Metallogenesis for the Boléo and Cananea Copper Mining Districts: a Contribution to the Understanding of Copper Ore Deposits in Northwestern México

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METALLOGENESIS FOR THE BOLÉO AND CANANEA COPPER MINING DISTRICTS: A CONTRIBUTION TO THE UNDERSTANDING OF COPPER ORE DEPOSITS IN NORTHWESTERN MÉXICO

Rafael Eduardo Del Rio Salas

A Dissertation Submitted to the Faculty of the

DEPARTMENT OF GEOSCIENCES

In Partial Fulfillment of the Requirements for the Degree of

DOCTOR OF PHILOSOPHY

In the Graduate College THE UNIVERSITY OF ARIZONA 2011

THE UNIVERSITY OF ARIZONA GRADUATE COLLEGE

As members of the Dissertation Committee, we certify that we have read the dissertation

prepared by Rafael Eduardo Del Rio Salas

entitled METALLOGENESIS FOR THE BOLÉO AND CANANEA COPPER MINING DISTRICTS: A CONTRIBUTION TO THE UNDERSTANDING OF COPPER ORE DEPOSITS IN NORTHWESTERN MÉXICO and recommend that it be accepted as fulfilling the dissertation requirement for the

Degree of Doctor of Philosophy

______Date: 04/13/2011 Joaquín Ruiz

______Date: 04/13/2011 George Gehrels

______Date: 04/13/2011 Eric Seedorff

______Date: 04/13/2011 Christopher J. Eastoe

______Date: 04/13/2011 Lucas Ochoa Landín

Final approval and acceptance of this dissertation is contingent upon the candidate’s submission of the final copies of the dissertation to the Graduate College.

I hereby certify that I have read this dissertation prepared under my direction and recommend that it be accepted as fulfilling the dissertation requirement.

______Date: 04/13/2011 Dissertation Director: Joaquín Ruiz 3

STATEMENT BY AUTHOR

This dissertation has been submitted in partial fulfillment of requirements for an advanced degree at the University of Arizona and is deposited in the University Library to be made available to borrowers under rules of the Library.

Brief quotations from this dissertation are allowable without special permission, provided that accurate acknowledgment of source is made. Requests for permission for extended quotation from or reproduction of this manuscript in whole or in part may be granted by the copyright holder.

SIGNED: Rafael Eduardo Del Rio Salas

ACKNOWLEDGMENTS

This research was accomplished thanks to the support and the scholarship (No. 166600) from the Consejo Nacional de Ciencia y Tecnología of México (CONACYT). I would like to express my gratitude to my advisor Joaquín Ruiz for his friendship, the support, the understanding, the patience, and the opportunity of being a member of his research group. I would like to thank also to my co-advisor Lucas Ochoa Landín for the support, the friendship, the patience, and for sharing his knowledge in geology, ore deposits, and life. I am really thankful to Christopher Eastoe for his friendship, understanding, for sharing his knowledge and discussions in geology and stable isotopes, and the opportunity of gaining experience in the Environmental Isotope Laboratory. I am really grateful to John Chesley, for sharing his friendship, the support, the thinking and different perspectives, and his experience in the analytical and chemical work, and life. I also express my gratitude to Mark Baker for his friendship, for the analytical support, for sharing his knowledge and experience in the laboratory. I am thankful to Jason Kirk and Ryan Mathur for the collaboration in this research. I thank to Diana Meza Figueroa for her friendship, the enthusiasm, the financial and analytical support. I also thank to Francisco Paz Moreno for the friendship and sharing his knowledge in geology. Thanks to David Dettman for the opportunity of gaining experience in the carbonate isotope laboratory. I also thank Ben McElhaney for his friendship and collaboration in the carbonate isotope lab. Thanks to Oscar Talavera for his friendship, enthusiasm, analytical support, and the support and sharing experience in geology and life fields. I thank and appreciate the support of the of my committee members. Thanks to George Gehrels for the help with the U-Pb zircon analysis, Eric Seedorff for his help and sharing field experience, Lucas Ochoa, Chris Eastoe, and Joaquín Ruiz. I thank my officemates for sharing the good and bad stuff (Victor Valencia, Fernando Barra, Sergio Salgado, Alyson Thiboudeau, Lisa Molofsky, Francisco Quintanar, Luis Zuñiga, Mauricio Ibañez). I thank Sergio Salgado Souto for his friendship, support, always being there, and listening. I thank Francisco Quintanar (el jefe) for his friendship, support, and sharing his knowledge in ore deposits and mining exploration. I thank Alyson Thiboudeau for her friendship, the team spirit in the office and lab, and enthusiasm and good vibes. I thank also Lisa Molofsky for her good vibes too and her friendship and good mood. I thank Fernando Barra for the early help and support during my research. Thanks to my colleagues for their support and collaboration, especially Hugo Zuñiga, Luis Zuñiga, Facundo Cazares, and Jose Alberto Campillo. Also thanks to Aimee Orci, Ramses Tarazón, Omar Noriega, Isidro Flores, Genaro Verdugo, Jesus Vidal, Hector Hinojosa, Saul Peña. I am grateful to Patricia Acuña for the love and support during this long stage. I thank my family for the support (Enrique, Graciela, Juan, Cristina, Lalo, el niño, Andrea, and los abuelos). Finally I like to thank my mother Martha Elena and my sister Martika for their love and unlimited support. Thanks to my wife Veronica Moreno for the love, encouragement, for believe in me and the unlimited support during the last and hard stage. 5

