (Cal) -5.9 (Cal) DDM 3-3-1 M-V 10.5 (Amph) -60 (Amph) DDM 3-3-2 M-V 4.8 (Cp) DDM 3-3-5 M-V 9.0 (Amph) -74 (Amph) DDM 3-3-7 M-V 3.6 (Py) 9.6 (Amph) -63 (Amph) DDM 3-3-8 M-V 0.8 (Py) 9.5 (Amph) -52 (Amph) MA3-8 M-V 5.0 (Py) 14.2 (Cal) -6.0 (Cal) MA3-35 M-V 4.6 (Py) 13.7 (Cal) -5.0 (Cal) MA2-7 M-V 13.8 (Cal) -5.0 (Cal)

152 Sample Stage T (oC) δ 34S δ 18O δ D δ 13C MA2-12 M-V 360 4.7 (Py); 3.6 (Cp) 16.5 (Cal) -5.9 (Cal) DDM 4-6-3 M-V 4.6 (Py) DDM 4-7-1 M-V 4.0 (Py) DDM 4-7-9 M-V 160 3.5 (Py); 2.7 (Cp) DDM 5-4-1 M-V 8.7 (Amph) -62 (Amph) DDM 5-4-2 M-V 2.1 (Py) 8.9 (Amph) -65 (Amph) DDM 5-4-7 M-V 9.0 (Amph) -60 (Amph) DDM 5-4-8 M-V 3.0 (Po) MA5-3 M-V 4.0 (Py) 18.7 (Cal) -6.8 (Cal) MA5-7-2 M-V 1.8 (Py) MA7-15 M-V 4.9 (Py) DDM 3-3-3 M-VII 8.6 (Mt); 14.5 (Cal) 1) -5.7 (Cal) MA3-25 M-VII 7.4 (Py ± Cp) MA3-35 M-VII 14.9 (Qtz) -42 (Qtz) 2) MA2-14 M-VII 12.3 (Cal) -5.1 (Cal) DDM 4-7-8 M-VII 13.1 (Cal) -7.7 (Cal) MA7-2 M-VII 13.7 (Cal) -6.8 (Cal) MA91-3 M11-I 9.4 (Amph) -61 (Amph) MA91-4 M11-II 430 2.5 (Py); 1.6 (Cp) 4.7 (Mt) MA91-3 Early M11-V 2.4 (Py) MA91-5 Early M11-V 2.6 (Mo) MA91-6 Early M11-V 3.9 (Py) MA91-3 Late M11-V 15.8 (Gy) MA91-5 Late M11-V 14.5 (Gy)

Ab-albite, Amph-amphibole, Apt-apatite, Bn-bornite, Bt-biotite, Cal-calcite, Cc-chalcocite, Cp-chalcopyrite, Cum-cummintonite, Gy-gypsum, Moly-molybdenite, Mt-magnetite, Phl-phlogopite, Po-pyrrhotite, Py-pyrite, Qtz-quartz; 1) magnetite and calcite in different veins; 2) from fluids hosted in fluid inclusions; underlined indicates mineral pair for geothermometer

153 Table 3-5. Stable Isotopic Compositions of Minerals from Mina Justa Sample Stage T (oC) δ 34S δ 18O δ D δ 13C MA17-5 J-V 2.6 (Py ±Cp) MA17-6 J-V 3.4 (Mt) MA35-1 J-V 2.7 (Py ±Cp) MA27-2 J-V 2.0 (Py ±Cp) MA27-5 J-V 3.0 (Mt) MA27-6 J-V 1.1 (Py) MA89-1 J-V 2.7 (Mt) 1); 13.5 (Qtz) MA89-2 J-V 580 1.1 (Py ± Cp) 2.9 (Mt); 8.0 (Apt) MA89-3 J-V 1.7 (Py ± Cp) MA89-3 J-V 3.9 (Py ± Cp) MA36-1 J-V 1.6 (Py ± Cp) MA45-2 J-V 540 1.7 (Py ± Cp) 3.3 (Mt); 9.0 (Apt) MA45-3 J-V 600 1.3 (Py ± Cp) 4.7 (Mt); 12.6 (Qtz) MA45-4 J-V 1.6 (Py) 4.4 (Mt) MA75-2 J-V 0.8 (Py ± Cp) MA64-4 J-VI 2.2 (Bn ± Cc) 12.3 (Cal) -4.4 (Cal) MA64-5 J-VI 3.2 (Bn) MA17-1 J-VI 1.6 (Bn) MA75-1 J-VI 2.0 (Bn) MA17-2 J-VI 13.7 (Cal) -6.0 (Cal) MA17-4 J-VI 2.6 (Cp) MA17-8 J-VI 3.7 (Cp ± Py) MA17-9 J-VI 3.3 (Cp) 12.2 (Cal) -7.0 (Cal) MA35-1 J-VI 14.7 (Cal) -6.9 (Cal) MA27-3 J-VI 2.5 (Cp ± Py) 14.3 (Cal) -6.3 (Cal) MA27-5 J-VI 2.7 (Cp ± Py) 14.6 (Cal) -7.2 (Cal) MA89-1 J-VI 1.5 (Cp) -40 (Qtz) 3) MA89-5 J-VI 2.1 (Cp) 13.3 (Cal); 12.6 (Ab) 2) -7.5 (Cal) MA54-2 J-VI 3.7 (Cp) MA36-1 J-VI 1.3 (Cp ± Py) MA45-3 J-VI 1.7 (Cp ± Py) -33 (Qtz) 3) MA45-6 J-VI 1.7 (Cp) 13.1 (Cal); 13.2 (Ab) 2) -6.8 (Cal)

Abbreviations of minerals as in Table 3-4. 1) minor hematization in magnetite; 2) albite is relatively late (see Fig. 3-3) and not in equilibrium with calcite; 3) from fluids hosted in fluid inclusions. Underlined indicates mineral pair for geothermometer

154 Table 3-6. Summary of Sulphur Isotopic Composition of Minerals and Ore-forming Fluids Marcona * Magnetite-sulphide stage Polymetallic sulphide stage Late vein stage (M-VII) (M-IV) (M-V) 34 δ Ssulphides +0.8 to +5.9‰ (n= 16) +1.8 to +5.0‰ (n=17) +7.4‰ (n=1) Estimated 430-600oC 160-360oC 70-170oC 34 34 Temperature Py-Cp and Po-Cp δ Sequil; Py-Cp and Po-Cp δ Sequil Fluid inclusions 18 Mt-Cal δ Oequil

fo2 indicator pyrrhotite pyrrhotite hematite or hematite/magnetie 34 34 Δ=δ SH2S-δ Sfluid 0‰ 0‰ ≥-30‰ 34 δ Sfluid +0.8 to +5.9‰ +1.8 to +5.0‰ ≥ +37.4‰ Mina Justa Magnetite-pyrite stage (J-V) Cu mineralization stage (J-VI) 34 δ Ssulphides +0.8 to +3.9‰ (n= 13) +1.3 to +3.7‰ (n= 16) Estimated 540-600oC 88-220oC 18 Temperature Mt-qtz and Mt-Apt δ Oequil Fluid inclusions

fo2 indicator magnetite-pyrite hematite 34 34 Δ=δ SH2S-δ Sfluid 0‰ ≥-28‰ 34 δ Sfluid +0.8 to +3.9‰ ≥29.3‰

* δ34S values of Mina 11 sulphides and sulphates not included (see text). Abbreviations as in Table 3-4.

mineralization stage (J-VI) have a range of +1.3 to +3.7‰. Both Stage J-V and Stage J-VI sulphides have δ34S values with modes at ~ +2.0‰ (Fig. 3-13B).

34 Given the fO2 and temperature during mineral precipitation, total δ S values of hydrothermal

34 fluids can be estimated from the δ S values of sulphides and sulfates (Ohmoto and Rye, 1979;

Ohmoto and Goldhaber, 1997). At Marcona, the temperatures determined from sulphur and oxygen isotope geothermometers range from 430 to 600oC for Stage M-IV and from 160 to 360oC for Stage M-V (Table 3-4). Pyrrhotite occurs in both stages (Fig. 3-3), evidence for a low fO2

(pyrite-pyrrhotite oxygen buffer), and the difference between the δ34S values of sulphide and

34 34 fluids (Δ=δ SH2S-δ Sfluid) would therefore be negligble (Ohmoto and Rye, 1979).

The local occurrence of hematite in late veins unaffected by supergene alteration is

o interpreted as evidence that high fO2 (log fO2 ≥ -34 at 250 C and 40 bar; Barton and Skinner, 155

Figure 3-13. δ34S values of sulphides and calculated ore-forming fluids from (A) Marcona and (B) Mina Justa mineralization and alteration stages.

1979) prevailed during Stage M-VII. If Stage M-VII pyrite was deposited under such conditions, the δ34S values of the fluids could have been as high as +37‰, but were probably lower (Table

3-6). The δ34S values of the gypsum from Mina 11 (+14.5 and +15.8‰) may represent the total

δ34S values of fluids precipitating anhydrite in this oxidized environment. However, sulphides from the Cu-rich Mina 11 mineralization stage were probably deposited from fluids having δ34S values similar to those in the other Marcona orebodies, on the basis of the temperature of 430oC

156 determined from sulphur isotope geothermometry (Table 3-4) and the coexistence of magnetite and pyrite (Ohmoto and Rye, 1979).

At Mina Justa, the difference in the δ34S values of sulphides and fluids for the magnetite -

o pyrite stage (J-V) would be negligible at 540 to 600 C (Table 3-5) and the moderate fO2 environment defined by the magnetite-pyrite oxygen buffer (Table 3-6, Ohmoto and Rye, 1979).

However, fine-grained hypogene hematite is abundant in the Cu mineralization stage (J-VI), evidence for a high oxygen fugacity (≥ hematite/magnetite buffer), and the δ34S conversion factor for fluids at the low inferred temperatures of 88-220oC is at least -28‰ (Table 3-6, Ohmoto and

34 Goldhaber, 1997), generating δ Sfluid values of ≥ 29.3‰ (Table 3-6 and Fig. 3-13B).

Oxygen isotopic compositions of minerals, melt and hydrothermal fluids: Stage M-III amphibole, biotite and phlogopite have δ18O values ranging narrowly between +8.0 and +8.9‰, whereas those for Stage M-III magnetite vary between +3.7 and +5.2‰ (Table 3-4). Because the main Stage M-III magnetite orebodies are interpreted as having crystallized from a high-T (≥

800oC, Table 3-4), iron oxide-dominant melt which formed through immiscibility from a parental andesitic magma, isotope exchange factors between minerals and melt, rather than between minerals and water, must be considered. Experiments show that 18O preferentially partitions into

Si-rich minerals relative to Si-poor minerals with less polymerized structures (Taylor and Epstein,

1962), and Kyser et al. (1998) demonstrated that a similar fractionation also occurs between

Si-rich and Fe-rich immiscible melts. Thus, in the system Fe2SiO4-KAlSi2O6-SiO2 at 1 bar and

o 18 1180 C, the δ O value of a melt with 66.7% SiO2 and 21.3% FeO was shown experimentally to be 0.6‰ higher than that of a coexisting immiscible Fe-rich melt with 50.1% SiO2 and 43.5%

FeO. The Marcona Stage M-III iron oxide-dominant melt would have had an extremely Fe-rich,

SiO2-poor composition (SiO2 ≈ 25%, FeO ≈ 70%) from the estimated percentages of minerals, and would therefore be predicted to be strongly 18O-depleted relative to a Si-rich melt (i.e.,

18 18 18 δ Osi-rich melt - δ OFe-rich melt » 0.6‰). Given the δ O values of the Neogene volcanic rocks of the

157 Central Andes, e.g., 6.8-7.4‰ for andesite (Harmon et al., 1984), the δ18O values of the Marcona oxide melts are inferred to have had a maximum range from 6.2 to 6.8‰. Further, on the basis of the extreme FeO content (~70%) of the Marcona melt, the oxygen fractionation factor between magnetite and melt would have been very small. At 700-800oC (Table 3-4), a maximum fractionation of 2.5‰ is inferred (Zhao and Zheng, 2003), so that on the basis of the δ18O values of Stage III magnetite, the δ18O values of Marcona oxide melt would have had a range from 5.2 to

7.7‰, in permissive agreement with the estimate based on the data from the Neogene volcanic rocks. The δ18O values of Stage M-III silicates (8.0 and 8.9‰) are ~ 2.5‰ higher than those inferred for parental oxide melt (5.2 to 7.7‰), evidence for a Fe-rich melt.

The δ18O values of silicates and magnetite from Stage M-IV have ranges similar to those of Stage M-III, but the associated calcites have higher δ18O values, from +11.9 to +14.5‰ (Table

3-4). At temperatures of 570-600oC, as estimated from oxygen isotope geothermometry (Table

3-4), the δ18O values of the Stage M-IV fluids are inferred to have ranged between +9.6 and

+12.2‰ (Table 3-7). The δ18O values of coarse-grained Stage M-V amphiboles are +8.7 to

+10.5‰, slightly higher than those of Stages M-III and M-IV (Table 3-4), with those for coexisting calcites varying from +13.7 to +18.7‰. At 360oC (Table 3-4), the calculated δ18O values of Stage M-V fluids range from +10.1 to +12.5‰ (Table 3-7). Four Stage M-VII (late vein) calcites have δ18O values of +12.3 to +14.5‰, and one quartz vein has a value of +14.9‰ (Table

3-4). At 120oC (Fig. 3-8), the δ18O values of fluids in equilibrium with the veins would have ranged from -3.6 to -1.4‰ (Table 3-7).

At Mina Justa, δ18O values of Stage J-V magnetite lie between +2.7 and +4.7‰, and two quartzes have values of +12.6 and +13.5‰ (Table 3-5). The calculated δ18O values of 540-600oC

(Table 3-5) fluids in Stage J-V range between +9.5 and +11.5‰ (Table 3-7). These δ18O values are similar to those of the fluids of the Marcona magnetite-sulphide stage (Stage M-IV). Calcites

158 from the Cu mineralization stage (J-VI) have δ18O values between +12.2 and +14.7‰. At 140oC

(Fig. 3-8), an average δ18O value of 0.1‰ is calculated for the ore-forming fluids (Table 3-7).

Table 3-7. δ18O, δD and δ13C Values of Minerals and Fluids From Various Stages at Marcona and Mina Justa 18 13 o 18 13 stages minerals δ O δ D δ C T ( C) δ Ofluid δ Dfluid δ Cfluid Marcona M-III Amph 8.6 (1) * -67 (1) 800 1) -54 Bt 8.7 (1) -71 (1) 800 1) -50 Phl 8.0 (1) -67 (1) 770 1) 5.2-7.7 5) -73 Amph 8.9 (1) -60 (1) 700 1) -43 Mt 4.2 (5) 770 2) M-IV Phl 9.3 (1) -62 (1) 600 1) 11.7 -60 Cal 13.2 (1) -6.0 (1) 570 1) 12.2 -3.3 Mt 3.9 (2) 590 2) 9.6 M-V Cal 16.5 (1) -5.9 (1) 360 1) 12.5 -3.4 Amph 9.3 (7) -60 (7) 360 3) 10.1 -8 M-VII Qtz 14.9 (1) 120 4) -3.6 -42 6) Cal 13.7 (4) -6.3 (4) 120 4) -1.4 -9.2 Mina Justa J-V Qtz 12.6 (1) 600 1) 11.5 Mt 3.7 (3) 570 2) 9.5 J-VI Cal 13.5 (8) -6.4 (8) 140 4) 0.1 -8.3 Qtz -37 (2) 6) Abbreviations of minerals as in Table 3-4. * value (number of analyses); 1) temperature estimated by isotope geothermometry from the sample; 2) average temperature estimated by geothermometry; 3) temperature estimated by geothermometry from other samples assigned to this stage; 4) temperature estimated by microthermometry; 5) δ 18O values are estimated for Stage M-III oxide melt (see text for discussion); 6) from fluid inclusions hosted by quartz.

Hydrogen isotopic compositions of silicates, melt and fluids: The δD values of Marcona

Stage M-III amphiboles range from -67 to -62‰, whereas M-III phlogopite gives a δD value of

-67‰, and a biotite a value of -71‰ (Table 3-4). These data are inferred to record the hydrogen

isotope chemistry of the oxide melt. At 700-800oC, the calculated δD values of the melt would

vary between -73 and -43‰ (Table 3-7). A phlogopite from later Stage M-IV has δD value of

-62‰ (Table 3-4), which at 600oC indicates that Stage M-IV fluids had a δD value of -60‰ 159 (Table 3-7). Stage M-V coarse-grained amphibole samples give a range of δD values between -74

o and -52‰ (Table 3-4), corresponding at 360 C to a δDfluid value of -8‰ (Table 3-7). The fluid trapped in inclusions hosted by a late Stage M-VII quartz vein has a δD value of -42‰ (Table

3-7).

At Mina Justa, δD values of -32 and -40‰ (average -37‰) were determined for fluid inclusions hosted by Stage J-V quartz (Table 3-7). However, paragenetic relationships and fluid inclusion petrography imply that these values record the hydrogen isotopic composition of the subsequent Cu mineralization fluids (Stage J-VI).

Carbon isotopic compositions of calcites and hydrothermal fluids: The δ13C values of

Stage M-IV calcites from Marcona range between -8.2 and -5.9‰ and those of Stage M-V calcites from -6.8 to -5.0‰ (Table 3-4). The fractionation factor for calcite-CO2 determined by

Ohmoto and Rye (1979) indicates that, at 570oC, the δ13C value of Marcona Stage M-IV fluids was -3.3‰. A similar value (-3.4‰) is estimated for Stage M-V 360oC fluids (Table 3-7). Calcite samples from late barren veins (Stage M-VII) have δ13C between -7.7 and -5.1‰, so that the average δ13C value of Stage M-VII ~120oC fluids was –9.2‰, generally lower than those of

Stages M-IV and M-V fluids (Table 3-7).

At Mina Justa, the δ13C values of eight calcites from the Cu mineralization stage fall between -7.5 and -4.4‰, similar to those of Marcona calcite (Table 3-5). However, at 140oC, the average δ13C value of the ore-forming fluid would have been -8.3‰, differing significantly from those of Marcona Stage M-IV and M-V fluids, but similar to those of Stage M-VII late vein fluids

(Table 3-7).

3.7 Discussion

3.7.1 Fluid evolution in the Marcona deposit

160 Oxygen isotope geothermometers give temperatures of 700oC-800oC for the sulphide-barren magnetite Stage III which dominates the Marcona deposit. The δ18O and δ D values estimated for the inferred Stage M-III melt plot within the magmatic water range, and close to those characteristic of Central Andean Neogene andesite (Fig. 3-14). A magmatic source for Stage

M-III is also indicated by the relationships between δ18O and temperature (Fig. 3-15), which are distinct from those of other documented, unambiguously hydrothermal, IOCG deposits (e.g.,

Olympic Dam, Ernest Henry and La Candelaria), but similar to those of Central Andean Neogene dacitic and andesitic magmas (Fig. 3-15). The δ18O values of Stage M-III magnetite (3.7-5.2‰,

Table 3-4) are comparable to those of magnetite from El Laco (3.6-4.6‰, Rhodes and Oreskes,

1999), an iron oxide deposit inferred to be of melt origin (Naslund et al., 2002).

Figure 3-14. Calculated δD and δ18O values of fluids at Marcona and Mina Justa. Fields for magmatic waters are from Taylor (1997) and for Neogene andesites of the Central Andes from Harmon et al. (1984). The trajectories for seawater (SMOW) exchange with host rocks are from Ripley and Ohmoto (1977). The deviation of the δD and δ18O values (dashed line) from the meteoric water line represents formation water of the Kettleman North Dome basin from California (Longstaffe, 1987).

161

18 Figure 3-15. δ Ofluid/melt-temperature relationships for the main mineralization and alteration stages at Marcona and Mina Justa and other IOCG deposits (data for Mantoverde from Benavides et al., 2007; other data from Williams et al., 2005 and references therein). Magmatic water and Neogene Andean dacite and andesite fields are from Taylor (1997) and Harmon et al. (1984).

162 The inferred primary fluid inclusions hosted by Marcona Stage M-IV calcite have moderate homogenization temperatures (up to 220oC), and are commonly highly saline, with varied cationic compositions. Sulphur and oxygen isotope geothermometers indicate a temperature range of from 430 to 600oC (Table 3-4). Stage M-IV fluids have δ18O and δD values that plot close to the magmatic water field (Fig. 3-14). Fluids at this stage also exhibit similarities in temperature and oxygen isotopic composition to those identified at Ernest Henry (Mark et al., 2000) and for the early, Cu (Au)-barren, magnetite stages at Olympic Dam (Oreskes and Einaudi, 1992), which are inferred to be of magmatic-hydrothermal origin (Fig. 3-15). The δ34S values of Stage M-IV fluids and sulphides fall in a narrow range from +0.8 to +5.9‰ (mode at ~2.5‰), indicating a magmatic source (Nielsen, 1979). These data are comparable to those documented from some

34 major IOCG deposits, which have δ Ssulphide values of 0 ± 5‰ (Fig. 3-16; Williams et al., 2005).

Further, the calculated δ13C and δ18O values of the fluids in Stage M-IV are similar to those of magmatic CO2 (Fig. 3-17) (Taylor, 1986; Javoy et al., 1988). Apparently, between Stage M-III and M-IV, an interval which saw the intital precipitation of sulphides and the stabilization of phlogopite (Fig. 3-3), physicochemical conditions at Marcona changed significantly (Fig. 3-18).

The fluids responsible for Stage M-V were apparently diverse (Fig. 3-9). Higher - temperature (Th

= 200 - ≥320oC), Na-rich/low-salinity, and Ca-rich/high-salinity groups were probably magmatic in origin, whereas the low-temperature (~120oC) populations record mixing of meteoric water

(low-salinity and Na-rich) with modified seawater (Mg-Fe- rich or Ca-rich, medium-to- high salinities), the latter being the product of reactions between seawater and Ca- rich andesite and metasediments (Fig. 3-18C). Incursion of external fluids during Stage M-V is also supported by oxygen and hydrogen isotopic compositions, which are distinct from those of magmatic Stage

M-III and M-IV (Fig. 3-14), and apparently record a shift away from magmatic compositions with decreasing temperature (Fig. 3-15). In the Callovian-Oxfordian, the andesites of the Upper

Río Grande Formation erupted in a shallow-marine environment (Caldas, 1978; Aguirre, 1998).

163

Figure 3-16. δ34S values of ore fluids at Marcona and Mina Justa and other major IOCG deposits (data from Ullrich and Clark, 1999; de Haller et al., 2002, Benavides et al.,. 2007, Williams et al., 2005, and references therein).

Metamorphic fluids, with temperatures of ca. 250oC, probably derived from the reaction of seawater with hot andesite (Aguirre, 1988), or intense heat flow during the initial development of a basin (i.e., diastathermal processes: Alt, 1999), generated extensive alteration. Modified seawater was also identified in the mineralization stage at Raúl-Condestable, 300 km NNW of

Marcona (Ripley and Ohmoto, 1977) and may have contributed to Stage M-V sulphide precipitation. However, the δ34S values of the fluid may have been strongly modified through water/rock reaction (Bowers, 1989). Jurassic seawater, which generally has δ34S values of 16‰

(Hoefs, 1997), is inferred to have interacted with andesite and basaltic andesite (δ34S values ~

0‰), generating a fluid with low δ34S values (1.8-5.0‰, Table 3-6). Calculated δ13C and δ18O values of fluids for Marcona Stage M-V at 360oC fall in the magmatic range (Fig. 3-17), but may record the involvement of seawater (δ13C=0‰, δ18O=0‰, Rotherhamet et al., 1998), with high

δ18O and low δ13C values as the result of reaction with intermediate to mafic volcanic rocks

(average δ13C values -10 - -5‰; δ18O values +8 - +10‰: Hoefs, 1997). 164

Figure 3-17. Calculated δ13C and δ18O values of hydrothermal fluids, Marcona and Mina Justa mineralization and alteration stages.

3.7.2 Fluid evolution in the Mina Justa deposit

The inferred primary fluid inclusions hosted by Mina Justa Stage J-V quartz have moderate -to-

o high homogenization temperatures (Th ≤ 400 C), high salinities and variable cationic compositions. Oxygen isotope geothermometers indicate temperatures of 540 to 600oC, similar to that of the Marcona magnetite-sulphide stage (M-IV). The δ18O and δ D values estimated for

Stage J-V plot close to the magmatic water field (Fig. 3-14), again comparable to Marcona Stage

M-IV fluids. A magmatic source for Stage J-IV is also indicated by the relationship between δ18O values and temperature (Fig. 3-15), wherein the Stage J-IV values are similar to those of the

Marcona Stage M-IV and other IOCG deposits inferred to be of magmatic-hydrothermal origin.

165

Figure 3-18. Cartoon showing the evolution of the main magnetite mineralization and major sulphide stage at Marcona (not to scale).

166 The δ34S values of Stage J-V sulphides and fluids fall in a narrow range of between +0.8 and

+3.9‰ (Fig. 3-16), similar to those of Marcona Stage M-IV and evidence for a dominant magmatic source (Nielsen, 1979).

The 40Ar/39Ar ages for Stage J-V hydrothermal microclines (101-104 Ma) coincide with the initial emplacement of the Coastal Batholith in the Acarí area to the east, i.e., ≤ 109±4 Ma (Vidal et al., 1990). Small granodiorite-diorite stocks crop out ~ 3km to the east of Mina Justa (Fig.

3-19A), supporting a genetic relationship between Stage J-V magnetite-pyrite alteration and the

Coastal Batholith. Magmatic-hydrothermal fluids may have risen along the ENE-striking Mina

Justa normal faults which controlled the emplacement of the Fe and Cu orebodies (Fig. 3-19A).

Fluids documented for the Cu mineralization stage are mainly low-temperature (~ 140oC), high-salinity and Ca-rich, but with a minor low-temperature, low-salinity and Na-dominant component. The former is inferred to be a Ca-rich basinal brine and the latter is probably meteoric water. TOF-ICP-MS analyses indicate that Cu and other ore metals are restricted to the

Ca-rich fluids, i.e., the exotic brines, both the high-temperature primary fluids at Stage J-V

(magnetite stage) and the meteoric water being Cu-barren. The mineralizing Stage J-VI fluids differ also markedly from those of magnetite stage in their lower temperatures and δ18O values

(Fig. 3-15), and are unlikely to have been of magmatic derivation. Such non-magmatic fluids were also involved in the hematite-rich mineralization stages at Olympic Dam (Oreskes and

Einaudi, 1992; Haynes et al., 1995), La Candelaria (Ullrich et al., 2001) and Mantoverde

(Benavides et al., 2007). A non-magmatic model for Cu mineralization is supported by the relationship between δ18O and δ D values (Fig. 3-14), which shows that the dominant fluids have the characteristics of basinal brines, especially those with high salinities occurring at middle latitudes, e.g., basins from California (Longstaffe, 1987). The δ34S values of sulphides from Stage

34 J-VI are concentrated in a narrow range around 0‰, but the calculated δ Sfluid values exceed

+29‰ (Fig. 3-16), strong evidence for a non-magmatic source and consistent with data for the

167

Figure 3-19. Genetic model for the magnetite-pyrite alteration and superimposed Cu mineralization at Mina Justa.

main Cu (-Au) ore-stages at Raúl-Condestable (Ripley and Ohmoto, 1977), La Candelaria

(Ullrich and Clark, 1999) and Mantoverde (Benavides et al., 2007). A non-magmatic source for the mineralizing fluid is also supported by the δ13C values (Fig. 3-17), which are similar to those of Marcona late veins (M-VII) and show close analogies to deep formation waters or basinal brines at low latitudes (Fig. 3-17).

168 During the Albian, faulting of the Mesozoic volcanic-sedimentary strata along the southwest margin of the Cañete Basin, could have provided pathways for fluid circulation (Fig. 3-19B).

During tectonic inversion of the basin, Ca-rich basinal brines, possibly derived from connate water in equilibrium with the permeable Copara Formation, would have been expelled from the basin, flowing westward to the Mina Justa area, probably driven by heat from the underlying

Coastal Batholith. During their rise to the surface, the basinal fluids replaced early magnetite-pyrite bodies to generate the Cu orebodies, with minor mixing with meteoric waters

(Fig. 3-19B).

3.7.3 Implications for Cu-mineralizing fluids in IOCG deposits

Both the Marcona and Mina Justa deposits share numerous alteration and mineralization characteristics with other major IOCG deposits. Table 3-8 compares the mineralizing fluids recognized at Mina Justa and Marcona (the hydrothermal polymetallic sulphide stage), with those in major IOCG deposits and several other deposit types elsewhere which are either iron oxide-rich or potentially affiliated, such as porphyry Cu-Au, stratabound Cu-Ag (mantos) and sediment-hosted Cu deposits.

Recent studies (Pollard, 2001) provide strong evidence that high-temperature (> 350oC) fluids of at least partial magmatic derivation are responsible for Cu-Au mineralization in some

IOCG deposits, e.g., Ernest Henry (Mark et al., 2000) and La Candelaria (early polymetallic and main ore stages: Ullrich and Clark, 1999). However, fluid inclusion data from Mina Justa and

Marcona indicate that the main sulphide deposition occurred at either low temperatures (<200oC in the case of the Mina Justa Cu mineralization) or over a wide temperature range (160-360oC for the Marcona polymetallic sulphide stage), which may imply the involvement of low-temperature external fluids. Similar conditions are documented for the mineralizing stages at the Olympic

169 Table 3-8. Copper Ore-forming Fluids Involved in IOCG Mineralization1) and Other Deposit Types Deposits Mineral Assemblages 2) Temperature (oC) Salinity3) Dominant Cations 4) Other Features References 5) 5) 6) IOCG: Mina Justa, Perú Cp-Bn-Cc-Cal-Hem 85-200 (140 ) 6-32 (24 ) Ca, Na (-K) CO2-poor This study

Marcona, Perú Cp-Py-Po-Cal-Amph 70-300 (120) 5-33 (20) Na, Mg (-Fe), Ca CO2-poor This study Raúl, Perú Cp-Py-Po-Qtz±Mt 320-360 (350) >20 Na (Ca ?) ? Ripley and Ohmoto (1977)

Mantoverde, Chile Cp-Hem(-Cal-Qtz) 150-360 (240) 14-40 Na (Ca?) CO2-poor Vila 1996 6) La Candelaria, Chile Cp-Po-Qtz-Amph ±Mt 300 - 450 high Na (Ca) CO2-rich Ullrich and Clark (1999) 6) Olympic Dam, Australia Cp-Bn-Cc-Ser-Fl-Qtz-Bar-Hem 130-280 (180) 7.3-42 Ca, Na CO2-bearing Oreskes and Einaudi (1992)

Ernest Henry, Australia Cp-Py-Mt-Cal-Qtz-Bt-Kfs >350 - 400 > 26 Na (Ca?) CO2-rich Mark et al. 2000

Eloise, Australia Cp-Po-Qtz-Cal-Mt-Chl 100-500 (220) 30-47 Ca, Na CO2-bearing Baker 1998

Starra, Australia Cp-Py-Qtz-Cal-Hem-Chl 220-360 29-42 Ca, Na CO2-poor Williams et al. 2001

Osborne, Australia Cp-Qtz-Cal-Chl±Mt ~300 20-37 Na, K, Ca CO2-bearing Adshead et al. 1998

Lightning Creek, Australia Cp-Py-Cal-Chl 120-180 15-28 Ca, Na CO2-poor Perring et al. 2000

Tennant Creek, Australia Cp-Py-Qtz-Chl-Hem 300-340 3-10 Na, Ca CO2-poor Skirrow and Walsh 2002

Aitik, Sweden Bn-Cc-Mt-Qtz 130-220 18-27 Ca, Na CO2-poor Wanhainen et al. 2003

Pahtohavare, Sweden Cp-Py-Po-Qtz-Cal±Mt 80-350 (120) 0.5-30 Ca, Na CO2-bearing Lindblom et al. 1996

Bidjovagge, Norway Cp-Py-Po-Qtz-Cal±Mt 300-370 30-45 Na, Ca CO2-CH4-rich Ettner et al. 1994

Wernecke, Canada Cp-Py-Qtz-Cal±Mt±Hem 70-350 (200) 24-42 Ca, Na CO2-poor Gillen et al. 2004; Hunt et al. 2005, 2007

Salobo, Brazil Cp-Bn-Cc-Qtz-Chl 173-485 39-52 Na (Ca?) CO2-bearing Requia et al. 2003; Pollard 2001

Porphyry Cu-Au: Panguna,Papua New Guinea Cp-Py-Qtz 350 – 700 (?) 46-76 Na, K CO2-bearing Eastoe 1978

Bajo de la Alumbrera, Argentina Cp-Py-Qtz 300 - 400 35-65 Na, K CO2-bearing Ulrich et al. 2001

Stratabound Cu-Ag: El Soldado, Chile Cp-Bn-Cc-Cal-Hem 140-180 21-34 Ca, Na CO2-poor Boric et al. 2002

Coastal cordillera of northern Chile Cp-Bn-Cc-Cal-Qtz-Hem 200-380 (250) 7-34 Na, Ca CO2-poor Kojima et al. 2002

Sediment-hosted Cu: Nchanga, Zambia Cp-Bn-Cc-Dol-Phl 105-300 (170) 31-38 Ca, Na CO2-bearing McGowan et al. 2006 Cp-chalcopyrite, Py-pyrite, Po-pyrrhotite, Cal-calcite, Amph-amphibole, Bn-bornite, Cc-chalcocite, Hem-hematite, Mt-magnetite, Qtz-quartz, Ser-sericite, Fl-fluorite, Bar-barite, Kfs-K-feldspar, chl chlorite, Phl-phlogopite, Dol-dolomite. 1) exclusively for Cu (Au, Ag, REE, U) mineralization stages. 2) italicized mineral used for fluid inclusion study. 3) wt% NaCl equiv or wt% NaCl + CaCl2 equiv. 4) first-cited cation dominant. 5) the majority. 6) Defined in this table: CO2-rich - fluid inclusions containing CO2 ≥ 20%;

CO2-bearing: a few fluid inclusions (<20%) contain CO2; CO2-poor: no CO2 identified in fluid inclusions.

171 Dam deposit (~180oC: Oreskes and Einaudi, 1992), Lightning Creek (late chalcopyrite veins,

120-180oC: Perring et al., 2000) and Eloise (100-500oC: Baker, 1998); Mantoverde (150-360oC:

Vila, 1996); Wernecke (70-350oC: Hunt et al., 2005, 2007); Aitik (130-240oC, bornite stage:

Wanhainen et al., 2003) and Salobo (173-485oC: Requia et al., 2003). In contrast, porphyry

Cu-Au deposits, which are generally magnetite-rich, commonly formed at consistently high temperatures (300->400oC) for Cu-Au mineralization (Table 3-8), whereas low-temperature fluids are commonly associated with stratabound Cu-Ag (El Soldado: Boric et al., 2002) and sediment-hosted Cu deposits (Nchanga: McGowan et al., 2006), which may reflect conditions similar to those for IOCG Cu (-Au) mineralization.

High-salinity fluids responsible for Cu-Au mineralization have been documented in many major IOCG deposits (Oreskes and Einaudi, 1992; Ullrich and Clark, 1999; Pollard, 2001;

Wanhainen et al., 2003). Whereas most porphyry Cu-Au deposits are products of high-salinity fluids dominated by Na and K, a high content of Ca has been widely demonstrated for the fluids responsible for Cu mineralization in IOCG deposits (Table 3-8). These Ca-rich fluids also commonly occur in stratabound Cu-Ag and sediment-hosted Cu deposits in which evaporitic brines may have contributed to mineralization (Boric et al., 2002; McGowan et al., 2006).

Whereas CO2 has been inferred to play an important role in many IOCG deposits through the unmixing of magmatic H2O-CO2-NaCl ± CaCl2-KCl fluids (Xu and Pollard, 1999; Pollard, 2001;

Fu et al., 2003), the relationship between this process and Cu (Au-Ag-REE-U) mineralization is not clear. At La Candelaria and Ernest Henry, the high-temperature mineralizing fluids contain considerable CO2 (Table 3-8) and are inferred to be responsible for Cu-Au mineralization

(Ullirich and Clark, 1999; Mark et al., 2000). However, only minor or negligible CO2 contributed to the Cu mineralization stages at many IOCG deposits, such as Mina Justa, Mantoverde, and

Wernecke (Table 3-8), which indicates that CO2-rich fluids are not a prerequisite for Cu mineralization in IOCG deposits.

172

3.8 Conclusions

The contiguous Jurassic Marcona Fe (162-156 Ma) and Early Cretaceous Mina Justa Cu (104-95

Ma) deposits were generated by fundamentally different magmatic-hydrothermal systems. At

Marcona, the major magnetite orebodies are interpreted as recording an evolution from high-temperature (700-800°C) and hydrous Fe oxide-dominated melts to rapidly cooling

(430-600°C) endogenous hydrothermal fluids with dominantly magmatic isotopic compositions.

The subsequent lower-temperature (160-360oC) polymetallic sulphide stage may record the invasion of seawater modified through reaction with the andesitic host rocks. However, the sulphide vein systems contain only uneconomic Cu, probably recording the restricted mineralization potential of the vesiculating Fe oxide melts (cf. Broman et al., 1999).

Whereas the early magnetite-pyrite assemblage at Mina Justa was deposited from 540-600°C,

Na-rich hydrothermal fluids with a magmatic signature, the Cu (-Ag) orebodies were entirely the product of low-temperature ( ≤ 200°C), Ca-rich brines with sulfur and oxygen isotopic compositions predicating a dominant evaporite-sourced basinal brine reservoir. Ore formation is inferred to have taken place during inversion of the contiguous Aptian-Albian Cañete back-arc basin, fluid migration being focused by the intrusion of flanking granitoid plutons of the Coastal

Batholith.

For those magnetite-dominant, “Kiruna-type” Fe deposits, such as Kiruna, El Laco, and the

Cretaceous Chilean iron deposits, a high-temperature, Fe-oxide dominant melt was probably responsible for the massive magnetite mineralization, whereas in probably the majority of Cu-rich

IOCG deposits, including Mina Justa, Olympic Dam, La Candelaria-Punta del Cobre,

Mantoverde and Raúl-Condestable, precursor ironstone formation was the product of high-temperature metasomatism and nonmagmatic, “exotic” fluids were a prerequisite for subsequent economic Cu (-Au, Ag) mineralization.

173 Chapter 4

CONCLUSIONS

4.1 Marcona – a unique “Kiruna-type”, magnetite deposit in the Middle Jurassic

metallogenetic sub-province of the Central Andes

New 40Ar/39Ar geochronological data demonstrate that the earliest hydrothermal alteration at

Marcona occurred in the Aalenian at ca. 177 Ma, but that the massive magnetite-calcic amphibole-trioctahedral mica orebodies did not form until the late Bathonian, during a 156-162

Ma episode of andesitic eruption which terminated the growth of the Middle Jurassic arc. The forms, paragenetic relationships and stable isotope geochemistry of the orebodies are interpreted as recording an evolution from intrusive high-temperature (700->800°C), intrusive, hydrous Fe oxide-dominated melts to rapidly cooling (400-600°C) endogenous hydrothermal fluids. In particular, the intimately interdigitated bodies of massive magnetite and dacite porphyry, as exemplified by the large Minas 2-3-4 orebody, are most plausibly explained as evidence for the commingling, at 159 to 162 Ma, of immiscible melts generated through the silicification of parental andesitic magma ponded in Marcona Formation siliciclastics. Sulphides initially precipitated at ≥ 400°C in the ensuing magmatic - hydrothermal stage. Finally, at 156-159 Ma, magnetite-free sulphide - calcic amphibole and calcite assemblages, modestly enriched in Cu, Zn and Pb, formed at markedly lower temperatures (160-360oC) in a relatively reduced environment

(pyrite-pyrrhotite buffer). Fluid inclusion and stable isotope analyses suggest that sulphide precipitation was triggered by the incursion of seawater.

The main Marcona magnetite (-actinolite) orebodies share numerous mineralogical, textural and geochemical similarities with the Kiruna, El Laco and Cretaceous Chilean iron belt (CIB) deposits, which are considered by most workers to be products of oxide melt crystallization

174 (Table 4-1; Nyström and Henríquez, 1994; Naslund et al., 2002; Lledó, 2005). However, the

Marcona mineralization is haloed by zones of intense K-metasomatism (biotite and K-feldspar alteration), in place of the strong Na- or Ca metasomatism surrounding other iron deposits, evidence for a K-rich parental magma. The significant, if uneconomic, average Cu content

(0.12%) at Marcona is ascribed herein to the involvement of seawater.

The 159-162 Ma Marcona magnetite orebodies constitute the major mineralization of the

Middle Jurassic metallogenetic sub-province of the Central Andes, unparalleled in scale elsewhere in the mid-Jurassic Cordillera de la Costa (Fig. 4-1). Thus, several of the larger individual magnetite orebodies are comparable in size to the major ore deposits of the CIB, and the Marcona district, only 25 km2 in extent, incorporates more economic magnetite than the entire

10,000 km2 CIB. Whereas Marcona formed in a newly-developed extensional submarine basin, coeval, 160-165 Ma, porphyry Cu (-Mo-Au) mineralization in the Cocachacra district in southern

Perú (Clark et al., 1990) was associated with granitoid intrusions emplaced during orogenic contraction and uplift (Quang, 2003). In contrast to central-south and southern Perú, transtensional regimes dominated northern Chile in the mid-Jurassic, controlling the emplacement of numerous iron oxide-rich Cu veins, e.g., Tocopilla, Julia and Las Animas (Maksaev, 1990;

Sillitoe, 2003). Largely hosted by submarine to subaerial Lower Jurassic andesitic strata (Cornejo et al., 2006), these small deposits are considered to be genetically related to Bathonian to

Callovian granitoid intrusions (Maksaev, 1990). Thereafter, with the increasing coupling of the convergent plates in northern Chile in the Valanginian (Jaillard et al., 2000), shallow dioritic and granodioritic plutons (Maksaev and Zentilli, 2002) were probably responsible for the

“manto-type” Mantos Blancos stratabound hematite-rich Cu deposit (500 Mt at 1.0% Cu;

141-142 Ma: Ramírez et al., 2006; Oliveros et al., 2006). However, only weak magmatism and magmatic-hydrothermal alteration occurred in central-south and southern Perú in the Late

175 Table 4-1. Characteristics of Major Cu-poor “Kiruna-type” Iron Deposits Deposits Kiirunavaara, Sweden Marcona, Perú El Romeral (CIB*), El Laco, Chile Chile Tonnage/grade 2600 Mt @ 60% Fe 1940 Mt @ 55% Fe; ~400 Mt @ 60% Fe 500 Mt @ >60% Fe 0.12% Cu Age of mineralization ~1900 Ma ~160 Ma ~110 Ma ~2 Ma Host rocks trachytic and syenitic Silicic metasediments, Andesite Andesite, dacite rocks andesite-dacite Major mineral magnetite+apatite+ magnetite+actinolite magnetite+actinolite+ magnetite+diopside assemblages in orebody actinolite (tremolite)+ biotite apatite +apatite+quartz Dominant alteration albitization- K-silicate (biotite – actinolitization Ca-metasomatism associated with magnetite scapolitization K-feldspar) (diopside) mineralization alteration magmatic features columnar, platy or mesoscopic contact Columnar, platy Columnar, platy dendritic magnetite; relationships; local magnetite; dendritic magnetite; dendritic vesicular; octahedral vesicles; local columnar amphibole; vesicles; pyroxene; vesicles; crystals, magnetite octahedral magnetite octahedral crystals; octahedral crystals; flow; “ore breccias”; crystals; local “ore magnetite flows; “ore magnetite flows; “ore droplets breccias” breccias” breccias” Magnetite composition High V, low Cr and High V, Low Ti High V, low Cr and Ti High V, low Cr and Ti Ti Ore-forming environment subaerial/submarine submarine subaerial ?/submarine subaerial Tectonomagmatic extensional, basin formation in an inversion of basins eruption of a subaerial setting intracratonic basin ?, extensional contiguous to an andesitic-dacitic arc in anorogenic andesitic-dacitic arc extensional arc an extensional magmatism environment Relationship with Cu-rich Spatially and Spatially with large Temporally with large Spatially and IOCG deposits temporally with small IOCG: Mina Justa IOCG, but only temporally with small IOCG’s: Pahtohavare locally overlapping in IOCG’s: Rio Grande and Rakkurijärvi space and Arizaro References Frietsch, 1978; This study; Bookstrom, 1977; Nyström and Nyström and Injoque et al., 1988 Nyström and Henríquez, 1994; Henríquez, 1994 Henríquez, 1994; Naslund et al., 2002 Naslund et al., 2002 *CIB-Chilean Iron Belt, with a total tonnage of ~2000 Mt and average grade of 60% Fe (Oyarzun et al., 2003).

Jurassic, e.g., the K-Fe metasomatism in the Mina Justa area (ca. 142 Ma). Apparently, with

respect to mineralization type, the Marcona magnetite deposit is unique in the later-Jurassic

evolution of the Central Andes. A conjunction of favourable conditions, viz. strongly oxidized 176 magmas, oxide-silicate melt immiscibility prompted by quartz-rich host rocks, and an extensional environment, was probably responsible for this anomaly. In contrast, the slightly younger, rhyolite-hosted Mantos Blancos mineralization has been ascribed to the repeated boiling of magmatic-hydrothermal fluids due to episodic decompression (Ramírez et al., 2006).

Figure 4-1. Model for the Middle-Late Jurassic Central Andean mineralization (N-S section)

4.2 The Mina Justa Cu (-Ag) deposit – a major Cu-rich IOCG deposit

in the Cretaceous Central Andes

40Ar/39Ar age data for actinolite and K-feldspar indicate that hydrothermal activity in the Mina

Justa area generated albite-actinolite alteration at 156-157 Ma, to be followed at ca. 142 Ma by

K-Fe metasomatism. However, the intense high-temperature (540-600oC) magnetite-pyrite alteration (Fe metasomatism) was not initiated until ca. 104 Ma, succeeding precursor actinolite alteration (Ca metasomatism) at ca. 109 Ma. The 95-99 Ma Mina Justa Cu (-Ag) mineralization, nucleated on bodies of the ca. 2.5 m.y. - older magnetite-pyrite alteration, was entirely the product of low-temperature (≤ 200°C) Ca-rich brines with isotopic compositions predicating a dominant evaporite-sourced, basinal brine reservoir. Ore formation is inferred to have taken place during inversion of the contiguous Aptian-Albian Cañete back-arc basin, fluid migration being focused by the intrusion of flanking granitoid plutons of the Coastal Batholith.

177 The Mina Justa IOCG mineralization shares many characteristics with other major Andean

Cu-rich IOCG deposits, e.g., Raúl-Condestable, Mantoverde and La Candelaria-Punta del Cobre

(Table 4-2). However, significant differences between these centres may be evidence for different ore-forming processes and/or hydrothermal fluid sources. Thus, both the Raúl-Condestable and

La Candelaria-Punta del Cobre districts are dominated by magnetite-rich Cu mineralization assemblages which formed at higher temperatures than those of the magnetite-absent Cu mineralization stages at Mina Justa and Mantoverde (Table 4-2). Syn-mineralization granitoid plutons were emplaced adjacent to the orebodies in the former two deposits, but not at Mina Justa or Mantoverde. This would support a more important contribution of magmatic fluids at

Raúl-Condestable and La Candelaria, although external fluids are still considered necessary for economic Cu mineralization. Fluid mixing between magmatic and “exotic” fluids may have been the dominant mechanism responsible for Cu mineralization in the higher temperature deposits, where the hydrothermal fluids were CO2-rich, whereas the reduction of Ca-rich, low-T and

CO2-poor external fluids by early magnetite bodies was a critical factor in the Cu mineralization at Mina Justa and Mantoverde. The ore metals, e.g., Cu, Au and Ag, were probably derived from both magmatic and non-magmatic fluids, but could have been largely introduced by external fluids at Mina Justa.

The mid-Cretaceous metallogenetic sub-province, incorporating the major Cu-rich IOCG deposits of the Andes, viz. Raúl-Condestable, Mina Justa, La Candelaria-Punta del Cobre and

Mantoverde, the Cu-poor CIB iron deposits, several “manto-type” stratabound Cu (-Ag) deposits and minor porphyry Cu (-Mo-Au) deposits (Fig. 4-2; Sillitoe, 2003), developed in an extensional arc in a transtensional tectonic regime (Polliand et al., 2005). Mineralization was largely coeval with the shallow emplacement of granitoid plutons along the margins of aborted mid-Cretaceous

178

Table 4-2. Salient Features of the Major Central Andean Cu-rich IOCG deposits

Deposit Raúl-Condestable 1) Mina Justa 2) Mantoverde 3) La Candelaria – Punta del Cobre 4) Magmatism broadly Coastal Batholith (~115 Ma): quartz Acarí Pluton (~110 Ma): Las Tazas Pluton (~130 Ma): San Gregorio Pluton (~112 Ma): coeval with diorite porphyry, tonalite (Y)* Granodiorite, tonalite, porphyritic granodiorite (Y)*; Sierra Dieciocho monzogranite (Y)* mineralization Volcanic rocks (~115 Ma): Casma dacite (N)* pluton (~120 Ma): quartz-diorite (N) Intramineral dacite porphyry dike (Y)* Formation - calc-alkaline dacitic to Volcanic rocks (<110 Ma): upper Volcanic rocks (~130 Ma): andesitic flows Copara Formation Intramineral andesitic and dacitic dikes, Punta del Cobre Formation Host rocks Copara Formation (~130 Ma): Upper Rio Grande Formation (~160 La Negra Formation (≥ 150 Ma): Punta del Cobre Formation (~130 Ma) : basalt, basaltic andesite and andesite Ma): andesite and volcanic andesite andesite sediments

Age of mineralization ca. 115 Ma (titanite-magnetite 101-104 Ma (Magnetite-pyrite stage) 1261-30 Ma (magnetite stage) Initiated at > 125 Ma (at Punta del Cobre) stage) 95-99 Ma (Cu stage) 118-120 Ma (part of Cu stage) 112 - 115 Ma (main ore stage) Structures Regional NW-striking normal Regional NW-striking strike-slip N (or NW)-striking strike-slip Atacama N (or NW)-striking strike-slip Atacama faults; orebodies (veins) controlled faults; orebodies are cotrolled by fault faults; orebodies controlled by by NW and NE-striking normal (?) extensional normal faults low-angle E-dipping extensional faults Candelaria shear-zone Alteration/mineralizati lateral; breccias and veins upward and lateral; breccias and upward; breccias and veins not clear, but hematite to south; breccias on zonation and style veins and veins Alteration types ** 1) biotite alteration 1) Na and K-Fe metasomatism 1) albitization (not in mine area) 1) Na-metasomatism 2) magnetite–actinolite / or chlorite 2) actinolite (-magnetite) alteration 2) K-feldspar - magnetite alteration 2) biotite–magnetite alteration – sericite alteration 3) Potassic alteration with 3) Chlorite–sericite–scapolite– pyrite 3) main-stage Cu mineralization with 3) Cu-mineralization with/without magnetite-pyrite mineralization alteration magnetite and amphibole alteration magnetite and hematite 4) Cu mineralization with 3) Hematite–chalcopyrite-calcite (Ca-metasomatism) calcite-hematite 4) Hematite-calcite-chalcopyrite

179 Table 4-2 (continued) Relationship between Cu-Au mineralization partially Cu-mineralization clearly later than Cu-mineralization clearly later than main-stage Cu-Au mineralization is iron oxides and Cu associated with magnetite/hematite magnetite-pyrite stage magnetite stage; closely associated with magnetite, but (-Au) mineralization followed main Fe metasomatism. Minor late hematite-Cu sulfide veins Mineralization cp-py-po-qtz±mt cp-bn-cc-cal-hm cp-hm (-cal-qtz) cp-po-qtz-amph ±mt assemblage o o o o Mineralization 320-360 C; CO2-? 85-200 (av. 140) C, CO2-poor 150-360 (av. 240) C; CO2-poor 300 – 450 C; CO2-rich conditions Ore-forming fluid Magmatic hydrothermal fluids for Magmatic hydrothermal fluids for Magmatic hydrothermal fluids for Magmatic hydrothermal fluids for early compositions magnetite stage; external magnetite stage; external magnetite stage; external low-medium magnetite-biotite stage; low-medium (including IO stage) high-salinity, high δ34S fluids low-temperature, high-salinity, temperature, high δ34S fluids (basinal temperature, Ca-rich, high δ34S-value (seawater or basinal fluids) involved Cu-Ca-rich, and high δ34S fluids fluids) dominant at Cu-mineralization fluids (basin brines) involved in Cu-mineralization (basinal fluids) dominant at stage increasingly in Cu-mineralization stage Cu-mineralization stage Mineralization Fluid mixing (as at Olympic Dam); Cu-rich fluid replacement of early Cu-rich (?) fluid mixing with magmatic Fluid mixing (similar to Olympic Dam); mechanism and hydrothermal replacement of magnetite-pyrite bodies fluids, and replacement of early and hydrothermal replacement of host host rocks magnetite stage rocks Possible Cu source Evidence of magmatic source No evidence of magmatic source No evidence of magmatic source Clear evidence for magmatic contribution Relationships to other Small Cu-rich IOCG deposits Giant IO (Marcona); small IO and Small Iron oxide and Cu-rich IOCG Small Iron oxide and “manto-type” Cu deposits Cu-rich IOCG deposits deposits (-Ag) deposits Common features 1) Regional strike-slip fault system (with local extensional faults controlling the orebodies); 2) hosted by low-grade metamorphic volcanic or volcanic - sedimentary rocks (partly Cu-rich); 3) commonly associated with small iron oxide deposits; 4) Cu-mineralization emplaced during the 95-120 Ma interval. *(Y): in the immediate mine area; (N): not in the immediate mine area; (?) indicate no enough evidence; ** numbers indicate the approximate paragenesis sequence. 1) de Haller et al., 2006; 2) This study; 3) Benavides et al., 2007; 4) Ullrich and Clark, 1999; Pop et al., 2000; Marschik and Fontobé, 2001; Arévalo et al., 2006

180

Figure 4-2. Model for the Early Cretaceous Central Andean mineralization (the deposits are not located geographically).

181 basins (Oyarzun et al., 2003; Sillitoe, 2003). The 95-104 Ma Mina Justa magnetite alteration and superimposed Cu mineralization, as well as the precursor 109 Ma actinolite alteration, may have been nucleated by plutons of the Arequipa Segment of the Coastal Batholith (≤109 Ma;

Vidal et al., 1990), which also host the magnetite mineralization at Acarí (Injoque, 1985), and iron oxide-rich Cu veins and Na- and K- metasomatism at Cobrepampa and La Argentina, 30-50 km to the southeast of Mina Justa (Injoque, 2002). The slightly older (114-115 Ma) IOCG mineralization at Raúl-Condestable and Eliana occurred simultaneously with the earliest intrusions of Coastal Batholith (Arequipa segment) in central-south Perú (de Haller et al., 2006).

All of the Peruvian IOCG deposits developed on the western margin of the Early Cretaceous

Cañete Basin. Small Albian iron-oxide Cu (-Au) deposits (e.g., Cerro Morritos: Clark et al.,

1990) and porphyry Cu deposits (e.g., Yaral: Quang, 2003) formed at this time in the Ilo-Tacna area of southernmost Perú. In contrast to southern Perú, the mid-Cretaceous metallogenic belt of northern Chile exhibits a complex and protracted history of mineralization. Initiated at ~ 128

Ma (Zentilli, 1974; Gelcich et al., 2002) and persisting to the late Albian (Wilson et al., 2003;

Sillitoe and Perelló, 2005), hydrothermal alteration and mineralization culminated at 110-125 Ma, represented by major Cu-Au mineralization at La Candelaria - Punta del Cobre (Ullrich and

Clark, 1999, Pop et al., 2000), the El Romeral magnetite deposit (Munizaga et al., 1985) and the

Andacollo porphyry Cu-Au deposit (Sillitoe and Perelló, 2005). Oxide melts may have generated some of the largest CIB deposits (Nyström and Henríquez, 1994), but magmatic-hydrothermal fluids were responsible for the porphyry Cu (-Mo-Au) mineralization, the Fe-(K-) metasomatism around some CIB deposits, and the early magnetite bodies in the IOCG deposits (Ullrich and

Clark, 1999; Ullrich et al., 2001; Sillitoe, 2003; Benavides et al., 2007). However, the significant but varying temporal separation of Fe-metasomatism and economic Cu mineralization in the

IOCG deposits, e.g., 2-3 m.y. at La Candelaria (Ullrich and Clark, 1999, but cf. Marschik and

Fontboté, 2001; Mathur et al., 2002), ~ 10 m.y. at Mantoverde (Gelcich et al., 2005), and 2-5 m.y.

182 at Mina Justa, probably records the involvement of different ore-forming fluid sources at these stages. The incursion of evaporite-sourced basinal brines may be a prerequisite for economic

Cu mineralization in these deposits. Basinal connate - metamorphic brines are also advocated as responsible for Cu mineralization in the El Soldado “manto-type” deposit in central Chile (Boric et al., 2002) and for several small stratabound copper deposits in the Copiapó area, 10-20 km east of the La Candelaria-Punta del Cobre IOCG centre and within the Neocomian back-arc basin

(Haggan et al., 2003; Cisternas et al., 2006).

4.3 The protracted history of alteration and mineralization in the Marcona-Mina Justa

district and other IOCG centres

Geochronological data reveal that the Marcona-Mina Justa district experienced a ~ 80 m.y. history of alteration and mineralization, extending from the Middle Jurassic to the mid-Cretaceous. At Marcona itself, magmatic-hydrothermal activity persisted for ca. 20 m.y., from the precursor Mg-silicate alteration (177 Ma) to the weak polymetallic sulphide stage (154

Ma), while at Mina Justa, IOCG-like early stage K-Fe metasomatism occurred at 142 Ma, ca. 50 m.y. before the termination of Cu mineralization at 95 Ma.

The protracted history of alteration and mineralization documented herein for each of the

Marcona and Mina Justa deposits has been recognized in other major IOCG centres in the Andes and elsewhere. In the La Candelaria – Punta del Cobre district, a U-Pb zircon age for intramineralization dacite indicates that Cu (-Au) mineralization had commenced by 125 Ma (Pop et al., 2000), whereas the major Cu-Au orebodies were emplaced between ca. 115 Ma (Ullrich and Clark, 1999; Mathur et al., 2002) and ca. 112 Ma (Ullrich and Clark, 1999; Marschik and

Fontboté, 2001). In the wider Mantoverde district, mineralization was initiated at 131-126 Ma, defined by a U-Pb isochron age for apatite from the probably magmatic Carmen iron deposit and a U-Pb age for titanite from Mantoverde (Gelcich et al., 2003, 2005), but K-Ar ages of sericite

183 from Mantoverde (Vila et al., 1996) and 40Ar/39Ar plateau ages of actinolite from the Todos Los

Santos Cu-Au veins and the Jerusalem magnetite-apatite deposit, although rejected by Gelcich et al. (2005), strongly suggest that IOCG activity persisted at least to 117 Ma. At Ernest Henry,

Queensland, early Na-Ca alteration and biotite-magnetite alteration took place at 1514-1529 Ma

(Mark et al., 2006), whereas economic Cu (-Au) mineralization probably persisted until ~1476

Ma (Perkins and Wyborn, 1998). In the Olympic Dam - Prominent Hill - Acropolis district, Cu

(-Au-U) mineralization extended from 1570 to 1600 Ma, overlapping with the intrusion of

Hiltaba Suite granitoids (Skirrow et al., 2002, 2007). Aitik, Norbotten, a porphyry Cu-Au deposit emplaced at 1890 Ma, experienced IOCG-type (Na-Ca and K) alteration and Cu and Au redistribution from 1800 to 1750 Ma (Wanhainen, 2005; Wanhainen et al., 2005). Whereas major magnetite mineralization at Kiirunavaara extended for ca. 20 m.y. from 1990 to 1880 Ma (Romer et al., 1994), hydrothermal activity, represented by disseminated and vein pyrite in the orebodies, probably persisted to ca. 1490 Ma (Cliff and Rickard, 1992). An extended duration of alteration and mineralization is therefore a characteristic feature of major IOCG centres and districts.

Similarly, both ore-controlling fault displacement and magmatism in IOCG provinces exhibit longlived activity, as exemplified by the Mesozoic Atacama fault system in northern Chile, where displacement was initiated in the Early Jurassic (Brown et al., 1993) and granitoid plutons were emplaced from Early Jurassic to Early Cretaceous (Grocott et al., 1994). In contrast, Tertiary magmatism and associated strictly magmatic-hydrothermal porphyry Cu mineralization in the

Central Andes was restricted to intervals of less than 10 m.y. and was concentrated in three discrete < 50 km-wide, longitudinal belts (Sillitoe and Perelló, 2005). Such “arc-movement” has been ascribed to slab flattening by Kay and Abruzzi (1996) or to tectonic erosion by Skewes and

Stern (1994). However, the Central Andes was dominated during the Mesozoic by an extentional arc overlying a steep subduction zone (Benavides-Cáceres, 1999). Fault systems developed along

184 the basin margins and were active during multiple basin opening and inversion, an environment favourable to the long-distance migration of external fluids (e.g., Benavides et al., 2007).

4.4 Implications for the genesis of IOCG deposits: a redefinition and reclassification of

the IOCG clan

4.4.1 The important role of external fluids in IOCG mineralization

The genetic models developed in this study for the Marcona and Mina Justa deposits imply that

Cu-poor “Kiruna-type” magnetite deposits form directly from oxide melt and are not genetically related to true, Cu-rich IOCG deposits in which an external fluid is a prerequisite for Cu

(-Au-Ag-U) mineralization, even though they are commonly spatially and temporally associated.

Small magnetite deposits, which are associated with the major Cu-rich deposits, e.g., the Acarí and Ferífera centres in the Central Andes and the Cu-barren “ironstones” of the Cloncurry district, are inferred to be poor in Cu because of the paucity or non-involvement of sulphur- and/or

Cu-rich external fluids (Hawkes et al., 2007; Oliver et al., 2004).

Although the sources of the ore metals remain problematic in most IOCG deposits, external sulphur, either from surficial basinal brines and seawater (e.g., Central Andean deposits, Ullrich and Clark, 1999; Ullrich et al., 2001; Benavides et al., 2007) or from formation water and metamorphic fluids (e.g., Olympic Dam and other IOCG deposits in the eastern Gawler Craton:

Bastrakov et al., 2007; the Cloncurry deposits: Mark et al., 2006; Kendrick et al., 2007), or introduced by magmatic assimilation of metasedimentary units (e.g., Phalaborwa: Mitchell and

Krouse, 1975; Harmer, 2000; Drüppel et al., 2006), has been documented in all major Cu-rich

IOCG centres. However, only the evaporite-sourced fluids yield diagnostically high δ34S values

(i.e., > 10‰: Thode, 1991), while sedimentary formation water or metamorphic fluids commonly have lower values and are less clearly distinguishable from magmatic fluids, as in the Cloncurry

185 deposits in which the involvement of external fluids is revealed by other evidence, such as noble gas isotopes (Kendrick et al., 2007).

The addition of sulphur, and hence the attainment of sulphide mineral solubility products in originally sulphur-poor systems, is considered to be responsible for the economic Cu (-Au) mineralization in most, if not all, IOCG centres. This takes place through mixing of external fluids and sulphur-poor magmatic fluids (e.g., Olympic Dam, Ernest Henry and La Candelaria), or by fluid reduction through the replacement of ironstones or other host rocks by late oxidized sulphur-rich external fluids (e.g., Mina Justa and Mantoverde). Alternatively, sulphur saturation in magmas with limited sulphur capacity, such as those of the carbonatites of the Phalaborwa complex, may have resulted directly from the assimilation of sulphur-rich country rocks, or the involvement of reduced and sulphur-rich magmatic fluids modified by Na (-K) metasomatism

(fenitization) of host rocks. Differing degrees of oxidation in such ore-forming systems may result in magnetite or hematite-dominant IOCG systems. In contrast, Cu-barren magnetite - dominated veins (the M veins of Arancibia and Clark (1996)) in sulphur-saturated porphyry

Cu-Au and skarn Cu-Au deposits are generated at high-temperature (e.g., 450 - ≥750oC at Bajo de la Alumbrera: Ullrich et al., 2002) and relatively higher fO2 (up to the magnetite-pyrite oxygen buffer) than later Cu-Au mineralization stages. In such an extreme environment, Cu sulphides do not attain saturation and Fe is commonly represented by iron oxides (Hezarkhani and

Williams-Jones, 1999).

4.4.2 A revised definition and classification of hydrothermal IOCG deposits

Recent definitions and classifications of the IOCG clan have been proposed by Williams et al.

(2005) and Hunt et al. (2007). The former defined IOCG deposits as a group of hydrothermal iron oxide-rich Cu (-Au) deposits which are not clearly associated with intrusions, a problematic criterion given the controversial relationships between granitoid bodies and mineralization at, e.g.,

186 La Candelaria and Raúl-Condestable, as well as at Olympic Dam itself. Although two major different ore-forming fluid sources, magmatic and non-magmatic, were considered to be responsible for IOCG mineralization, Williams et al. (2005) did not advocate to a systematic classification of IOCG clan on the basis of these important differences. A classification based on fluid sources has, however, been proposed by Hunt et al. (2007), viz. magmatic, hybrid magmatic

– non-magmatic and non-magmatic. In this classification, however, most IOCG deposits, and all large examples such as Olympic Dam, Ernest Henry, Salobo, La Candelaria and Aitik, have been assigned to a non-magmatic-associated subclans, and the magmatic-source examples are either magnetite-dominant with subeconomic Cu (-Au), e.g., Lightning Creek, or small magnetite-rich

Cu-Au deposits in which non-magmatic fluids may still have been involved, e.g., Eloise. Further, a classification based on the covert nature of the ore-forming fluids is difficult to utilize in mineral exploration.

On the basis of these arguments, IOCG deposits are redefined as a clan of Cu (-Au-Ag-U) deposits containing abundant hypogene iron oxide (magnetite and/or hematite), in which externally-derived sulphur is probably a prerequisite for the Cu (-Au-Ag-U) mineralization. In this definition, all “Kiruna-type” magnetite deposits, hydrothermal iron deposits and magnetite-rich porphyry Cu-Au and skarn Cu-Au deposits are excluded. A classification and the defining characteristics of this newly-defined IOCG clan are presented in Table 4-3, the salient criteria being ore and alteration mineralogy and the nature of the dominant ore-forming fluids.

Common Features: In addition to the abundant iron oxide and external sulphur involvement, most IOCG deposits exhibit several common features, viz. (1) with the exception of Phalaborwa, orebodies are commonly controlled by faults (shear zones and normal faults); (2) arc located on basin margins; (3) hydrothermal breccias; (4) regional Na (±Ca) alteration; (5) Cu (-Au) mineralization associated with K (±Cl±Ca) alteration; (6) alteration zonation unrelated to local intrusions; and (7) low-Ti magnetite.

187 Table 4-3. Proposed Classification of IOCG Deposits

IOCG subtype magnetite subclan hematite subclan Ore-forming fluids dominantly magmatica) hybrid magmatic – non-magmatic hybrid magmatic – non-magmatic and non-magmatic Examples Phalaborwa1-2 , Glenover2 (South Africa); Sossego4, Salobo5 (Carajás); Ernest Henry6-7, Eloise?-8 Olympic Dam23, Prominent Hill24, Redbank25 Swartbooisdrif3 (Nambia) (Cloncurry); West Peko9 (Tennant Creek); La Candelaria-Punta (Australia); Mantoverde26, Mina Justa27, Mantos del Cobre10-11, Raúl-Condestable12 (Central Andes); Tjårrojåkka13, Blancos?-28-29 (Central Andes); Mont-de-l’Aigle30, Aitik(c)-14 (Norbotten); Khetri15 (NW India); Boss-Bixby16 (SW Sue-Dianne19 (Canada); Salton Sea31 (USA); Missouri); Wernecke17-18, Nico19 (Canada); Guelb Moghrein20 Malundae?-32 (Lufilian arc); Boleo33 (Mexico) (Mauritania); Ossa Morena21-22 (Spain) Economic elements Cu (±Au±P±REE±Ni±F±Mg) Cu-Au (±Co±Bi±U±REE±Ni) Cu±Ag±Au±Co±Zn (U-REE : Olympic Dam) Ore assemblages Cpy±Bn±Cub±Mt±Apt±Phl Commonly: Mt-Cpy-Py±Po±Hem Hem-Cpy-Bn-Cc Locally: Mt-Cpy-Bn-Cc (Salobo) and Hem-Cpy-Py Hydrothermal Regional/Mineralization-related: Regional: Na±Ca Regional: weak Na±Cl±Ca alteration Na±F±K±Mg Mineralization-related: K±Fe±Ca±Cl±Na Mineralization-related: K-Cl±Na±Ca Ore morphology Pipe-like. Disseminated grains or massive Veins and breccias; local lenses (Salobo) Breccias, lenses (manto) and veins blebs of sulphides in carbonatite Ore precipitation Immiscibility and cooling Fluid mixing and cooling, local fluid reduction Fluid mixing and fluid reduction mechanism 34 34 Non-magmatic no Basinal brine/seawater (Candelaria); δ Sfluids > +10‰; Basinal brine/seawater: δ Sfluids commonly > fluids Metamorphic fluids/formation water (Ernest Henry; Guelb +10‰ 34 Moghrein): δ Sfluids < +10‰ o b o o Characteristics of High T (>650 C); CO2-rich High-T (Ernest Henry: > 350 C) or wide range of T (Salobo: Moderate-low T (<300 C), and moderate-high ore-forming fluids 100-500oC); moderate-high salinity (15-50 wt.% NaCl eq.); Na-Ca salinity (>15 wt.% NaCl eq.); Ca-Na dominant

dominant and Na>Ca; CO2-rich or CO2-bearing and Ca>Na; CO2-bearing or CO2-poor Mineralization level Deep (>3 km) Deep (> 6 km, Ernest Henry) to shallow (< 3 km, Candelaria) Shallow (generally < 3 km)

188 Subtype magnetite subclan IOCG hematite subclan IOCG Ore-controlling Dilation zone (?) in rift Shear zone Normal fault-detachement-dilational zone / Shear structure zone Host rocks Archean greenstones: gneiss, amphibolite Sedimentary and volcanic rocks and intrusions Sedimentary and volcanic rocks and intrusions and schist Distance to Hosted by carbonatite Close (< 3 km) – Sossego, Eloise, La Candelaria; Close (< 3 km) – Olympic Dam; Mantos Blancos syn-mineralization Far (> 3 km) – Ernest Henry, Wernecke Far (> 3 km) – Mina Justa, Mantoverded) magmatism Mineralization age Proterozoic Archean to Tertiary Mesoproterozoic to Pleistocene Tectonomagmatic Rift at Precambrian cratoon edges; (A) Precambrian intracratonic basins, anorogenic; A-type (A) (Olympic Dam, Prominent Hill) setting decompression melting of metasomatised magmatism (Carajás, Norbotten, SW Missouri) mantle, alkaline carbonatitic magmatism. (B) Inversion of basins in extensional arc on subduction-related (B) (Central Andean deposits) continental margin; calc-alkaline arc magmatism (Central Andes) (C) Inversion of post-collisional orogenic basins; calc-alkaline (C) (Mont-de-l’Aigle) magmatism postdating metamorphic peak (Cloncurry) (D) extension following retrograde metamorphism (Guelb Moghrein, Wernecke)

a) immiscible melt and hydrothermal fluids ?; b) Defined in this table: CO2-rich - fluid inclusions containing CO2 ≥ 20%; CO2-bearing: a few fluid inclusions (<20%) contain CO2;

CO2-poor: no CO2 identified in fluid inclusions; c) late IOCG-like magnetite-bornite-chalcopyrite stage; d) intra-minerlaization dykes locally developed. ? – evidence for external fluids not clear. References: 1) Groves and Vielreicher, 2001; 2) Harmer, 2000; 3) Drüppel et al., 2006; 4) Monteiro et al., 2008; 5) Requia and Fontboté, 1999; 6) Mark et al., 2006; 7) Kendrick et al., 2007; 8) Baker, 1998; 9) Skirrow and Walsh, 2002; 10) Ullrich and Clark, 1999; 11) Barton et al., 2007; 12) de Haller et al., 2006; 13) Edflet et al., 2005; 14) Wanhainen et al., 2003; 15) Knight et al., 2002; 16) Day et al., 2001; 17) Hunt et al., 2005; 18) Hunt et al., 2007; 19) Goad et al., 2000; 20) Kolb et al., 2006; 21) Tornos and Casquet, 2005; 22) Tornos et al., 2005; 23) Haynes et al., 1995; 24) Skirrow et al., 2002; 25) Knutsen et al., 1979; 26) Benavides et al., 2007; 27) this study; 28) Maksaev and Zentilli, 2002; 29) Ramírez et al., 2006; 30) Simard et al., 2006; 31) McKibben and Elders, 1985; 32) Nisbet et al., 2000; 33) Conly et al., 2001 Mineral abbreviations: Apt-apatite, Bn-bornite, Cc-chalcocite, Cpy-chalcopyrite, Cub-cubanite, Hem-hematite, Mt-magnetite, Phl-phlogopite, Po-pyrrhotite, Py-pyrite.

189 Two subtypes of IOCG deposits are recognized on the basis of the predominant iron oxide directly associated with the Cu- (Au-) mineralization, whether magnetite or hematite.