DEDICATION

Con amor y cariño a

mis abuelos María y Enrique,

Martha Salas y Martha Del Rio,

Verónica Moreno,

Mara y Leo Del Rio Moreno

TABLE OF CONTENTS

LIST OF FIGURES ...... 9 LIST OF TABLES ...... 14 ABSTRACT ...... 16 CHAPTER 1: INTRODUCTION TO THE PRESENT STUDY ...... 18 CHAPTER 2: ORIGIN OF THE MINERALIZATION OF THE CU-CO-ZN BOLÉO DISTRICT, BAJA CALIFORNIA SUR: INSIGHT FROM ISOTOPIC METHODS .... 28 2.1, ABSTRACT ...... 28 2.2, INTRODUCTION ...... 29 2.3, GEOLOGICAL SETTING ...... 31 2.4, MINERALIZATION FROM THE BOLÉO DISTRICT ...... 35 2.4.1, The Boléo district ...... 36 2.4.2, Neptuno area ...... 39 2.4.3, Lucifer deposit ...... 40 2.4.4, Paragenesis of the Boléo district ...... 41 2.5, CONCEPCIÓN PENINSULA MN DEPOSITS ...... 44 2.6, ANALYTICAL PROCEDURES ...... 46 2.6.1, Rare earths and other trace elements ...... 46 2.6.2, Sulfur and oxygen isotopes ...... 47 2.6.3, Oxygen and carbon isotopes ...... 48 2.6.4, Copper isotopes ...... 49 2.6.5, Lead and strontium isotopes ...... 50 2.7, RESULTS ...... 51 2.7.1, Major and trace elements in manganese oxides ...... 51 2.7.2, Sulfur and oxygen isotopes ...... 53 2.7.3, Oxygen and carbon isotopes ...... 54 2.7.4, Copper isotopes ...... 54 2.7.5, Pb and Sr isotopes ...... 55 2.8, DISCUSSION ...... 57 2.8.1, Mineralization and hydrothermal activity ...... 57 2.8.2, Mineralization age ...... 60 2.8.3, Geochemistry of manganese oxides from the Boléo mantos ...... 62 2.8.4, Sulfur and oxygen isotopes ...... 67 2.8.5, Oxygen and carbon isotopes ...... 70 2.8.6, Pb and Sr isotopes ...... 75 2.8.7, Source of metals ...... 77 2.8.8, Copper isotopes in nature...... 80 2.8.9, Copper isotope fractionation ...... 83 7

TABLE OF CONTENTS - Continued

2.8.10, Copper isotope data for the Boléo mineralization ...... 84 2.8.11, Metal budget ...... 88 2.8.12, Origin of the Cu-Co-Zn and Mn mineralization in Santa Rosalía region ...... 89 2.9, CONCLUSIONS...... 92

CHAPTER 3: GEOLOGY, GEOCHEMISTRY, AND U-PB GEOCHRONOLOGY OF THE MARIQUITA PORPHYRY COPPER AND LUCY CU-MO DEPOSITS, CANANEA DISTRICT, MÉXICO ...... 127 3.1, ABSTRACT ...... 127 3.2, INTRODUCTION ...... 128 3.3, PREVIOUS STUDIES OF THE CANANEA DISTRICT ...... 129 3.4, REGIONAL GEOLOGIC SETTING ...... 130 3.4.1, Cananea district stratigraphy ...... 130 3.4.2, Structural geology ...... 134 3.5, GEOLOGY OF MARIQUITA DEPOSIT ...... 135 3.5.1, Geology ...... 135 3.5.2, Structure ...... 138 3.5.3, Alteration and mineralization ...... 138 3.6, GEOLOGY OF LUCY DEPOSIT ...... 143 3.7, ANALYTICAL PROCEDURES ...... 144 3.7.1, Stable isotopes ...... 144 3.7.2, U-Pb method ...... 146 3.8, RESULTS ...... 148 3.8.1, Oxygen and hydrogen isotopes ...... 148 3.8.2, Sulfur isotopes ...... 149 3.8.3, U-Pb ages ...... 149 3.9, DISCUSSION ...... 150 3.9.1, Alteration and mineralization ...... 152 3.9.2, Supergene events ...... 153 3.9.3, Sulfur isotopes ...... 154 3.9.4, Isotope geothermometry ...... 155 3.9.5, Ore fluids ...... 156 3.9.6, Magmatic-hydrothermal geochronology ...... 159 3.10, CONCLUSIONS...... 161

TABLE OF CONTENTS - Continued

CHAPTER 4: GEOCHRONOLOGY OF THE PORPHYRY COPPER AND RELATED DEPOSITS IN THE CANANEA DISTRICT, NORTHWESTERN MEXICO ...... 182 4.1, ABSTRACT ...... 182 4.2, INTRODUCTION ...... 183 4.3, CANANEA DISTRICT ...... 184 4.3.1, Cananea district geology ...... 185 4.4, ANALYTICAL PROCEDURES ...... 189 4.4.1, Re-Os method ...... 189 4.4.2, U-Pb method ...... 191 4.5, RESULTS ...... 193 4.5.1, Re-Os geochronological data ...... 193 4.5.2, U-Pb zircon data ...... 194 4.6, DISCUSSION ...... 195 4.6.1, Molybdenite mineralization events in the Cananea district ...... 196 4.6.2, Mineralizing porphyritic intrusions ...... 197 4.6.3, Southeastern migration of the mineralization ...... 199 4.7, CONCLUSIONS...... 201

REFERENCES ...... 219

APPENDIX A: GEOLOGY, GEOCHEMISTRY AND RE–OS SYSTEMATICS OF MANGANESE DEPOSITS FROM THE SANTA ROSALÍA BASIN AND ADJACENT AREAS IN BAJA CALIFORNIA SUR, MÉXICO...... 242

LIST OF FIGURES

Figure 1.1, Map showing the Basin and Range (BR) and the Sierra Madre Occidental (SMO) provinces in northwestern Mexico. The western, central, and eastern belts represent the different metallogenetic provinces for the orogenic gold deposits (squares), porphyry copper deposits (circles), and the epithermal deposits (triangles) respectively. The small province in Baja California Sur represents the copper (open diamonds) and manganese (solid diamonds) Miocene mineralization. The dashed line represents the Cananea Lineament (Hollister 1978) ...... 27

Figure 2.1, Simplified geological map showing the location of Cu-Co-Zn Boléo district, and other copper and manganese deposits in Baja California Sur, Mexico (modified after Conly et al. 2006). Localities: (1) Lucifer, (2) Neptuno area, (3) Boléo, (4) San Alberto, (5) Rosario, (6) Caracol, (7) Gavilán, (8) Mantitas, (9) Trinidad, (10) Pilares, (11) Minitas, (12) Santa Teresa, (13) Santa Rosa, (14) Las Delicias ...... 96

Figure 2.2, Generalized stratigraphic column of the Santa Rosalía region (modified after Conly et al. 2006). The Cu-Co-Zn mantos and Mn oxide mineralization are located at the beginning of each sedimentary cycle; less important mantos are indicated by italics. Age data from (1) Schmidt 1975, (2) Sawlan and Smith 1984, (3) Holt et al. 2000, and (4) Conly 2003...... 97

Figure 2.3, Schematic geological section showing the major faults that affected the ASL volcanic rocks, previous to the formation of the Santa Rosalía basin and the deposition of the Boléo Formation (after Wilson and Rocha 1955). Symbols as in Fig. 2...... 98

Figure 2.4, (a) Cross-section and (b) schematic stratigraphic column of the Lucifer manganese oxide deposit, northern the Boléo district ...... 99

Figure 2.5, Paragenetic sequence for the Boléo district (after Conly 2003) ...... 100

Figure 2.6, NASC-normalized REE patterns for the manganese oxide mineralization from the Boléo region and the Mn deposits from Concepción Peninsula. The REE data of Mn deposits from Concepción Peninsula taken from Rodríguez Díaz (2009). REE data from hydrothermal and hydrogenetic fields, and average fossil and modern deposits from Usui and Someya (1997) ...... 101