Magnetite subclan: Magnetite is the major iron oxide associated with the main Cu-Au mineralization stage(s) in these deposits. Hematite formed at a late paragenetic stage and, if it formed earlier, is replaced by magnetite (“mushketovite”; e.g., La Candelaria). The mineralization comprises economic Cu and Au (> 0.2 g/t), with minor Co, Bi and U-REE (e.g.,

Tennant Creek), locally rich in P, Ni, F and Mg (e.g., Phalaborwa). The main ore assemblage includes magnetite, chalcopyrite, pyrite and/or pyrrhotite, but locally includes the magnetite - chalcopyrite - bornite - chalcocite association (e.g., Salobo and Phalaborwa). Regional Na±Ca

(albite - scapolite ± actinolite) alteration is common but, except at Phalaborwa, Cu-Au mineralization is closely associated with potassic (biotite and/or K-feldspar) alteration, and locally with calcic (actinolite or clinopyroxene) alteration (e.g., La Candelaria and Eloise). The shear zone-controlled orebodies are composed of veins and breccias, but are locally lensoid (e.g.,

Salobo). At Phalaborwa, the ore-controlling structures are not clear and mineralization occurs in a carbonatite pipe, in which disseminated grains and massive blebs of sulphides are hosted by fine- to coarse-grained carbonate matrix. Evidence of the incursion of non-magmatic fluids is unambiguous in most magnetite subclan IOCG’s, and metal precipitation is inferred in most examples to have resulted from fluid mixing and concomitant cooling. The main Cu mineralization at Phalaborwa and Swartbooisdrif is intimately associated with and hosted by carbonatite and lacks evidence of non-magmatic fluids, but the sulphur isotopic values (≤ +5 ‰:

Mitchell and Krouse, 1975; Drüppel et al., 2006) and εNd and εSr (Harmer, 2000; Yuhara et al.,

2005) of the carbonatite indicate either that metasedimentary, probably evaporitic, units were was assimilated by the carbonatitic magma, or that magmatic fluids modified through fenitization invaded the carbonatitic magma (Harmer, 2000). Such reduced, sulphur-rich fluids may have triggered sulphide saturation and sulphide melt immiscibility in the carbonatite magmas at

190 temperatures exceeding 650oC (Helz and Wyllie, 1979; Eriksson, 1989). The external fluids involved in other hydrothermal, magnetite-dominant IOCG systems were probably derived from either basinal brine/seawater with high δ34S values (commonly > +10‰: La Candelaria and Raúl-

Condestable), or metamorphic fluids/formation water from the middle crust or from sedimentary units with lower δ34S values (commonly < +10 ‰: Ernest Henry and Guelb Moghrein). The ore-forming fluids were high-temperature (e.g., Ernest Henry: 350-440oC) or record a wide range of temperatures (e.g., Salobo: 100-500oC, La Candelaria: 275-450oC), which may imply fluid mixing. Na is the dominant cation in the fluids, but Ca was invariably present and locally abundant (e.g., Eloise and Aitik). CO2 is commonly a major fluid constituent, particularly in the

Cloncurry deposits and at Phalaborwa. Hot, CO2-rich and hypersaline brines have been identified in many magnetite-rich IOCG deposits (Pollard, 2006). However, the areal separation of mineralization and syn-mineralization intrusions is commonly considerable (e.g., ~ 15 km at

Ernest Henry), in strong contrast to magnetite-rich porphyry Cu-Au and skarn deposits. Magnetite subclan IOCG deposits were emplaced over a wide depth interval, from more than 6 km at Ernest

Henry and Guelb Moghrein, to relatively shallow levels (e.g., Candelaria and Sossego). In contrast, the hematite-series IOCG deposits were all emplaced at shallow levels (i.e., < 3 km;Fig.

4-3).

Hematite subclan: Hematite- and sulphide- cemented hydrothermal breccias are common in this subtype, and the Cu sulphides are dominated by Fe-poor chalcopyrite- bornite-chalcocite

(and/or digenite) assemblages, although only chalcopyrite occurs in some deposits (e.g.,

Mantoverde). In contrast to magnetite-rich IOCG deposits, Au is only locally enriched in the hematite subtype (e.g., Olympic Dam), whereas Ag and Zn are more abundant. Vertical and lateral sulphide zonation, from chalcopyrite-pyrite to bornite-chalcocite, is documented in many hematite IOCG’s (e.g., Olympic Dam, Mantos Blancos and Mina Justa). Hydrothermal alteration

191

Figure 4-3. Carton illustrating the settings of IOCG deposits.Alteration zoning in IOCG deposits is summarized in Fig. 1-7. Regional Na-Ca or Na alteration commonly precedes mineralization, while potassic alteration and hydrolytic (sericite-chlorite) alteration are usually mineralization-related. 1) Ca-metasomatism dominates the La Candelaria-Punta del Cobre district. 2) Na-Ca alteration widespread in the Cloncurry district, whereas Na-alteration dominates the Central Andean IOCG centres. NB. Under conditions of high thermal gradient/heat-flow, the Y-axis will be compressed. The thicknesses of arrows indicate the relative contributions of different fluid sources. Basinal brines or seawater are probably more oxidized than magmatic and metamorphic fluids.

associated with Cu mineralization is dominated by sericitic or K-feldspar - chlorite alteration, rather than the biotite±amphibole assemblages characteristic of magnetite-rich IOCG deposits, and regional Na-Ca alteration is only weakly developed around most hematitic IOCG systems.

The ore-forming fluids in hematite IOCG’s also differ from those of the magnetite subclan. They

o are usually cooler (< 300 C), and have higher contents of Ca (Ca > Na), but less CO2. External fluids, mainly basinal brines and modified seawater with high δ34S values, commonly > +10‰, unambiguously played a major role in the hematitic IOCG ore-forming systems. Fluid mixing

192 was directly responsible for Cu mineralization at Olympic Dam, Mantos Blancos and Mantoverde, all dominated by breccias, whereas at Mina Justa, manto orebodies formed through the replacement of ironstone or other host- rocks by heated basinal brines or seawater, with fluid reduction as the dominant mechanism for Cu mineralization.

Neither magnetite- nor hematite-rich IOCG deposits show any preference for specific host rocks, and both range in age from Neoarchean to Pleistocene. The tectonomagmatic settings include: Precambrian intracratonic rifts with A-type (Olympic Dam, Sossego), or even carbonatitic (Phalaborwa) magmatism; inverted basins within extensional arcs along convergent continental margins with calc-alkaline arc magmatism (Central Andean IOCG’s); and inverted post-collisional orogenic basins with calc-alkaline magmatism (Cloncurry, the Lufilian arc).

Extension following retrograde metamorphism may accompany IOCG mineralization at Guelb

Moghrein and Wernecke.

193 References

Adshead, N.D., Voulgaris, P., and Muscio, V.N., 1998, Osborne copper–gold deposit, in

Berkman, D.A., and Mackenzie, D.H., eds., Geology of Australian and Papua New

Guinean Mineral Deposits: The Australasian Institute of Mining and Metallurgy,

Melbourne, p. 793–799.

Aguirre, L., 1988, Chemical mobility during low-grade metamorphism of a Jurassic lava flow:

Río Grande Formation, Perú: Journal of South American Earth Sciences, v. 1, p.

343–361.

Aguirre, L., and Offler, R., 1985. Burial metamorphism in the Western Peruvian Trough: Its

relation to Andean magmatism and tectonics, in Pitcher, W.S., Atherton, M.P., Cobbing,

E.J., and Beckinsale, R.D., eds., Magmatism at a Plate Edge: The Peruvian Andes:

Glasgow, Blackie and Son Ltd., p. 59-71.

Alt, J.C., 1999, Very low-grade hydrothermal metamorphism of basic igneous rocks, in Frey, M.,

and Robinson, D., eds., Low-grade Metamorphism: Oxford, U.K., Blackwell Science, p.

169–201.

Arancibia, O.N., and Clark, A.H., 1996, Early magnetite-amphibole-plagioclase

alteration-mineralization in the Island Copper porphyry copper-gold-molybdenum

deposit, British Columbia: Economic Geology, v. 91, p. 402-438.

Atchley, F.W., 1956, Geology of the Marcona Iron Deposits, Peru: Unpublished Ph.D. thesis,

Stanford University, California, 150 p.

Atherton, M.P., 1990, The Coastal Batholith of Peru: the product of rapid recycling of new crust

formed within a rifted continental margin: Geological Journal v. 25, p. 331-349.

Atherton, M.P., and Aguirre, L., 1992, Thermal and geotectonic setting of Cretaceous volcanic

rocks near Ica, Perú, in relation to Andean crustal thinning: Journal of South American

Earth Sciences, v. 5, p. 47-69.

194 Atherton, M.P., and Webb, S., 1989, Volcanic facies, structure, and geochemistry of the marginal

basin rocks of central Peru: Journal of South American Earth Sciences, v. 2, p. 241–261.

Atkin, B.P., Injoque-Espinoza, J., and Harvey, P.K., 1985, Cu-Fe amphibole mineralization in the

Arequipa segment, in Pitcher, W.S., Atherton, M.P., Cobbing, E.J., and Beckinsale, R.D.,

eds., Magmatism at a Plate Edge: The Peruvian Andes: Glasgow, Blackie and Son Ltd., p.

261–270.

Baker, T., 1998, Alteration, mineralization, and fluid evolution at the Eloise Cu-Au deposit,

Cloncurry District, Northwest Queensland, Australia: Economic Geology, v. 93, p.

1213-1236.

Baker, T., Perkins, C., Blake, K.L., and Williams, P.J., 2001, Radiogenic and stable isotope

constraints on the genesis of the Eloise Cu–Au deposit, Cloncurry district, NW

Queensland: Economic Geology, v. 96, p. 723–742.

Balcerzak, M., 2003, An overview of analytical applications of time of flight mass spectrometric

(TOF-MS) analysers and an inductively coupled plasma-TOF-MS technique: Analytical

Sciences, v. 19, p. 979-989.

Baldassaro, P.M., Bodnar, R.J., 2000, Low temperature phase relations in the system H2O-NaCl-

FeCl2: application to fluid inclusion studies [abs.]: Geological Society of America,

Abstracts with Program, v. 32, no. 6, p. A-4.

Barton, M.D., and Johnson, D.A., 1996, Evaporite-source model for igneous-related Fe

oxide-(REE-Cu-Au-U) mineralization: Geology, v. 24, p. 259-262.

Barton, M.D., and Johnson, D.A., 2000, Alternative brine sources for Fe-oxide (-Cu-Au) systems:

Implications for hydrothermal alteration and metals, in Porter, T.M., ed., Hydrothermal

Iron Oxide Copper-Gold and Related Deposits: A Global Perspective: Adelaide, Porter

Geoscience Consultancy Publishing, v. 1, p. 43-60.

195 Barton, M.D., and Johnson, D.A., 2004, Footprints of Fe-oxide (-Cu-Au) systems: The University

of Western Australia Special Publication, v. 33, p.112-116.

Barton, M.D., Jensen, E.P., Johnson, D.A., and Ducea, M., 2007, A Cordilleran perspective on

fluid sources for Fe-oxide (-Cu-Au) (IOCG) systems: geologic and isotopic constrains fro

the Copiapo area, Chile [abs.]: Ores and Orogenesis, Program with Abstracts, Arizona

Geological Society, p. 155-156.

Bastrakov, E.N., Skirrow, R.G., and Davidson, G.J., Fluid Evolution and Origins of Iron Oxide

Cu-Au Prospects in the Olympic Dam District, Gawler Craton, South Australia:

Economic Geology, v. 102, p. 1415-1440.

Baxter, R., Meder, K., Cinits, R., and Berezowski, M., 2005, The Marcona copper project – Mina

Justa prospect geology and mineralisation: Congreso Internacional de Prospectores y

Exploradores, Lima, Conferencias, 4th, Instituto de Ingenieros de Minas del Perú, Lima,

Actas [CD-ROM].

Belevtsev, Y.N., 1973, Genesis of high-grade iron ores of Krivoyrog type: Symposium on the

Genesis of Pre-Cambrian Iron and Manganese Deposits, Kiev, 1970, United Nations

Educational Scientific and Cultural Organization UNESCO [Paris], Proceedings, p.

167–180.

Benavides, J., 2006, Iron Oxide-Copper-Gold Deposits of the Mantoverde Area, Northern Chile:

Ore Genesis and Exploration Guidelines: Unpublished Ph.D. thesis, Queen’s University,

Kingston, 269 p.

Benavides, J., Kyser, T.K., Clark, A.H., Oates, C., Zamora, R., Tarnovschi, R., and Castillo, B.,

2007, The Mantoverde iron oxide-copper-gold district, III Región, Chile: the role of

regionally-derived, non-magmatic fluid contributions to chalcopyrite mineralization:

Economic Geology, v. 102, p. 415-440.

196 Benavides-Cáceres, V., 1999, Orogenic evolution of the Peruvian Andes: the Andean cycle, in

Geology and Ore Deposits of the Central Andes: Society of Economic Geologists Special

Publication, v. 7, p. 61-107.

BHP Billiton, 2007, Annual Report. http://www.bhpbilliton.com/bb/investorsMedia/reports/

2007/2007BhpBillitonAnnualReportPage.jsp

Bodnar, R.J., and Vityk, M.O., 1994, Interpretation of microthermometric data for H2O-NaCl

fluid inclusions, in de Vivo, B., and Frezzotti, M.L., eds., Fluid Inclusions in Minerals:

Methods and Applications: International Mineralogical Association, Short Course of the

Working Group, p. 117-130.

Bodnar, R., 2003, Introduction to fluid inclusions, in Samson, I., Anderson, A., Marshall, D., eds.,

Fluid Inclusions - Analysis and Interpretation: Mineralogical Association of Canada,

Short Course Series 32, p. 1-9.

Bookstrom, A.A., 1977, The magnetite deposits of El Romeral, Chile: Economic Geology, v. 72,

p. 1101–1130.

Boric, R., Holmgren, C., Wilson, N.S.F., and Zentilli, M., 2002, The geology of the El Soldado

manto type Cu (Ag) deposit, central Chile, in Porter, T.M., ed., Hydrothermal Iron Oxide

Copper-Gold and Related Deposits: A global perspective: Adelaide, Porter Geoscience

Consultancy Publishing, v. 2, p. 163-184.

Bottinga, Y., and Javoy, M., 1973, Comments on oxygen isotope geothermometry: Earth &

Planetary Science Letters, v. 20, p. 250-265.

Bouzari, F., 2003, Hypogene and Supergene Evolution of the Cerro Colorado Porphyry Copper

(-Molybdenum) Deposit, I Región, Chile: Unpublished Ph.D. thesis, Queen's University,

Kingston, 276 p.

197 Bouzari, F., and Clark, A.H., 2006, Prograde evolution and geothermal affinities of a major

porphyry copper deposit: the Cerro Colorado hypogene protore, I Región, Northern Chile:

Economic Geology, v. 101, p. 95-134.

Bowers, T.S., 1989, Stable isotope signatures of water-rock interaction in mid-ocean ridge

hydrothermal systems: Sulphur, oxygen, and hydrogen: Journal of Geophysical Research,

v. 94, p. 5775-5786.

Brand, L.R., Esperante, R., Chadwick, A.V., Porras, O.P., and Alomia, M., 2004, Fossil whale

preservation implies high diatom accumulation rate in the Miocene–Pliocene Pisco

Formation of Peru: Geology, v. 32, p. 165-168.

Brett, R., 1964, Experimental data from the system Cu-Fe-S and their bearing on exsolution

textures in ores: Economic Geology, v. 59, p. 1241-1269.

Broggi, J. A., 1946. Las terazas marinas de la bahia de San Juan: Boletin de la Sociedad

Geólogica del Perú, v.19, p. 21-33.

Broman, C., Nyström, J.O., Henríquez, F., and Elfman, M., 1999, Fluid inclusions in

magnetite-apatite ore from a cooling magmatic system at El Laco, Chile: Geologiska

Föreningen i Stockholm Förhandlingar (GFF), v. 121, p. 253–267.

Brown, M., 1991, Comparative geochemical interpretation of Permian-Triassic plutonic

complexes of the Coastal Range and Altiplano (25°30' to 26°30'S), northern Chile:

Geological Society of America Special Paper 265, p.157–177.

Caldas, J., 1978, Geologia de los cuadrangulos de San Juan, Acari y Yauca: Instituto Geológico

Minero y Metalúrgico del Perú, Lima, Perú, Boletín 30.

Candela, P.A., 1989a, Magmatic ore-forming fluids: thermodynamic and mass transfer

calculations of metal concentrations: Reviews in Economic Geology, v. 4, p. 203-221.

Candela, P.A., 1989b, Felsic magmas, volatiles, and metallogenesis: Reviews in Economic

Geology, v. 4, p. 223-233.

198 Cardero Resource Corp., 2005, News Release, September 6 2005. http://www.cardero.com/s/

NewsReleases.asp

Chapin, C.E., and Lindley, J.I., 1986, Potassium metasomatism of igneous and sedimentary rocks

in detachment terranes and other sedimentary basins: Economic implications: Arizona

Geological Society Digest, v. 16, p. 118–126.

Chariot Resources, 2006, Mining Journal, November 24 2006, p. 8.

Chew, D., Kirklan, C., Schaltegger, U., and Goodhue, R., 2007, Neoproterozoic glaciation in the

proto-Andes: Tectonic implications and global correlation: Geology, v. 35, p. 1095-1098.

Cisternas, M.E., and Hermosilla, J., 2006, The role of bitumen in strata-bound copper deposit

formation in the Copiapó area, Northern Chile: Mineralium Deposita, v. 41, p. 339-355.

Clark, A.H., 1993, Are outsize porphyry copper deposits either anatomically or environmentally

distinctive?: Society of Economic Geologists Special Publication, v. 2, p. 213–282.

Clark, A.H., Farrar, E., Kontak, D.J., Langridge, R.J., Arenas F., M.J., France, L.J., McBride,

S.L., Woodman, P.L., Wasteneys, H.A., Sandeman, H.A., and Archibald, D.A., 1990,

Geologic and geochronologic constraints on the metallogenic evolution of the Andes of

southeastern Perú: Economic Geology, v. 85, p. 1520–1583.

Clark, A.H., and Kontak, D.J., 2004, Fe-Ti-P oxide melts generated through magma mixing in the

Antauta subvolcanic centre, Peru: Implications for the origin of nelsonite and iron

oxide-dominated hydrothermal deposits: Economic Geology, v. 99, p. 377-395.

Clayton, R., and Mayeda, T.K., 1963, The use of bromine pentafluoride in the extraction of

oxygen from oxides and silicates for isotopic analysis: Geochimica et Cosmochimica

Acta, v. 27, p. 43-52.

Clayton, R.N., O'Neil, J.R., and Mayeda, T.K., 1972, Oxygen isotope exchange between quartz

and water: Journal of Geophysical Research, v. 77, p. 3057-3067.

199 Clayton, R.N., and Keiffer, S.W., 1991, Oxygen isotopic thermometer calibrations, in Taylor,

H.P., O'Neil, J.R., and Kaplan, I.R., eds., Stable Isotope Geochemistry: A Tribute to

Samuel Epstein: The Geochemical Society Special Publication, v. 3, p. 3-10.

Cobbing, E.G., 1998, The Coastal Batholith and other aspects of Andean magmatism in Perú:

Boletin de la Sociedad Geólogica del Perú, v. 88, p. 5-20.

Cole, D.R., Horita, J., Eniamin, V., Polyakov, V.B., Valley, J.W., Spicuzza, M.J., and Coffey,

D.W., 2004, An experimental and theoretical determination of oxygen isotope

o fractionation in the system magnetite-H2O from 300 to 800 C: Geochimica et

Cosmochimica Acta, v. 68, p. 3569-3585.

Conly, A.G., Scott, S.D., Bellón, H., and Beaudoin, G., 2001, The Boléo Cu-Co-Zn deposit, Baja

California Sur, Mexico: the role of magmatic fluids in the genesis of a synsedimentary

deposit, in Piestrzyñski et al., eds., Mineral Deposits at the Beginning of the 21st Century:

Proceedings of the joint 6th Biennial SGA-SEG meeting, Kraków, Poland, 26-29 August,

2001. P. 223-226.

Cornejo, P., Latorre, J.J., Matthews, S., Marquardt, C., Toloza, R., Basso, M., Rodríguez, J., and

Ulloa, C., 2006, U/Pb and 40Ar/39Ar geochronology of volcanic and intrusive events at the

Mantos Blancos copper deposit, II Region, Chile [abs.]: Congreso Geológico Chileno,

11th, Antofagasta, Actas, v. 2, p. 223-226.

Correa, A., 2003, El Espino, un Nuevo Depósito del Tipo Fe-Cu (Au) en Chile. Comuna de Illapel,

Región de Coquimbo : Tésis de Magister, Universidad Católica del Norte, Antofagasta,

Chile, 69 p.

Corriveau, L., 2005, Iron Oxide Copper Gold (±Ag ±Nb, ±P ±REE ±U) Deposits: a Canadian

perspective: Geological Survey of Canada. http://gsc.nrcan.gc.ca/mindep/synth_dep/

iocg/index_e.php

Cox, L. R., 1956, Jurassic mollusca from Perú: Journal of Paleontology, v. 30, p. 1179-1186.

200 Cruise, M., Hitzman, M., Lopez, G., 2007, Baja California, Mexico - new IOCG discoveries in a

frontier district [abs.]: Ores and Orogenesis, Program with Abstracts, Arizona

Geological Society, p. 138-139.

Dalrymple, G. B., and Lanphere, M.A., 1974, 40Ar/39Ar age spectra of some undisturbed

terrestrial samples: Geochimica et Cosmochimica Acta, v. 38, p. 715-738.

Dalstra, H., and Guedes, S., 2004, Giant hydrothermal hematite deposits with Mg-Fe

metasomatism: a comparison of the Carajás, Hamersley, and other iron ores: Economic

Geology, v. 99, p. 1793-1800.

Davis, D.W., Lowenstein, T.K., and Spencer, R.J., 1990, Melting behavior of fluid inclusions in

laboratory-grown halite crystals in the systems NaCl-H2O, NaCl-KCl-H2O,

NaCl-MgCl2-H2O, and NaCl-CaCl2-H2O: Geochimica et Cosmochimica Acta, v. 54, p.

591-601.

Day, W.C., Seeger, C.M., and Rye, R.O., 2001, Review of the iron oxide deposits of Missouri:

magmatic end members of the iron oxide-Cu-Au-U-REE deposit family [abs.]:

Geological Society of America Annual Meeting, Boston, MA, Abstracts with Programs,

v.33, no. 6, p. A-4.

Deckart, K., Clark, A.H., Aguilar, C.A., and Vargas, R.R., 2005, Magmatic and hydrothermal

chronology of the supergiant Río Blanco porphyry copper deposit, Central Chile:

implications of an integrated U–Pb and Ar–Ar database: Economic Geology, v. 100, p.

905–934.

Deer, W.A., Howie, R.A., and Zussman, J., 1997, Rock-Forming Minerals, Volume 2B,

Double-Chain Silicates: Geological Society, London, 2nd edition, 764 p. de Haller, A., Corfu, F., Fontboté, L, Schaltegger, U, Barra, F., Chiaradia, M., Frank, M., and

Alvarado, J.Z., 2006, Geology, geochronology, and Hf and Pb isotope data of the

Raúl-condestable iron oxide-copper-gold deposit, central coast of Perú: Economic

Geology, v. 101, p. 281-310.

201 de Haller, A., Zúñiga, A.J., Corfu, F., and Fontboté, L., 2002, The iron oxide-Cu-Au deposit of

Raúl-Condestable, Mala, Lima, Perú [abs.]: Congreso Geológico Peruano, 11th, Abstracts

volume, p. 80.

Devries, T.J., 1998, Oligocene deposition and Cenozoic sequence boundaries in the Pisco Basin

(Perú): Journal of South American Earth Sciences, v. 11, p. 217–231.

Drüppel, K., Wagner, T., and Boyce, A.J., 2006, Evolution of sulphide mineralization in

ferrocarbonatite, Swartbooisdrif, Northwestern Namibia: constraints from mineral

compositions and sulfur isotopes: Canadian Mineralogist, v. 44, p. 877-894.

Dunin-Borkowski, E., 1970, The Acarí pluton (Perú) as an example of the differentiation of the

tonalitic magma: Geologische Rundschau, v. 59, p. 1141–1180.

Eagle Plains Resources Ltd., 2005, News release, 13 June, 2005. http://www.eagleplains.ca

/news/documents/IRNRjun132005.pdf

Eastoe, C.J., 1978, A fluid inclusion study of the Panguna porphyry copper deposit, Bougainville,

Papua New Guinea: Economic Geology, v. 73, p. 721–748.

Edflet, A., Armstrong, R.N., Smith, M., and Martinsson, O., 2005, Alteration paragenesis and

mineral chemistry of the Tjärrojäkka apatite–iron and Cu (-Au) occurrences, Kiruna area,

northern Sweden: Mineralium Deposita, v. 40, p. 409-434.

Equinox Minerals Limited, 2007, Sulphide Reserves and Resources within Designed Pits.

http://www.equinoxminerals.com/development/lumwana-project

Ettner, D.C., Bjolykke, A., and Andersen, T., 1994, A fluid inclusion and stable isotope study of

the Proterozoic Bidjovagge Au–Cu deposit, Finnmark, northern Norway: Mineralium

Deposita, v. 29, p. 16-29.

Evans, B.W., and Ghiorso, M.S., 1995, Thermodynamics and petrology of cummingtonite:

American Mineralogist, v. 80, p. 649-663.

Far West Mining Ltd., 2007, News Release, September 06, 2007. http://www.farwestmining.

com/s/NewsReleases.asp

202 Frietsch, R., 1978, On the magmatic origin of iron ores of the Kiruna type: Economic Geology, v.

73, p. 478-485.

Frietsch, R., Tuiska, P., Martinsson, O., and Perdahl, J.-A., 1997, Early Proterozoic Cu(-Au) and

Fe ore deposits associated with regional Na-Cl metasomatism in northern Fennoscandia:

Ore Geology Reviews, v. 12, p. 1–34.

Force, E.R. 1991, Geology of Titanium-Mineral Deposits: Geological Society of America, Special

Paper 259, 113 p.

Fortune Minerals Limited, 2007. http://www.fortuneminerals.com/s/NICO.asp.

Friehauf, K.C., Smith, R.C., and Volkert, R.A., 2002, Comparison of the geology of Proterozoic

iron oxide deposits in the Adirondack and mid-Atlantic belt of Pennsylvania, New Jersey

and New York, in Porter, T.M., ed., Hydrothermal Iron Oxide Copper-Gold and Related

Deposits: A Global Perspective: Adelaide, Porter Geoscience Consultancy Publishing, v.

2, p. 247–252.

Fu, B., Williams, P.J., Oliver, N.H.S., Dong, G., Pollard, P.J., and Mark, G., 2003, Fluid mixing

versus unmixing as an ore-forming process in the Cloncurry Fe-oxide–Cu–Au District,

NW Queensland, Australia: evidence from fluid inclusions: Journal of Geochemical

Exploration, v. 78/79, p. 617–622.

Gander, M.J., French, G.M., and Ohlin, H., 2007, The geology and mineralization of the Pumpkin

Hollow Cu-Fe-Au deposit, Yerington, Lyon County, Nevada [abs.]: Geological Society of

America Annual Meeting, Denver, CO, Abstracts with Programs, v.39, no. 6, p. 626.

Gauthier, M., Chartrand, F., Cayer, A., and David, J., 2004, The Kwyjibo Cu-REE-U-Au-Mo- F

property, Quebec: a Mesoproterozoic polymetallic iron oxide deposit in the northeastern

Grenville Province: Economic Geology, v. 99, p. 1177–1196.

Geijer, P., 1910, Igneous Rocks and Iron Ores of Kirunavaara, Luossavaara and Tuolluvaara:

Scientific and Practical Researches in Lapland Arranged by Luossuvaara - Kiirunavaara

Aktiebolag, Stockholm, 278 p.

203 Geijer, P., 1931, The iron ores of the Kiruna type: Sveriges Geologiska Undersökning, series C,

no. 367, 39 p.

Gelcich, S., Davis, D.W., and Spooner, E.T.C., 2002, New U-Pb ages for host rocks,

mineralization and alteration of iron oxide (Cu-Au) deposits in the Coastal Cordillera of

northern Chile: South American Symposium on Isotope Geology, 4th, Salvador de Bahia,

Brazil, Proceedings, p. 63–65.

Gelcich, S., Davis, D.W., and Spooner, E.T.C., 2005, Testing the apatite-magnetite

geochronometer: U-Pb and 40Ar/39Ar geochronology of plutonic rocks, massive

magnetite-apatite tabular bodies, and IOCG mineralization in Northern Chile:

Geochimica et Cosmochimica Acta, v. 69, p. 3367-3384.

Gillen, D., Baker, T., Hunt, J., Ryan, C., and Win, T.T., 2004, PIXE analysis of hydrothermal

fluids in the Wernecke Mountains, Canada: Predictive mineral discovery CRC

Conference, Barossa Valley, Australia, p. 69-73.

Goad, R.E., Mumin, A.H., Duke, N.A., Neale, K.L., and Mulligan, D.L., 2000, Geology of the

Proterozoic iron oxide-hosted NICO cobalt-gold-bismuth and Sue-Dianne copper-silver

deposits, Southern Great Bear magmatic zone, Northwest Territories, Canada, in Porter,

T.M., ed., Hydrothermal Iron Oxide Copper-Gold and Related Deposits: A Global

Perspective: Adelaide, Porter Geoscience Consultancy Publishing, v. 1, p. 249–267.

Goldstein, R.H., 2003, Petrographic analysis of fluid inclusions, in Samson, I., Anderson, A., and

Marshall, D., eds., Fluid Inclusion-Analysis and Interpretation: Mineralogical

Association of Canada, Short Course Series 32, p. 9-54.

Goldstein, R.H., and Reynolds, T.J., 1994, Systematics of fluid inclusions in diagenetic minerals:

Society for Sedimentary Geology Short Course 31, 199 p.

Graham, C.M., Harmon, R.S., and Sheppard, S.M.F., 1984, Experimental hydrogen isotope

exchange between amphibole and water: American Mineralogist, v. 69, p. 128-138.

204 Grip, E., 1978, Sweden - Extract on the Bergslagen district of central Sweden: in Bowie S.H.U.,

Kvalheim, A., and Haslam, H.W., eds., Mineral Deposits of Europe, v. 1: Institution of

Mining and Metallurgy, London, p. 97-138.

Grocott, J., Brown, M., Dallmeyer, R.D., Taylor, G.K., and Treolar, P.J., 1994, Mechanisms of

continental growth in extensional arcs: an example from the Andean plate-boundary zone:

Geology, v. 22, p. 391–394.

Gross, G.A., Gower, C.F., and Lefebure, D.V., 1997, Magmatic Ti-Fe±V oxide deposits, in

Geological Fieldwork 1997: British Columbia Ministry of Employment and Investment,

Paper 1998-1, p. 24J-1 - 24J-3.

Groves, D.I., and Vielreicher, N.M., 2001, The Phalabowra (Palabora) carbonatite-hosted

magnetite-copper sulphide deposit, South Africa: an end-member of the iron

oxide-copper-gold-rare earth element deposit group?: Mineralium Deposita, v. 36, p.

189-194.

Gustafson, L.B., 1978, Some major factors of porphyry copper genesis: Economic Geology, v. 73,

p. 600-607.

Haggan, T., Parnell, J., and Cisternas, M.E., 2003, Fluid history of andesite-hosted CuS-bitumen

mineralization, Copiapó district, North Central Chile: Journal of Geochemical

Exploration, v. 78-79, p. 631-635.

Harmer, R.E., 2000, Mineralization of the Phalaborwa complex and the carbonatite connection in

iron oxide-Cu-Au-U-REE deposits, in Porter, T.M., ed., Hydrothermal Iron Oxide

Copper-Gold and Related Deposits: A Global Perspective: Adelaide, Porter Geoscience

Consultancy Publishing, v. 1, p. 331–340.

Harmon, R.S., Barreiro, B.A., Moorbath, S., Hoefs, J., Francis, P.W., Thorpe, R.S., Déruelle, B.,

McHugh, J., and Viglino, J.A., 1984, Regional O-, Sr-, and Pb-isotope relationships in

late Cenozoic calc-alkaline lavas of the Andean Cordillera: Journal of the Geological

Society, London, v. 141, p. 803-822.

205 Harrison, T.M., 1981, Diffusion of 40Ar in hornblende: Contributions to Mineralogy and

Petrology, v. 78, p. 324–331.

Harrison, T.M., Duncan, I., and McDougall, I., 1985, Diffusion of 40Ar in biotite: Temperature,

pressure and compositional effects: Geochimica et Cosmochimica Acta, v. 49, p.

2461–2468.

Harrison, T.M., Heizler, M.T., and Lovera, O.M., 1993, In vacuo crushing experiments and

K-feldspar thermochronometry: Earth & Planetary Science Letters, v. 117, p. 169-180.

Harrison, T.M., and McDougall, I., 1982, The thermal significance of potassium feldspar K-Ar

ages inferred from 40Ar/39Ar age spectrum results: Geochimica et Cosmochimica Acta, v.

46, p. 1811–1820.

Hawkes, N., Clark, A.H., and Moody, T.C., 2002, Marcona and Pampa de Pongo: giant Mesozoic

Fe-(Cu, Au) deposits in the Peruvian coastal belt, in Porter, T.M., ed., Hydrothermal Iron

Oxide Copper-Gold and Related Deposits: A Global Perspective: Adelaide, Porter

Geoscience Consultancy Publishing, v. 2, p. 115-130.

Hawkes, N., Moody, T., Loader, S., and Abbott, C., 2003, Marcona JV Exploration Report, Rio

Tinto-Shougang Hierro Perú for the Period 16 January to 15 June, 2003, Unpublished

Report, Rio Tinto Mining and Exploration Limited, 42 p.

Haynes, D., 2006, The Olympic Dam ore deposit discovery – a personal view: Society of

Economic Geologists Newsletter, v. 66, p. 7-15.

Haynes, D.W., Cross, K.C., Bills, R.T., and Reed, M.H., 1995, Olympic Dam ore deposit: A

fluid-mixing model: Economic Geology, v. 90, p. 281- 307.

Hedenquist, J.W., Arribas, A.J., and Reynolds, T.J., 1998, Evolution of an intrusion-centreed

hydrothermal system: Far Southeast-Lepanto porphyry and epithermal Cu-Au deposits,

Philippines: Economic Geology, v. 93, p. 373-404.

206 Helz, G. R., and Wyllie, P. J., 1979, Liquidus relationships in the system CaCO3-Ca(OH)2-CaS

and the solubility of sulfur in carbonatite magmas: Geochimica et Cosmochimica Acta, v.

43, p. 259-265.

Hemley, J.J., Cygan, G.L., Fein, B.J., Robinson, G.R., and D’Angelo, W.M., 1992, Hydrothermal

ore-forming processes in the light of studies in rock-buffered systems: I.

Iron-copper-zinc-lead sulphide solubility relations: Economic Geology, v. 87, p. 1-22.

Henríquez, F., and Martin, R.F., 1978, Crystal-growth textures in magnetite flows and feeder

dykes, El Laco, Chile: Canadian Mineralogist, v. 16, p. 581-589.

Henríquez, F., Naslund, H.R., Nyström, J.O., Vivallo, W., Aguirre, R., Dobbs, F.M., and Lledó,

H., 2003, New field evidence bearing on the origin of the El Laco magnetite deposit,

northern Chile - a discussion: Economic Geology, v. 98, p. 1497–1500.

Hey, M.H., 1954, A new review of the chlorites: Mineralogical Magazine, v. 30, p. 277-292.

Hezarkhani, A., Williams-Jones, A.E., and Gammons, 1999, Factors controlling copper solubility

and chalcopyrite deposition in the Sungun porphyry copper deposit, Iran: Mineralium

Deposita, v. 34, p. 770-783.

Hitzman, M.W., Oreskes, N., and Einaudi, M.T., 1992, Geological characteristics and tectonic

setting of Proterozoic iron oxide (Cu-U-Au-REE) deposits: Precambrian Research, v. 58,

p. 241-288.

Hitzman, M.W. and Valenta, R.K., 2005, Uranium in iron oxide-copper-gold (IOCG) systems:

Economic Geology, v. 100, p. 1657-1661.

Hoefs, J., 1997, Stable Isotope Geochemistry, 4th Edition: Springer-Verlag, Berlin, 214 p.

Hosmer, H.L., 1959, Geology and Structural Development of the Andean System of Perú:

Unpublished Ph.D. Thesis, Ann Arbor, University of Michigan, 291 p.

Hooper, D. and Correa, A., 2000, The Panulcillo and Teresa de Colmo copper deposits: two

contrasting examples of Fe-ox Cu-Au mineralization from the coastal Cordillera of Chile,

in Porter, T.M., ed., Hydrothermal Iron Oxide Copper-Gold and Related Deposits: A

207 Global Perspective: Adelaide, Porter Geoscience Consultancy Publishing, v. 1, p.

177–189.

Hudson, C., 1974, Metallogenesis as Related to Crustal Evolution in South-west Central Perú:

Unpublished Ph.D. thesis, University of Liverpool, 150 p.

Huhn, S.R.B., Soares, A.D.V., Souza, C.I.J., Albuquerque, M.A.C., Leal, E.D., Vieira, E.A.P.,

Masotti, F.S., Brustolin, V., 2000, The Cristalino copper–gold deposit, Serra dos Carajás,

Pará. IUGS, International Geological Congress, 31 [CD-ROM].

Hunt, J., Baker, T., and Thorkelson, D., 2005, Regional-scale Proterozoic IOCG-mineralized

breccia systems: examples from the Wernecke Mountains, Yukon, Canada: Mineralium

Deposita, v. 40, p. 492-514.

Hunt, J., Baker, T., and Thorkelson, D., 2007, A review of iron oxide copper-gold deposits, with

focus on the Wernecke Breccias, Yukon, Canada, as an example of a non-magmatic end

member and implications for IOCG genesis and classification: Exploration and Mining

Geology; v. 16, p. 209-232.

Injoque, J., 1985, Geochemistry of the Cu-Fe-Amphibole Skarn Deposits of the Peruvian Central

Coast: Unpublished Ph.D. thesis, University of Nottingham, 597 p.

Injoque, J., 2002, Fe oxide-Cu-Au deposits in Peru: an integrated view, in Porter, T.M., ed.,

Hydrothermal Iron Oxide Copper-Gold and Related Deposits: A Global Perspective:

Adelaide, Porter Geoscience Consultancy Publishing, v. 2, p. 97–113.

Jaillard, E., Hérail, G., Monfret, T. Díaz-Martínez, E., Baby, P., Lavenu, A., and Dumont, J.F.,

2000, Tectonic evolution of the Andes of Ecuador, Perú, Bolivia and northernmaost Chile,

in Cordani, U., Milani, E.J., Thomas Filho, A., and Campos, D.A., eds., Tectonic

Evolution of South America: International Geological Congress, Rio de Janeiro, p.

481–559.

208 Javoy, M., Pineau, F., Cheminee, J.L., and Kraft, M., 1988, The gas-magma relationship in the

1988 eruption of Oldoinyo Lengai (Tanzania): America Geophysical Union, EOS

Transactions, v. 69, p. 1466

Jenkin, G.R.T., Ellam, R.M., Rogers, G., and Stuart, F.M., 2001, An investigation of the closure

temperature of the biotite Rb-Sr system: The importance of cation exchange: Geochimica

et Cosmochimica Acta, v. 65, p. 1141-1160.

Johnson, J.P., and McCulloch, M.T., 1995, Sources of mineralizing fluids for the Olympic Dam

deposit (South Australia): Sm-Nd isotopic constraints: Chemical Geology, v. 121, p.

177-199.

Kajiwra, Y., and Krouse, H.R., 1971, Sulphur isotope partitioning in metallic sulphide systems:

Canadian Journal of Earth Sciences, v. 8, p. 1397-1408.

Kay, S.M., and Abruzzi, J.M., 1996, Magmatic evidence for Neogene lithospheric evolution of

the Central Andean ‘flat-slab’ between 30° and 32°S: Tectonophysics, v. 259, p. 5–28.

Kay, R.W., and Kay, S.M., 1993, Delamination and delamination magmatism: Tectonophysics, v.

29, p. 117-189.

Kelley, S.P., and Wartho, J.A., 2000, Rapid kimberlite ascent and the significance of Ar-Ar ages

in xenolith phlogopites: Science, v. 289, p. 609-611.

Kendrick, M.A., Mark, G., and Phillips, D., 2007, Mid-crustal fluid mixing in a Proterozoic Fe

oxide-Cu-Au deposit, Ernest Henry, Australia: evidence from Ar, Kr, Xe, Cl, Br and I:

Earth & Planetary Science Letters, v. 256, p. 328-343.

Knight, J., Lowe, J., Joy, S., Cameron, J., Merrillees, J., Nag, S., Shah, N., Dua, G., and Jhala, K.,

2002, The Khetri copper belt, Rajasthan: iron oxide copper-gold terrane in the

Proterozoic of NW India, in Porter, T.M., ed., Hydrothermal Iron Oxide Copper-Gold

and Related Deposits: A Global Perspective: Adelaide, Porter Geoscience Consultancy

Publishing, v. 2, p. 321–341.

209 Knutson, J., Fergurson, J., Roberts, W.M.B., and Donnelly, T.H., 1979, Petrogenesis of

copper-bearing breccia pipes, Redbank, Northern Territory, Australia: Economic Geology,

v. 74, p. 814–826.

Kojima, S., Astudillo, J., Rojo, J., Tristá, D., Hayashi, K., 2003, Ore mineralogy, fluid inclusion,

and stable isotopic characteristics of stratiform copper deposits in the coastal Cordillera

of northern Chile: Mineralium Deposita, v. 38, p. 208-216.

Kolb, J., Sakellaris, G.A., and Meyer, F.M., 2006, Controls on hydrothermal Fe oxide-Cu-Au-Co

mineralization at the Guelb Moghrein deposit, Akjoujt area, Mauritania: Mineralium

Deposita, v. 41, p. 68-81.

Kyser, T.K., and O’Neil, J., 1984, Hydrogen isotope systematics of submarine basalts:

Geochimica et Cosmochimica Acta, v. 48, p. 48-53.

Kyser, T.K., Lesher, C.E., and Walker, D., 1998, The effects of liquid immiscibility and thermal

diffusion on oxygen isotopes in silicate liquids: Contribution to Mineralogy and

Petrology, v. 133, p. 373-381.

Larson, R.L., 1991, Latest pulse of Earth: evidence for a mid-Cretaceous superplume: Geology, v.

19, p. 547-550.

Leach, A.M., and Hieftje, G.M., 2001, Standardless semiquantitative analysis of metals using

single-shot laser ablation inductively coupled plasma time-of-flight mass spectrometry:

Analytical Chemistry, v. 73, p. 2959-2967.

Leake, B.E., et al., Nomenclature of amphiboles: report of the subcommittee on amphiboles of the

international mineralogical association, commission on new minerals and mineral names:

Canadian Mineralogist, v. 235, p. 219-246.

Lester, G.W., 2002, The Effects of Excess H2O, and H2O in Combination with F, Cl, S, or P on

Liquid Immiscibility in the System Si-Fe-Al-K-O, at 2 kbar: Implications for the Genesis

of Fe-oxide Magmas: Unpublished M.Sc. thesis, Binghamton, New York, State

University of New York, 66 p.

210 Liipo, J., Laajoki, K., 1991, Mineralogy, geochemistry and metamorphism of the early

Proterozoic Vähäjoki iron ores, northern Finland: Bulletin of the Geological Society of

Finland, v. 63, p. 69-85.

Lindblom, S., Broman, C., and Martinsson, O., 1996, Magmatic-hydrothermal fluids in the

Pahtohavare Cu-Au deposit in greenstone at Kiruna, Sweden: Mineralium Deposita, v. 31,

p. 307-318.

Lledó, L.H., 2005, Experimental Studies on the Origin of Iron Deposits; and Mineralization of

Sierra La Bandera, Chile: Unpublished Ph.D. thesis, Binghamton, New York, State

University of New York, 200 p.

Loewy, S.L., Connelly, J.N., and Dalziel, I.W.D., 2004, An orphaned basement block: The

Arequipa-Antofalla Basement of the central Andean margin of South America:

Geological Society of America Bulletin, v. 116, p. 171-187.

Longstaffe, F.J., 1987, Stable isotope studies of diagenetic processes, in Kyser, T.K., ed., Stable

Isotope Geochemistry of Low Temperature Fluids: Mineralogical Association of Canada,

Short Course Series 13, p. 187-257.

Longstaffe, F.J., 2000, An introduction to stable oxygen and hydrogen isotopes and their use as

fluid tracers in sedimentary systems, in Kyser, T.K., ed., Fluids and Basin Evolution:

Mineralogical Association of Canada, Short Course Series 28, p. 115-150.

Lortie, R. B., and Clark, A . H., 1976, Stratabound fumarolic copper deposits in rhyolitic lavas

and ash-flow tuffs, Copiapó district, Atacama, Chile: in Problems of Ore Deposition vol.

1, Volcanogenous Ore Deposits, Stuttgart, E . Schweizerbart'sche Verlagsbuchhandlung,

p. 256-264.

Lundberg, B., and Smellie, J.A.T., 1979, Painirova and Mertainen iron ores: two deposits of the

Kiruna iron ore type in northern Sweden: Economic Geology, v. 74, p. 1131-1152.

Lyons, J.I., 1988, Volcanogenic iron oxide deposits, Cerro de Mercado and vicinity, Durango:

Economic Geology, v. 83, p. 1886-1906.

211 Mahoney, P., Li, G., and Hieftje, G.M., 1996, Laser ablation-inductively coupled plasma mass

spectrometry with a time-of-flight analyzer: Journal of Analytical Atomic Spectrometry, v.

11, p. 401-405.

Maiden, K.J., Innes, A.H., King, M.J., Master, S., and Pettitt, I., 1984, Regional controls on the

localization of stratabound copper deposits: Proterozoic examples from Southern Africa

and South Australia: Precambrian Research, v. 25, p. 99-118.

Maiden, K.J., and Master, S., 1986, The Mangula copper-silver-gold deposit, Zimbabwe: a Lower

Proterozoic Olympic Dam-type deposit?: Canadian Mineralogist, v. 24, p. 195-196.

Maksaev, V., Zentilli, M., 2002, Chilean strata-bound Cu- (Ag) deposits: an overview, in Porter,

T.M., ed., Hydrothermal Iron Oxide Copper-Gold and Related Deposits: A Global

Perspective: Adelaide, Porter Geoscience Consultancy Publishing, v. 2, p. 185-205.

Marcona Mining Company, 1968, Geologic Map of Marcona.

Mark, G., and Oliver, N.H.S., 2006, Mineralogical and chemical evolution of the Ernest Henry Fe

oxide–Cu–Au ore system, Cloncurry district, northwest Queensland, Australia:

Mineralium Deposita, v. 40, p. 769-801.

Mark, G., Oliver, N.H.S., Williams, P.J., Valenta, R.K., and Crookes, R.A., 2000, The evolution

of the Ernest Henry Fe-oxide-(Cu-Au) hydrothermal system, in Porter, T.M., ed.,

Hydrothermal Iron Oxide Copper-Gold and Related Deposits: A Global Perspective:

Adelaide, Porter Geoscience Consultancy Publishing, v. 1, p. 123-136.

Marschik, R., and Fontboté, L., 2001, The Candelaria-Punta del Cobre iron oxide Cu-Au (-Zn, Ag)

deposits, Chile: Economic Geology, v. 96, p. 1799-1826.

Marrett, R.A., Allmendinger, R.W., Alonso, R.N., and Drake, R.E., 1994, The late Cenozoic

tectonic evolution of the Punta plateau and adjacent foreland, northwest Argentina Andes:

Journal of South American Earth Sciences, v. 7, p. 179-207.

212 Marshall, L.J., and Oliver, N.H.S., 2006, Monitoring fluid chemistry in iron

oxide-copper-gold-related metasomatic processes, eastern Mt Isa Block, Australia:

Geofluids, v. 6, p. 45-66.

Mathur, R.D., Marschik, R., Ruiz, J., Munizaga, F., and Martin, W., 2002, Age of mineralization

of the Candelaria iron oxide Cu-Au deposit, and the origen of the Chilean iron belt based

on Re-Os isotopes: Economic Geology, v. 97, p. 1799–1825.

Matthews, S.J., Sparks, R.S.J., and Gardeweg, M.C., 1995, The relationships between magma

mixing and volatile behaviour at Lascar Volcano (23°22'S; 67°44'W), northern Chile:

significance for the formation of copper sulphide and magnetite-apatite orebodies, in

Clark, A.H., ed., Giant Ore Deposits II: Queen’s University, Kingston, Ontario, p.

169–196.

Matthews, S.J., Sparks, R.S.J., and Gardeweg, M.C., 1999, The Piedras Grandes-Soncor

eruptions, Lascar volcano, Chile; evolution of a zoned magma chamber in the central

Andean upper crust: Journal of Petrology, v. 40, p. 1891-1919.

McGowan, R.R., Roberts, S., Boyce, A.J., 2006, Origin of the Nchanga copper–cobalt deposits of

the Zambian Copperbelt: Mineralium Deposita, v. 40, p. 617-638.

McDougall, I., and Harrison, T.M., 1999, Geochronology and Thermochronology by the 40Ar/39Ar

Method: Oxford University Press, New York, 261 p.

McKibben, M.A., and Elders, W.A., 1985, Fe-Zn-Cu-Pb mineralization in the Salton Sea

geothermal system, Imperial Valley, California: Economic Geology, v. 80, p. 539-559.

McLlaren, S., Dunlap, W.J., and Powell, R., 2007, Understanding K-feldspar 40Ar/39Ar data:

reconciling models, methods and microtextures: Journal of the Geological Society,

London, v. 164, p. 941-944.

McLean, R.C., 2002, The Sin Quyen IOCG-rare earth oxide mineralization of North Vietnam, in

Porter, T.M., ed., Hydrothermal Iron Oxide Copper-Gold and Related Deposits: A

213 Global Perspective: Adelaide, Porter Geoscience Consultancy Publishing, v. 2, p.

293–302.

McLelland, J., Ashwal, Z., and Moore, L., 1994, Composition and petrogenesis of oxide-,

apatite-rich gabbronorites associated with Proterozoic anorthosite massifs: examples from

the Adirondack Mountains, New York: Contributions to Mineralogy and Petrology, v.

116, p. 225–238.

Mitchell, R .H., and Krouse, H .R., 1975, Sulfur isotope geochemistry of carbonatites:

Geochimica et Cosmochimica Acta, v. 39, p. 1505-1513.

Monteiro, L.V.S., Xavier, R.P., de Carvalho, E.R., Hitzman, M.W., Johnson, C.A., de Souza,

C.R., Torresi, I., 2008, Spatial and temporal zoning of hydrothermal alteration and

mineralization in the Sossego iron oxide–copper–gold deposit, Carajás Mineral Province,

Brazil: paragenesis and stable isotope constraints: Mineralium Deposita, v. 43, p.

129-159.

Moody, T.C., Hawkes, N., Ramos, D., Loader, S., Pañez, R., Abbott, C., Carbońell, J., and

Sillitoe, R.H., 2003, The Marcona iron oxide-copper deposits, Perú: Congreso

Internacional de Prospectores y Exploradores, Lima, Conferencias, 3o, Instituto de

Ingenieros de Minas del Perú, Lima, Actas [CD-ROM].

Mukasa, S.B., and Henry, D.J., 1990, The San Nicolás Batholith of coastal Perú: early Paleozoic

continental arc or continental rift magmatism?: Journal of the Geological Society,

London, v. 147, p. 27-39.

Munoz, J. L., 1984, F-OH and CI-OH exchange in micas with application to hydrothermal ore

deposits: Reviews in Mineralogy, Mineralogical Society of America, v. 13, p. 469-491.

Naslund, H.R., 1983, The effects of oxygen fugacity on liquid immiscibility in iron-bearing

silicate melts: American Journal of Science, v. 283, p. 1034–1059.

214 Naslund, H.R., Aguirre, R., and Lledó, H., 2004, Evidence for the formation of massive

magnetite ores by liquid immiscibility [abs.]: IAVCEI General Assembly, 14-19

November, 2004, Pucón, Chile, Abstract volume [CD-ROM].

Naslund, H.R., Henríquez, F., Nyström, J.O., Vivallo, W., and Dobbs, F.M., 2002, Magmatic iron

ores and associated mineralization: examples from the Chilean High Andes and Coastal

Cordillera, in Porter, T.M., ed., Hydrothermal Iron Oxide Copper-Gold and Related

Deposits: A Global Perspective: Adelaide, Porter Geoscience Consultancy Publishing, v.

2, p. 207–228.

Nielsen, H., 1979, Sulfur isotopes: in Jager, E., and Hunziker, J., eds., Lectures in Isotope

Geology, Springer, Berlin Heidelberg New York, p. 283-312.

Niiranen, T., Mänttäri, I., Poutiainen, M., Oliver, N.H.S., and Miller, J.A., 2005, Genesis of

Palaeoproterozoic iron skarns in the Misi region, northern Finland: Mineralium Deposita,

v. 40, p. 192-217.

Nisbet, B., Cooke, J., Richards, M., and Williams, C., 2000, Exploration for iron oxide -

copper-gold deposits in Zambia and Sweden: comparison with Australian experience, in

Porter, T.M., ed., Hydrothermal Iron Oxide Copper-Gold and Related Deposits: A

Global Perspective: Adelaide, Porter Geoscience Consultancy Publishing, v. 1, p.

297–308.

Noble, D.C., McKee, E.H., Mourier, T., and Megard. F., 1990, Cenozoic stratigraphy, magmatic

activity, compressive deformation, and uplift in northern Perú: Geological Society of

American Bulletin, v. 102, p. 1105-1113.

Nyström, J.O., 1985, Apatite iron ores of the Kiruna Field, northern Sweden: magmatic textures

and carbonatitic affinity: Geologiska Föreningens i Stockholm Förhandlingar (GFF), v.

107, p. 133-141.

215 Nyström, J.O., and Henríquez, F., 1994, Magmatic features of iron ores of the Kiruna type in

Chile and Sweden: ore textures and magnetite geochemistry: Economic Geology, v. 89, p.

820-839.

Oakes, C.S., Bodnar, R.J., and Simonson, J.M., 1990, The system NaCl–CaCl2–H2O: I. The ice

liquidus at 1 atm total pressure: Geochimica et Cosmochimica Acta, v. 54, p. 603–610.

Ohmoto, H., Goldhaber, M., 1997, Sulphur and carbon isotopes, in Barnes, H.L., ed.,

Geochemistry of Hydrothermal Ore Deposits, 3rd edn. : Wiley and Sons, New York, p.

517-611.

Ohmoto, H., and Rye, R.O., 1979, Isotopes of sulphur and carbon, in Barnes, H.L., ed.,

Geochemistry of Hydrothermal Ore Deposits, 2nd edn. : Wiley and Sons, New York, p.

509-567.

Oliver, N.H.S., Cleverley, J.S., Mark, G., Pollard, P.J., Fu, B., Marshall, L.J., Rubenach, M.J.,

Williams, P.J., and Baker, T., 2004, Modeling the role of sodic alteration in the genesis of

iron oxide-copper-gold deposits, Eastern Mount Isa block, Australia: Economic Geology,

v. 99, p. 1145-1176.

Oliveros, V., Féraud, G., Aguirre, L., Tristá, D., Morata, D., Fornari, M., Ramírez, L., Palacios,

C., Ferraris, F., 2006, Método 40Ar/39Ar en minerals de ganga: una aproximación a la

edad de la mineralización de cobre en los yacimientos Mantos Blancos y Michilla,

Cordillera de la Costa, norte de Chile [abs.]: 11o Congreso Geológico Chileno,

Antofagasta, Actas, v.2, p. 323-326.

O'Neil, J.R., Clayton, R.N., and Mayeda, T.K., 1969, Oxygen isotope fractionation in divalent

metal carbonates: Journal of Chemical Physics, v. 51, p. 5547-5558.

Oreskes, N., and Einaudi, M.T., 1992, Origin of hydrothermal fluids at Olympic Dam:

preliminary results from fluid inclusion and stable isotopes: Economic Geology, v. 87, p.

64-90.

216 Orrego, M., Urrutia, J., Sanhueza, A., Boric, R., and Cornejo, P., 2006, Mineralización tipo

OxFe-Cu-Au (IOCG) en el distrito Mantos Blancos (Cu-Ag), Antofagasta, norte de Chile

[abs.]: 11o Congreso Geológico Chileno, Antofagasta, Actas, v.2, p. 331-334.

Ortlieb, L. and Machare, J., 1990, Quaternary marine terraces on the Peruvian coast and Recent

vertical motions: in Laubacher, G., Olivier R. A., and Vatin P. N., eds., Symposium

international; Geodynamique andine; resumes des communications Editions de l’Office

de la Recherche Scientifique et Technique d’Outre-mer, Paris, p. 95–98.

Oxiana Resources, 2006, Mineral Resources Statement as at 30 June 2006: Oxiana Limited

Annual Report, p. 28-30.

Oyarzun, R., Oyarzún, J., Ménard, J.J., and Lillo, J., 2003, The Cretaceous iron belt of northern

Chile: role of oceanic plates, a superplume event, and a major shear zone: Mineralium

Deposita, v. 38, p. 640-646.

Pálfy, J., Smith, P.L., and Mortensen, J.K., 2000, A U–Pb and 40Ar/39Ar time scale for the

Jurassic: Canadian Journal of Earth Sciences, v. 37, p. 923-944.

Pan, Y., and Dong, P., 1999, The lower Changjiang (Yangzi/Yangtze River) metallogenic belt,

east central China: intrusion-and wall rock-hosted Cu-Fe-Au, Mo, Zn, Pb, Ag deposits:

Ore Geology Reviews, v. 15, p. 177-242.

Park, C.F., Jr., 1961, A magnetite “flow” in Northern Chile: Economic Geology, v. 56, p.

431-436.

Park, C.F., Jr., 1972, The iron ore deposits of the Pacific Basin: Economic Geology, v. 67, p.

339-349.

Perring, C.S., Pollard, P.J., Dong, G., Nunn, A.J., Blake, K.L., 2000, The Lightning Creek Sill

Complex, Cloncurry District, Northwestern Queensland: a source of fluids for the Fe

oxide Cu–Au mineralization and sodic–calcic alteration: Economic Geology, v. 95, p.

1037–1089.

217 Philpotts, A.R., 1967, Origin of certain iron-titanium oxide and apatite rocks: Economic Geology,

v. 62, p. 303-315.

Philpotts, A.R., 1982, Compositions of immiscible liquids in volcanic rocks: Contributions to

Mineralogy and Petrology, v. 80, p. 201–218.

Pitcher, W.S., Cobbing, E.J., 1985, A model for the Coastal batholith, in Pitcher, W.S., Atherton,

M.P., Cobbing, E.J., and Beckingsale, R.D., eds., Magmatism at a Plate Edge: the

Peruvian Andes: Glasgow, Blackie and Son Ltd., p. 19–25.

Pollard, P.J., 2000, Evidence of a magmatic fluid and metal source for Fe-oxide Cu-Au

mineralization, in Porter, T.M., ed., Hydrothermal Iron Oxide Copper-Gold and Related

Deposits: A Global Perspective: Adelaide, Porter Geoscience Consultancy Publishing, v.

1, p. 27-41.

Pollard, P.J., 2001, Sodic (-calcic) alteration associated with Fe-oxide-Cu-Au deposits: an origin

via unmixing of magmatic-derived H2O-CO2-salt fluids: Mineralium Deposita, v. 36, p.

93-100.

Pollard, P.J., 2006, An intrusion-related origin for Cu-Au mineralization in iron

oxide-copper-gold (IOCG) provinces: Mineralium Deposita, v. 41, p. 179-187.

Pollard, P.J., Taylor, R.G., and Peters, L., 2005, Ages of intrusion, alteration, and mineralization

at the Grasberg Cu-Au deposit, Papua, Indonesia: Economic Geology, v. 100, p.

1005–1020.

Polliand, M., Schaltegger, U., Frank, M., and Fontboté, L., 2005, Formation of intra-arc

volcano-sedimentary basins in the western flank of the central Peruvian Andes during

Late Cretaceous oblique subduction-field evidence and constraints from U-Pb ages and

Hf isotopes: International Journal of Earth Sciences, v. 94, p. 231–242.

Pop, N., Heaman, L., Edelstein, O., Isache, C., Zentilli, M., Pecskay, Z., Valdman, S., and Rusu,

C., 2000, Geocronología de las rocas ígneas y los productos de alteración hidrotermal

relacionados con la mineralización de Cu-Fe(Au) del Sector Adriana-Carola-Cobriza

218 (Parte Este del Distrito Punta del Cobre-Candelaria), en base a dataciones U-Pb (en

circón), 40Ar/39Ar y K-Ar: IX Congreso Geológico Chileno, Puerto Varas, Chile, Actas, p.

155–160.

Pope, A., 2003, Structural Controls on Copper Mineralization, Mina Justa Prospect and

Marcona District: Unpublished Report to Rio Tinto Mining and Exploration Limited, 16

Quang, C.X., 2003, An Expanded Geochronological and Geomorphological Framework for

Hypogene and Supergene Mineralization in the Southern Peru Porphyry Copper Belt:

Unpublished M.Sc. thesis, Kingston, Ontario, Queen’s University, 201 p.

Quang, C.X., Clark, A.H., and Lee, J.K.W., 2001, Cenozoic Landform Development and

Associated Supergene Processes in Southern Perú and Northernmost Chile: Unpublished

report to Rio Tinto Mining and Exploration Ltd., 59 p.

Ramírez, L.E., Palacios, C., Townley, B., Parada, M.A., Sial, A.N., Fernandez-Turiel, J.L.,

Gimeno, D., Garcia-Valles, M., and Lehmann, B., 2006, The Mantos Blancos copper

deposit: an upper Jurassic breccia-style hydrothermal system in the Coastal Range of

northern Chile: Mineralium Deposita, v. 41, p. 246-258.

Requia, K., and Fontboté, 1999, Hydrothermal alkali metasomatism in the Salobo iron oxide Cu

(-Au) deposit, Carajás mineral province, northern Brazil, in Stanley, C.J. et al., eds.,

Mineral Deposits: Processes to Processing: Rotterdam, Balkema, p. 1025-1028.

Requia, K., Stein, H., Fontboté, L., and Chiaradia, M., 2003, Re–Os and Pb–Pb geochronology of

the Archean Salobo iron oxide copper–gold deposit, Carajás mineral province, northern

Brazil: Mineralium Deposita, v. 38, p. 727-738.

Reynolds, I.M., 2002 a, Mineralogy and Petrography of a Third Suite of Oxide and Sulphide Ore

Samples from the Mina Justa Deposit, Marcona Area, Perú: Unpublished Report

BR3081, Rio Tinto Mining and Exploration Limited, 20 p.

219 Reynolds I.M., 2002 b, Descriptions of Individual Samples-Appendix: Unpublished Report

BR3046, Rio Tinto Mining and Exploration Limited, 50 p.

Rhodes, A.L., Oreskes, N., and Sheets, S., 1999, Geology and rare earth element geochemistry of

magnetic deposits at El Laco, Chile: Society of Economic Geologists Special Publication,

v. 7, p. 299-332.

Richards, J.P., and Spooner, E.T.C., 1989, Evidence for Cu-(Ag) mineralization by

magmatic-meteoric fluid mixing in Keweenawan fissure veins, Mamainse Point, Ontario:

Economic Geology, v. 84, p. 360-385.

Ripa, M., 1999, A review of the Fe oxide deposits of Bergslagen, Sweden and their connection to

Au mineralization, in Stanley, C.J. et al., eds., Mineral Deposits: Processes to Processing:

Rotterdam, Balkema, p. 1349-1352.

Ripley, E.M., Ohmoto, H., 1977, Mineralogic, sulphur isotope, and fluid inclusion studies of the

stratabound copper deposits at the Raúl Mine, Perú: Economic Geology, v. 72, p.

1017-1041.

Robert, D.E., and Hudson, G.R.T., 1983, The Olympic Dam copper-uranium-gold deposit, Roxby

Downs, South Australia: Economic Geology, v. 78, p. 799-822.

Roddy, M.S., Reynolds, S.J., Smith, B.M., and Ruiz, J., 1988, K-metasomatism and

detachment-related mineralization, Harcuvar Mountains, Arizona: Geological Society of

America Bulletin, v. 100, p. 1627-1639.

Roedder, E., 1951, Low-temperature liquid immiscibility in the system K2O-FeO-Al2O3-SiO2:

American Journal of Science, v. 35, p. 282–286.

Roedder, E., 1984, Fluid Inclusions: Mineralogical Society of America, Reviews in Mineralogy, v.

12, 646 p.

Romeuf, N., Aguirre, L., Carlier, G., Soler, P., Bonhomme, M., Elmi, S., and Salas, G., 1993,

Present knowledge of the Jurassic volcanogenic formations of the southern coastal Perú:

220 Second International Symposium Andean Geodynamics (ISAG), Oxford, 21-23

September, p. 437-440.

Roperch, P. and Carlier, G., 1992, Paleomagnetism of mesozoic rocks from the Central Andes of

southern Perú: Importance of rotations in the development of the Bolivian orocline:

Journal of Geophysical Research, v. 97 (B12), p. 17233-17249.

Rose, A.W., Herrick, D.C., Deines, P., 1985, An oxygen and sulfur isotope study of skarn-type

magnetite deposits of the Cornwall type, southeastern Pennsylvania: Economic Geology,

v. 80, p. 418-443.

Rotherham, J.F., Blake, K.L., Cartwright, I., and Williams, P.J., 1998, Stable isotope evidence for

the origin of the Mesoproterozoic Starra Au-Cu deposit, Cloncurry District, Northwest

Queensland: Economic Geology, v. 93, p. 1435-1449.

Rüegg, W., 1956, Geologie Zwischen Cañete-San Juan 13°00 °- 15°24'Sudperu: Geologische

Rundschau, v. 45, p. 775-858.

Rüegg, W., 1961, Hallazgo y posición estragráfico-tectónica del Titoniano en la costa sur del

Peru. Anales, II Congreso Nacional de Geologia (Sociedad Geológica del Peru), Lima,

Parte I, v. 36, p. 203-208.

Ruiz, C., Aguirre, L., Corvalán, J., Klohn, C., Klohn, E., Levi, B., 1965, Geología y Yacimientos

Metalíferos de Chile: Instituto de Investigaciones Geológicas, Santiago (Chile), 305 p.

Ruiz, F.C., and Peebles, L.F., 1988, Geología, Distribución y Génesis de los Yacimientos

Metalíferos Chilenos: Editorial Universitaria, Santiago, 334 p.

Ryan, P.J., Lawrence, A.L., Jenkins, R.A., Matthews, J.P., Zamora, J.C., Marino, E., and Urqueta,

I., 1995, The Candelaria copper-gold deposit, Chile, in Pierce, F.W. and Bolm, J.G., eds.,

Porphyry Copper Deposits of the American Cordillera: Arizona Geological Society

Digest, v. 20, p. 625–645.

221 Sánchez, A. W., 1982, Edades Rb-Sr en los segmentos Arequipa-Toquepala del batolito de la

costa del Perú: Quinto Congreso Latinoamericano de Geologia, Argentina, Actas, v. 111,

p. 487–504.

Seedorff, E., and Einaudi, M., 2004, Henderson porphyry molybdenum system, Colorado: II.

Decoupling of introduction and deposition of metals during geochemical evolution of

hydrothermal fluids: Economic Geology, v. 99, p. 39-72.

Sharma, T., Clayton, R.N., 1965, Measurement of 18O/16O ratios of total oxygen of carbonates:

Geochimica et Cosmochimica Acta, v. 29, p. 1347-1353.

Sherwood Copper Corp., 2007, Minto Project. http://www.sherwoodcopper.com/s/

MintoProject.asp

Shi, P., 1992, Fluid fugacities and phase equilibria in the Fe-Si-O-H-S system: American

Mineralogist, v. 77, p. 1050-1066.

Shougang Hierro Perú SA., 2003, Resource Estimate of the Marcona Iron Mine, 5 p. (in

Chinese).

Sillitoe, R.H., 2003, Iron oxide-copper-gold deposits: an Andean view: Mineralium Deposita, v.

38, p. 787-812.

Sillitoe, R.H., and Burrows, D.R., 2002, New field evidence bearing on the origin of the El Laco

magnetite deposit, northern Chile: Economic Geology, v. 97, p. 1101–1109.

Sillitoe, R.H., and Burrows, D.R., 2003, New field evidence bearing on the origin of the El Laco

magnetite deposit, northern Chile - a reply: Economic Geology, v. 98, p. 1501–1502.

Sillitoe, R., and Perelló, J., 2005, Andean copper province: tectonomagmatic settings, deposit

types, metallogeny, exploration, and discovery: Economic Geology 100th Anniversary

Volume, p. 845–890.

Simard, M., Beaudoin, G., Bernard, J., and Hupé, A., 2006, Metallogeny of the Mont-de-l’ Aigle

IOCG deposit, Gaspé Peninsula, Québec, Canada: Mineralium Deposita, v. 41, p.

607-636.

222 Skewes, A., Stern, C.R., 1994, Tectonic trigger for the formation of Late Miocene Cu-rich breccia

pipes in the Andes of central Chile: Geology, v. 22, p. 551-554.

Skirrow, R.G., and Walsh, J.L., 2002, Reduced and oxidized Au-Cu-Bi deposits of the Tennant

Creek inlier, Australia: An integrated geologic and chemical model: Economic Geology,

v. 97, p. 1167-1202.

Skirrow, R.G., Bastrakov, E.N., Barovich, K, Fraser, G.L, Creaser, R.A., Fanning, C.M.,

Raymond, O.L., and Davidson, G.J., 2007, Timing of iron oxide Cu-Au-(U)

hydrothermal activity and Nd isotope constraints on metal sources in the Gawler craton,

South Australia: Economic Geology, v. 102, p. 1441–1470.

Skirrow, R.G., Bastrakov, E., Davidson, G., Raymond, O.L., and Heithersay, P., 2002, The

Geological Framework, distribution and controls of Fe-oxide Cu-Au mineralization in the

Gawler Carton, South Australia: Part II – alteration and mineralization, in Porter, T.M.,

ed., Hydrothermal Iron Oxide Copper-Gold and Related Deposits: A Global Perspective:

Adelaide, Porter Geoscience Consultancy Publishing, v. 2, p. 33–48.

Smith, M.P., and Henderson, P., 2000, Preliminary fluid inclusion constraints on fluid evolution

in the Bayan Obo Fe-REE-Nb deposit, Inner Mongolia, China: Economic Geology, v. 95,

p. 1371-1388.

Smith, M. and Wu, C., 2000, The geology and genesis of the Bayan Obo Fe-REE-Nb deposit: A

review, in Porter, T.M., ed., Hydrothermal Iron Oxide Copper-Gold and Related

Deposits: A Global Perspective: Adelaide, Porter Geoscience Consultancy Publishing, v.

1, p. 271–283.

Smith, M., Coppard, J., Herrington, R., and Stein, H., 2007, The geology of the Rakkurijärvi

Cu-(Au) prospect, Norrbotten: A new iron oxide-copper-gold deposit in Northern

Sweden: Economic Geology, v. 102, p. 393-414.

223 Souza, L.J., Vieira, E.A., 2000, Salobo 3 Alpha deposit: geology and mineralization, in Porter,

T.M., ed., Hydrothermal Iron Oxide Copper-Gold and Related Deposits: A Global

Perspective: Adelaide, Porter Geoscience Consultancy Publishing, v. 1, p. 213–224.

Stumpfl, E.F., Clifford, T.N., Burger, A.J., and Zyl, D.V., 1976, The copper deposits of the

O’Okiep district, South Africa: new data and concepts: Mineralium Deposita, v. 11, p.

46-70.

Suzuoki, T., and Epstein, S., 1976, Hydrogen isotope fractionation between OH-bearing minerals

and water: Geochimica et Cosmochimica Acta, v. 40, p. 1229-1240.

Tafti, R. and Mortensen, J.K., 2004, Early Jurassic porphyry(?) copper (-gold) deposits at Minto

and Williams Creek, Carmacks Copper Belt, western Yukon, in Emond, D.S., and Lewis,

L.L., eds., Yukon Exploration and Geology 2003: Yukon Geological Survey, p. 289-303.

Tallarico, F.H.B., Figueiredo, B.R., Groves, D.I., Kositcin, N., McNaughton, N.J., Fletcher, I.R.,

Rego, J.L., 2005, Geology and SHRIMP U–Pb geochronology of the Igarapé Bahia

deposit, Carajás copper–gold belt, Brazil: an Archean (2.57 Ga) example of iron–oxide

Cu–Au–(U–REE) mineralization: Economic Geology, v. 100, p. 7–28.

Taylor, B.E., 1986, Magmatic volatiles: Isotopic variation of C, H and S: Reviews in Mineralogy,

v. 16, p. 185-271.

Taylor, B.E., 1987, Stable isotope geochemistry of ore-forming fluids, in Kyser, T.K., ed., Stable

Isotope Geochemistry of Low Temperature Fluids: Mineralogical Association of Canada,

Short Course Series 13. p. 337-445.

Taylor, H.P., 1997, Oxygen and hydrogen isotope relationships in hydrothermal mineral deposits,

in Barnes, H.L., ed., Geochemistry of Hydrothermal Ore Deposits, 3rd edn. : Wiley and

Sons, New York, p. 229-302.

Taylor, H.P., and Epstein, S., 1962, Relationship between 18O/16O ratios in coexisting minerals of

igneous and metamorphic rocks, part I: Geological Society America Bulletin, v. 73, p.

461-480.

224 Teck Cominco Limited, 2007, News Release April 24, 2007. http://www.teckcominco.com/

Generic.aspx?PAGE=Investors+Pages%2fQuarterly+Reports&portalName=tc

Thode, H.G., 1991, Sulphur isotopes in nature and the environment: An overview, in Krouse,

H.R., and Grinenko, V.A., eds., Stable Isotopes: Natural and Anthropogenic Sulphur in

the Environment, SCOPE 43: John Wiley, New York, p 1-26.

Thomas, H., 1958, Geología de la Cordillera de la Costa entre el valle de La Ligua y la Cuesta de

Barriga: Instituto de Investigaciones Geológicas Bolettín, v. 2, p. 1-86.

Torab, F.M., and Lehmann, B., 2006, Iron oxide-apatite deposits of the Bafq district, Central Iran:

an overview from geology to mining: World of Mining, v. 58, p. 392-399.

Tornos, F., and Casquet, C., 2005, A new scenario for related IOCG and Ni-(Cu) mineralization:

the relationship with giant mid-crustal mafic sills, Variscan Iberian Massif: Terra Nova, v.

17, no. 3, p. 236-241.

Tornos, F., Casquet, C., Pevida, L.R., and Velasco, F., 2005, The iron oxide – (Cu–Au) Deposits

of SW Iberia, Fregenal–Burguillos–Cala District: Lat. 38°18' N, Long. 6°40' W: Ore

Geology Reviews, v. 27, p. 166-167.

Ullrich, T.D., and Clark, A.H., 1997, Paragenetic Sequence of Mineralization in the Main

Orebody, Candelaria Copper-Gold Deposit, Chile: Unpublished Report to Phelps Dodge

Exploration Corp., 36 p.

Ullrich, T.D., and Clark, A.H., 1999, The Candelaria Cu-Au deposit, III Región, Chile:

paragenesis, geochronology and fluid composition, in Stanley, C.J. et al., eds., Mineral

Deposits: Processes to Processing: Rotterdam, Balkema, p. 201-204.

Ullrich, T.D., Clark, A.H., and Kyser, T.K., 2001, The Candelaria Cu-Au deposit, III Región,

Chile: product of long-term mixing of magmatic-hydrothermal and evaporite-sourced

fluids[abs.]: Geological Society of America Annual Meeting, Boston, MS, Abstracts with

Programs, v.33, no. 6, p. A-3.