Figure 2.7, Trace element discrimination diagram for manganese oxides deposits between supergene (or hydrogenous) and hydrothermal deposits (Nicholson 1992) .... 102

LIST OF FIGURES - Continued

Figure 2.8, Histogram showing the sulfur isotope data for the sulfide and sulfate samples from the Boléo district (Ortlieb and Colleta 1984; Ochoa Landín 1998; Conly et al., 2006; present study) ...... 103

Figure 2.9, Oxygen and sulfur isotope data of sulfates from the Boléo district. Samples located within the dotted field correspond to gypsum veinlets cross-cutting the indicated mantos; sulfate data outside dotted field from manto 3 and 4 taken from Conly et al. (2006). Gray square represents Miocene evaporite deposits precipitated from Miocene seawater (Claypool et al., 1980) ...... 104

Figure 2.10, Carbon and oxygen isotope data of carbonates from the Boléo district, Lucifer, and Gavilán deposits. The gray-shaded square represents the field for marine carbonates precipitated in equilibrium with seawater followed Conly et al. (2006) ..... 105

Figure 2.11, 207Pb/204Pb vs 206Pb/204Pb diagram after Rollinson (1993). The mantle reservoirs of Zindler and Hart (1986) are as follows: DM - depleted mantle; BSE - bulk silicate earth; EMI and EMII - enriched mantle. EMII coincides with the field of oceanic pelagic sediments; PREMA - prevalent mantle composition; MORB - mid- ocean ridge basalts. Note that the Miocene volcanic rocks from the Boléo district plots within the lower continental crust ...... 106

Figure 2.12, Lead isotope diagram showing in detail the Miocene volcanic rocks from the Santa Rosalía region. NHRL - Northern Hemisphere Reference Line; MORB gray field ...... 107

Figure 2.13, Lead isotope diagram showing the isotope data for the copper and manganese mineralization from the different mineralized mantos and the manganese deposits around Santa Rosalia. Also is shown the lead data fields for the Miocene volcanic rocks from the Boléo district and the peninsular batholith. NHRL - Northern Hemisphere Reference Line ...... 108

Figure 2.14, 206Pb/204Pb vs. 87Sr/86Sr diagram for the rocks and Cu and Mn mineralization from the Boléo district ...... 109

Figure 2.15, Copper isotope variations for continental and marine environments. (*) Present study; (1) Vance et al., 2008; (2) Jiang et al., 2002; (3) Asael et al., 2009; (4) Markl et al., 2006; (5) Botfield 1999; (6) Li et al., 2009; (7 ) Larson et al., 2003; (8) Maréchal et al., 1999; (9) Larson et al., 2003; (10) Graham et al., 2004; (11) Mathur et al., 2005; (12) Mathur et al., 2009 ...... 110

LIST OF FIGURES – Continued

Figure 2.16, Schematic model for the copper isotope fractionation in the mineralized mantos from the Boléo district. a) Mineralizing fluids ascended along the fault system and encountered the biogenic pyrite reduced horizons within the fine facies at the beginning of the sedimentary cycle of the Boléo Formation. b) Proposed systematic for the copper isotope fractionation. (1) Seawater and meteoric water; (2) Oxidation of the sulfide ores producing fluids with heavier δ65Cu; (3) Mineralization relicts with lighter δ65Cu; (4) Continental flow of meteoric water through the pre-Gulf of California ...... 111

Figure 2.17, Histogram showing the copper isotope data for the Cu-Co-Zn Boléo district and adjacent manganese oxide localities. a) Copper data for secondary copper mineralization and manganese oxides from manto 2; b) Copper isotope data for Copper isotope data for Cu-sulfides, secondary copper mineralization, and Mn mineralization from the Boléo manto 3; c) Copper isotope data for the Mn mineralization from the Boléo manto 4; d) Copper isotope data for manganese oxides from Lucifer, Neptuno area, and Gavilán deposits ...... 112

Figure 2.18, Model showing the mineralization for the Boléo district. (a) Schematic section showing the fine-grained sediments of the first sedimentary cycle of the Boléo Formation, and the formation of the mineralized manto 4. (b) formation of the second sedimentary cycle and manto 3 ...... 113

Figure 3.1, Regional map showing the porphyry copper deposit belt northwestern Mexico and southeastern Arizona ...... 164

Figure 3.2, Geologic map of the Cananea district modified after Wodzicki (1995) and Noguez-Alcántara 2008 ...... 165

Figure 3.3, Stratigraphic columns of (a) the Cananea district (modified after Wodzicki 1995), and (b) the Mariquita deposit. Geochronologic data: (1) Anderson and Silver 1977; (2) Wodzicki 1995; (3) Cox et al., 2000; (4) Noguez-Alcántara 2008; (5) Carreón-Pallares 2002; (6) Valencia et al., 2006; (7) McDowell and Clabaugh 1979; McDowell et al., 1997; (9) Varela 1972; (10) Damon and Mauger 1966; (11) Wodzicki 2001 ...... 166

Figure 3.4, a) Geologic map of Mariquita PCD area; b) structural framework of Mariquita area showing the faulting stages ...... 167

Figure 3.5, Map showing the different alteration zones from Mariquita area ...... 168

Figure 3.6, Schematic cross section of Lucy deposit showing the different alteration zones ...... 169 12

LIST OF FIGURES – Continued

Figure 3.7, Histogram showing the sulfur isotope data from Mariquita (gray columns) and Lucy (black bars) deposits ...... 170

Figure 3.8, U-Pb zircon ages from the mineralizing porphyritic units in the Mariquita PCD ...... 171

Figure 3.9, U-Pb zircon age from the Cuitaca granodiorite that hosts the Cu-Mo mineralization at Lucy deposit ...... 172

Figure 3.10, Schematic cross section showing the Cuitaca half-graben filled by the sediments of Sonora Group. Also shown are the Mariquita and Lucy deposits; in this case, Lucy is located further north, and a porphyritic body is shown in dotted line as the mineralizing system, even though it has not been seen (see text) ...... 173

Figure 3.11, Oxygen and hydrogen isotope composition of water in equilibrium involved during the hydrothermal stages from Mariquita and Lucy deposits. (1) Present study oxygen isotope data of quartz from stage I and III from Mariquita and Lucy deposits; (2) oxygen isotope rage of magmatic water from Cananea district (Wodzicki 2001); (3) Primary magmatic water field from Taylor (1974); (4) Water in arcs and crustal melts from Taylor (1992); (5) volcanic fumaroles and vapor from convergent volcanoes (Giggenbach 1992). Winter and summer water samples from Sierra Vista Arizona (Coes and Pool, 2007) ...... 174