225 Ulrich, T., Guenther, D., and Heinrich, C.A., 2001, The evolution of a porphyry Cu-Au deposit,

based on LA-ICP-MS analysis of fluid inclusions; Bajo de la Alumbrera, Argentina:

Economic Geology, v. 96, p. 1743-1774.

Ulrich, T., and Heinrich, C.A., 2002, Geology and alteration geochemistry of the porphyry Cu-Au

deposit at Bajo de la Alumbrera, Argentina: Economic Geology, v. 97, p. 1865–1888.

Valley, J.W., 2003, Oxygen isotopes in zircon, in Hanchar, J.M., and Hoskin, P.W.O., eds.,

Zircon: Reviews in Mineralogy and Geochemistry, v. 53, p. 343-385.

Vidal, C.E., Injoque-Espinoza, J.L., Sidder, G.B., and Mukasa, S.B., 1990, Amphibolitic Cu-Fe

skarn deposits in the central coast of Perú: Economic Geology, v. 85, p. 1447–1461.

Vielreicher, N.M., Groves, D.I., and Vielreicher, R.M., 2000, The Phalaborwa (Palabora) deposit

and its potential connection to iron-oxide copper-gold deposits of Olympic Dam type, in

Porter, T.M., ed., Hydrothermal Iron Oxide Copper-Gold and Related Deposits: A

Global Perspective: A global perspective: Adelaide, Porter Geoscience Consultancy

Publishing, v. 1, p. 321–331.

Vila, T., Lindsay, N., Zamora, R., 1996, Geology of the Mantoverde Copper Deposit, Northern

Chile: a specularite-rich hydrothermal-tectonic breccia related to the Atacama Fault Zone,

in Camus, F., Sillitoe, R.H., Petersen, R., eds., Andean Copper Deposits, New

Discoveries, Mineralization, Styles and Metallogeny: Society of Economic Geologists

Special Publication, v. 5, p. 157-170.

Vila, T., and Sillitoe, R.H., 1991, Gold-rich porphyry systems in the Maricunga Belt, northern

Chile: Economic Geology, v. 86, p. 1238-1260.

Visser, W., and Koster Van Groos, A.F., 1979, Effects of P2O5 and TiO2 on liquid-liquid

equilibria in the system K2O-FeO-Al2O3-SiO2: American Journal of Science, v. 279, p.

970–988.

226 Vivallo, W., Henríquez, F., 1997. Relación genética entre los yacimientos estratoligados de Cu

(“Tipo Manto”), de Cu-Fe±Au y de hierro tipo Kiruna [abs.]: 8o Congreso Geológico

Chileno, Antofagasta, Actas, v. 2, p. 1189-1193.

Wanhainen, C., 2005, On the Origin and Evolution of the Paleoproterozoic Aitik Cu-Au-Ag

Deposit, Northern Sweden: Unpublished Ph.D. thesis, Luleå University of Technology,

Sweden, 151 p.

Wanhainen, C., Billström, K., Martinsson, O., Stein, H., and Nordin, R., 2005, 160 Ma of

magmatic/hydrothermal and metamorphic activity in the Gällivare area: Re-Os dating of

molybdenite and U-Pb dating of titanite from the Aitik Cu-Au-Ag deposit, northern

Sweden: Mineralium Deposita, v. 40, p. 435-447.

Wanhainen, C., Broman, C., and Martinsson, O., 2003, The Aitik Cu-Au-Ag deposit in northern

Sweden: a product of high salinity fluids: Mineralium Deposita, v. 38, p. 715-726.

Wasteneys, H.A., Clark, A.H., Farrar, E., and Langridge, R.J., 1995, Grenvillian granulite facies

metamorphism in the Arequipa massif, Perú: a Laurentia-Gondwana link: Earth &

Planetary Science Letters, v. 132, p. 63-73.

White, W.S., 1968, The native copper deposits of northern Michigan, in Ridge, J. D., ed., Ore

Deposits of the United States 1933-1967 (Graton-Sales vol.): The American Institute of

Mining, Metallurgical, and Petroleum Engineers, New York, p. 303-325.

Williams, P.J., Barton, M.D., Johnson, D.A., Fontboté, L., Halter, A.D., Mark, G., Oliver, N.H.S.,

and Marschik, R., 2005, Iron-oxide copper-gold deposits: geology, space-time

distribution, and possible modes of origin: Economic Geology 100th Anniversary Volume,

p. 371-405.

Williams, P.J., Dong, G., Ryan, C.G., Pollard, P.J., Rotherham, J.F., Mernagh, T.P., and

Chapman, L.H., 2001, Geochemistry of hypersaline fluid inclusions from the Starra (Fe

oxide)-Au- Cu deposit, Cloncurry District, Queensland: Economic Geology, v. 96, p.

875–883.

227 Williams, P.J., and Pollard, P.J., 2003, Australian Proterozoic iron oxide-Cu-Au deposits: an

overview with new metallogenic and exploration data from the Cloncurry district,

northwest Queensland: Exploration and Mining Geology, v. 10, p. 191–213.

Wilson, N.S.F., Zentilli, M., Reynolds, P.H., and Boric, R., 2003, Age of mineralization by

basinal fluids at the El Soldado manto-type Cu deposit, Chile: 40Ar/39Ar geochronology

of K-feldspar: Chemical Geology, v. 197, p. 161-176.

Wones, D.R., Eugster, H.P., 1965, Stability of biotite: experiment, theory and application:

American Mineralogist, v. 50, p. 1228–1272.

Woodall, R., 1993, The multidisciplinary team approach to successful mineral exploration:

Society of Economic Geologists Newsletter, v. 14, p. 1-6.

Wright, T.L., 1968, X-ray and optical study of alkali feldspar: II. An X-ray method for

determining the composition and structural state from measurement of 2θ values for three

reflections: American Mineralogist, v. 53, p. 88-101.

Xu, G., and Pollard, P.J., 1999, Origin of CO2-rich fluid inclusions in synorogenic veins from the

Eastern Mount Isa Fold Belt, NW Queensland, and their implications for mineralization:

Mineralium Deposita, v. 34, p. 395-404.

Yuhara, M., Hirahara, Y., Nishi, N., and Kagami, H., 2005, Rb-Sr, Sm-Nd ages of the

Phalaborwa carbonatite complex, South Africa: Polar Geosciences, v. 18, p. 101-113.

Zamora, R., and Castillo, B., 2001, Mineralización de Fe-Cu-Au en el distrito Mantoverde,

Cordillera de la Costa, III Región de Atacama, Chile: 2o Congreso Internacional de

Prospectores y Exploradores, Lima, Conferencias, Instituto de Ingenieros de Minas del

Perú, Lima, Actas, 13 p.

Zentilli, M., 1974, Geological Evolution and Metallogenetic Relationships in the Andes of

Northern Chile Between 26° and 29° S: Unpublished Ph.D. thesis, Kingston, Ontario,

Queen’s University, 394 p.

228 Zhao, Z.F., Zheng, Y.F., 2003, Calculation of oxygen isotope fractionation in magmatic rocks:

Chemical Geology, v. 193, p. 59-80.

Zheng, Y.F., 1991, Calculation of oxygen isotope fractionation in metal oxides: Geochemica et

Cosmochimica Acta, v. 55, p. 2299-2307.

Zheng, Y.F., 1993a, Calculation of oxygen isotope fractionation in anhydrous silicate minerals:

Geochimica et Cosmochimica Acta, v. 57, p. 1079-1091.

Zheng, Y.F., 1993b, Calculation of oxygen isotope fractionation in hydroxyl-bearing silicates:

Earth & Planetary Science Letters, v. 120, p. 247-263.

Zheng, Y.F., and Simon, K., 1991, Oxygen isotope fractionation in hematite and magnetite: a

theoretical calculation and application to geothermometry of metamorphic

iron-formations: European Journal of Mineralogy, v. 3, p. 877-886.

229 Appendix A

Analytical Techniques

Electron microprobe techniques

Quantitative analyses were made on an automated 4 spectrometer Camebax MBX electron probe by the wavelength dispersive X-ray analysis method (WDX). Operating conditions were: 15kv accelerating potential and a beam current of 20 nano-amperes (nA). Peak counting times for each analyzed element were: 15-30 seconds or 40,000 accumulated counts. Peak counting time for fluorine was 60 seconds. Background measurements were made at 50% peak counting time on each side of the analyzed peak. Raw X-ray data were converted to elemental weight % by the

Cameca PAP matrix correction program. A suite of well characterized natural and synthetic minerals and compounds were used as calibration standards: Si (wollastonite), Al (spinel syn.),

Mg (olivine), Na (albite), K (orthoclase), Ca (wollastonite), Ti (MnTiO3 syn.), Mn (MnTiO3 syn.),

Fe (fayalite syn.), Cl (tugtupite), F (lithium fluoride syn.).

Digital BSE (back-scattered electron) images were collected at 512 × 512 resolution with a

Lamont 4-element solid state BSE detector and BSE Quad Summing Amplifier interfaced to a 4Pi

Analysis Inc. digital imaging system and EDX X-ray system and Power Macintosh computer.

Sulphides were analysed at 20 kv accelerating potential and a beam current of 20 nÅ. A defocused beam approximately 10 microns in diameter was used for Cu sulphides. The following standards and X-ray lines were used: Fe (pyrrhotite syn.), S (pyrrhotite syn.), Cu (chalcopyrite),

Ag (Ag metal syn.), As (NiAs)

230 Sulfur and oxygen isotope analytical techniques

Analyses were performed in the stable isotope laboratory at Queen’s University (QFIR) on specimens collected from outcrop, drill core, and the open pits. Mineral identification was through standard petrographic analysis of polished thin sections using both reflected- and transmitted-light microscopy and X-ray diffraction (XRD). The minerals for stable isotope analysis were extracted from a crushed and washed fraction of the sample and from drilling on the selected sector for sulphides and iron oxides. Sulphur was extracted online with continuous-flow technology: 0.2-0.3 mg sulphide samples were converted to SO2 in a Carlo Erba Element

Analyzer NCS 2500, with CuO as oxidant. δ34S values are reported relative to Cañón Diablo

Troilite (CDT) standard. The oxygen isotope analysis of silicates and iron oxides was carried out

18 using the BrF5 method of Clayton and Mayeda (1963). δ O values are reported relative to Vienna

Standard Mean Ocean Water (V-SMOW). All isotopes were measured using a Finnigan MAT

252 isotope-ratio mass spectrometer in QFIR and reported in units of per mil (‰). Replicate δ34S and δ18O analyses were reproducible to ±0.3 and ±0.2 per mil, respectively.

40Ar/ 39Ar techniques

Mineral separates were hand-picked, washed in acetone, dried, wrapped in aluminum foil and stacked in an irradiation capsule with similar-aged samples and neutron flux monitors (Fish

Canyon Tuff sanidine, 28.02 Ma, Renne et al., 1998). The samples were irradiated at the

McMaster Nuclear Reactor in Hamilton, Ontario, for 90 MWH, with a neutron flux of approximately 3x1016 neutrons/cm2. Analyses (n=57) of 19 neutron flux monitor positions produced errors of <0.5% in the J value. The samples were analyzed at the Noble Gas Laboratory,

Pacific Centre for Isotopic and Geochemical Research, University of British Columbia,

Vancouver, BC, Canada. The mineral separates were step-heated at incrementally higher powers in the defocused beam of a 10W CO2 laser (New Wave Research MIR10) until fused. The gas

231 evolved from each step was analyzed by a VG5400 mass spectrometer equipped with an ion-counting electron multiplier. All measurements were corrected for total system blank, mass spectrometer sensitivity, mass discrimination, radioactive decay during and subsequent to irradiation, as well as interfering Ar from atmospheric contamination and the irradiation of Ca, Cl and K.

The plateau and correlation ages were calculated using Isoplot ver.3.09 (Ludwig, 2003).

Errors are quoted at the 2-sigma (95% confidence) level and are propagated from all sources except mass spectrometer sensitivity and age of the flux monitor. The best statistically-justified plateau and plateau age were picked based on the following criteria: (1) Three or more contiguous steps comprising more than 50% of the 39Ar; (2) Probability of fit of the weighted mean age greater than 5%; (3) Slope of the error-weighted line through the plateau ages equals zero at 5% confidence; (4) Ages of the two outermost steps on a plateau are not significantly different from the weighted-mean plateau age (at 1.8ơ six or more steps only); (5) Outermost two steps on either side of a plateau must not have nonzero slopes with the same sign (at 1.8 ơ nine or more steps only).

Staining Procedure for feldspars (for slabbed samples, by R. Foster, July 1988 at Queen’s

University)

1) Immerse the clean, dry face of the rock slab into the hydrofluoric acid for 30 seconds.

2) Rinse the sample in water to remove the excess hydrofluoric acid.

3) Immerse the etched face of the sample in the sodium cobaltinitrite solution for 60

seconds. Gently agitate the solution throughout the immersion.

4) Rinse the sample with tap water to remove the excess sodium cobaltinitrite.

5) Using a wash bottle squirt a generous amount of barium chloride over the etched surface

of the sample. Rinse the sample with tap water.

232 6) Immerse the etched face of the sample in the amaranth red solution for three minutes.

Quickly wash away the excess amaranth red and dry the sample with compressed air.

233 Appendix B

Summarized 40Ar/39Ar Analytical Data for Hydrothermal Minerals

from Marcona and Mina Justa

Laser

Step power %40Ar Age (Ma) no. (W) 40Ar/39Ar 38Ar/39Ar 37Ar/39Ar 36Ar/39Ar Ca/K Cl/K atm f 39Ar 40Ar*/39ArK 2ơ errow MA5-9A Cummingtonite (±Greenalite±Chlorite)

J=0.010055 ± 0.000010 Volume 39ArK = 742.63 × 1E-13, Integrated age = 161.81 ± 1.29, Plateau age = 177.0 ± 1.5 Ma (61% of 39ArK)

1 2 78.577 ± 0.006 2.009 ± 0.015 0.223 ± 0.025 0.260 ± 0.019 0.847 0.449 96.86 0.99 2.407 ± 1.420 43.14 ± 25.15

2 2.3 18.800 ± 0.005 4.413 ± 0.010 0.227 ± 0.015 0.046 ± 0.021 0.891 1.012 71.66 7.77 5.257 ± 0.290 92.93 ± 5.00

3 2.5 21.195 ± 0.006 2.457 ± 0.010 0.125 ± 0.020 0.046 ± 0.019 0.487 0.561 62.9 11.83 7.795 ± 0.263 136.13 ± 4.42

4 2.6 21.212 ± 0.011 1.761 ± 0.014 0.083 ± 0.024 0.042 ± 0.021 0.323 0.401 57.91 8.24 8.838 ± 0.295 153.60 ± 4.91

5 2.7 19.733 ± 0.012 1.767 ± 0.016 0.081 ± 0.026 0.034 ± 0.027 0.313 0.402 49.61 8.18 9.842 ± 0.313 170.24 ± 5.17

6* 2.8 17.345 ± 0.011 1.723 ± 0.014 0.082 ± 0.027 0.025 ± 0.022 0.321 0.393 40.92 9.65 10.145 ± 0.211 175.25 ± 3.47

7* 2.9 15.172 ± 0.013 1.739 ± 0.014 0.080 ± 0.026 0.017 ± 0.022 0.31 0.397 31.39 10.46 10.303 ± 0.204 177.83 ± 3.36

8* 3 14.522 ± 0.013 1.714 ± 0.015 0.087 ± 0.024 0.014 ± 0.025 0.34 0.391 28.04 9.88 10.336 ± 0.183 178.38 ± 3.01

9* 3.1 13.585 ± 0.014 1.688 ± 0.016 0.108 ± 0.032 0.012 ± 0.024 0.422 0.385 23.4 8.4 10.274 ± 0.189 177.36 ± 3.11

10* 3.3 13.225 ± 0.014 1.682 ± 0.016 0.155 ± 0.024 0.011 ± 0.029 0.607 0.384 22.15 9.3 10.172 ± 0.182 175.69 ± 2.99

11* 3.5 13.012 ± 0.013 1.671 ± 0.017 0.194 ± 0.018 0.010 ± 0.035 0.757 0.381 20.17 7.54 10.241 ± 0.174 176.82 ± 2.86

12* 3.9 14.241 ± 0.010 1.660 ± 0.015 0.547 ± 0.017 0.014 ± 0.031 2.147 0.379 26.79 5.78 10.263 ± 0.170 177.18 ± 2.80

13 4.5 15.159 ± 0.025 1.619 ± 0.025 0.901 ± 0.028 0.021 ± 0.045 3.566 0.369 33.7 1.98 9.678 ± 0.386 167.53 ± 6.38

MA5-9A (rerun) cummingtonite (±Greenalite±Chlorite)

J=0.010216±0.000012 Volume 39ArK = 194.26 × 1E-13, Integrated age = 162.05 ± 2.89, plateau age: 175.2± 2.3 Ma (73.8% 39ArK) .

1 2 38.637 ± 0.020 3.084 ± 0.025 0.152 ± 0.093 0.123 ± 0.040 0.449 0.702 94.1 5.02 2.210 ± 1.271 40.27 ± 22.91

2 2.3 23.124 ± 0.010 4.032 ± 0.015 0.322 ± 0.027 0.052 ± 0.037 0.966 0.924 65.64 17.72 7.851 ± 0.552 139.19 ± 9.42

3* 2.5 23.910 ± 0.005 1.754 ± 0.014 0.089 ± 0.038 0.048 ± 0.025 0.264 0.399 58.25 15.7 9.863 ± 0.348 173.19 ± 5.82

4* 2.7 17.778 ± 0.005 1.708 ± 0.010 0.097 ± 0.045 0.026 ± 0.052 0.291 0.389 42 20.22 10.183 ± 0.402 178.55 ± 6.71

5* 2.9 13.726 ± 0.006 1.647 ± 0.013 0.101 ± 0.040 0.013 ± 0.039 0.303 0.376 26.6 23.11 9.938 ± 0.161 174.44 ± 2.69

6* 3.1 14.387 ± 0.011 1.639 ± 0.020 0.250 ± 0.043 0.015 ± 0.115 0.747 0.374 29.24 8.39 9.879 ± 0.535 173.46 ± 8.96

7* 3.4 14.950 ± 0.015 1.662 ± 0.026 0.740 ± 0.041 0.016 ± 0.077 2.219 0.379 28.66 6.4 10.285 ± 0.398 180.25 ± 6.64

8 4 15.185 ± 0.015 1.547 ± 0.029 1.835 ± 0.026 0.022 ± 0.113 5.515 0.353 37.52 3.44 8.900 ± 0.767 156.99 ± 12.96

MA5-2 Phlogopite (±Chlorite±Talc)

J=0.010048 ± 0.000010 Volume 39ArK = 1233.12 × 1E-13, Integrated age = 155.24 ± 0.53, Plateau age = 171.5 ± 1.1 Ma (35.5% of 39ArK)

1 2 146.177 ± 0.009 1.342 ± 0.020 0.057 ± 0.128 0.489 ± 0.022 0.107 0.284 97.37 0.2 3.775 ± 2.954 67.17 ± 51.60

2 2.2 16.663 ± 0.008 0.247 ± 0.018 0.038 ± 0.080 0.052 ± 0.033 0.1 0.051 88.08 0.94 1.917 ± 0.497 34.42 ± 8.84

234 Laser

Step power %40Ar Age (Ma) no. (W) 40Ar/39Ar 38Ar/39Ar 37Ar/39Ar 36Ar/39Ar Ca/K Cl/K atm f 39Ar 40Ar*/39ArK 2ơ errow

3 2.4 10.174 ± 0.006 0.178 ± 0.012 0.014 ± 0.101 0.028 ± 0.031 0.036 0.037 79.37 2.15 2.039 ± 0.255 36.59 ± 4.53

4 2.6 11.456 ± 0.013 0.218 ± 0.013 0.003 ± 0.123 0.020 ± 0.021 0.006 0.046 51.34 6.27 5.517 ± 0.166 97.33 ± 2.86

5 2.8 10.423 ± 0.006 0.244 ± 0.012 0.006 ± 0.079 0.010 ± 0.027 0.016 0.053 27.65 7.38 7.476 ± 0.093 130.66 ± 1.57

6 2.9 10.936 ± 0.010 0.251 ± 0.014 0.002 ± 0.108 0.008 ± 0.024 0.006 0.054 20.63 8.98 8.616 ± 0.113 149.79 ± 1.88

7 3 10.649 ± 0.007 0.253 ± 0.011 0.008 ± 0.041 0.005 ± 0.023 0.023 0.055 14.04 11.8 9.095 ± 0.075 157.76 ± 1.25

8 3.1 10.535 ± 0.005 0.272 ± 0.010 0.002 ± 0.083 0.003 ± 0.035 0.005 0.059 8.01 13.79 9.633 ± 0.058 166.68 ± 0.95

9 3.2 10.413 ± 0.007 0.271 ± 0.011 0.002 ± 0.126 0.002 ± 0.035 0.005 0.059 5.03 13.03 9.829 ± 0.076 169.92 ± 1.25

10* 3.3 10.385 ± 0.006 0.260 ± 0.010 0.003 ± 0.081 0.002 ± 0.038 0.009 0.057 4.14 13.51 9.896 ± 0.063 171.02 ± 1.04

11* 3.4 10.372 ± 0.007 0.258 ± 0.012 0.004 ± 0.066 0.001 ± 0.055 0.012 0.056 3.57 11.73 9.939 ± 0.072 171.73 ± 1.18

12* 3.5 10.394 ± 0.008 0.265 ± 0.014 0.004 ± 0.111 0.002 ± 0.042 0.011 0.058 3.45 9.84 9.968 ± 0.083 172.22 ± 1.37

13* 3.7 11.368 ± 0.009 0.292 ± 0.040 0.016 ± 0.185 0.010 ± 0.115 0.012 0.063 8.76 0.38 9.628 ± 0.369 166.60 ± 6.09

MA5-2 (rerun) Phlogopite (±Chlorite±Talc)

J=0.010048 ± 0.000010 Volume 39ArK = 693.11 × 1E-13, Integrated age = 164.39 ± 1.35, no plateau age, excess A r; inverse isochron age: 171.0 ± 1.9 Ma

1 2 389.567 ± 0.018 1.319 ± 0.056 0.161 ± 0.161 1.096±0.093 0.825 0.254 82.1 0.16 68.003±29.861 939.48±321.58

2 2.2 13.384 ± 0.010 0.177 ± 0.023 0.033 ± 0.092 0.035 ± 0.108 0.255 0.036 74.89 1.92 3.082 ± 1.131 55.03 ± 19.89

3 2.5 15.686 ± 0.006 0.178 ± 0.028 0.013 ± 0.127 0.041 ± 0.085 0.1 0.036 75.39 5.32 3.739 ± 1.013 66.54 ± 17.70

4 2.8 11.640 ± 0.005 0.229 ± 0.013 0.003 ± 0.147 0.012 ± 0.038 0.024 0.049 29.73 14.78 8.050 ± 0.140 140.33 ± 2.36

5 3 11.005 ± 0.005 0.270 ± 0.024 0.002 ± 0.118 0.003 ± 0.045 0.008 0.059 6.25 17.1 10.172 ± 0.063 175.56 ± 1.03

6 3.2 10.809 ± 0.004 0.281 ± 0.011 0.002 ± 0.131 0.001 ± 0.049 0.008 0.061 2.76 17.11 10.361 ± 0.048 178.67 ± 0.79

7 3.4 10.669 ± 0.004 0.285 ± 0.039 0.003 ± 0.105 0.001 ± 0.068 0.017 0.062 2.22 16 10.275 ± 0.053 177.26 ± 0.87

8 3.6 10.509 ± 0.004 0.277 ± 0.025 0.003 ± 0.120 0.001 ± 0.078 0.023 0.061 1.38 11.58 10.162 ± 0.054 175.41 ± 0.88

9 3.8 10.359 ± 0.004 0.272 ± 0.043 0.011 ± 0.090 0.001 ± 0.126 0.091 0.059 0.92 9.92 10.033 ± 0.062 173.28 ± 1.02

10 4.1 10.388 ± 0.010 0.257 ± 0.035 0.007 ± 0.129 0.001 ± 0.085 0.045 0.056 0.48 6.11 9.990 ± 0.106 172.58 ± 1.75

MA5-9B Biotite

J=0.010041 ± 0.000010 Volume 39ArK = 528.36 × 1E-13, Integrated age = 160.21 ± 0.47, Plateau age = 161.42 ± 0.89 Ma (80% of 39ArK)

1 2 26.321 ± 0.079 1.202 ± 0.079 0.011 ± 0.239 0.075 ± 0.082 0.004 0.27 79.56 1.07 5.217 ± 0.819 92.11 ± 14.11

2 2.2 13.117 ± 0.009 1.163 0.020 0.009 0.353 0.021 0.094 0 0.264 35.32 1.13 8.058 ± 0.580 140.37 ± 9.72

3 2.4 11.344 ± 0.006 0.907 0.012 0.003 0.361 0.009 0.079 0 0.205 19.2 3.53 8.972 ± 0.220 155.61 ± 3.66

4 2.6 10.496 ± 0.006 0.879 ±0.013 0.002 ± 0.352 0.004 ± 0.060 0.001 0.199 8.73 5.39 9.427 ± 0.098 163.16 ± 1.62

5* 2.8 10.021 ± 0.005 0.878 ± 0.010 0.001 ± 0.309 0.002 ± 0.044 0.001 0.199 5.79 13.16 9.354 ± 0.056 161.96 ± 0.93

6* 3 9.803 ± 0.006 0.886 ± 0.011 0.001 ± 0.288 0.002 ± 0.090 0 0.201 3.85 14.08 9.341 ± 0.072 161.74 ± 1.18

7* 3.2 9.677 ± 0.006 0.964 ± 0.010 0.001 ± 0.298 0.001 ± 0.056 0 0.219 3.2 15.98 9.288 ± 0.058 160.86 ± 0.96

8* 3.4 9.641 ± 0.004 1.048 ± 0.010 0.001 ± 0.261 0.001 ± 0.072 0 0.238 2.31 13.27 9.330 ± 0.051 161.55 ± 0.85

9* 3.6 9.622 ± 0.005 1.073 ± 0.011 0.001 ± 0.497 0.001 ± 0.075 0 0.244 2.21 12.74 9.318 ± 0.053 161.37 ± 0.88

10* 3.9 9.561 ± 0.004 1.057 ± 0.010 0.001 ± 0.216 0.001 ± 0.111 0 0.24 1.65 10.74 9.303 ± 0.057 161.10 ± 0.94

11* 4.2 9.571 ± 0.004 1.069 ± 0.010 0.001 ± 0.221 0.002 ± 0.120 0 0.243 2.04 6.89 9.241 ± 0.072 160.08 ± 1.19

12* 4.5 9.840 ± 0.007 1.050 ± 0.014 0.004 ± 0.454 0.004 ± 0.108 0 0.238 2.89 2.01 9.187 ± 0.149 159.19 ± 2.48

235 Laser

Step power %40Ar Age (Ma) no. (W) 40Ar/39Ar 38Ar/39Ar 37Ar/39Ar 36Ar/39Ar Ca/K Cl/K atm f 39Ar 40Ar*/39ArK 2ơ errow

MA3-30 Microcline

J=0.009489 ± 0.000010 Volume 39ArK = 595.95 × 1E-13, Integrated age = 107.83 ± 0.47, Plateau age = 109.18 ± 0.64 Ma (76% of 39ArK)

1 2 56.575 ± 0.007 0.454 ± 0.016 0.0058 ± 0.1511 0.174 ± 0.018 0.012 0.094 90.83 1.76 5.152 ± 0.869 86.11 ± 14.18

2 2.2 8.538 ± 0.004 0.049 ± 0.016 0.004 ± 0.088 0.008 ± 0.029 0.011 0.008 28.1 8.32 6.066 ± 0.078 100.97 ± 1.26

3 2.4 6.941 ± 0.005 0.024 ± 0.047 0.004 ± 0.117 0.002 ± 0.031 0.011 0.002 8.71 6.75 6.239 ± 0.040 103.77 ± 0.65

4 2.6 6.902 ± 0.006 0.021 ± 0.042 0.005 ± 0.074 0.001 ± 0.082 0.012 0.002 4.02 5.74 6.513 ± 0.051 108.19 ± 0.82

5* 2.9 7.170 ± 0.005 0.024 ± 0.021 0.006 ± 0.059 0.002 ± 0.063 0.016 0.002 6.97 9.61 6.591 ± 0.048 109.44 ± 0.77

6* 3.2 7.482 ± 0.005 0.031 ± 0.021 0.006 ± 0.025 0.003 ± 0.037 0.017 0.004 11.74 14.78 6.541 ± 0.046 108.64 ± 0.75

7* 3.5 7.782 ± 0.005 0.042 ± 0.021 0.009 ± 0.052 0.004 ± 0.039 0.024 0.006 14.43 9.67 6.585 ± 0.056 109.34 ± 0.91

8* 3.8 7.751 ± 0.004 0.040 ± 0.013 0.008 ± 0.025 0.004 ± 0.022 0.021 0.006 14.78 28.12 6.557 ± 0.038 108.89 ± 0.61

9* 4 7.539 ± 0.005 0.042 ± 0.032 0.009 ± 0.087 0.003 ± 0.062 0.021 0.006 10.66 5.99 6.633 ± 0.065 110.12 ± 1.04

10* 4.3 7.483 ± 0.018 0.049 ± 0.027 0.006 ± 0.103 0.003 ± 0.057 0.016 0.008 10.05 7.72 6.643 ± 0.136 110.29 ± 2.19

11 4.8 15.468 ± 0.008 0.179 ± 0.030 0.012 ± 0.079 0.029 ± 0.046 0.028 0.037 52.97 1.53 7.121 ± 0.395 117.96 ± 6.34

MA3-24 Microcline

J=0.009481 ± 0.000010 Volume 39ArK = 1747.12 × 1E-13, Integrated age = 100.61 ± 0.34, Plateau age = 101.04 ± 0.56 Ma (61% of 39ArK)

1 2 263.157 ± 0.0070 1.731 ± 0.017 0.008 ± 0.448 0.863 ± 0.019 0.01 0.363 96.97 0.2 8.031 ± 4.721 132.38 ± 75.04

2 2.2 12.447 ± 0.005 0.110 ± 0.020 0.005 ± 0.094 0.021 ± 0.019 0.012 0.021 49.31 2.81 6.249 ± 0.122 103.85 ± 1.97

3 2.4 8.744 ± 0.005 0.061 ± 0.031 0.004 ± 0.082 0.009 ± 0.035 0.011 0.011 30.2 2.06 6.018 ± 0.099 100.11 ± 1.61

4 2.6 8.496 ± 0.004 0.058 ± 0.019 0.005 ± 0.103 0.009 ± 0.023 0.012 0.01 28.69 2.85 5.985 ± 0.063 99.58 ± 1.02

5 2.8 7.737 ± 0.004 0.045 ± 0.019 0.005 ± 0.042 0.006 ± 0.029 0.013 0.007 22.32 3.78 5.944 ± 0.059 98.90 ± 0.96

6 3 6.788 ± 0.004 0.031 ± 0.022 0.003 ± 0.056 0.003 ± 0.035 0.007 0.004 13.35 9.17 5.833 ± 0.041 97.10 ± 0.67

7 3.2 6.788 ± 0.004 0.034 ± 0.016 0.002 ± 0.062 0.002 ± 0.022 0.005 0.005 10.09 14.79 6.059 ± 0.031 100.77 ± 0.50

8 3.3 6.885 ± 0.005 0.036 ± 0.025 0.003 ± 0.112 0.003 ± 0.059 0.008 0.005 11.78 3.09 5.995 ± 0.060 99.73 ± 0.97

9* 3.4 7.022 ± 0.006 0.038 ± 0.028 0.004 ± 0.131 0.003 ± 0.025 0.01 0.005 12.64 2.96 6.053 ± 0.047 100.68 ± 0.75

10* 3.5 7.015 ± 0.004 0.040 ± 0.014 0.007 ± 0.023 0.003 ± 0.020 0.017 0.006 13.09 16.81 6.053 ± 0.032 100.67 ± 0.52

11* 3.7 6.836 ± 0.004 0.037 ± 0.011 0.004 ± 0.048 0.002 ± 0.018 0.01 0.005 10.21 13.19 6.093 ± 0.028 101.32 ± 0.45

12* 3.9 6.940 ± 0.005 0.036 ± 0.014 0.004 ± 0.148 0.003 ± 0.036 0.011 0.005 11.37 4.66 6.086 ± 0.043 101.21 ± 0.70

13* 4 7.043 ± 0.005 0.034 ± 0.026 0.006 ± 0.047 0.003 ± 0.023 0.014 0.004 12.84 4.94 6.076 ± 0.037 101.05 ± 0.60

14* 4.2 6.776 ± 0.007 0.036 ± 0.016 0.008 ± 0.031 0.002 ± 0.022 0.02 0.005 9.5 18.68 6.089 ± 0.050 101.25 ± 0.81

MA3-19 Phlogopite

J=0.010033 ± 0.000010 Volume 39ArK = 1562.11 × 1E-13, Integrated age = 158.82 ± 0.30, Plateau age = 159.69 ± 0.84 Ma (97.8% 39ArK)

1 2 44.684 ± 0.010 0.159 ± 0.127 0.029 ± 0.792 0.177 ± 0.037 0 0.022 98.15 0.06 0.728 ± 1.997 13.13 ± 35.88

2 2.3 4.644 ± 0.009 0.136 ± 0.038 0.015 ± 0.340 0.020 ± 0.076 0.024 0.027 85.16 0.3 0.490 ± 0.448 8.86 ± 8.07

3 2.6 10.571 ± 0.007 0.138 ± 0.062 0.009 ± 0.336 0.028 ± 0.055 0.012 0.027 62.44 0.42 3.656 ± 0.448 64.99 ± 7.83

4 2.9 12.018 ± 0.005 0.150 ± 0.019 0.003 ± 0.295 0.012 ± 0.051 0.005 0.031 23.79 1.4 8.950 ± 0.181 155.14 ± 3.00

5* 3.2 9.618 ± 0.004 0.128 ± 0.012 0.001 ± 0.200 0.002 ± 0.039 0.001 0.026 3.34 5.42 9.202 ± 0.045 159.31 ± 0.74

6* 3.4 9.421 ± 0.004 0.131 ± 0.013 0.000 ± 0.409 0.001 ± 0.097 0.001 0.027 1.04 9.46 9.252 ± 0.045 160.15 ± 0.74

236 Laser

Step power %40Ar Age (Ma) no. (W) 40Ar/39Ar 38Ar/39Ar 37Ar/39Ar 36Ar/39Ar Ca/K Cl/K atm f 39Ar 40Ar*/39ArK 2ơ errow

7* 3.6 9.519 ± 0.004 0.144 ± 0.016 0.001 ± 0.359 0.001 ± 0.088 0.001 0.03 0.86 5.85 9.344 ± 0.048 161.67 ± 0.79

8* 3.8 9.359 ± 0.004 0.142 ± 0.011 0.000 ± 0.347 0.000 ± 0.114 0 0.029 0.44 9.72 9.247 ± 0.042 160.06 ± 0.69

9* 4 9.273 ± 0.005 0.137 ± 0.011 0.000 ± 0.286 0.000 ± 0.169 0.001 0.028 0.03 12.67 9.211 ± 0.051 159.46 ± 0.85

10* 4.1 9.321 ± 0.004 0.140 ± 0.013 0.001 ± 0.352 0.000 ± 0.210 0.001 0.029 0.01 8.7 9.248 ± 0.047 160.07 ± 0.78

11* 4.2 9.317 ± 0.004 0.137 ± 0.013 0.001 ± 0.323 0.000 ± 0.105 0.001 0.028 0.27 7.62 9.213 ± 0.043 159.49 ± 0.72

12* 4.3 9.320 ± 0.004 0.139 ± 0.012 0.001 ± 0.391 0.000 ± 0.151 0.001 0.029 0.03 7.3 9.236 ± 0.043 159.87 ± 0.72

13* 4.4 9.392 ± 0.004 0.138 ± 0.019 0.001 ± 0.488 0.001 ± 0.136 0.001 0.029 0.35 4.79 9.255 ± 0.050 160.18 ± 0.82

14* 4.6 9.324 ± 0.004 0.139 ± 0.009 0.000 ± 0.442 0.001 ± 0.177 0 0.029 0.23 6.8 9.217 ± 0.048 159.57 ± 0.79

15* 4.8 9.338 ± 0.004 0.144 ± 0.013 0.001 ± 0.315 0.001 ± 0.148 0.001 0.03 0.2 6.27 9.230 ± 0.047 159.78 ± 0.77

16* 5.1 9.332 ± 0.004 0.143 ± 0.014 0.001 ± 0.197 0.001 ± 0.125 0.001 0.03 0.51 5.99 9.193 ± 0.045 159.16 ± 0.74

17* 5.5 9.283 ± 0.005 0.139 ± 0.015 0.001 ± 0.468 0.001 ± 0.290 0.001 0.029 0.33 7.24 9.170 ± 0.062 158.78 ± 1.04

DDM3-3-8 Tremolite

J=0.010099 ± 0.000014 Volume 39ArK = 105.78 × 1E-13, Integrated age = 159.69 ± 2.31, Plateau age = 158.5 ± 1.9 Ma (98.7% 39ArK)

1 2 1043.520 ± 0.077 2.897 ± 0.095 28.382 ± 0.082 3.819 ± 0.086 125.44 0.525 100.74 0.07 (-)8.08±50.71 (-)153.5±1006.0

2 2.2 336.853 ± 0.041 1.777 ± 0.055 35.883 ± 0.046 1.209 ± 0.052 153.92 0.371 97.1 0.25 9.68 ± 12.81 168.2 ± 212.6

3 2.5 181.775 ± 0.018 1.873 ± 0.037 43.493 ± 0.023 0.616 ± 0.035 186.54 0.429 90.78 0.69 17.04 ± 6.03 286.5 ± 93.7

4 2.8 202.106 ± 0.033 1.600 ± 0.069 32.610 ± 0.037 0.694 ± 0.052 138.44 0.35 88.63 0.33 21.99 ± 8.95 361.8 ± 133.5

5* 3.1 173.848 ± 0.029 0.963 ± 0.065 28.775 ± 0.034 0.618 ± 0.042 120.64 0.197 95.66 0.56 7.36 ± 6.00 129.3 ± 101.7

6* 3.4 35.624 ± 0.010 1.071 ± 0.027 23.927 ± 0.017 0.117 ± 0.047 98.962 0.246 72.01 1.78 9.15 ± 1.67 159.5 ± 27.8

7* 3.8 11.204 ± 0.010 1.160 ± 0.013 18.956 ± 0.016 0.020 ± 0.027 77.626 0.271 17.8 19.8 9.13 ± 0.19 159.2 ± 3.2

8* 4.2 10.335 ± 0.016 1.210 ± 0.019 17.655 ± 0.021 0.016 ± 0.034 71.365 0.282 12.1 32.38 9.09 ± 0.22 158.5 ± 3.7

9* 4.6 10.590 ± 0.006 1.147 ± 0.011 18.202 ± 0.014 0.018 ± 0.036 73.642 0.267 13.16 21 9.11 ± 0.20 158.8 ± 3.4

10* 5 10.218 ± 0.008 1.077 ± 0.012 17.968 ± 0.014 0.017 ± 0.045 72.673 0.251 10.42 18.5 9.01 ± 0.24 157.1 ± 4.0

11* 5.5 11.440 ± 0.007 0.947 ± 0.014 20.084 ± 0.014 0.025 ± 0.091 81.553 0.22 10 4.63 9.12 ± 0.70 158.9 ± 11.7

DDM5-4-2 Actinolite

J=0.010061 ± 0.000010 Volume 39ArK = 62.61 × 1E-13, Integrated age = 148.75 ± 4.94, Plateau age = 156.6 ± 4.2 Ma (80.5% 39ArK)

1 2 1385.898 ± 0.073 23.179 ± 0.079 37.261 ± 0.076 4.913 ± 0.077 154.07 5.457 101 0.15 (-)14.170±41.196 (-)277.49±872.06

2 2.3 344.175 ± 0.031 13.908 ± 0.036 37.338 ± 0.036 1.171 ± 0.041 155.42 3.344 95.67 0.72 15.217 ± 9.996 256.99 ± 157.35

3 2.6 200.408 ± 0.020 8.638 ± 0.023 43.539 ± 0.024 0.710 ± 0.036 183.24 2.096 98.35 1.43 3.407 ± 6.785 60.80 ± 119.08

4 2.9 175.983 ± 0.032 4.719 ± 0.050 45.280 ± 0.035 0.723 ± 0.068 190.7 1.137 105.08 0.33 (-)8.228±13.883 (-)155.89±274.74

5 3.2 78.665 ± 0.008 1.888 ± 0.019 42.979 ± 0.014 0.275 ± 0.029 180.81 0.449 91.5 4.43 6.941 ± 2.442 121.78 41.43

6 3.5 37.700 ± 0.008 1.239 ± 0.024 42.853 ± 0.015 0.138 ± 0.043 180.25 0.295 83.16 4.53 6.417 ± 1.851 112.87 31.55

7 3.8 18.286 ± 0.010 1.195 ± 0.016 43.835 ± 0.015 0.068 ± 0.031 184.69 0.288 59.74 7.92 7.386 ± 0.655 129.32 11.06

8* 4.1 14.025 ± 0.018 1.183 ± 0.020 39.254 ± 0.021 0.044 ± 0.039 164.18 0.284 38.93 21.73 8.816 ± 0.558 153.31 9.30

9* 4.4 12.966 ± 0.011 1.107 ± 0.014 36.117 ± 0.016 0.037 ± 0.031 150.3 0.264 32.16 36.4 9.101 ± 0.371 158.06 6.17

10* 4.8 12.214 ± 0.011 1.020 ± 0.018 35.369 ± 0.018 0.035 ± 0.037 147 0.243 27.57 22.36 9.019 ± 0.416 156.69 6.91

237 Laser

Step power %40Ar Age (Ma) no. (W) 40Ar/39Ar 38Ar/39Ar 37Ar/39Ar 36Ar/39Ar Ca/K Cl/K atm f 39Ar 40Ar*/39ArK 2ơ errow

DDM3-3-1 Tremolite

J=0.010086 ± 0.000012 Volume 39ArK = 118.95 × 1E-13, Integrated age = 155.26 ± 3.52, Plateau age = 156.2 ± 2.4 Ma (100% 39ArK)

1* 2 749.766 ± 0.030 1.536 ± 0.080 27.772 ± 0.033 2.587 ± 0.035 119.93 0.262 99.51 0.35 3.839 ± 15.797 68.54 ± 276.72

2* 2.2 265.227 ± 0.013 0.924 ± 0.046 30.922 ± 0.019 0.918 ± 0.029 131.68 0.182 98.37 1.01 4.490 ± 7.649 79.91 ± 133.15

3* 2.5 143.263 ± 0.008 0.949 ± 0.028 33.388 ± 0.015 0.496 ± 0.036 142.48 0.207 95.12 1.26 7.220 ± 5.445 126.81 ± 92.35

4* 2.8 106.620 ± 0.025 1.044 ± 0.045 25.692 ± 0.031 0.367 ± 0.047 108.85 0.236 89.87 0.72 10.723 ± 4.665 185.27 ± 76.60

5* 3.2 30.139 ± 0.007 1.212 ± 0.013 25.775 ± 0.014 0.095 ± 0.052 108.07 0.284 70.58 2.97 8.750 ± 1.535 152.58 ± 25.67

6* 3.6 13.113 ± 0.014 1.228 ± 0.020 25.419 ± 0.020 0.033 ± 0.039 106.24 0.29 32.72 12.01 8.880 ± 0.397 154.76 ± 6.64

7* 4 12.388 ± 0.018 1.684 ± 0.020 26.398 ± 0.022 0.029 ± 0.029 110.44 0.4 28.19 38.93 9.149 ± 0.285 159.23 ± 4.75

8* 4.5 14.719 ± 0.018 1.555 ± 0.020 27.735 ± 0.022 0.039 ± 0.030 117.42 0.369 41.16 24 8.898 ± 0.405 155.05 ± 6.76

9* 5 10.996 ± 0.013 1.458 ± 0.016 28.092 ± 0.018 0.028 ± 0.030 119.08 0.347 21.19 12.7 8.728 ± 0.270 152.21 ± 4.52

10* 4.5 11.177 ± 0.006 1.439 ± 0.015 28.240 ± 0.014 0.029 ± 0.028 119.89 0.343 16.13 6.06 9.105 ± 0.258 158.51 ± 4.30

MA91-2 Actinolite

J=0.009466 ± 0.000010 Volume 39ArK = 63.5 × 1E-13, Integrated age = 157.9 ± 3.7, Plateau age = 157.3 ± 3.2 Ma (99.9% 39ArK)

1 2 1088.341 ± 0.097 61.887 ± 0.098 44.783 ± 0.320 3.682 ± 0.101 58.271 16.38 97.36 0.11 30.456 ± 39.870 457.02 ± 528.51

2* 2.4 295.871 ± 0.015 33.702 ± 0.019 25.527 ± 0.071 0.974 ± 0.026 63.461 8.033 95.15 0.98 14.259 ± 6.520 228.42 ± 98.11

3* 2.8 58.113 ± 0.021 6.956 ± 0.027 18.476 ± 0.074 0.183 ± 0.051 46.781 1.633 84.44 1.7 8.085 ± 2.596 133.05 ± 41.18

4* 3.2 18.795 ± 0.007 2.954 ± 0.011 23.702 ± 0.016 0.042 ± 0.033 66.283 0.693 45.12 9.1 9.781 ± 0.429 159.75 ± 6.71

5* 3.6 19.507 ± 0.016 2.345 ± 0.014 23.859 ± 0.018 0.045 ± 0.026 67.264 0.549 50.83 19.33 9.483 ± 0.369 155.09 ± 5.78

6* 4 17.905 ± 0.012 1.964 ± 0.015 22.469 ± 0.018 0.039 ± 0.027 63.491 0.459 47.22 46.23 9.499 ± 0.342 155.34 ± 5.36

7* 4.3 26.706 ± 0.016 1.529 ± 0.017 22.204 ± 0.025 0.069 ± 0.043 61.697 0.354 61.6 7.41 9.816 ± 0.927 160.29 ± 14.49

8* 4.6 29.200 ± 0.018 1.492 ± 0.031 23.028 ± 0.028 0.078 ± 0.053 63.638 0.346 64.52 5.39 9.760 ± 1.256 159.43 ± 19.64

9* 4.9 24.309 ± 0.014 1.587 ± 0.029 23.106 ± 0.024 0.060 ± 0.051 64.368 0.369 54.44 5.08 10.193 ± 0.964 166.18 ± 15.01

10* 5.5 23.621 ± 0.017 1.758 ± 0.024 23.213 ± 0.031 0.059 ± 0.054 64.538 0.41 53.41 4.66 9.991 ± 0.995 163.03 ± 15.53

MA91-2 (rerun) Actinolite

J=0.009466 ± 0.000010 Volume 39ArK = 81.14 × 1E-13, Integrated age = 155.61 ± 3.64, Plateau age = 156.8 ± 2.9 Ma (99.9% 39ArK)

1 2 952.2868±0.0748 53.7575±0.0768 34.2192±0.4249 3.2359±0.0855 101.27 12.95 99.52 0.12 4.554±42.299 76.14±692.50

2* 2.4 579.9541 0.0170 53.7268 0.0201 28.5100 0.0671 1.9648 0.0245 83.171 12.74 99.41 0.54 3.476 10.796 58.40 178.50

3* 2.8 245.0068 0.0340 21.9726 0.0392 31.7853 0.0639 0.8098 0.0441 93.087 5.226 95.45 0.54 11.073 7.067 179.83 109.24

4* 3.2 37.4338 0.0142 3.7780 0.0180 26.3103 0.0241 0.1089 0.0512 76.448 0.889 75.38 3.13 9.002 1.647 147.54 25.91

5* 3.6 23.7048 0.0065 2.7871 0.0130 25.1387 0.0145 0.0593 0.0229 72.933 0.654 59.36 10.9 9.654 0.409 157.76 6.41

6* 4 23.7730 0.0097 2.4326 0.0137 23.4471 0.0160 0.0593 0.0237 68.475 0.569 60.73 31.82 9.473 0.428 154.93 6.70

7* 4.2 28.6572 0.0157 1.7374 0.0216 22.8505 0.0223 0.0758 0.0232 66.71 0.404 67.42 14.89 9.413 0.553 153.99 8.68

8* 4.4 30.3480 0.0143 1.7925 0.0210 22.9767 0.0215 0.0806 0.0273 67.108 0.417 68.04 9.67 9.724 0.699 158.86 10.93

9* 4.7 22.7067 0.0159 1.7509 0.0211 22.6680 0.0205 0.0551 0.0335 66.204 0.408 57.5 8.85 9.585 0.578 156.68 9.05

10* 5.1 21.1838 0.0138 2.1476 0.0205 22.5029 0.0180 0.0500 0.0288 65.729 0.502 54.14 7.88 9.591 0.469 156.78 7.34

11* 6 27.1146 0.0092 1.8748 0.0150 22.5834 0.0158 0.0694 0.0201 65.993 0.437 64.15 11.67 9.754 0.415 159.33 6.49

238 Laser

Step power %40Ar Age (Ma) no. (W) 40Ar/39Ar 38Ar/39Ar 37Ar/39Ar 36Ar/39Ar Ca/K Cl/K atm f 39Ar 40Ar*/39ArK 2ơ errow

MA45-2 Actinolite

J=0.009458 ± 0.000010 Volume 39ArK = 58.31 × 1E-13, Integrated age = 156.0 ± 5.7, Plateau age = 157.3 ± 3.5 Ma (95.2% 39ArK)

1 2 4638.381 ± 0.179 59.753 ± 0.181 33.906 ± 0.803 15.584 ± 0.181 0 16.22 98.59 0.07 78.39 ± 153.21 1000.8 ± 1501.2

2 2.4 1293.244 ± 0.019 38.443 ± 0.022 23.863 ± 0.116 4.362 ± 0.026 55.506 9.038 99.35 0.8 8.623 ± 24.394 141.44 ± 384.85

3 2.8 374.194 ± 0.036 12.942 ± 0.040 39.330 ± 0.080 1.252 ± 0.044 103.37 3.094 96.76 0.91 12.307 ± 10.027 198.65 ± 153.26

4 3.2 80.944 ± 0.010 6.144 ± 0.021 37.103 ± 0.026 0.266 ± 0.026 103.77 1.464 89.71 3.03 8.195 ± 1.969 134.68 ± 31.17

5* 3.6 29.092 ± 0.016 5.967 ± 0.017 28.949 ± 0.021 0.081 ± 0.031 82.048 1.412 68.69 16.34 9.100 ± 0.761 148.95 ± 11.95

6* 3.9 23.463 ± 0.007 6.038 ± 0.011 33.096 ± 0.014 0.063 ± 0.024 94.559 1.436 59.81 26.95 9.530 ± 0.467 155.70 ± 7.32

7* 4.2 19.675 ± 0.019 5.874 ± 0.022 31.697 ± 0.025 0.049 ± 0.034 90.408 1.396 52.16 27.44 9.462 ± 0.544 154.64 ± 8.53

8* 4.5 17.331 ± 0.009 5.719 ± 0.012 34.224 ± 0.015 0.043 ± 0.040 97.089 1.364 40.17 8.06 9.689 ± 0.539 158.18 ± 8.42

9* 4.9 17.567 ± 0.011 6.006 ± 0.015 35.349 ± 0.018 0.045 ± 0.034 100.36 1.435 40.87 7.46 9.649 ± 0.474 157.56 ± 7.42

10* 5.5 20.069 ± 0.011 6.184 ± 0.015 38.425 ± 0.017 0.053 ± 0.039 109.21 1.482 44.71 5.72 10.182 ± 0.660 165.88 ± 10.27

11* 6.2 23.077 ± 0.014 6.389 ± 0.018 39.985 ± 0.021 0.065 ± 0.056 112.65 1.536 48.21 3.22 10.201 ± 1.129 166.17 ± 17.56

MA45-2 (rerun) Actinolite

J=0.009458 ± 0.000010 Volume 39ArK = 35.92 × 1E-13, Integrated age = 154.06 ± 8.5, Plateau age = 154.3 ± 5.5 Ma (99.8% 39ArK)

1 2 2427.664 ± 0.104 25.051 ± 0.105 12.380 ± 1.202 8.236 ± 0.107 35.213 5.637 100.09 0.18 -2.231±63.468 -38.47±1106.30

2* 2.4 1363.377 ± 0.031 29.045 ± 0.034 29.532 ± 0.111 4.540 ± 0.0364 86.309 6.745 98.22 0.92 24.959 ± 28.069 382.34 ± 387.47

3* 2.8 355.132 ± 0.037 9.905 ± 0.043 42.534 ± 0.082 1.187 ± 0.045 126.15 2.346 97.05 1.26 10.665 ± 9.610 173.37 ± 148.95

4* 3.2 59.114 ± 0.012 6.520 ± 0.017 32.372 ± 0.021 0.188 ± 0.035 94.7 1.546 86.66 10.72 7.975 ± 1.930 131.19 ± 30.63

5* 3.6 35.535 ± 0.014 6.574 ± 0.016 34.046 ± 0.021 0.105 ± 0.021 99.769 1.565 74.37 32.63 9.317 ± 0.712 152.36 ± 11.17

6* 4 34.623 ± 0.023 6.805 ± 0.027 39.019 ± 0.027 0.104 ± 0.029 114.99 1.629 73.68 27.29 9.353 ± 0.879 152.92 ± 13.79

7* 4.4 23.914 ± 0.010 6.277 ± 0.018 36.276 ± 0.017 0.065 ± 0.041 106.59 1.5 58.84 13.34 9.807 ± 0.811 160.03 ± 12.66

8* 4.8 23.133 ± 0.020 6.382 ± 0.020 38.964 ± 0.022 0.066 ± 0.033 114.86 1.53 57.8 7.02 9.370 ± 0.660 153.19 ± 10.34

9* 5.5 33.817 ± 0.015 6.988 ± 0.019 41.699 ± 0.021 0.101 ± 0.036 123.31 1.68 70.5 6.64 9.833 ± 1.116 160.42 ± 17.42

MJ-6 Microcline + Albite

J=0.010028 ± 0.000010 Volume 39ArK = 33.01 × 1E-13, Integrated age = 135.05 ± 12.35, Plateau age = 142.4 ± 6.7 Ma (89.4% 39ArK)

1 2 474.135 ± 0.019 2.807 ± 0.025 0.221 ± 0.111 1.607 ± 0.027 0.53 0.574 99.47 4.31 2.455 ± 9.679 43.87 ± 170.90

2 2.2 311.894 ± 0.013 2.222 ± 0.020 0.833 ± 0.033 1.051 ± 0.021 2.363 0.463 98.82 6.34 3.618 ± 5.332 64.30 ± 93.09

3* 2.4 84.038 ± 0.007 0.819 ± 0.027 7.176 ± 0.014 0.262 ± 0.020 21.138 0.175 89.73 16 8.599 ± 1.517 149.22 ± 25.27

4* 2.6 46.170 ± 0.010 0.458 ± 0.034 9.655 ± 0.017 0.140 ± 0.026 28.519 0.097 82.46 10.04 7.961 ± 1.019 138.57 ± 17.07

5* 2.9 21.687 ± 0.007 0.112 ± 0.037 2.095 ± 0.017 0.048 ± 0.034 6.13 0.02 61.84 38.96 8.151 ± 0.485 141.75 ± 8.11

6* 3.3 57.626 ± 0.008 0.812 ± 0.019 4.261 ± 0.016 0.174 ± 0.032 12.494 0.177 85.73 14.75 8.121 ± 1.626 141.25 ± 27.20

7* 3.8 88.913 ± 0.011 1.394 ± 0.019 7.153 ± 0.019 0.280 ± 0.025 21.054 0.308 89.72 9.6 9.059 ± 1.897 156.87 ± 31.47

MA64-3 Actinolite

J=0.009474 ± 0.000010 Volume 39ArK = 286.71 × 1E-13, Integrated age = 117.2 ± 0.7, No Plateau age

1 2 281.322 ± 0.029 2.725 ± 0.060 4.651 ± 0.047 0.939 ± 0.066 15.627 0.728 96.79 0.05 10.794 ± 20.740 175.66 ± 321.60

239 Laser

Step power %40Ar Age (Ma) no. (W) 40Ar/39Ar 38Ar/39Ar 37Ar/39Ar 36Ar/39Ar Ca/K Cl/K atm f 39Ar 40Ar*/39ArK 2ơ errow

2 2.3 58.784 ± 0.014 0.946 ± 0.025 3.571 ± 0.020 0.177 ± 0.028 9.867 0.211 87.49 0.77 7.336 ± 1.312 121.22 ± 20.97

3 2.6 12.926 ± 0.006 0.189 ± 0.043 1.260 ± 0.017 0.024 ± 0.046 3.435 0.039 52.11 2.46 6.041 ± 0.329 100.40 ± 5.32

4 2.9 9.066 ± 0.006 0.243 ± 0.014 0.920 ± 0.017 0.011 ± 0.045 2.502 0.053 31.6 4.07 6.059 ± 0.147 100.70 ± 2.38

5 3.2 11.108 ± 0.005 1.464 ± 0.011 4.369 ± 0.013 0.016 ± 0.029 11.929 0.335 37.58 5.73 6.859 ± 0.146 113.57 ± 2.35

6 3.5 11.187 ± 0.007 2.855 ± 0.012 9.432 ± 0.014 0.015 ± 0.029 25.871 0.661 28.12 9.43 8.035 ± 0.141 132.36 ± 2.24

7 3.8 11.350 ± 0.009 2.962 ± 0.016 10.321 ± 0.016 0.015 ± 0.036 28.327 0.687 27.53 12.23 8.239 ± 0.181 135.60 ± 2.88

8 4.1 9.361 ± 0.006 1.855 ± 0.011 7.312 ± 0.014 0.009 ± 0.031 19.946 0.427 17.04 6.25 7.600 ± 0.099 125.44 ± 1.58

9 4.3 8.330 ± 0.009 1.160 ± 0.015 5.091 ± 0.022 0.007 ± 0.041 13.803 0.266 9.16 3.57 7.207 ± 0.104 119.16 ± 1.67

10 4.6 8.414 ± 0.008 1.176 ± 0.015 5.051 ± 0.031 0.007 ± 0.084 13.688 0.269 11.07 3.43 7.115 ± 0.188 117.69 ± 3.02

11 5 8.088 ± 0.007 0.974 ± 0.016 4.455 ± 0.017 0.006 ± 0.063 12.194 0.222 9.86 5.54 7.056 ± 0.118 116.74 ± 1.88

12 5.6 7.935 ± 0.004 0.463 ± 0.016 2.046 ± 0.020 0.005 ± 0.036 5.577 0.103 14.49 9.77 6.639 ± 0.065 110.04 ± 1.05

13 8 7.779 ± 0.004 0.636 ± 0.009 3.311 ± 0.013 0.005 ± 0.023 9.088 0.143 13.39 36.71 6.692 ± 0.046 110.89 ± 0.74

MA64-3 (rerun) Actinolite

J=0.0094748 ± 0.000010 Volume 39ArK = 569.1 × 1E-13, Integrated age = 121.2 ± 0.6, Plateau age = 109.9 ± 1.0 Ma (28.2% 39ArK)

1 2 241.606 ± 0.036 2.303 ± 0.059 4.027 ± 0.885 0.835 ± 0.045 8.115 0.518 100.27 0.07 -0.691±7.226 -11.85±124.32

2 2.4 45.508 ± 0.006 0.659 ± 0.011 2.968 ± 0.034 0.135 ± 0.022 8.01 0.143 86.13 1.51 6.224 ± 0.879 103.37 ± 14.18

3 2.8 9.778 ± 0.007 0.171 ± 0.021 1.228 ± 0.047 0.015 ± 0.033 3.307 0.035 41.5 3.96 5.551 ± 0.148 92.47 ± 2.40

4 3.1 10.841 ± 0.005 0.902 ± 0.012 2.772 ± 0.022 0.016 ± 0.023 7.499 0.205 39.52 5.41 6.432 ± 0.114 106.72 ± 1.83

5 3.3 11.377 ± 0.005 2.566 ± 0.010 8.345 ± 0.015 0.016 ± 0.025 22.797 0.593 28.95 4.33 7.960 ± 0.121 131.16 ± 1.92

6 3.5 11.962 ± 0.008 2.398 ± 0.012 8.878 ± 0.016 0.018 ± 0.038 24.256 0.554 32.51 2.98 7.895 ± 0.215 130.14 ± 3.41

7 3.7 14.134 ± 0.007 3.393 ± 0.010 10.505 ± 0.014 0.025 ± 0.024 28.787 0.786 40.81 6 8.328 ± 0.189 137.00 ± 2.99

8 3.9 13.020 ± 0.008 3.488 ± 0.012 12.404 ± 0.015 0.020 ± 0.026 34.071 0.81 32.27 6.46 8.796 ± 0.172 144.41 ± 2.71

9 4.1 10.912 ± 0.007 2.897 ± 0.011 10.794 ± 0.014 0.013 ± 0.028 29.586 0.672 21.01 5.1 8.531 ± 0.127 140.21 ± 2.00

10 4.3 10.333 ± 0.008 2.280 ± 0.011 8.542 ± 0.017 0.012 ± 0.037 23.345 0.527 22.19 4.62 7.913 ± 0.149 130.42 ± 2.36

11 4.6 10.017 ± 0.007 2.122 ± 0.011 7.531 ± 0.014 0.011 ± 0.024 20.557 0.489 21.87 5.08 7.704 ± 0.099 127.09 ± 1.58

12 4.9 9.900 ± 0.006 1.764 ± 0.010 6.954 ± 0.015 0.011 ± 0.023 18.965 0.406 22.64 4.87 7.525 ± 0.092 124.23 ± 1.47

13 5.2 10.154 ± 0.006 1.676 ± 0.011 6.597 ± 0.017 0.012 ± 0.030 17.98 0.385 24.43 4.88 7.541 ± 0.119 124.49 ± 1.90

14 5.6 10.168 ± 0.007 1.267 ± 0.011 6.566 ± 0.015 0.013 ± 0.029 17.902 0.29 26.94 5.69 7.320 ± 0.122 120.96 ± 1.96

15* 6 8.750 ± 0.005 0.561 ± 0.012 2.979 ± 0.019 0.008 ± 0.033 8.08 0.126 22.87 9 6.651 ± 0.089 110.24 ± 1.44

16* 6.5 8.913 ± 0.005 0.437 ± 0.011 2.435 ± 0.013 0.009 ± 0.023 6.603 0.097 25.13 12.67 6.598 ± 0.070 109.39 ± 1.12

17* 7 9.521 ± 0.005 0.557 ± 0.013 2.840 ± 0.016 0.011 ± 0.034 7.691 0.125 28.56 6.55 6.681 ± 0.116 110.72 ± 1.86

18 8 10.983 ± 0.005 1.279 ± 0.010 6.510 ± 0.013 0.015 ± 0.027 17.763 0.293 30.98 10.8 7.535 ± 0.123 124.39 ± 1.97

MA17-7 Microcline

J=0.010073 ± 0.000012 Volume 39ArK = 2671.12 × 1E-13, Integrated age = 107.59 ± 0.32, Plateau age = 103.73 ± 0.64 Ma (33.9% 39ArK)

1 2 18.009 ± 0.005 0.113 ± 0.018 0.007 ± 0.060 0.041 ± 0.024 0.019 0.021 65.52 1.41 6.148 ± 0.285 108.40 ± 4.87

2 2.2 7.532 ± 0.005 0.028 ± 0.025 0.007 ± 0.101 0.006 ± 0.028 0.019 0.003 22.42 2.54 5.774 ± 0.058 101.99 ± 1.00

3 2.4 7.739 ± 0.004 0.029 ± 0.023 0.008 ± 0.099 0.005 ± 0.035 0.024 0.003 18.37 2.58 6.246 ± 0.062 110.08 ± 1.05

240 Laser

Step power %40Ar Age (Ma) no. (W) 40Ar/39Ar 38Ar/39Ar 37Ar/39Ar 36Ar/39Ar Ca/K Cl/K atm f 39Ar 40Ar*/39ArK 2ơ errow

5 2.9 7.927 ± 0.004 0.027 ± 0.017 0.009 ± 0.053 0.004 ± 0.025 0.027 0.003 13.35 4.84 6.812 ± 0.042 119.74 ± 0.71

6 3.2 7.748 ± 0.004 0.030 ± 0.014 0.005 ± 0.053 0.004 ± 0.021 0.015 0.003 15.23 6.03 6.517 ± 0.038 114.70 ± 0.65

7 3.5 7.727 ± 0.004 0.029 ± 0.016 0.002 ± 0.078 0.005 ± 0.020 0.005 0.003 19.2 11.16 6.199 ± 0.040 109.28 ± 0.68

8 3.6 7.558 ± 0.004 0.028 ± 0.012 0.001 ± 0.095 0.005 ± 0.020 0.004 0.003 19.48 10.75 6.042 ± 0.040 106.59 ± 0.69

4 2.6 7.717 ± 0.010 0.028 ± 0.023 0.014 ± 0.059 0.004 ± 0.072 0.042 0.003 12.26 1.81 6.680 ± 0.114 117.49 ± 1.94

9 3.7 7.583 ± 0.008 0.028 ± 0.017 0.002 ± 0.056 0.005 ± 0.024 0.007 0.003 19.21 8.92 6.081 ± 0.061 107.26 ± 1.05

10 3.8 7.554 ± 0.004 0.028 ± 0.015 0.003 ± 0.077 0.005 ± 0.021 0.009 0.003 17.5 7.47 6.184 ± 0.039 109.01 ± 0.67

11 3.9 7.597 ± 0.004 0.027 ± 0.024 0.002 ± 0.063 0.005 ± 0.024 0.004 0.003 18.71 8.57 6.129 ± 0.044 108.08 ± 0.76

12* 4 7.407 ± 0.005 0.024 ± 0.015 0.003 ± 0.048 0.005 ± 0.025 0.01 0.002 20.28 9.84 5.861 ± 0.051 103.48 ± 0.87

13* 4.1 7.315 ± 0.004 0.024 ± 0.018 0.003 ± 0.061 0.005 ± 0.026 0.008 0.002 19.01 9.39 5.879 ± 0.044 103.79 ± 0.76

14* 4.2 7.462 ± 0.006 0.025 ± 0.021 0.003 ± 0.075 0.005 ± 0.023 0.008 0.002 20.89 7.74 5.856 ± 0.050 103.40 ± 0.86

15* 4.3 7.456 ± 0.004 0.026 ± 0.015 0.002 ± 0.084 0.005 ± 0.027 0.005 0.002 20.1 3.27 5.894 ± 0.050 104.06 ± 0.86

16* 4.4 7.407 ± 0.004 0.024 ± 0.026 0.001 ± 0.238 0.005 ± 0.031 0.001 0.002 19.71 3.65 5.887 ± 0.055 103.93 ± 0.94

MA14-3 Microcline

J=0.010092 ± 0.000014 Volume 39ArK = 986.07 × 1E-13, Integrated age = 101.16 ± 0.51, Plateau age = 101.49 ± 0.67 Ma (83.3% 39ArK)

1 2 10.814 ± 0.005 0.024 ± 0.018 0.004 ± 0.158 0.020 ± 0.022 0.009 0.001 51.93 5.63 5.138 ± 0.127 91.20 ± 2.20

2 2.2 7.445 ± 0.005 0.016 ± 0.063 0.008 ± 0.091 0.007 ± 0.052 0.021 0 21.84 3.56 5.717 ± 0.105 101.20 ± 1.80

3 2.4 8.362 ± 0.005 0.017 ± 0.039 0.003 ± 0.330 0.009 ± 0.038 0.002 0 27.71 2.41 5.919 ± 0.111 104.66 ± 1.90

4* 2.6 7.968 ± 0.004 0.019 ± 0.032 0.001 ± 0.241 0.008 ± 0.026 0.002 0.001 26.88 8.59 5.765 ± 0.065 102.02 ± 1.12

5* 2.8 7.431 ± 0.005 0.018 ± 0.052 0.007 ± 0.097 0.006 ± 0.054 0.017 0.001 20.87 3.08 5.767 ± 0.107 102.06 ± 1.84

6* 3 7.962 ± 0.004 0.020 ± 0.018 0.002 ± 0.168 0.008 ± 0.034 0.005 0.001 27.03 7.1 5.743 ± 0.083 101.65 ± 1.43

7* 3.2 8.225 ± 0.004 0.021 ± 0.013 0.001 ± 0.219 0.009 ± 0.024 0.002 0.001 30.01 18.23 5.710 ± 0.066 101.08 ± 1.14

8* 3.4 8.198 ± 0.004 0.021 ± 0.018 0.000 ± 0.338 0.008 ± 0.018 0.001 0.001 30.04 22.98 5.691 ± 0.052 100.75 ± 0.90

9* 3.5 8.118 ± 0.004 0.022 ± 0.020 0.001 ± 0.274 0.008 ± 0.025 0.002 0.001 28.49 9.92 5.748 ± 0.066 101.72 ± 1.13

10* 3.6 8.487 ± 0.004 0.022 ± 0.050 0.003 ± 0.406 0.011 ± 0.038 0.002 0.001 31.02 1.97 5.713 ± 0.124 101.13 ± 2.14

11* 3.8 9.872 ± 0.005 0.023 ± 0.026 0.000 ± 0.253 0.014 ± 0.019 0 0.002 40.87 11.43 5.787 ± 0.082 102.40 ± 1.40

12 4 10.011 ± 0.005 0.024 ± 0.041 0.001 ± 0.204 0.014 ± 0.042 0.001 0.002 39.12 3.71 6.011 ± 0.179 106.24 ± 3.06

13 4.3 9.641 ± 0.006 0.024 ± 0.035 0.004 ± 0.216 0.014 ± 0.046 0.001 0.001 36.35 1.38 5.962 ± 0.197 105.42 ± 3.39

MA45-6 Albite (± Microcline)

J=0.010064 ± 0.000010 Volume 39ArK = 1400.58 × 1E-13, Integrated age = 96.67 ± 1.16, Plateau age = 99.05 ± 0.92 Ma (84.9% 39ArK)

1 2 49.971 ± 0.005 0.073 ± 0.016 0.009 ± 0.057 0.160 ± 0.017 0.022 0.007 94.22 2.2 2.850 ± 0.782 51.01 ± 13.80

2 2.2 12.524 ± 0.012 0.030 ± 0.024 0.009 ± 0.072 0.026 ± 0.020 0.026 0.002 61.72 6.16 4.744 ± 0.177 84.14 ± 3.07

3 2.4 9.564 ± 0.004 0.021 ± 0.027 0.012 ± 0.072 0.014 ± 0.018 0.034 0.001 43.19 6.73 5.377 ± 0.078 95.07 ± 1.35

4* 2.6 10.568 ± 0.004 0.025 ± 0.031 0.024 ± 0.043 0.017 ± 0.024 0.07 0.002 46.55 3.86 5.578 ± 0.126 98.54 ± 2.17

5* 2.8 12.161 ± 0.005 0.036 ± 0.036 0.053 ± 0.033 0.022 ± 0.023 0.156 0.004 51.82 2.42 5.773 ± 0.149 101.89 ± 2.56

6* 3 14.060 ± 0.004 0.050 ± 0.025 0.110 ± 0.027 0.029 ± 0.026 0.329 0.007 58.83 2.37 5.709 ± 0.220 100.79 ± 3.78

7* 3.2 21.041 ± 0.005 0.068 ± 0.016 0.064 ± 0.026 0.053 ± 0.019 0.192 0.01 73.6 3.93 5.504 ± 0.288 97.26 ± 4.96

241 Laser

Step power %40Ar Age (Ma) no. (W) 40Ar/39Ar 38Ar/39Ar 37Ar/39Ar 36Ar/39Ar Ca/K Cl/K atm f 39Ar 40Ar*/39ArK 2ơ errow

10* 3.6 13.310 ± 0.004 0.038 ± 0.019 0.025 ± 0.039 0.026 ± 0.016 0.074 0.004 57.43 7.54 5.617 ± 0.129 99.21 ± 2.22

11* 3.7 16.037 ± 0.004 0.046 ± 0.024 0.035 ± 0.026 0.035 ± 0.018 0.104 0.006 64.44 5.31 5.650 ± 0.191 99.78 ± 3.28

12* 3.8 17.384 ± 0.004 0.049 ± 0.017 0.039 ± 0.019 0.040 ± 0.017 0.116 0.006 68.59 7.85 5.416 ± 0.201 95.74 ± 3.46

13* 3.9 17.284 ± 0.004 0.045 ± 0.023 0.032 ± 0.034 0.040 ± 0.019 0.095 0.005 67.86 4.81 5.503 ± 0.221 97.25 ± 3.80

14* 4 20.268 ± 0.004 0.053 ± 0.022 0.041 ± 0.038 0.050 ± 0.020 0.121 0.007 71.63 3.11 5.691 ± 0.295 100.48 ± 5.06

8* 3.4 11.803 ± 0.004 0.031 ± 0.016 0.016 ± 0.039 0.021 ± 0.017 0.046 0.003 52.56 12.09 5.554 ± 0.106 98.13 ± 1.82

9* 3.5 11.689 ± 0.004 0.028 ± 0.021 0.018 ± 0.028 0.021 ± 0.019 0.055 0.002 51.65 10.68 5.605 ± 0.116 99.00 ± 2.00

15* 4.1 25.375 ± 0.005 0.066 ± 0.017 0.061 ± 0.025 0.068 ± 0.019 0.183 0.009 78.55 4.38 5.396 ± 0.375 95.41 ± 6.47

16* 4.2 20.196 ± 0.006 0.057 ± 0.018 0.039 ± 0.027 0.049 ± 0.018 0.115 0.008 71.56 4.58 5.694 ± 0.274 100.53 ± 4.70

17* 4.3 25.782 ± 0.005 0.084 ± 0.020 0.063 ± 0.026 0.068 ± 0.019 0.188 0.013 77.53 2.38 5.734 ± 0.384 101.21 ± 6.59

18* 4.5 27.439 ± 0.004 0.081 ± 0.024 0.065 ± 0.032 0.073 ± 0.018 0.194 0.012 78.88 5 5.751 ± 0.383 101.50 ± 6.57

19* 4.7 25.593 ± 0.004 0.089 ± 0.018 0.058 ± 0.021 0.068 ± 0.019 0.173 0.014 77.79 3.68 5.634 ± 0.371 99.49 ± 6.38

20* 5 43.791 ± 0.006 0.145 ± 0.034 0.120 ± 0.031 0.128 ± 0.020 0.351 0.024 85.66 0.94 6.205 ± 0.721 109.28 ± 12.32

MA17-9 Microcline

J=0.010080 ± 0.000012 Volume 39ArK = 1040.26 × 1E-13, Integrated age = 93.77 ± 0.41, Plateau age = 95.04 ± 0.60 Ma (72.6% 39ArK)