Figure 4.1, Map showing the Basin and Range and the Sierra Madre Occidental provinces in northwestern Mexico. The western, central, and eastern belts represent the different metallogenetic provinces for the orogenic gold deposits (squares), porphyry copper deposits (circles), and the epithermal deposits (triangles) respectively ...... 202

Figure 4.2, Simplified geological map of northern Sonora and southern Arizona showing the Cananea Lineament and the porphyry copper deposits along the trace (modified after Hollister 1978) ...... 203

Figure 4.3, Geologic map of the Cananea district modified after Wodzicki (1995) and Noguez-Alcántara 2008 ...... 204

Figure 4.4, Stratigraphic column of the Cananea district (modified after Wodzicki 1995). Geochronologic data: (1) Anderson and Silver 1977; (2) Wodzicki 1995; (3) Cox et al., 2000; (4) Noguez-Alcántara 2008; (5) Carreón-Pallares 2002; (6) Valencia et al., 2006; (7) McDowell and Clabaugh 1979; McDowell et al., 1997; (9) Varela 1972; (10) Damon and Mauger 1966; (11) Wodzicki 2001 ...... 205

LIST OF FIGURES - Continued

Figure 4.5, Evolution of molybdenite mineralization and mineralizing porphyritic pulses in the Cananea district. Geochronological data from: 1) Present study; 2) Valencia et al., 2006; 3) Barra et al., 2005 ...... 206

Figure 4.6, U-Pb zircon ages for the mineralizing porphyries of the Cananea mine; a) and b) granodiorite porphyries, c) quartz monzonite porphyry, and d) monzodiorite porphyry ...... 207

Figure 4.7, U-Pb zircon ages from the Alacrán mineralizing porphyry (a), and the hosting rocks from the Pilar deposit (b and c) ...... 208

LIST OF TABLES

Table 2.2, Rare earth element concentration in ppm from manganese oxides from the Boleo district and adjacent areas in Baja California Sur, México ...... 116

Table 2.3, Sulfur isotope data for the sulfide phases in the Cu-Co-Zn Boléo District...... 118

Table 2.4, Sulfur and oxygen isotope data for sulfate phases in the Cu-Co-Zn Boléo District, Baja California Sur ...... 119

Table 2.5, Carbon and oxygen isotope data for the Cu-Co-Zn Boléo District and adjacent areas ...... 120

Table 2.6, Copper stable isotope data from the Boléo district, Lucifer, and the Gavilán deposit in Concepcion peninsula ...... 123

Table 2.7, Lead and strontium isotope data of the Boléo district and adjacent deposits ...... 124

Table 2.8, Rare earth element average concentrations from hydrothermal and hydrogenous manganese oxide deposits including those from Baja California ...... 126

Table 3.1, Oxygen and hydrogen stable isotope data for the Mariquita PCD and Cu- Mo Lucy deposit ...... 175

Table 3.2, Oxygen and sulfur stable isotopes and fluid inclusion data for the Mariquita and Lucy deposits ...... 177

Table 3.3, U-Pb geochronologic analyses of the mineralizing porphyry 104 from Mariquita PCD ...... 179

Table 3.4, U-Pb geochronologic analyses of the mineralizing porphyry 604 from Mariquita PCD ...... 180

Table 3.5, U-Pb geochronologic analyses of the Cuitaca granodiorite from Lucy deposit ...... 181

Table 4.1, General geologic features of the Porphyry copper deposits from the Cananea district, northwestern Mexico...... 209 15

LIST OF TABLES - Continued

Table 4.2, Re-Os geochronologic data of molybdenite mineralization from the Pilar, Mariquita, and Lucy copper deposits from de Cananea district ...... 210

Table 4.3, U-Pb geochronologic analyses of granodiorite porphyry from Cananea mine...... 211

Table 4.4, U-Pb geochronologic analyses of granodiorite porphyry from Cananea mine...... 212

Table 4.5, U-Pb geochronologic analyses of quartz monzonite porphyry from Cananea mine ...... 213

Table 4.6, U-Pb geochronologic analyses of monzodiorite porphyry from Cananea mine...... 214

Table 4.7, U-Pb geochronologic analyses of the mineralizing porphyry from Alacrán PCD ...... 215

Table 4.8, U-Pb geochronologic analyses of the granodiorite hosting rock in the Pilar Cu deposit ...... 216

Table 4.9, U-Pb geochronologic analyses of the granodiorite hosting rock in the Pilar Cu deposit ...... 217

Table 4.10, Geochronologic compilation of the lithology from the Cananea District, Sonora ...... 218

ABSTRACT

Northwestern Mexico is characterized by different metallogenic provinces that are

included along the Basin and Range, the Sierra Madre Occidental, and the Baja

California geological provinces. With the purpose of contribute to the current understanding of the mineralizing processes, the present study focused on two important copper metallogenic provinces: the Cananea Porphyry District in Sonora, and the

Sediment-hosted Stratiform Copper- and Mn-deposits in Baja California Sur.

The U-Pb zircon ages from the mineralizing porphyries from Cananea district suggest a continued magmatic activity period of ~6 Ma. Also suggests a period of ~20

Ma for the entire magmatic activity in the district. The Re-Os molybdenite ages demonstrate five well-constrained mineralization events in the district; the main mineralization is constrained over a short period of time (~4 Ma). The new molybdenite- age from the Pilar deposit documents the oldest mineralizing pulse, suggesting possibly the initiation of the Laramide mineralization in northern Sonora.

A detailed study of Mariquita porphyry Cu and Lucy Cu-Mo deposits in the

Cananea district was performed. Four hydrothermal stages were defined in Mariquita, whereas a single hydrothermal pulse characterizes Lucy. Emplacement depths between 1-

1.2 km, and temperatures between 430-380ºC characterized the mineralization from

Mariquita, whereas deeper emplacement depths and higher mineralization temperatures characterized Lucy. The stable isotope systematic and fluid inclusion data determined that the mineralizing fluids in Mariquita deposit are essentially magmatic during the earlier hydrothermal stages, whereas the last stage is the mixing between magmatic and 17

winter meteoric-waters. The mineralizing fluids from Lucy deposit are magmatic in

origin.

A comprehensive study was performed in the Cu-Co-Zn-Mn mineralization of the

Boléo District, and Mn-oxide mineralization along the eastern coast Baja California Sur.

The REE and trace element in the Mn-oxides demonstrated the exhalative nature of the

mineralizing hydrothermal fluids, and exclude the hydrogenous nature. The stable isotope systematic in ore and gangue minerals, along with the Cu-isotope data helped to decipher the nature of mineralizing and non-mineralizing fluids. The application of Pb, Sr and Re-

Os isotope systems was applied to constrain the nature of the fluids involved during the mineralization processes and that the metal sources.