1 2 13.574 ± 0.005 0.044 ± 0.028 0.022 ± 0.051 0.032 ± 0.020 0.064 0.006 67.74 4.8 4.326 ± 0.190 77.01 ± 3.31

2 2.2 7.340 ± 0.004 0.029 ± 0.032 0.097 ± 0.022 0.008 ± 0.026 0.299 0.003 30.35 5.8 5.044 ± 0.068 89.47 ± 1.17

3 2.4 7.602 ± 0.005 0.037 ± 0.022 1.228 ± 0.013 0.009 ± 0.025 3.817 0.005 31.01 6.03 5.185 ± 0.073 91.91 ± 1.27

4 2.6 7.118 ± 0.008 0.030 ± 0.062 1.242 ± 0.017 0.008 ± 0.102 3.854 0.003 16.69 1.05 5.674 ± 0.248 100.34 ± 4.27

5* 2.7 6.639 ± 0.005 0.022 ± 0.028 0.568 ± 0.015 0.005 ± 0.030 1.761 0.002 19.08 5.82 5.301 ± 0.053 93.91 ± 0.91

6* 2.8 6.612 ± 0.005 0.021 ± 0.030 0.278 ± 0.019 0.005 ± 0.060 0.859 0.001 15.44 3.85 5.495 ± 0.085 97.25 ± 1.46

7* 3 6.932 ± 0.005 0.026 ± 0.032 0.241 ± 0.016 0.006 ± 0.036 0.746 0.003 22.42 6.23 5.309 ± 0.068 94.04 ± 1.17

8* 3.2 6.970 ± 0.005 0.029 ± 0.023 0.086 ± 0.021 0.006 ± 0.022 0.266 0.003 22.47 16.69 5.357 ± 0.048 94.87 ± 0.83

9* 3.3 7.122 ± 0.004 0.028 ± 0.020 0.054 ± 0.027 0.006 ± 0.022 0.168 0.003 24.07 11.81 5.355 ± 0.047 94.84 ± 0.81

10* 3.4 7.343 ± 0.004 0.027 ± 0.019 0.032 ± 0.043 0.007 ± 0.022 0.097 0.003 26.02 8.7 5.373 ± 0.051 95.15 ± 0.88

11* 3.5 7.899 ± 0.004 0.024 ± 0.017 0.026 ± 0.046 0.008 ± 0.027 0.08 0.002 30.14 8.72 5.460 ± 0.073 96.65 ± 1.25

12* 3.6 7.639 ± 0.005 0.026 ± 0.034 0.061 ± 0.035 0.008 ± 0.054 0.184 0.002 27.6 2.64 5.417 ± 0.137 95.91 ± 2.37

13* 3.8 7.551 ± 0.005 0.025 ± 0.036 0.028 ± 0.041 0.008 ± 0.038 0.085 0.002 28.07 3.82 5.343 ± 0.096 94.63 ± 1.65

14* 4.4 7.757 ± 0.004 0.023 ± 0.020 0.084 ± 0.026 0.008 ± 0.025 0.259 0.002 30.37 9.7 5.346 ± 0.066 94.69 ± 1.14

15* 4.5 7.901 ± 0.005 0.024 ± 0.036 0.042 ± 0.032 0.009 ± 0.022 0.129 0.002 30.73 4.32 5.392 ± 0.063 95.49 ± 1.09 Measured volumes are × 1E-13 cm3 Isotope production ratios: All ages are Ma and errors are 2σ standard error (40Ar/39Ar) K = 0.0302 ± 0.00006 % 40Ar atm : percentage of atmospheric 40Ar (37Ar/39Ar) Ca = 1416.4 ± 0.5 f 39Ar : fraction of 39Ar (36Ar/39Ar) Ca = 0.3952 ± 0.0004 40Ar* : radiogenic 40Ar Ca/K = 1.83±0.01 × 37ArCa/39ArK) 39ArK : 39Ar from K * : steps used to determine plateau age

242 Appendix C

Fluid Inclusion Database

Sample No. Host Origin Type Size μm Filling % Te Tm(ice) Thh Th(vapor) Tm(halite) mineral MA7-23-1 cal A S. 4*3 60 -41.5 -9.1 164.1 MA7-23-2 cal A S. 7*4 65 -37.9 -10.4 173.9 MA7-23-3 cal A S. 7*3 65 -46.4 -10.8 170.5 MA7-23-4 cal A S. 7*4 60 -49.5 -8.2 185.6 MA7-23-5 cal A U.O. 4*2 60 114.3 MA7-23-6 cal A S. 6*4 60 -6.0 125.1 MA7-23-7 cal A S. 7*3 65 156.0 MA7-23-8 cal A U.O. 5*3 60 172.1 MA7-23-9 cal A U.O. 10*4 65 168.5 MA7-23-10 cal A U.O. 5*2 60 166.1 MA7-23-11 cal A U.O. 4*3 60 MA7-23-12 cal A U.O. 10*5 60 MA7-23-13 cal A U.O. 4*2 60 MA7-23-14 cal A U.O. 6*3 65 MA7-23-15 cal A U.O. 6*3 65 139.7 MA7-23-16 cal A U.O. 15*5 65 MA7-23-17 cal A U.O. 8*4 70 115.1 MA7-23-18 cal A U.O. 10*6 65 125.6 MA7-23-19 cal A U.O. 8*5 70 121.1 MA7-23-20 cal A U.O. 6*4 70 131.2 MA7-23-21 cal A U.O. 5*4 65 133.5 MA7-23-29 cal A U.O. 12*4 70 -61.5 -45.0 140.9 MA7-23-30 cal A U.O. 6*4 60 -54.3 -26.5 157.6 MA7-23-31 cal A U.O. 6*3 65 142.1 MA7-23-32 cal B-2 U.O. 7*4 60 -41.2 142.8 MA7-23-33 cal B-1 U.O. 8*4 60 -65.3 -37.6 125.6 130.1 MA7-23-34 cal B-1 U.O. 15*5 70 -43.1 -31.5 105.1 128.7 MA7-23-35 cal A U.O. 12*4 70 179.1 MA7-23-36 cal A U.O. 3*3 60 -10.0 158.6

243 Sample No. Host Origin Type Size μm Filling % Te Tm(ice) Thh Th(vapor) Tm(halite) mineral MA7-23-37 cal A U.O. 12*5 60 -33.1 -12.0 163.8 MA7-23-38 cal A U.O. 4*3 60 -45.6 -22.9 158.6 MA7-23-39 cal A U.O. 4*2 65 121.6 MA7-23-40 cal A U.O. 3*3 60 120.7 MA7-23-41 cal A U.O. 8*6 60 -43.2 -24.8 MA7-23-42 cal A U.O. 18*9 60 -65.1 -31.2 142.4 MA7-23-43 cal A U.O. 20*10 75 140.0 MA7-23-44 cal A U.O. 8*7 60 -63.8 -35.4 161.1 MA7-23-45 cal B-2 U.O. 10*6 60 102.1 MA7-23-46 cal A U.O. 6*4 70 -40.1 -6.5 112.0 MA7-23-47 cal A U.O. 8*5 70 -60.4 -16.7 100.1 MA7-23-48 cal B-2 U.O. 9*4 70 110.1 MA7-23-49 cal A U.O. 8*4 70 -50.1 -9.1 100.1 MA7-23-50 cal A U.O. 6*5 60 100.9 MA7-23-51 cal A U.O. 7*4 65 102.1 MA7-23-52 cal A U.O. 8*5 65 -41.1 -27.2 163.1 MA7-23-53 cal A U.O. 15*6 65 -10.0 112.3 MA7-23-54 cal A U.O. 6*4 60 -35.9 -25.5 160.5 MA7-23-55 cal A U.O. 7*3 65 -45.0 -20.1 161.1 MA7-23-56 cal A U.O. 10*5 70 -40.0 -5.1 115.4 MA7-23-57 cal A U.O. 8*3 65 -39.0 -24.8 160.0 MA7-23-58 cal A U.O. 5*3 60 -23.2 159.1 MA7-23-59 cal A S. 7*7 75 -53.2 -41.1 121.2 MA7-23-60 cal A S. 5*3 70 -51.6 -40.1 120.1 MA7-23-61 cal A S. 5*3 60 -54.2 -37.1 119.9 MA7-23-62 cal A S. 8*5 65 -51.0 -38.9 122.4 MA7-23-63 cal A S. 6*4 60 -64.6 -38.9 131.2 MA7-23-64 cal A S. 6*4 60 -39.5 -23.3 125.1 MA7-23-65 cal A U.O. 10*5 75 -63.1 -26.9 -49.1 122.9 MA7-23-66 cal A U.O. 10*9 70 -8.5 136.9 MA7-23-67 cal A U.O. 7*5 65 -37.6 -24.4 135.7 MA7-23-68 cal A U.O. 8*6 65 -47.3 -18.9 120.0 MA7-23-69 cal A S. 6*6 70 -49.0 -24.5 118.7 MA7-23-70 cal A S. 7*5 70 -51.0 -25.1 122.1 MA7-23-71 cal A S. 7*5 60 -54.0 -25.0 120.5

244 Sample No. Host Origin Type Size μm Filling % Te Tm(ice) Thh Th(vapor) Tm(halite) mineral MA7-23-72 cal A S. 4*3 60 -53.1 -25.3 119.1 MA7-23-73 cal A U.O. 10*5 65 -72.1 -48.6 105.1 MA7-23-74 cal A U.O. 7*5 70 -70.5 -41.0 106.2 MA7-23-75 cal A U.O. 6*5 65 -71.1 -40.5 105.7 MA7-23-76 cal A U.O. 10*5 60 -64.1 -25.6 106.3 MA7-23-77 cal A U.O. 8*4 65 -63.1 -10.8 162.1 MA7-23-78 cal A U.O. 7*5 65 -51.0 -26.7 136.9 MA7-23-79 cal A U.O. 9*5 65 -53.2 -25.2 135.1 MA7-23-80 cal A U.O. 8*6 65 -50.0 -26.0 134.2 MA7-23-81 cal A U.O. 5*4 60 -26.7 135.4 MA7-23-82 cal A U.O. 4*3 60 133.9 MA7-23-83 cal A U.O. 6*4 65 -23.7 131.5 MA7-23-84 cal A U.O. 6*4 65 -31.2 -11.0 165.7 MA7-23-85 cal A U.O. 8*6 75 -30.9 -9.8 178.6 MA7-23-86 cal A U.O. 8*5 70 -65.1 -47.8 123.1 MA7-23-87 cal A U.O. 7*5 65 -59.8 -41.2 126.5 MA7-23-88 cal A U.O. 10*7 75 -64.1 -38.4 118.9 MA7-23-89 cal A U.O. 5*3 60 -57.3 -37.5 121.2 MA7-23-90 cal A S. 4*3 70 -52.1 -26.5 117.6 MA7-23-91 cal A S. 5*4 70 -50.1 -25.0 126.5 MA7-23-92 cal A S. 6*3 60 -49.8 -26.7 120.0 MA7-23-93 cal A S. 4*3 65 -51.0 -25.1 121.5 MA7-23-94 cal A S. 3*3 55 -49.1 -24.2 123.4 MA7-23-95 cal A U.O. 20*10 75 -73.2 -48.2 114.1 MA7-23-96 cal A U.O. 15*8 75 -46.5 -25.0 153.2 MA7-23-97 cal A U.O. 8*4 65 -31.1 114.0 MA7-23-98 cal A U.O. 10*6 70 -48.6 -28.6 122.7 MA7-23-99 cal A U.O. 3*7 60 -69.7 -48.9 100.1 MA7-23-100 cal A U.O. 7*3 60 -71.0 -52.2 143.2 MA7-23-101 cal A U.O. 4*4 70 -65.3 -46.9 103.2 MA7-23-102 cal A U.O. 4*3 60 -69.7 -51.8 98.9 MA7-23-103 cal A U.O. 6*5 65 92.3 MA7-23-104 cal A U.O. 10*8 65 -24.3 95.1 MA7-23-105 cal A U.O. 8*4 65 -70.1 -49.8 94.9 MA7-23-106 cal A U.O. 8*3 65 -55.6 -41.2 95.0

245 Sample No. Host Origin Type Size μm Filling % Te Tm(ice) Thh Th(vapor) Tm(halite) mineral MA7-23-107 cal A U.O. 7*5 65 -54.1 -36.0 98.9 MA7-23-108 cal A U.O. 20*10 80 -66.5 -39.6 -46.8 106.7 MA7-23-109 cal A U.O. 20*15 80 -65.3 -40.1 108.9 MA7-23-110 cal A U.O. 9*6 65 -37.6 105.1 MA7-23-111 cal A U.O. 7*4 65 -9.8 126.7 MA7-23-112 cal A U.O. 11*4 70 -32.1 110.2 MA7-23-113 cal A U.O. 8*3 60 -61.3 -32.0 107.1 MA7-23-114 cal A U.O. 9*5 70 105.6 MA7-23-115 cal A U.O. 10*5 75 106.1 MA7-23-116 cal A U.O. 9*3 65 -48.7 -9.1 125.6 MA7-23-117 cal A U.O. 12*8 50 -56.7 -36.3 115.0 MA7-23-118 cal A U.O. 10*8 70 -58.5 -33.3 109.7 MA7-23-119 cal A U.O. 7*6 65 -57.6 -34.5 95.6 MA7-23-120 cal A U.O. 10*6 70 -55.1 -37.1 87.8 MA7-23-121 cal A U.O. 10*6 70 -26.7 90.1 MA7-23-122 cal A U.O. 7*3 60 -41.0 94.5 MA7-23-123 cal A U.O. 9*6 65 -66.7 -43.4 96.7 MA7-23-124 cal A U.O. 8*4 65 -40.0 92.3 MA7-23-125 cal A U.O. 8*3 60 -77.3 -48.5 95.1 MA7-23-126 cal A U.O. 4*4 55 -70.1 -36.0 95.0 MA7-23-127 cal A U.O. 9*6 70 -68.9 -35.4 93.2 MA7-23-128 cal A U.O. 5*5 65 -41.1 152.7 MA7-23-129 cal A U.O. 4*3 65 -61.2 -36.0 149.6 MA7-23-130 cal A U.O. 7*5 60 -63.1 -31.0 138.6 MA7-23-132 cal A U.O. 7*4 70 -61 -40.5 142.0 MA7-23-133 cal A U.O. 6*4 70 -36.0 141.1 MA7-23-134 cal A U.O. 6*5 70 -50.1 -23.7 139.9 MA7-23-135 cal A U.O. 7*4 65 -60.5 -32.0 140.0 MA7-23-136 cal A U.O. 3*2 60 -61.0 -31.1 141.2 MA7-23-137 cal B-1 U.O. 10*4 60 -63.1 92.3 195.4 MA7-23-138 cal C U.O. 9*7 65 -54.1 -25.6 94.3 213.2 MA7-23-139 cal B-1 U.O. 8*5 70 91.2 193.1 MA7-23-140 cal B-1 U.O. 12*7 70 87.6 178.6 MA7-23-141 cal B-1 U.O. 15*5 75 68.6 201.1 MA7-23-142 cal B-1 U.O. 9*6 70 -24.3 75.1 191

246 Sample No. Host Origin Type Size μm Filling % Te Tm(ice) Thh Th(vapor) Tm(halite) mineral MA7-23-143 cal A U.O. 12*4 75 -65.3 -25.1 110.5 MA7-23-144 cal A U.O. 8*5 60 -47.8 -28.7 91.6 MA7-23-145 cal A U.O. 9*6 60 -46.1 29.2 153.6 MA7-23-146 cal B-1 U.O. 10*6 60 -54.6 -36.5 113.4 117.3 MA7-23-147 cal B-1 U.O. 10*7 60 -55.0 -36.7 118.2 121.3 MA7-23-148 cal B-1 U.O. 9*4 65 -57.2 -36.1 117.9 118.5 MA7-23-149 cal A U.O. 8*5 65 -45.0 -23.7 124.1 MA7-23-150 cal A U.O. 12*8 70 -18.9 119.6 MA7-23-151 cal A U.O. 9*4 60 -24.2 113.1 MA7-23-152 cal B-1 U.O. 7*5 60 -60 -36.5 115.6 127.3 MA7-23-153 cal A U.O. 6*5 60 -48.7 -20.1 127.3 MA7-23-154 cal A U.O. 15*10 75 -58.7 -34.2 124.1 MA7-23-155 cal B-1 U.O. 9*6 65 -59.8 -37.0 121.3 135.4 MA7-23-156 cal A U.O. 12*8 75 -47.0 -19.6 119.8 MA7-23-157 cal B-1 U.O. 14*9 75 -60.1 -36.9 114.7 125.6 MA7-23-158 cal A U.O. 13*7 70 -51.9 -27.3 140.1 MA7-23-159 cal A U.O. 6*3 65 -62.1 -39.3 98.1 MA7-23-160 cal A S. 9*4 65 -54.5 -28.7 117.1 MA7-23-161 cal A S. 8*3 65 -64.8 -38.3 117.3 MA7-23-162 cal A S. 7*3 60 -51.5 -26.3 123.6 MA7-23-163 cal A S. 5*4 65 -63.2 -33.7 118.2 MA7-23-164 cal A S. 4*3 65 -61.0 -39.6 119.6 MA7-23-165 cal A S. 4*4 60 -64.3 -38.9 119.7 MA7-23-166 cal A S. 4*3 60 -57.1 -39.6 126.2 MA7-23-167 cal A S. 4*3 60 -38.5 120.5 MA7-23-168 cal B-2 U.O. 9*8 20 114.5 MA7-23-169 cal B-1 U.O. 20*5 65 119.5 220.1 MA7-23-170 cal B-1 U.O. 10*5 70 118.6 218.7 MA7-23-171 cal C U.O. 25*5 70 119.0 MA7-23-172 cal A U.O. 9*7 70 137.2 MA7-23-173 cal B-1 U.O. 10*8 65 120.1 205.6 MA7-23-174 cal B-1 U.O. 9*6 70 118.6 208.5 MA7-23-175 cal B-2 U.O. 11*7 65 119.5 MA7-23-176 cal B-2 U.O. 11*5 75 117.5 DDM4-7-4-1 cal A U.O. 4*4 55 -50.0 -29.8 100.4

247 Sample No. Host Origin Type Size μm Filling % Te Tm(ice) Thh Th(vapor) Tm(halite) mineral DDM4-7-4-2 cal A U.O. 5*4 65 -47.9 -33.2 72.8 DDM4-7-4-3 cal A U.O. 9*5 60 -48.6 -33.5 126.7 DDM4-7-4-4 cal A U.O. 7*5 65 -49.7 -34.1 89.7 DDM4-7-4-5 cal A U.O. 6*5 60 -42.5 -15.6 151.2 DDM4-7-4-6 cal A U.O. 5*4 60 -39.6 -23.1 147.5 DDM4-7-4-7 cal A U.O. 4*2 60 -18.9 145.1 DDM4-7-4-8 cal A U.O. 6*6 60 -31.0 -4.2 171.3 DDM4-7-4-9 cal A U.O. 6*4 65 -30.5 -3.1 156.8 DDM4-7-4-10 cal A U.O. 8*5 65 -29.0 -4.7 165.0 DDM4-7-4-11 cal A U.O. 9*6 60 -28.6 -5.2 193.3 MA2-12-1 cal A U.O. 6*3 70 -6.6 136.9 MA2-12-2 cal A U.O. 10*4 70 -26.1 -14.2 129.0 MA2-12-3 cal A U.O. 10*5 70 -66.5 -15.0 -47.5 118.9 MA2-12-4 cal A U.O. 15*10 70 -52.1 -25.2 -39.0 120.1 MA2-12-5 cal A U.O. 12*5 65 -37.1 -14.3 122.3 MA2-12-6 cal A U.O. 10*4 70 140.7 MA2-12-7 cal A U.O. 14*6 55 220.7 MA2-12-8 cal A U.O. 5*3 45 MA2-12-9 cal A U.O. 10*5 70 100.1 MA2-12-10 cal A U.O. 20*8 80 113.4 MA2-12-11 cal A U.O. 14*6 70 118.3 MA2-12-12 cal A U.O. 25*10 80 115.1 MA2-12-13 cal A U.O. 18*7 75 130.6 MA2-12-14 cal A U.O. 8*4 50 MA2-12-15 cal A U.O. 18*10 65 155.5 MA2-12-17 cal A U.O. 12*5 75 105.0 MA2-12-18 cal A U.O. 12*3 75 101.5 MA2-12-19 cal A U.O. 12*5 60 -7.8 257.1 MA2-12-20 cal A U.O. 5*4 55 220.0 MA2-12-21 cal A U.O. 8*4 45 MA2-12-22 cal A U.O. 12*4 70 113.0 MA2-12-23 cal A U.O. 15*6 70 120.1 MA2-12-24 cal A U.O. 5*4 65 161.2 MA2-12-25 cal A U.O. 5*4 50 278.5 MA2-12-26 cal A U.O. 5*5 70 98.1

248 Sample No. Host Origin Type Size μm Filling % Te Tm(ice) Thh Th(vapor) Tm(halite) mineral MA2-12-27 cal A U.O. 6*5 65 -15.0 -6.0 127.6 MA2-12-28 cal A U.O. 10*4 65 -11.6 138.9 MA2-12-29 cal A U.O. 10*5 55 MA2-12-30 cal A U.O. 8*5 30 MA2-12-31 cal A U.O. 4*4 75 100.1 MA2-12-32 cal A U.O. 12*5 70 -20.1 -6.5 119.8 MA2-12-33 cal A U.O. 10*5 70 -22.0 -6.1 140.1 MA2-12-34 cal A U.O. 14*7 75 -61.5 -24.1 -45.3 121.0 MA2-12-35 cal A U.O. 10*4 70 -40.1 -5.8 125.3 MA2-12-36 cal A U.O. 6*4 70 -38.7 -22.1 135.0 MA2-12-37 cal A U.O. 20*8 70 126.5 MA2-12-38 cal A U.O. 6*3 50 MA2-12-39 cal A U.O. 15*8 60 -65 -20.4 -48 138.7 MA2-12-40 cal A U.O. 6*2 70 117.5 MA2-12-41 cal A U.O. 12*5 60 MA2-12-42 cal A U.O. 10*5 65 -39.5 -22.7 129.2 MA2-12-43 cal A U.O. 10*5 65 -38.6 -10.8 151.5 MA2-12-44 cal A U.O. 12*7 65 MA2-12-45 cal A U.O. 8*2 65 105.6 MA2-12-46 cal A U.O. 6*2 70 74.8 MA2-12-47 cal A U.O. 10*5 35 MA2-12-48 cal A U.O. * * MA2-12-49 cal A U.O. 18*9 70 MA2-12-50 cal A U.O. 8*4 60 119.1 MA2-12-51 cal A U.O. 20*10 70 105.7 MA2-12-52 cal A U.O. 10*8 70 115.4 MA2-12-53 cal A U.O. 8*5 70 103.1 MA2-12-54 cal A U.O. 15*10 75 108.9 MA2-12-55 cal A U.O. 20*10 60 -44.1 -8.9 144.1 MA2-12-56 cal A U.O. 8*5 65 -46.4 -31.0 109.8 MA2-12-57 cal A U.O. 7*3 70 128.6 MA2-12-58 cal A U.O. 12*5 70 -34 -18.5 115.1 MA2-12-59 cal A U.O. 10*4 70 114.5 MA2-12-60 cal A U.O. 8*4 20 MA2-12-61 cal A U.O. 8*7 60 -32.1 -15.7 147.4

249 Sample No. Host Origin Type Size μm Filling % Te Tm(ice) Thh Th(vapor) Tm(halite) mineral MA2-12-62 cal A U.O. 10*7 50 MA2-12-63 cal A U.O. 10*4 60 -4.8 212.3 MA2-12-64 cal A U.O. 4*2 70 -6.0 137.2 MA2-12-65 cal A U.O. 15*7 80 MA2-12-66 cal A U.O. 12*8 85 85.1 MA2-12-67 cal A U.O. 7*7 75 -37.0 -22.7 109.6 MA2-12-68 cal A U.O. 10*5 65 -35.6 -24.4 105.7 MA2-12-69 cal A U.O. 18*10 65 -33.2 -5.2 149.9 MA2-12-70 cal A U.O. 12*5 65 -23.1 -6.8 125.2 MA2-12-71 cal A U.O. 4*3 70 -37.1 -22.9 112.0 MA2-12-72 cal A U.O. 20*10 70 -23.1 -8.9 132.6 MA2-12-73 cal A U.O. 8*5 65 -48.3 -16.8 102.1 MA2-12-74 cal A U.O. 12*4 70 -39 -13.7 138.5 MA2-12-75 cal A U.O. 8*7 70 -33.9 -11.3 152.9 MA2-12-76 cal A U.O. 7*5 30 MA2-12-77 cal A U.O. 20*10 55 -50.2 -20.6 -43.1 320.6 MA2-12-78 cal A U.O. 8*7 60 -44.5 -22.7 258.6 MA2-12-79 cal A U.O. 12*7 65 -38.0 -23.9 304.5 MA2-12-80 cal A U.O. 14*6 65 -53.1 -20.1 -43.8 220.1 MA2-12-81 cal A U.O. 7*5 50 MA2-12-82 cal A U.O. 15*8 70 270.8 MA2-12-83 cal A U.O. 10*5 75 -35.6 -17.8 210.4 MA2-12-84 cal A U.O. 5*3 70 -36.7 -25.7 124.8 MA2-12-85 cal A U.O. 15*8 80 -53.2 -19.8 -38.6 115.7 MA2-12-86 cal A U.O. 10*5 70 -42.1 -13.7 115 MA2-12-87 cal A U.O. 30*20 80 -33.1 -6.8 119.8 MA2-12-88 cal A U.O. 15*10 75 -50.1 -6.0 -35 150.4 MA2-12-89 cal A U.O. 8*6 65 -37.7 -7.6 155.6 MA2-12-90 cal A U.O. 12*7 70 154.5 MA2-12-91 cal A U.O. 10*8 70 -39.8 -3.5 180.2 MA2-12-92 cal A U.O. 7*5 70 103.5 MA2-12-93 cal A U.O. 15*10 75 95.1 MA2-12-94 cal A U.O. 12*9 80 -7.4 112.7 MA2-12-95 cal A U.O. 10*5 75 98.9 MA2-12-96 cal A U.O. 8*6 65 -53.5 -20.2 178.6

250 Sample No. Host Origin Type Size μm Filling % Te Tm(ice) Thh Th(vapor) Tm(halite) mineral MA2-12-97 cal A U.O. 11*5 70 -68.1 -17.8 -40.4 86.5 MA2-12-98 cal A U.O. 20*8 85 -65.2 -27.1 -41.2 78.7 MA2-12-99 cal A U.O. 8*5 80 -63.4 -26.2 -42 100.7 MA2-12-100 cal A U.O. 6*3 65 -3.3 131.0 MA2-12-101 cal A U.O. 8*7 65 -39.7 -23.5 88.9 MA2-12-102 cal A U.O. 10*5 70 -20.1 101.2 MA2-12-103 cal A U.O. 12*7 80 -25.1 -2.6 124.5 MA2-12-104 cal A U.O. 20*15 50 -38.9 -5.1 358.9 MA2-12-105 cal A U.O. 5*4 70 -46.9 -11.1 130.5 MA2-12-107 cal A U.O. 9*5 70 -33.2 87.3 MA2-12-108 cal A U.O. 15*8 65 113.9 MA2-12-109 cal A U.O. 8*7 65 -56 -16.2 -45.4 107.8 MA2-12-110 cal A U.O. 6*4 70 -44 -12.6 105.7 MA2-12-111 cal A U.O. 6*5 70 -43.1 -14.9 99.6 MA2-12-112 cal A U.O. 7*5 65 -37.1 -11.8 95.5 MA2-12-113 cal A U.O. 8*7 65 -50.2 -15.4 119.0 MA2-12-114 cal A U.O. 25*8 60 -19.6 -0.8 227.1 MA2-12-115 cal A U.O. 8*5 65 -32.1 -3.8 121.2 MA2-12-116 cal A U.O. 8*4 65 -47.5 -30.9 106.2 MA2-12-117 cal A U.O. 8*7 70 -35 -18.1 122.1 MA2-12-118 cal A U.O. 8*5 65 -37.1 -11.9 142.8 MA2-12-119 cal A U.O. 6*3 70 -38.7 -12.0 111.5 MA2-12-120 cal A U.O. 8*6 75 -44.1 -23.0 99.5 MA2-12-121 cal A U.O. 7*6 75 -38.6 -24.0 119.8 MA2-12-122 cal A U.O. 4*3 75 -38.9 -25.2 100.1 MA2-12-123 cal A U.O. 9*5 70 -48.7 -25.6 133.4 MA2-12-124 cal A U.O. 8*4 65 -32.1 -6.8 100.7 MA2-12-125 cal A U.O. 15*9 75 -27.1 -4.9 124.5 MA2-12-126 cal A U.O. 5*2 70 117.0 MA2-12-127 cal A U.O. 7*3 70 -26.5 -10.5 117.3 MA2-12-128 cal A U.O. 11*6 70 110.0 MA2-12-129 cal A U.O. 12*6 65 144.1 MA2-12-130 cal A U.O. 9*6 70 -52.4 -30.1 105.9 MA2-12-131 cal A U.O. 8*5 70 -57.6 -21.9 -43.8 104.1 MA2-12-132 cal A U.O. 12*8 70 -23.1 -4.2 90.8

251 Sample No. Host Origin Type Size μm Filling % Te Tm(ice) Thh Th(vapor) Tm(halite) mineral MA2-12-133 cal A U.O. 8*4 65 -57.9 -23.8 -42 110.5 MA2-12-134 cal A U.O. 12*8 65 -68.1 -22.6 -44 115.6 MA2-12-135 cal A U.O. 10*7 65 -17.5 -3.6 154.3 MA2-12-136 cal A U.O. 7*5 65 -33.1 -11.4 136.1 MA2-12-137 cal A U.O. 7*3 60 -48.2 -18.8 148.6 MA2-12-138 cal A U.O. 7*2 70 -41.1 -14.0 142.5 MA2-12-139 cal A U.O. 5*3 75 -8.7 144.7 MA2-12-140 cal A U.O. 8*5 65 -31.2 -4.5 138.9 MA2-12-141 cal A U.O. 25*3 70 -19.8 -4.0 153.5 MA2-12-142 cal A U.O. 8*5 75 -42.3 -17.5 112.7 MA2-12-143 cal A U.O. 8*5 70 -34.5 -5.2 156.7 MA2-12-144 cal A U.O. 7*5 70 -35.6 -16.7 130.1 MA2-14-1 cal A U.O. 9*8 75 -12.6 -4.1 158.2 MA2-14-2 cal A U.O. 8*8 70 -18.9 -5.3 107.1 MA2-14-3 cal A U.O. 7*5 70 -15.0 -4.0 130.1 MA2-14-4 cal A U.O. 5*4 60 -16.1 -4.2 127.5 MA2-14-5 cal A U.O. 7*7 65 -10.8 -5.8 134.2 MA2-14-6 cal A U.O. 7*5 65 -11.0 -5.6 135.1 MA2-14-7 cal A U.O. 14*8 65 -23.2 -13.8 121.0 MA2-14-8 cal A U.O. 10*8 60 -21 -8.3 117.5 MA2-14-9 cal A U.O. 9*8 65 -21.9 -10.1 147.6 MA2-14-10 cal A U.O. 8*7 60 -22 -7.6 125 MA2-14-11 cal A U.O. 9*7 60 -21.2 -8.4 169.7 MA2-14-12 cal A U.O. 9*6 65 -28.2 -15.7 140.5 MA2-14-13 cal A U.O. 10*7 65 -21.5 -15.1 149.7 MA2-14-14 cal A U.O. 15*5 75 -28.3 -16.8 131.1 MA2-14-15 cal A U.O. 12*7 75 128.7 MA2-14-16 cal A U.O. 9*6 70 -14.7 114.7 MA2-14-17 cal A U.O. 10*8 70 -27.6 -15.1 127.5 MA2-14-18 cal A U.O. 8*7 70 -8.6 -4.0 143.7 MA2-14-19 cal A U.O. 8*5 70 -9.9 -3.9 128.6 MA2-14-20 cal A U.O. 7*4 65 -29.0 -14.1 125.6 MA2-14-21 cal A U.O. 7*5 60 -29.5 -14.5 125.1 MA3-35-1 qtz A U.O. 9*6 75 -27.4 MA3-35-2 qtz A U.O. 8*5 70 -45.6 -31.7 110.2

252 Sample No. Host Origin Type Size μm Filling % Te Tm(ice) Thh Th(vapor) Tm(halite) mineral MA3-35-3 qtz A U.O. 4*4 75 -51.2 -28.9 105.7 MA3-35-4 qtz A U.O. 6*4 60 -57.5 -30.1 111.3 MA3-35-5 qtz A U.O. 5*2 65 -50 -30.5 109.8 MA3-35-6 qtz A U.O. 6*3 65 -49.7 -29.8 109.6 MA3-35-7 qtz A U.O. 8*4 65 -50.0 -30.0 110.0 MA3-35-8 qtz A U.O. 5*5 60 -53.6 -32.9 107.6 MA3-35-9 qtz A U.O. 9*6 70 -40.1 -25.6 131.6 MA3-35-10 qtz A U.O. 6*3 60 -51.5 -30.7 113.5 MA3-35-11 qtz A U.O. 5*4 60 -52.7 -31.4 112.6 MA3-35-12 qtz A U.O. 5*3 60 -54.1 -29.9 93.8 MA3-35-13 qtz A U.O. 4*4 60 -49.7 -30.5 100.5 MA3-35-14 qtz A U.O. 9*6 65 -49.8 -25.6 113.8 MA3-35-15 qtz A U.O. 7*5 60 -50.0 -23.7 100.5 MA3-35-16 qtz A U.O. 5*5 60 -51.2 -27.5 70.1 MA3-35-17 qtz A U.O. 9*5 65 -51.0 -30.3 121.7 MA3-35-18 qtz A U.O. 5*5 60 -54.5 -29.8 110.2 MA3-35-19 qtz A U.O. 6*4 60 -51.2 -27.6 121.5 MA3-35-20 qtz A U.O. 4*2 60 -27.0 135.6 MA45-3-1 qtz A U.O. 10*7 85 -9.8.0 174.5 MA45-3-2 qtz A U.O. 10*6 80 -34.7 -14.1 134.9 MA45-3-3 qtz A U.O. 5*4 75 -30.5 -16.7 174.5 MA45-3-4 qtz A U.O. 5*3 65 -13.2 132.7 MA45-3-5 qtz A U.O. 10*4 75 151.8 MA45-3-6 qtz A U.O. 4*3 65 -24.2 -5.1 174.9 MA45-3-7 qtz A U.O. 10*5 75 MA45-3-8 qtz A U.O. 10*10 70 -26 -4.4 234.7 MA45-3-9 qtz A U.O. 10*5 85 -23.1 -8.2 153.8 MA45-3-10 qtz A U.O. 12*3 75 -40.1 -22.4 145.6 MA45-3-11 qtz A U.O. 7*4 70 -40.1 MA45-3-12 qtz A U.O. 8*3 70 -44.6 -11.6 131.0 MA45-3-13 qtz A U.O. 9*3 80 -4.6 176.1 MA45-3-14 qtz A U.O. 6*3 80 -44.8 -23.6 137.3 MA45-3-15 qtz A U.O. 10*4 70 -37.3 -19.3 147.9 MA45-3-16 qtz A U.O. 4*3 70 -14.2 134.1 MA45-3-17 qtz A U.O. 3*3 70 -15.1 123.6

253 Sample No. Host Origin Type Size μm Filling % Te Tm(ice) Thh Th(vapor) Tm(halite) mineral MA45-3-18 qtz A U.O. 10*3 65 -33.0 -9.9 230.2 MA45-3-19 qtz A U.O. 5*3 70 -1.5 137.6 MA45-3-20 qtz A U.O. 4*3 65 -31.0 -12.7 123.1 MA45-3-21 qtz A U.O. 5*3 65 -30.0 -15.1 148.4 MA45-3-22 qtz A U.O. 7*3 75 -30.8 -12.0 175.5 MA45-3-23 qtz A U.O. 8*3 75 -32.1 -13.1 166.2 MA45-3-24 qtz A U.O. 8*2 70 -34.2 -6.5 226.9 MA45-3-25 qtz A U.O. 9*4 70 -36.8 -11 250.1 MA45-3-26 qtz A U.O. 15*6 75 -39.9 -24.2 173.6 MA45-3-27 qtz A U.O. 9*4 70 -30.1 -1.8 148.3 MA45-3-28 qtz A U.O. 6*2 75 -5.9 169.9 MA45-3-29 qtz A U.O. 7*4 65 -40.0 -1.5 323.8 MA45-3-30 qtz A U.O. 8*2 75 -40.5 148.1 MA45-3-31 qtz A U.O. 12*3 65 -36.2 -22.8 290.8 MA45-3-32 qtz A U.O. 5*4 70 -40.1 -6.2 144.1 MA45-3-33 qtz A U.O. 4*4 70 -40.0 -24.4 157.1 MA45-3-34 qtz A U.O. 8*4 70 -2.5 172.2 MA45-3-35 qtz A U.O. 8*3 70 -58.0 -22.1 136.4 MA45-3-36 qtz A U.O. 7*5 70 -54.5 -19.8 -47.8 124.7 MA45-3-37 qtz A U.O. 6*4 55 -28.8 -1.4 214.5 MA45-3-38 qtz A U.O. 7*5 60 -10.7 162.3 MA45-3-39 qtz A U.O. 7*4 70 -11.4 139.5 MA45-3-40 qtz A U.O. 7*4 60 145.8 MA45-3-41 qtz A U.O. 8*2 70 -40.0 -28.1 135.6 MA45-3-42 qtz A U.O. 4*3 45 -58.9 -22.1 -33.8 217.1 MA45-3-43 qtz A U.O. 3*2 70 147.8 MA45-3-44 qtz A U.O. 4*3 65 -23.0 232.1 MA45-3-45 qtz A U.O. 4*4 65 -66.7 -10.0 171.5 MA45-3-46 qtz A U.O. 9*5 60 -23.0 -5.1 223 MA45-3-47 qtz A U.O. 12*3 70 -28.0 -4.9 197.1 MA45-3-48 qtz A U.O. 5*3 60 -24.1 -5.2 236.1 MA45-3-49 qtz A U.O. 10*2 75 -19.8 -4.5 181.8 MA45-3-51 qtz A U.O. 5*2 75 -0.3 220.2 MA45-3-52 qtz A U.O. 6*3 70 -21.3 -8.5 149.5 MA45-3-53 qtz A U.O. 15*4 75 -13.8 138.1