CHAPTER 1: INTRODUCTION TO THE PRESENT STUDY

The Mexican territory is known for a long history of mining, which has been practiced since pre-colonial times. The country is still among the world’s largest metal producers. The mineral wealth of Mexico, world-class and world famous mines and mining districts have played a crucial role in the history of Mexico, during the Spanish conquest and after Mexican independence (Ordoñez Cortés 2009).

The mineral endowment of Mexico includes a variety of metals and industrial minerals. These commodities are distributed among a wide variety of ore deposits and different mineralization styles. The mineralization ages in Mexico range from the

Proterozoic to the Pleistocene. The variety of mineralization styles has led different authors to classify the ore deposits according to metallogenic provinces and mineralizing epochs (Gonzalez-Reyna 1956; Salas 1976; Damon 1978; Campa and Coney 1983;

Staude and Barton 2001; Camprubí 2009; Clark 2009).

Northwestern Mexico is characterized by different metallogenic provinces and several mineralization styles in the Basin and Range and the Sierra Madre Occidental geological provinces (Titley 2001; Staude and Barton 2001). Precious and base metal epithermal deposits are spatially associated with the Sierra Madre Occidental (SMO) province. They are mostly of Mid-Tertiary age, although scarce older epithermal systems have been reported (Clark et al., 1982; Montaño 1988; Bennett 1993; Pérez-Segura 1993;

Staude 1995; Staude and Barton 2001; Zawada et al,. 2001; Valencia et al., 2005;

Camprubí and Albinson 2006). Some of the outstanding epithermal deposits in 19

northwestern Mexico are El Tigre, Mulatos, Magallanes, Ocampo, Palmarejo, etc.

(Knowling 1975; Montaño 1988; Arriaga et al., 1993; Staude 1995; Murray et al., 2008).

The Basin and Range province includes both orogenic gold and porphyry copper

deposit belts (Fig. 1.1). The orogenic gold deposits in Sonora are located along a NW-SE

belt about 300 km long and 90 km wide. The gold mineralization is mainly hosted in

metamorphic rocks, mostly of greenschist facies. The host rock ages range from

Proterozoic to Eocene, and the mineralization age appears to be restricted to the Laramide

orogeny around 60 Ma (Silberman et al., 1988; Albinson 1989; Pérez-Segura et al., 1996;

Araiza Martínez 1998; de la Garza et al. 1998; Araux-Sanchéz 2000; Iriondo and

Atkinson 2000; Quintanar Ruiz 2008). The prominent mineral deposits from this orogenic belt in Sonora are Quitovac, La Choya, La Herradura, Tajitos-San Francisco, El

Chanate and Sierra Pinta (Durgin and Teran 1996; Summers et al. 1998; Araux Sánchez

2000; Iriondo and Atkinson 2000; Pérez-Segura et al., 1996; Quintanar Ruiz 2008).

The porphyry copper deposit (PCD) province in the North American southwest is

well known and is a world-class copper producer (Fig. 1.1). The Cananea and Nacozari

PCDs in Sonora represent the southeastern extension of the province (Titley 1982), also

known as the “great cluster” (Keith and Swan 1996). The Cananea and Nacozari districts

are the most important porphyry copper mining districts in Mexico.

The Cananea district lies along a ~350 km northwest-trending regional line

defined as the Cananea Lineament (Hollister 1978), from the Silver Bell PCD at the

northwestern end in Arizona, through La Caridad PCD at the southeastern end in Sonora

(Fig. 1.1). The Cananea district includes the famous world-class porphyry copper 20

mineralization of the Cananea mine, but also includes other smaller PCDs along with other mineralization styles like those of skarn, manto, and breccia pipe deposits. This mineralization occurs along a NW-SE belt, and some of the outstanding mineral deposits in the district include those from the Cananea mine, Maria, Mariquita, Milpillas, Alacrán,

Puertecitos, Lucy and Pilar (Velasco 1966; Meinert 1982; Wodzicki 1995; Virtue 1996; de la Garza et al., 2003; Arellano 2004; Ochoa Landín et al., 2007; Noguez-Alcántara

2008; Aponte-Barrera 2009).

A younger mineralization episode is recorded on the eastern coast of Baja

California Sur. These deposits are hosted within Miocene sedimentary and volcanic rocks. These deposits have a hydrothermal signature and are related to magmatic activity associated with the opening of the Gulf of California. The most important are the Boléo

Cu-Co-Zn and Mn deposits, along with Lucifer, San Alberto, Gavilán, Las Delicias,

Santa Teresa, Santa Rosa and Minitas (Noble 1950; Wilson and Rocha 1955; González-

Reyna 1956; Ochoa Landín 1998; Conly 2003; Conly et al 2006; Camprubí et al., 2008;

Rodríguez Díaz et al., 2010).

In order to contribute with the understanding of the mineralizing processes in northwestern Mexico, the present study focuses on the porphyry copper mineralization from the Cananea district and the Cu-Mn mineralization from the Boléo region and surroundings areas. This work contributes with new geochemical data (REE, trace elements, stable and radiogenic isotopes), and focuses on improving the understanding of metallogenesis in these important mining districts. This work is divided in two sections.

The first section (Chapters II and Appendix A) comprises the metallogenesis of the Boléo 21

mining district and surroundings mineralizing regions in Baja California Sur, and the

second section (Chapters III and IV) deals with the metallogenesis of the Cananea mining district in Sonora.

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Figure 1.1. The Basin and Range (BR) and the Sierra Madre Occidental (SMO) provinces in northwestern Mexico. The western, central, and eastern belts represent the different metallogenetic provinces for the orogenic gold deposits (squares), porphyry copper deposits (circles), and the epithermal deposits (triangles) respectively. The small province in Baja California Sur represents the copper (open diamonds) and manganese (solid diamonds) Miocene mineralization. The dashed line represents the Cananea Lineament (Hollister 1978). 28

CHAPTER 2: ORIGIN OF THE MINERALIZATION OF THE CU-CO-ZN BOLÉO DISTRICT, BAJA CALIFORNIA SUR: INSIGHT FROM ISOTOPIC METHODS

2.1, ABSTRACT

The Cu-Co-Zn Boléo district is the only district known in northwestern Mexico that belongs to the sediment-hosted stratiform copper deposits type. The Cu-Co-Zn mineralization is hosted within the Boléo Formation, a marine-clastic sequence of Upper

Miocene age. The mineralization occurs as stratiform horizons or mantos, mostly associated to the fine-grained sediments of the Boléo Formation. In addition to the Cu-

Co-Zn mineralization, Mn-oxide mineralization is related to the mineralized mantos, and also is present as isolated deposits along the eastern coast of Baja California Sur.

The sulfur isotope data in ore and gangue minerals indicate that marine sulfate

(pore sediments and gypsum member) is the most important source of sulfur for the mineralization processes in the Boléo mantos. The C and O isotope data in the ore and gangue carbonate indicate the mixing of two end-members for the source of carbonate in the district (seawater and meteoric water-organic material); the mixing of fluids is recorded at each stratigraphic level, suggesting a systematic mixing occurring at in mineralized manto.

The δ65Cu data of the Cu and Mn mineralization from the Boléo district and

adjacent areas is slightly negative and around zero per mil. The copper and manganese

mineralization from the Boléo mantos display slightly differences in the δ65Cu values.

The copper mineralization from the mantos is characterized by slightly lower Cu isotope

compositions (–1.62 to +0.13‰), while slightly higher isotope compositions are found in

the manganese mineralization in the mantos (–0.73 to +0.16‰). Similar δ65Cu values are

found in the manganese mineralization from Lucifer (–0.86 to –0.17‰) and Gavilán

(+0.48‰) deposits. The copper isotope fractionation takes place in each manto, during the redox changes within the lacustrine conditions at the beginning of each sedimentary cycle of the Boléo Formation.

The lead isotope data demonstrate that metal sources for the Boléo mineralized mantos and the Mn mineralization along the eastern coast Baja California Sur is mostly the Miocene Andesite of Sierra Santa Lucia volcanic and the Peninsular batholith rocks.

The strontium isotope data indicate mixing between the Sierra Santa Lucia volcanic rocks and the Gypsum member of the Boléo Formation.

The geological observations and geochemical data indicate that hydrothermal and exhalative nature for the mineralization in the Boléo district and Mn deposits of the eastern coast of Baja California Sur, related to the Miocene rifting and the opening of the

Gulf of California.

2.2, INTRODUCTION

Sediment-hosted stratiform copper deposits (SSC) are a world-class mineral deposit type, economically important, and represent around 23% of the world’s copper production and known reserves (Singer 1995). These deposits can also represent important sources of Ag, Co, Pb, Zn, U, and a few have important concentrations of Au and platinum group elements (Hitzman et al., 2005). Classic example of this deposit type includes the super-giant mineral deposits of the Central African Copper Belt, the

Kupferschiefer in Europe, and the Udokan in Siberia (Hitzman et al., 2005).

Although the Cu-Co-Zn Boléo district of Baja California Sur is not a super-giant

mineral deposit, it is the only district known in Mexico that belongs to the SSC type (Fig.

2.1). Several authors have studied the Boléo district and surrounding areas, and most of

them have described the basic geology, structure, tectonics, geochronology, and volcanic activity. Even though the mineralization has been exploited since the XIX century, few

studies have attempted to constrain the genesis of the Cu-Co-Zn mineralization (Wilson

and Rocha 1955; Ochoa Landín 1998; Conly et al., 2003). In addition to the Cu-Co-Zn

mineralization, there is a clear spatial and temporal relationship with manganese oxide

mineralization in the Boléo district, and elsewhere in the Santa Rosalía region.

Considering the geological context at the Boléo district (active tectonics and

proto-rift related basin), the transitional marine-continental environment, along with the oxidation-reduction processes within the basin, and the nature of the mineralization (the low temperature, several mantos at different stratigraphic levels), the Boléo district and surroundings areas offer an ideal opportunity to improve the understanding of the SSC deposit type.

The purpose of this study is to trace the metal sources of the primary mineralization and examine the chemistry of the ore-fluids involved in the primary and

secondary mineralization stages, along with the understanding of the different

geochemical processes during the evolution of the Cu-Co-Zn and Mn mineralization. In

order to elucidate the genesis of the mineralization, the present study uses the isotopic

systematics of Pb and Sr, along with Cu isotopes and the traditional stable isotopes (S, C,

and O). Finally, the geochemistry of the rare earth elements (REE) and other trace

elements in the manganese mineralization is used to characterize the genesis of the

manganese mineralization present at the different mineralized mantos in the Boléo

district and surroundings areas.

2.3, GEOLOGICAL SETTING

The Gulf of California formed as a result of continental rifting and the slow transfer of the Baja California Peninsula from North America to the Pacific plate

(Lonsdale 1989; Spencer and Normark 1989; Stock and Hodges 1989). The Santa Rosalía

basin is an incipient rift basin, formed as a result of northwest- to southeast-trending

continental rifting during Late Miocene during the Pre-Gulf of California (Karig and

Jensky 1972; Stock and Hodges 1989). This basin hosts the Cu-Co-Zn deposits from the

Boleo district as well as the manganese oxide deposits of the Lucifer deposit and adjacent

areas (Fig. 2.1; Del Rio Salas et al., 2008a). The basin is bounded to the north-northwest

by the Plio-Quaternary Tres Vírgenes volcanic field and La Reforma Caldera, and to the

west-southwest by the 24-12 Ma andesite of Sierra Santa Lucia (ASL) volcanic rocks, and to the east by the Gulf of California (Fig. 2.1).

The oldest rock in the Santa Rosalía area is biotite quartz-monzonite dated at 91.2

 2.1 Ma using K-Ar geochronology (Schmidt 1975). This intrusive rock corresponds to the southeastern extension of the Mesozoic Peninsula Batholith complex that represents the crystalline basement of Baja California (Gastil et al., 1975). The quartz-monzonite crops out locally in Las Palmas Creek and in La Reforma Caldera, 15 and 35 km north

and northwest of the town of Santa Rosalía, respectively, and it appears also to the south and to the west of the Concepción Peninsula (McFall 1968).

The ASL suite lies unconformably over the biotite quartz-monzonite; it is more than 1 km thick and consists mostly of andesite, basaltic andesite and basalt flows, tuffs, breccias, agglomerates and tuffaceous sandstones, predominantly of andesitic composition (Sawlan and Smith 1984). Previous K-Ar geochronology in the ASL volcanic rocks yielded ages between 24 and 13 Ma (Sawlan and Smith 1984), in agreement with K-Ar geochronological data from the volcanic rocks from Santa Rosalía region (Conly 2003; Conly et al., 2005).

The ASL suite is the result of the oblique subduction of the Farallon-Guadalupe plate under the North American plate along the western margin of Baja California

(Atwater 1989; Stock and Hodges 1989; Londsdale 1989). These volcanic rocks are medium-K calc-alkaline and are widely exposed through the Boléo district and

Concepción Peninsula (Sawlan and Smith 1984). The final stage of the arc volcanism is represented by the Santa Rosalía Dacite unit, which consists of lavas that erupted between 13 and 12 Ma (Conly 2003; Conly et al., 2005).