254 Sample No. Host Origin Type Size μm Filling % Te Tm(ice) Thh Th(vapor) Tm(halite) mineral MA45-3-54 qtz A U.O. 6*2 75 -17.4 -9.8 106.8 MA45-3-55 qtz A U.O. 4*4 75 -41.0 -9.6 167.6 MA45-3-56 qtz A U.O. 8*6 70 -24.1 -9.0 187.1 MA45-3-57 qtz A U.O. 7*2 75 -31.7 -12.0 135.1 MA45-3-58 qtz A U.O. 4*4 75 -40.1 -11.6 -25.5 192 MA45-3-59 qtz A U.O. 9*6 65 -16.7 -8.5 147.5 MA45-3-60 qtz A U.O. 10*5 75 -30.1 -6.1 213.4 MA45-3-61 qtz A U.O. 6*4 70 -31.1 -9.0 153.2 MA45-3-62 qtz A U.O. 10*2 70 -22.8 -2.6 183.1 MA45-3-63 qtz A U.O. 8*2 70 -30.7 -17.1 116.8 MA45-3-64 qtz A U.O. 6*4 65 -18.0 91.9 MA45-3-65 qtz A U.O. 6*3 70 128.1 MA45-3-66 qtz A U.O. 12*4 70 -16.5 151.9 MA89-1-1 qtz A U.O. 18*8 65 -19.2 -0.1 280.2 MA89-1-2 qtz B-1 U.O. 12*4 70 -63.6 -37.4 127.1 165.1 MA89-1-3 qtz B-1 U.O. 14*6 70 146.3 182.0 MA89-1-4 qtz B-1 U.O. 15*6 70 138.1 168.4 MA89-1-5 qtz B-2 U.O. 15*7 70 -66.6 -24.5 145.7 MA89-1-6 qtz A U.O. 5*3 60 -64.7 -18.1 119.7 MA89-1-7 qtz A U.O. 13*7 65 167.5 MA89-1-8 qtz A U.O. 25*8 65 -66.0 -27.6 149.3 MA89-1-9 qtz A U.O. 10*4 70 143.1 MA89-1-10 qtz A U.O. 5*5 70 -20.1 150.8 MA89-1-11 qtz A U.O. 20*4 70 208.9 MA89-1-12 qtz A U.O. 12*5 65 -9.9 146.4 MA89-1-13 qtz A U.O. 10*4 65 -6.1 137.5 MA89-1-14 qtz A U.O. 15*8 65 -36.5 -10.1 154.8 MA89-1-15 qtz A U.O. 15*6 75 -10.7 122.3 MA89-1-16 qtz A U.O. 15*5 65 -10.7 123.1 MA89-1-17 qtz A U.O. 8*2 70 -35.7 -14.2 86.5 MA89-1-18 qtz A U.O. 8*2 70 -32.0 -4.5 115.6 MA89-1-19 qtz A U.O. 7*4 70 -33.1 -3.8 175.2 MA89-1-20 qtz A U.O. 5*2 75 -30.1 -4.8 168.3 MA89-1-21 qtz A U.O. 9*4 70 -44.1 -19.8 83.3 MA89-1-22 qtz B-1 U.O. 12*5 75 -12.8 130.5 218.1

255 Sample No. Host Origin Type Size μm Filling % Te Tm(ice) Thh Th(vapor) Tm(halite) mineral MA89-1-23 qtz A U.O. 15*2 75 101.5 MA89-1-24 qtz A U.O. 8*4 70 105.6 MA89-1-25 qtz A U.O. 5*4 60 -41.1 -21.3 -33.1 110.3 MA89-1-26 qtz B-2 U.O. 5*3 65 -51.0 -21.6 139.6 MA89-1-27 qtz B-2 U.O. 8*5 70 138.6 MA89-1-28 qtz A U.O. 7*4 75 178.9 MA89-1-29 qtz A U.O. 3*3 70 219.1 MA89-1-30 qtz A U.O. 3*2 75 -6.3 222.2 MA89-1-31 qtz B-1 U.O. 18*4 75 -59.8 -9.0 138.6 123.2 MA89-1-32 qtz A U.O. 8*4 65 -49.6 -9.1 110.5 MA89-1-33 qtz A U.O. 12*5 70 -51.1 -12.7 120.1 MA89-1-34 qtz A U.O. 7*5 70 -55.4 -9.5 56.8 MA89-1-35 qtz A U.O. 5*2 70 -13.3 158.6 MA89-1-36 qtz A U.O. 10*6 70 -49.1 -9.1 179.5 MA89-1-37 qtz A U.O. 5*4 70 -36.5 -10.0 MA89-1-38 qtz A U.O. 10*4 70 > 0 MA89-1-39 qtz A U.O. 10*4 70 -33.9 -11.2 153.9 MA89-1-40 qtz A U.O. 6*4 65 -12.5 130.8 MA89-1-41 qtz A U.O. 6*5 75 -35.1 -13.9 140.1 MA89-1-42 qtz A U.O. 12*6 75 -49.8 -11.8 110.9 MA89-1-43 qtz A U.O. 7*5 70 -10.0 135.4 MA89-1-44 qtz A U.O. 10*5 70 -36.8 > 0 116.5 MA89-1-45 qtz A U.O. 5*5 65 > 0 117.1 MA89-1-46 qtz A U.O. 12*8 75 -38.0 -11.5 75.6 MA89-1-47 qtz A U.O. 4*4 60 -35.1 -14.3 100.6 MA89-1-48 qtz A U.O. 10*4 65 -60.1 -18.7 117.5 MA89-1-49 qtz A U.O. 8*4 60 -34.1 119.9 MA89-1-50 qtz A U.O. 6*5 -20.7 111.4 MA89-1-51 qtz A U.O. 7*4 60 -20.1 95.7 MA89-1-52 qtz A U.O. 5*3 75 -39.6 -20.5 132.1 MA89-1-53 qtz A U.O. 10*3 70 -21.0 136.8 MA89-1-54 qtz A U.O. 10*4 70 -61.0 -33.5 146.7 MA89-1-55 qtz A U.O. 6*4 65 -59.4 -36.5 125.3 MA89-1-56 qtz A U.O. 6*3 70 -59.5 -36.3 108.4 MA89-1-57 qtz A U.O. 6*4 65 -55.1 -23.8 149.0

256 Sample No. Host Origin Type Size μm Filling % Te Tm(ice) Thh Th(vapor) Tm(halite) mineral MA89-1-58 qtz A U.O. 15*8 70 -62.5 -38.3 140.5 MA89-1-59 qtz A U.O. 12*8 75 -66.2 -38.3 88.3 MA89-1-60 qtz A U.O. 5*4 80 -80.0 -36.5 104.4 MA89-1-61 qtz A U.O. 9*5 60 -57.8 -35.2 130.0 MA89-1-62 qtz A U.O. 5*3 65 -55.3 -33.5 128.9 MA89-1-63 qtz B-1 U.O. 8*4 65 -54.0 -34.1 112.1 161.5 MA89-1-64 qtz B-1 U.O. 25*6 70 -58.6 -30.2 107.8 194.4 MA89-1-65 qtz A U.O. 12*7 70 -22.8 128.2 MA89-1-66 qtz B-2 U.O. 12*10 65 -63.1 -25.2 126.9 MA89-1-67 qtz A U.O. 10*6 65 -23.1 128.5 MA89-1-68 qtz A U.O. 5*5 60 -54.0 -31.1 122.1 MA89-1-69 qtz B-2 U.O. 8*5 60 -30.0 96.1 MA89-1-70 qtz C U.O. 20*10 70 -66.7 -37.2 201.1 375.1 MA89-1-71 qtz B-1 U.O. 8*4 65 -55.7 -38.1 138.5 190.7 MA89-1-72 qtz A U.O. 12*4 70 -65.5 -41.1 134.5 MA89-1-73 qtz B-1 U.O. 7*3 75 -56.1 -16.7 138.5 191.6 MA89-1-74 qtz A U.O. 7*4 70 -51.0 -36.0 121.6 MA89-1-75 qtz A U.O. 12*7 70 -54.1 -18.5 165.4 MA89-1-76 qtz A U.O. 12*5 80 -51.0 -16.4 154.0 MA89-1-77 qtz B-2 U.O. 7*7 65 -50.0 -23.8 127.5 MA89-1-78 qtz A U.O. 8*4 70 -21.0 -5.4 155.5 MA89-1-79 qtz A U.O. 9*4 70 -51.3 -24.1 153.4 MA89-1-80 qtz A U.O. 4*3 55 -28.1 -7.6 286.1 MA89-1-82 qtz A U.O. 9*3 70 -51.0 -21.5 177.1 MA89-1-83 qtz A U.O. 14*8 65 -51.5 -26.3 195.7 MA89-1-84 qtz A U.O. 15*6 65 -24.6 153.6 MA89-1-85 qtz A U.O. 8*4 65 -52.1 -23.3 175.1 MA89-1-86 qtz A U.O. 10*8 70 -42.0 -23.1 178.4 MA89-1-87 qtz A U.O. 12*4 70 -64.1 -27.4 178.6 MA89-1-88 qtz A U.O. 8*5 70 -31.4 141.3 MA89-1-89 qtz A U.O. 25*10 65 -68.7 -41.8 206.1 MA89-1-90 qtz A U.O. 28*12 70 -61.0 -31.0 184.5 MA89-1-91 qtz A U.O. 10*5 55 -41.2 -15.4 159.9 MA89-1-92 qtz A U.O. 9*7 55 -49.1 -22.5 -29.8 159.7 MA89-1-93 qtz C U.O. 15*9 50 -72.1 -35.6 -58.9 142.3 390

257 Sample No. Host Origin Type Size μm Filling % Te Tm(ice) Thh Th(vapor) Tm(halite) mineral MA89-1-94 qtz A U.O. 6*5 75 -21.2 136.3 MA89-1-95 qtz A U.O. 9*7 65 -34.3 -23.7 133.8 MA89-1-96 qtz A U.O. 10*8 65 -31.1 -19.7 113.7 MA89-1-97 qtz A U.O. 10*9 60 -66.7 -25.4 -43.8 123.2 MA89-1-98 qtz A U.O. 7*5 65 -41.2 -18.6 134.1 MA89-1-99 qtz A U.O. 9*6 60 -57.8 -19.5 135.3 MA89-1-100 qtz A U.O. 10*8 70 -55.1 -25.6 130.5 MA89-1-101 qtz B-1 U.O. 10*9 70 -53.9 -35.2 -35.2 137.9 400.2 MA89-1-102 qtz A U.O. 15*7 60 -63.8 -18.7 170.2 MA89-1-103 qtz A U.O. 15*10 70 -39.8 -21.0 147.1 MA89-1-104 qtz A U.O. 14*8 70 -76.7 -22.9 155.1 MA89-1-105 qtz A U.O. 13*10 75 -46.2 -22.0 136.5 MA89-1-106 qtz A U.O. 14*7 65 -43.1 -21.8 137.2 MA89-1-107 qtz A U.O. 8*6 70 -46.2 -23.8 119.7 MA89-1-108 qtz B-1 U.O. 10*9 60 -66.7 -28.1 143.2 313.1 MA89-1-109 qtz A U.O. 9*8 70 -45.7 -26.5 139.2 MA89-1-110 qtz A U.O. 8*6 55 -53.2 -26.0 142.1 MA89-1-111 qtz A U.O. 12*4 60 -20.0 157.8 MA89-1-112 qtz A U.O. 9*8 60 -63.1 -30.3 -46.2 135.6 MA89-1-113 qtz A U.O. 12*10 70 -87.4 -26.1 -46.7 110.5 MA89-1-114 qtz A U.O. 9*5 60 -59.8 -22.8 -37.1 138.7 MA89-1-115 qtz A U.O. 7*6 60 -46.4 -26.7 106.7 MA89-1-116 qtz A U.O. 12*5 65 -21.2 127.6 MA89-1-117 qtz A U.O. 9*4 60 -56.7 -30.1 -49.7 108.6 MA89-1-118 qtz A U.O. 9*6 60 -57.5 -24.2 -38.6 152.1 MA89-1-119 qtz A U.O. 14*10 60 -23.7 -7.8 156.7 MA89-1-120 qtz A U.O. 6*5 60 -63.1 -23.7 -47.5 110 MA89-1-121 qtz A U.O. 12*6 65 -63.2 -27.1 -44.6 108.7 MA89-1-122 qtz A U.O. 10*8 60 -55.6 -18.9 -42.3 125.7 MA89-1-123 qtz A U.O. 8*5 65 -14.7 115.6 MA89-1-124 qtz A U.O. 10*4 65 -34.1 -8.2 171.5 MA89-1-125 qtz A U.O. 9*8 60 -7.6 195.1 MA89-1-126 qtz A U.O. 8*6 60 -41.5 -15.7 86.1 MA89-1-127 qtz A U.O. 9*6 70 -51.1 -26.0 -35.6 82.4 MA89-1-128 qtz A U.O. 9*6 65 -60.6 -19.3 115.6

258 Sample No. Host Origin Type Size μm Filling % Te Tm(ice) Thh Th(vapor) Tm(halite) mineral MA89-1-129 qtz A U.O. 9*5 60 -45.0 -14.6 128.9 MA89-1-130 qtz A U.O. 10*6 60 -46.5 -19.8 159.1 MA89-1-131 qtz A U.O. 13*5 70 -48.7 -19.6 116.1 MA89-1-132 qtz A U.O. 7*6 65 -47.1 -17.2 117.2 MA89-1-133 qtz B-1 U.O. 15*10 50 -65 -33.1 -43.2 154.6 374.2 MA89-5-1 cal A U.O. 10*10 70 -17.1 -0.6 141.7 MA89-5-2 cal A U.O. 15*6 70 -11.3 -2.0 114.1 MA89-5-3 cal A U.O. 20*4 65 153.8 MA89-5-4 cal A U.O. 7*3 70 -12.1 -1.2 147.8 MA89-5-5 cal A U.O. 20*8 65 -34.1 -14.5 161.6 MA89-5-6 cal A U.O. 15*6 65 -19.1 -1.3 197.1 MA89-5-7 cal A U.O. 15*7 60 -43.2 -15.0 155.1 MA89-5-8 cal A U.O. 10*5 60 -62.1 -30.7 177.7 MA89-5-9 cal A U.O. 7*2 70 -39.8 -19.2 161.5 MA89-5-10 cal A U.O. 7*3 65 130.8 MA89-5-11 cal A U.O. 20*10 80 132.5 MA89-5-12 cal A U.O. 15*8 80 165.3 MA89-5-13 cal A U.O. 9*4 75 136.1 MA89-5-14 cal A U.O. 15*5 75 147.6 MA89-5-15 cal A U.O. 20*10 80 137.1 MA89-5-16 cal A U.O. 8*4 65 135.4 MA89-5-17 cal A U.O. 15*4 70 145.8 MA89-5-18 cal A U.O. 6*2 70 149.5 MA89-5-19 cal A U.O. 12*8 70 154.0 MA89-5-20 cal A U.O. 10*6 70 154.1 MA89-5-21 cal A U.O. 12*5 70 154.1 MA89-5-22 cal A U.O. 12*5 70 143.5 MA89-5-23 cal A U.O. 30*20 70 161.2 MA89-5-24 cal B-2 U.O. 10*4 70 153.0 MA89-5-25 cal A U.O. 15*8 75 152.7 MA89-5-26 cal A U.O. 9*5 70 176.4 MA89-5-27 cal A U.O. 14*8 70 158.2 MA89-5-28 cal A U.O. 7*4 70 163.4 MA89-5-29 cal A U.O. 6*4 65 193.6 MA89-5-30 cal A U.O. 4*3 65 194.7

259 Sample No. Host Origin Type Size μm Filling % Te Tm(ice) Thh Th(vapor) Tm(halite) mineral MA89-5-31 cal A U.O. 5*4 65 197.7 MA89-5-32 cal A U.O. 10*5 65 157.1 MA89-5-33 cal A U.O. 14*8 65 157.4 MA89-5-34 cal A U.O. 4*4 60 176.1 MA89-5-35 cal B-2 U.O. 9*8 60 218.9 MA89-5-36 cal B-2 U.O. 4*3 60 187.5 MA89-5-37 cal A U.O. 5*4 60 -25.6 -4.0 128.4 MA89-5-38 cal A U.O. 15*8 70 -22.1 -11.4 128.1 MA89-5-39 cal A U.O. 15*4 75 146.1 MA89-5-40 cal A U.O. 16*5 75 -50.0 135.6 MA89-5-41 cal A U.O. 8*4 70 > 0 153.2 MA89-5-42 cal A U.O. 10*7 65 -24.2 -11.6 110.0 MA89-5-43 cal A U.O. 10*5 70 -27.7 -11.5 142.0 MA89-5-44 cal A U.O. 8*4 70 -21.0 -12.1 141.5 MA89-5-45 cal A U.O. 5*2 70 -2.6 133.4 MA89-5-46 cal A U.O. 10*3 70 -12.1 -7.7 136.0 MA89-5-47 cal A U.O. 7*3 70 125.1 MA89-5-48 cal A U.O. 6*3 70 -0.5 142.1 MA89-5-49 cal A U.O. 7*4 70 -0.4 130.5 MA89-5-50 cal A U.O. 8*3 70 -1.0 140.1 MA89-5-51 cal A U.O. 8*3 70 -1.2 145.1 MA27-5-1 cal A U.O. 12*8 75 -65.7 -23.4 129.1 MA27-5-2 cal A U.O. 12*7 70 -35.6 -19.4 117.4 MA27-5-3 cal A U.O. 14*8 70 -45.3 -21.6 133.1 MA27-5-4 cal A U.O. 15*4 70 -57.5 -24.2 144.7 MA27-5-5 cal A U.O. 12*5 65 -65.3 -23.1 135.5 MA27-5-6 cal A U.O. 12*4 65 -22.0 107.4 MA27-5-7 cal A U.O. 18*10 75 -69.3 -34.5 117.1 MA27-5-8 cal A U.O. 8*5 70 -45.6 -26.7 110.2 MA27-5-9 cal A U.O. 11*8 70 -73.1 -44.7 145.8 MA27-5-10 cal A U.O. 15*4 70 -75.1 -44.5 137.6 MA27-5-11 cal A U.O. 8*5 65 -67.1 -45.8 147.3 MA27-5-12 cal A U.O. 14*8 70 -70.0 -37.4 153.2 MA27-5-13 cal A U.O. 8*5 70 -69.1 -49.2 154.6 MA27-5-14 cal A U.O. 9*5 70 -70.3 -46.3 150.8

260 Sample No. Host Origin Type Size μm Filling % Te Tm(ice) Thh Th(vapor) Tm(halite) mineral MA27-5-15 cal A U.O. 10*5 70 -64.8 -28.8 120.0 MA27-5-16 cal A U.O. 12*5 70 -60.1 -37 141.1 MA27-5-17 cal A U.O. 12*4 70 -61.0 -34.5 124.1 MA27-5-18 cal A U.O. 8*7 65 -65.1 -30.5 120.3 MA27-5-19 cal A U.O. 14*4 70 -56 -28.1 143.6 MA27-5-20 cal A U.O. 7*5 65 -45.7 -25.6 140.1 MA27-5-21 cal A U.O. 7*5 65 -48.9 -25.0 141.2 MA27-5-22 cal A U.O. 7*4 70 -47.1 -24.8 139.7 MA17-2-1 cal A U.O. 10*6 80 -51.0 -27.0 87.4 MA17-2-2 cal A U.O. 15*10 75 -47.8 -35.6 106.8 MA17-2-3 cal A U.O. 10*5 75 -7.1 168.9 MA17-2-5 cal A U.O. 5*3 70 -7.9 105.7 MA17-2-6 cal A U.O. 7*3 70 -43.0 -8.0 158.6 MA17-2-7 cal B-2 U.O. 17*8 65 -60.1 -33.5 182.1 MA17-2-8 cal A U.O. 18*9 70 -53.6 -38.4 126.9 MA17-2-9 cal A U.O. 9*3 65 -62.7 -38.9 115.2 MA17-2-10 cal A U.O. 10*4 70 -59.1 -27.0 180.1 MA17-2-11 cal A U.O. 10*5 70 -65.3 -50.4 128.1 MA17-2-12 cal A U.O. 12*5 75 -52.7 -27.1 115.6 MA17-2-13 cal A U.O. 12*5 70 -51.0 -25.0 123.1 MA17-2-14 cal A U.O. 7*4 70 -17.0 -6.5 144.1 MA17-2-15 cal A U.O. 12*4 70 -47.9 -21.0 138.7 MA17-2-16 cal A U.O. 12*7 75 -55.8 -15.7 117.5 MA17-2-17 cal A U.O. 20*10 75 -48.0 -18.1 135.6 MA17-2-18 cal A U.O. 12*5 75 -44.4 -16.2 120.1 MA17-2-19 cal A U.O. 8*4 75 -49 -19.3 112.1 MA17-2-20 cal A U.O. 5*4 65 -44.5 -10.0 126.0 MA17-2-21 cal A U.O. 9*4 65 -46.7 -18.9 110.0 MA17-2-22 cal A U.O. 12*8 75 -46.2 -11.2 121.2 MA17-2-23 cal A U.O. 8*4 70 -43.4 -11.5 123.1 MA17-2-24 cal A U.O. 11*4 70 -52.0 -15.3 122.9 MA17-2-25 cal A U.O. 13*4 70 -56.4 -15.2 117.9 MA17-2-26 cal A U.O. 15*10 70 -46.5 -17.9 119.3 MA17-2-27 cal B-2 U.O. 12*5 70 -75.6 -49.7 87.5 MA17-2-28 cal A U.O. 18*10 70 -39.7 -28.7 122.7

261 Sample No. Host Origin Type Size μm Filling % Te Tm(ice) Thh Th(vapor) Tm(halite) mineral MA17-2-29 cal A U.O. 15*5 70 -43.5 -26.3 123.0 MA17-2-30 cal A U.O. 10*5 70 -59.8 -27.5 122.9 MA17-2-31 cal A U.O. 6*4 65 -46.5 -24.7 115.0 MA17-2-32 cal A U.O. 10*4 70 -60.1 -26.9 123.3 MA17-2-33 cal A U.O. 12*7 70 -46.1 -24.0 145.6 MA17-2-34 cal B-2 U.O. 25*12 60 -52.1 -22.1 210.1 MA17-2-35 cal A U.O. 6*3 65 -42.1 -21.0 147.8 MA17-2-36 cal A U.O. 13*6 60 -52.4 -36.9 138.9 MA17-2-37 cal A U.O. 10*7 70 -60.8 -42.9 130.5 MA17-2-38 cal A U.O. 6*4 65 -45.0 -27.1 156.5 MA17-2-39 cal A U.O. 6*4 65 -32.5 138.7 MA17-2-40 cal A U.O. 5*4 60 -28.1 135.6

262

Appendix D

LA-TOF-ICP-MS Database

Sample No. Li B Na Mg Al P K Ca Sc Ti V

Mina Justa Magnetite-pyrite stage (Stage J-V): high temperature and high salinity, Na-Ca-contained MA89-1-93 375 3664 113724 19 1118 n.d. 1475 22809 187 n.d. n.d. MA89-1-101 n.d. 363 119202 81 n.d. n.d. 640 11211 81 208 n.d. MA89-1-108 n.d. 132 100139 273 565 3407 912 12204 177 1187 184 MA89-1-133 3 n.d. 117735 14 20 n.d. 1985 3622 18 n.d. 1 Mina Justa Cu-mineralization stage (Stage J-VI): low temperature and high salinity, Ca-Na contained (Stage J-VI-1) MA89-1-94 193 2361 78868 80 1278 n.d. 1575 19398 260 90 29 MA89-1-95* 87 1462 90027 56 624 1320 816 11106 n.d. 61 38 MA89-1-98* 593 1901 75637 57 n.d. n.d. 2775 12699 6 n.d. 9 MA89-1-99 340 2347 65793 58 1306 6672 708 34227 46 2260 104 MA89-1-102 4 638 77396 74 136 630 774 10124 n.d. n.d. n.d. MA89-1-103* 98 590 78487 42 1299 2901 1409 19148 122 941 128 MA89-1-104 n.d. n.d. 92265 24 n.d. n.d. 350 3821 n.d. n.d. 8 MA89-1-105 1089 386 73305 75 2141 4702 2653 32707 38 n.d. 230 MA89-1-106 118 n.d. 80155 97 846 1776 1305 19915 6 816 39 MA89-1-109 459 1374 77380 222 5565 1085 7892 45052 200 1430 339 MA89-1-110 451 n.d. 70320 216 13998 22293 4155 55198 364 n.d. 161 MA89-1-111 879 n.d. 78152 305 11552 2918 4774 15082 21 2797 175 MA89-1-112 431 267 108843 99 1799 751 1073 6701 210 791 47 MA89-1-113 n.d. 26 94873 23 n.d. n.d. 794 12892 n.d. n.d. n.d. MA89-1-114 11 149 89848 38 144 2025 2366 7496 7 655 32 MA89-1-115 57 100 99298 43 327 198 1004 7691 8 n.d. n.d. MA89-1-116 777 3156 75195 203 1553 8486 138 25806 426 5783 130 MA89-1-118 528 2040 85972 25 6099 4425 94 20304 18 20694 108 MA89-1-120 876 n.d. 79990 192 7779 8944 196 43003 703 21680 169 MA89-1-121 633 n.d. 80134 n.d. 3626 n.d. n.d. 28288 n.d. 13751 n.d. MA89-1-122 60 1371 62840 76 3773 n.d. 250 36450 81 17308 151 MA89-1-123 n.d. 1874 50218 n.d. 2984 n.d. 643 36181 196 n.d. 154 MA89-1-126* 200 440 57800 114 3438 313 79 28633 135 605 36 MA89-1-127 463 145 79125 57 3564 712 1490 39887 119 107 71 MA89-1-128 670 462 57563 540 2981 167 1168 47720 6 n.d. 127 MA89-1-131 314 n.d. 63289 45 444 2295 n.d. 39039 17 n.d. n.d. Mina Justa Cu-mineralization stage (Stage J-VI): low temperature and low salinity, Na-contained (Stage J-VI-2) MA89-1-119 397 n.d. 38479 14 n.d. 384 52 9525 85 6373 3

263

Sample No. Cr Mn Fe Co Ni Cu Zn As Rb Sr Y Zr

Mina Justa Magnetite-pyrite stage (Stage J-V): high temperature and high salinity, Na-Ca-contained MA89-1-93 185 97 5721 33 67 n.d. 175 46 n.d. 106 n.d. 27 MA89-1-101 n.d. 254 1639 12 n.d. n.d. 55 n.d. n.d. 6 4 n.d. MA89-1-108 n.d. 277 4007 59 464 n.d. 205 235 103 133 77 144 MA89-1-133 34 306 2762 n.d. 144 n.d. 67 n.d. 17 46 n.d. 3 Mina Justa Cu-mineralization stage (Stage J-VI): low temperature and high salinity, Ca-Na-contained (Stage J-VI-1) MA89-1-94 163 154 2380 29 542 83 515 177 91 52 28 71 MA89-1-95* 56 74 1049 49 263 103 223 14 69 42 9 52 MA89-1-98* 14 142 1996 n.d. 388 n.d. 159 583 57 48 15 80 MA89-1-99 10 237 1838 58 235 79 374 385 99 61 5 65 MA89-1-102 204 92 2253 27 87 196 105 54 18 37 7 5 MA89-1-103* 521 118 4905 64 164 107 252 292 74 80 19 19 MA89-1-104 n.d. 88 1101 7 35 n.d. 14 48 n.d. 13 n.d. n.d. MA89-1-105 210 410 15510 80 496 n.d. 704 550 n.d. 38 n.d. n.d. MA89-1-106 228 217 870 35 156 19 178 65 34 42 8 22 MA89-1-109 2085 937 14023 94 445 291 888 162 141 221 67 157 MA89-1-110 n.d. 379 10333 218 938 187 1049 81 204 242 165 97 MA89-1-111 1537 531 4562 167 875 144 462 425 76 108 48 49 MA89-1-112 155 220 1047 48 12 82 179 48 43 48 32 22 MA89-1-113 23 189 6700 n.d. 30 5 66 26 16 89 n.d. n.d. MA89-1-114 17 325 5443 14 115 32 127 72 27 45 n.d. 27 MA89-1-115 87 279 3856 5 n.d. n.d. 173 n.d. 2 14 4 32 MA89-1-116 1116 2297 22301 153 1693 225 817 591 69 52 56 124 MA89-1-118 n.d. n.d. 2865 11 n.d. 98 493 249 n.d. 123 38 5 MA89-1-120 5515 n.d. 101 141 1130 94 705 243 107 120 n.d. 254 MA89-1-121 n.d 113 5350 32 n.d. 94 929 130 9 106 30 n.d. MA89-1-122 n.d. 97 3479 n.d. n.d. 70 561 n.d. n.d. 29 n.d. 53 MA89-1-123 n.d. 494 7947 n.d. 339 63 437 502 42 120 n.d. 96 MA89-1-126* 796 50 6057 62 648 114 660 276 106 43 58 117 MA89-1-127 n.d. 299 4285 106 298 128 777 130 35 97 n.d. 72 MA89-1-128 470 225 18458 72 111 162 1013 70 n.d. 114 7 75 MA89-1-131 n.d. 21 2452 n.d. n.d. 231 691 149 n.d. 31 n.d n.d. Mina Justa Cu-mineralization stage (Stage J-VI): low temperature and low salinity, Na-contained (Stage J-VI-2) MA89-1-119 n.d. 158 485 39 233 n.d. 336 n.d. n.d. 52 2 50

264

Sample No. Nb Mo Ag Cd Sn Sb Ba Ce Nd Sm Eu Gd

Mina Justa Magnetite-pyrite stage (Stage J-V): high temperature and high salinity, Na-Ca-contained MA89-1-93 n.d. n.d. n.d. n.d. 26 20 95 14 n.d. n.d. 5 15 MA89-1-101 18 n.d. 11 n.d. n.d. n.d. n.d. 9 15 n.d. n.d. n.d. MA89-1-108 28 227 16 n.d. 52 88 289 47 n.d. 75 n.d. 38 MA89-1-133 3 27 2 104 8 n.d. 44 4 5 6 4 2 Mina Justa Cu-mineralization stage (Stage J-VI): low temperature and high salinity, Ca-Na-contained (Stage J-VI-1) MA89-1-94 20 461 9 125 62 103 275 31 78 15 44 78 MA89-1-95* 23 188 n.d. 206 59 22 198 n.d. 20 61 17 58 MA89-1-98* 28 n.d. 11 17 9 83 n.d. n.d. n.d. 36 n.d. 5 MA89-1-99 14 7 67 95 46 26 148 21 35 50 21 21 MA89-1-102 14 12 7 28 n.d. 16 42 n.d. n.d. n.d. 3 26 MA89-1-103* 32 49 86 208 20 88 132 10 26 20 10 78 MA89-1-104 n.d. 1 3 n.d. 5 8 18 n.d. n.d. 2 n.d. n.d. MA89-1-105 8 147 2 289 24 6 352 13 20 n.d. n.d. 84 MA89-1-106 6 4 n.d. 110 9 38 82 18 30 12 2 54 MA89-1-109 59 403 185 272 148 82 531 38 120 151 89 152 MA89-1-110 146 635 211 778 326 508 494 90 242 n.d. 80 742 MA89-1-111 3 152 241 n.d. 277 178 478 60 104 121 n.d. 363 MA89-1-112 26 37 47 137 64 101 40 9 58 9 12 42 MA89-1-113 1 18 n.d. 23 n.d. 6 26 8 n.d. n.d. 5 n.d. MA89-1-114 19 37 26 174 26 10 62 4 11 4 2 14 MA89-1-115 1 n.d. n.d. n.d. 23 n.d. 37 n.d. 4 n.d. 1 14 MA89-1-116 106 331 139 419 102 243 542 28 79 71 45 122 MA89-1-118 85 n.d. n.d. 1162 72 26 232 14 15 71 19 n.d. MA89-1-120 n.d. 421 n.d. 706 135 184 253 50 162 137 60 n.d. MA89-1-121 39 n.d. n.d. 605 185 n.d. 129 16 2 71 n.d. 59 MA89-1-122 n.d. n.d. n.d. 126 n.d. n.d. n.d. n.d. 10 22 n.d. n.d. MA89-1-123 n.d. 271 57 n.d. 57 2 125 n.d. n.d. 8 n.d. n.d. MA89-1-126* 97 78 145 537 111 43 123 7 40 29 33 40 MA89-1-127 22 92 45 92 168 17 149 25 1 23 12 20 MA89-1-128 n.d. n.d. 144 n.d. 380 90 161 4 6 46 16 n.d. MA89-1-131 n.d. n.d. n.d. n.d. 126 15 109 n.d. n.d. n.d. 23 5 Mina Justa Cu-mineralization stage (Stage J-VI): low temperature and low salinity, Na-contained (Stage J-VI-2) MA89-1-119 30 26 20 244 22 n.d. 37 7 7 11 20 72

265

Sample No. Tb Dy Ho Er Tm Yb Lu W Au Pb Th U

Mina Justa Magnetite-pyrite stage (Stage J-V): high temperature and high salinity, Na-Ca-contained MA89-1-93 13 13 n.d. 3 n.d. 13 1 108 9 49 n.d. 16 MA89-1-101 3 n.d. n.d. n.d. n.d. 6 n.d. n.d. n.d. 26 n.d. 8 MA89-1-108 16 20 26 116 8 166 40 271 122 81 51 60 MA89-1-133 3 4 7 3 1 4 4 9 6 n.d. 3 n.d. Mina Justa Cu-mineralization stage (Stage J-VI): low temperature and high salinity, Ca-Na-contained (Stage J-VI-1) MA89-1-94 39 52 33 142 18 24 20 85 n.d. 36 14 10 MA89-1-95* 3 13 6 6 7 16 n.d. 129 71 24 13 20 MA89-1-98* 5 136 13 47 17 61 10 161 126 57 60 n.d. MA89-1-99 12 76 15 114 2 41 13 104 83 49 20 24 MA89-1-102 n.d. n.d. n.d. n.d. 2 4 1 26 34 8 n.d. n.d. MA89-1-103* 12 28 12 29 6 67 14 149 76 35 6 10 MA89-1-104 1 n.d. 1 2 3 4 n.d. n.d. 5 8 n.d. 4 MA89-1-105 24 59 7 98 12 54 22 n.d. n.d. 56 n.d. n.d. MA89-1-106 6 24 n.d. 8 8 36 n.d. 52 3 54 3 n.d. MA89-1-109 53 78 47 86 35 220 35 127 261 55 42 71 MA89-1-110 31 61 58 287 31 32 115 430 293 173 213 103 MA89-1-111 n.d. 15 8 125 33 330 34 301 n.d. 60 n.d. 5 MA89-1-112 2 151 14 64 34 86 n.d. 61 42 17 17 35 MA89-1-113 1 n.d. n.d. n.d. n.d. n.d. 1 n.d. n.d. 28 1 6 MA89-1-114 2 26 n.d. 13 2 15 6 18 22 36 3 4 MA89-1-115 n.d. n.d. 2 n.d. 2 n.d. 4 n.d. n.d. 16 n.d. n.d. MA89-1-116 41 168 7 130 11 50 22 156 75 463 28 31 MA89-1-118 23 123 19 75 n.d. 57 29 131 151 66 39 13 MA89-1-120 n.d. 103 22 33 n.d. n.d. 55 213 253 112 23 33 MA89-1-121 n.d. n.d. 4 25 n.d. n.d. 21 n.d. 108 19 n.d. n.d. MA89-1-122 n.d. n.d. 28 n.d. n.d. n.d. 2 105 n.d. 4 n.d. n.d. MA89-1-123 n.d. 29 n.d. n.d. 39 n.d. n.d. n.d. n.d. 6 4 12 MA89-1-126* 26 76 38 64 13 33 7 128 99 102 7 17 MA89-1-127 15 48 20 n.d. 7 n.d. 8 81 73 333 5 22 MA89-1-128 16 19 22 62 n.d. 123 6 76 n.d. 657 10 17 MA89-1-131 n.d. 8 3 n.d. n.d. n.d. n.d. n.d. 13 264 n.d. n.d. Mina Justa Cu-mineralization stage (Stage J-VI): low temperature and low salinity, Na-contained (Stage J-VI-2) MA89-1-119 3 32 11 11 19 35 10 15 35 5 6 14

266