Subsequent volcanic activity records the transition from arc to rift volcanism associated with the initial opening of the Gulf of California (Conly 2003; Conly et al.,

2005). The rift-related suite was emplaced unconformably over the ASL volcanic rocks

(Fig. 2.2) and consists of three volcanic groups (Conly 2003; Conly et al., 2005): 1) the

11-9 Ma lava flows of the Boléo basalts and the basaltic andesites; 2) El Morro tuff, a 9-8

Ma felsic lapilli tuff to ignimbrite unit; and 3) the 9.5 to 7.7 Ma high-K andesitic lava

flow of the Cerro San Lucas unit.

The Miocene ASL and rift-related volcanic rocks in the Santa Rosalía region are

overlain by a series of sedimentary marine and non-marine formations including the

Boléo Formation (Fig. 2.2, Wilson and Rocha 1955). The thickness of the Boléo

Formation varies from 250 to 350 m. The formation has been divided into four members:

the basal conglomerate member, the limestone member followed by the gypsum member,

and the upper clastic member (Fig.2.2; Wilson and Rocha 1955). The basal conglomerate

has a maximum thickness of 10 m. This member is composed of angular to sub-angular

boulders and pebbles derived from the ASL volcanic rocks, supported by a brownish sandy matrix (Wilson and Rocha 1955). Locally, the basal conglomerate and the ASL suite are overlain by El Morro tuff, whose thickness varies from 0.1 to 6 m, and which consists mostly of pumice fragments (<1 to 3 cm in diameter), and minor lithic fragments probably from the ASL rocks and the Boléo basalts. Otherwise, where El Morro tuff is absent, the basal conglomerate is directly overlain by extensive, bedded, 0 to 4 m thick, reddish to brownish, locally fossiliferous marine limestone. The gypsum member occurs overlying either the limestone member or the ASL volcanic rocks. The gypsum member ranges in thickness from a few meters to 60 m, and forms massive to distinctly banded beds. The size and shape of these gypsum bodies are variable and range from large horizontal massive beds (hundreds of meters) to small lenses (tens of meters). Overlying

the gypsum is the clastic member, which consists of at least three well-organized, upward- coarsening sedimentary cycles, where each cycle ranges between 90 to 100 m in

thickness (Ochoa-Landín 1998). The clastic cycles consist of siltstones and sandstones at the base, grading to conglomerates towards the top (Fig. 2.2). The clastic section has been subdivided into five distinctive lithofacies, which correspond to distinct alluvial fan-delta deposits of a series of progradational episodes. Each deposit formed in response to a period of basin floor subsidence related to the initial stages of the opening of the Gulf of

California (Ochoa-Landín, 1998; Ochoa-Landín et al., 2000). Holt et al. (2000) constrained the age of the Boléo Formation between 7.1 Ma and 6.9 Ma at the base and from 6.3 to 6.1 Ma at the top. An important stratigraphic marker is present between the sedimentary cycles 3 and 2, and is locally known as ‘cinta colorada’. This unit consists of a thin layer (~0.5 m) of a coarse grained lithic tuff, containing coarse ash to lapilli, whose composition is mostly andesitic (Wilson and Rocha 1955). Additionally, pyroclastic volcanism of dacitic to rhyolitic composition is recorded during the deposition of the

Boléo Formation, particularly during the deposition of the finest facies (claystone- siltstone) at the beginning of each sedimentary cycle (Ochoa Landín 1998).

The Early to Middle Pliocene Gloria Formation (Ortlieb and Colleta 1984) lies unconformably over the Boléo Formation. The thickness of the Gloria Formation is estimated at ~60 m along the coastal area, and it thins and pinches out inland. The

Pliocene Gloria Formation grades inland from shallow marine to non-marine conglomerates and sandstones, which locally rest on a basal conglomerate (Wilson and

Rocha 1955). The Gloria Formation is unconformably overlain by the 20-30 m thick

Upper Pliocene Infierno Formation (Wilson and Rocha 1955), which consists of fossiliferous marine sandstone grading southwest to a continental conglomerate. The

Infierno Formation is overlain by the Pleistocene Santa Rosalía Formation (Wilson and

Rocha 1955), which has a thickness of 10-15 m. It consists of fossiliferous sandstones and non-marine conglomerates, grading to continental breccias and conglomerates landward (Ortlieb and Colleta 1984).

2.4, MINERALIZATION FROM THE BOLÉO DISTRICT

The copper mineralization in Santa Rosalía region was discovered in 1868 by a rancher, who found small, green, copper-bearing nodules or balls; hence the term “boléo”

(Bailes et al., 2001). After small scale mining production, a French mining company named the Compagnie du Boléo was founded in 1885, and operated until 1953. The following year operations were taken over by the Compañía Minera Santa Rosalía, S.A., jointly owned by Federal and State Governments until 1985. Subsequently, the bulk of the district was held in the Mexican Strategic National Mining Reserve until 1991.

Between 1992 and 1997, Minera Curator S.A. de C.V., a subsidiary of International

Curator Resources, Ltd., undertook an extensive exploration campaign (Bailes et al.,

2001). At the present time, the mineral concessions covering El Boléo deposit are 100% owned by Minera y Metalúrgica del Boleo S.A. de C.V., a Mexican company involved in mineral exploration and development and a wholly owned subsidiary of Baja Mining

Corp.

Early studies recognized the controversial nature of the Boléo deposit and provided a variety of genetic models. Wilson and Veytia (1949) and Wilson and Rocha

(1955) provided the first and most complete geologic maps of the Santa Rosalía region,

along with more detailed description of the Boléo Cu–Co–Zn and manganese oxide

mineralization. The hydrothermal activity in the Boléo district has been documented

decades ago in several reports and papers (Bouglise and Cumenge 1885; Tinoco 1885;

Fuchs 1886; Saladin 1892; Martinez and Servin 1896; Touwaide 1930; Bellanger 1931;

Peña 1931; Locke 1935; Wilson and Veytia 1949; Wilson and Rocha 1955; Nishihra

1957; Guilloux and Pélissonier 1974; Freiberg 1983; Ochoa Landín 1998; Conly 2003).

In addition, manganese mineralization has been described in detail at the Lucifer deposit

(Freiberg 1983), located northwest of the Boléo district. More manganese oxide mineralization has been described in the Concepción Peninsula and the Cerro

Mencenares volcanic center, southeast of the Boléo district (Fig. 2.1; Antúnez-Echegaray

1944; Noble 1950; Mapes-Vázquez 1956; González-Reyna 1956; McFall 1968).

Within the last decade, several studies have focused on constraining the nature

and genesis of the Cu-Co-Zn and the Mn mineralization at the Boléo district and the Mn

mineralization of the Concepcion Peninsula area (Ochoa-Landín 1998; Bailes et al.,

2001; Conly 2003; Canet et al., 2005a,b; Conly et al., 2006; Del Rio Salas et al., 2008a,b;

Camprubí et al., 2008). The following section provides a brief description of the ore

deposits of the Santa Rosalía region and Concepción Peninsula.

2.4.1, The Boléo district

The mineralization in the Boléo district consists of laterally extensive and

stratiform ore bodies (known as mantos) of disseminated Cu-Co-Zn sulfides and related

manganese oxides, constrained within the fine-grained facies (claystone and claystone

breccia) at the base of each cycle of the Boléo Formation (Fig. 2.2; Wilson and Rocha

1955; Ochoa-Landín 1998; Conly 2003). The “mantos” are numbered from 4 to 0, with

manto 4 being the lowest in the stratigraphic column and manto 0 the uppermost (Wilson

and Rocha 1955).

Manto 4 is commonly related to faults that affected the ASL volcanic rocks (Figs.

2.2 and 2.3). These fault zones also show evidence of Cu and Mn mineralization (Wilson and Rocha 1955; Ochoa-Landín 1998). Manto 4 occurs either above the limestone or the

ASL volcanic rocks, within the fine-grained facies of the first sedimentary cycle of the

Boléo formation, and consists of 1 m thick laminar calcareous mudstone overlain by a 2

m thick monomictic breccia with high Mn and Fe oxide content (Wilson and Rocha 1955;

Ochoa-Landín 1998). Locally, the Mn and Fe oxides are mixed with jasper and have

replaced the limestone. The Mn-oxide mineralization in manto 4 occurs directly over the

ASL volcanic rocks and shows dendritic textures within the fine-grained sediments of the

Boléo Formation. Manto 4 appears to correlate geologically and temporally with the

Lucifer manganese deposit (Wilson and Rocha 1955; Freiberg 1983), which is located at the margin of the Santa Rosalía basin.

Manto 3 is the more extensive and can be continuously traced over an area of 6 km x 3 km. Detailed mapping shows that manto 3 pinches out gulfward with a notable decrease in ore grade (Wilson and Rocha 1955). The fine-grained facies that host manto 3 consist of a 0.25 to 0.5 m thick interval of calcareous mudstone at the bottom and a 1 to

20 m thick of a fine-laminar claystone-siltstone interval at the top. The later contains a

chaotic breccia zone with claystone-siltstone fragments ranging in diameter from 1 to 5

cm, in a claystone-siltstone matrix (Ochoa-Landín 1998). The mineralization occurs

along the laminar structures of the brecciated fragments and rarely in the matrix (Ochoa-

Landín 1998). This manto is important because of its high content of copper minerals,

such as chalcocite, covellite, chalcopyrite, bornite, native copper, and minor cuprite

(Wilson and Rocha 1955; Pérez-Segura 1995; Ochoa-Landín 1998). Manganese oxides

are present above the copper-rich zone, usually as thin horizons along with chrysocolla,

as veinlets, and as small nodules. Locally, the manganese oxide mineralization in manto 3

occurs as nodules, with diameters ranging from 8 to 15 cm, hosted within the fine-grained

sediments.

Manto 2 is less extensive, and is also hosted within a sedimentary facies loke the

host of manto 3. The base of the fine-grained sequence consists of a 1 m thick mudstone,

and a 2.5 m thick breccia zone with siltstone-sandstone fragments in a siltstone-sandstone matrix. The mineralization consists of Cu-sulfides, mostly chalcocite, along with

disseminated pyrite. The manganese content is usually higher at the top of the

mineralized manto as seen in manto 3. Many of the ores in manto 2 are manganiferous

and ferruginous, and are associated with NW-SE silicified structures (Wilson and Rocha

1955).

A few manganiferous horizons with small quantities of copper are found between

mantos 1 and 2, and between mantos 2 and 3. These horizons are commonly

discontinuous, thin, and in general, low-grade (Wilson and Rocha 1955). Manto 1 is the

second most important manto after manto 3 as a producer, with ore grade ~4.5% Cu. This

manto is of the most extensive in the district but has been productive only in the

southeastern portion of the district (Wilson and Rocha 1955). Manto 0 is the least

mineralized of all the mantos in the Boléo district. It is commonly manganiferous and ferruginous, and the copper grade is as high as 1% Cu (Wilson and Rocha 1955). The manto thickness ranges between 1 and 1.5 m, and it grades inland to a conglomerate

(Touwaide 1930).

2.4.2, Neptuno area

The Neptuno area is located on a small hill 8 km northwest of the town of Santa

Rosalía, included near the boundary of the Boléo district, and has been exploited by small-scale miners using rudimentary methods (Wilson and Rocha 1955). The manganese mineralization correlates with manto 4. It is mostly located near the base of the first sedimentary cycle of the Boléo Formation and is hosted by siltstones and sandstones.

These manganese oxide bodies were deposited in lenticular basins (10×15 m) with depths

of ~2 m and directly on top of the ASL volcanic rocks. These lenticular basins were

geomorphic traps that favored the preservation of manganese oxides. In these basins, the

Mn oxides, Fe oxides, and jasper facies show a strong zonation similar to that observed at

the Lucifer deposit. Other manganese oxide outcrops are located directly above the

limestone, restricted to N–S fault structure depressions. Some manganese oxides with

botryoidal morphology are intercalated within the fine-grained sediments below manto 3.

In addition to the Mn mineralization in manto 4, the Neptuno area also has

evidence of mineralization corresponding to manto 3 just a few meters above manto 4

(Wilson and Rocha 1955). The mineralization in manto 3 consists of a thin horizon less than 30 cm thick, hosted within claystone-siltstones of the sedimentary cycle 2.

2.4.3, Lucifer deposit

The Lucifer deposit was the most important source of manganese in the Santa

Rosalía region. The Lucifer mine produced more than 300,000 tonnes with grades of over

40% Mn between 1941 and the 1950s (Wilson 1956). The stratigraphy in the Lucifer deposit is similar to the lower section of the Boléo district, except for the absence of the gypsum member. The ASL volcanic rocks in the Lucifer deposit are at least 600 m thick, and consist mostly of massive lava flows of andesitic, basaltic andesitic, and basaltic composition, and andesitic agglomerates. The limestone member rests unconformably above the ASL volcanic rocks (Freiberg 1983). This member is partly recrystallized with unidirectional fractures, with manganese oxides and jasper within notable stratification structures, replacing and filling open spaces within the limestone.

The Lucifer Mn deposit is a stratiform manto deposit hosted by the first sedimentary cycle of the Boléo Formation (manto 4, Fig. 2.2), although the ASL rocks are crosscut by 1 to 5 mm wide manganese oxide veinlets, filling a NW–SE fracture trend and localized inside a weak argillization zone as pointed out by Freiberg (1983).

The lower zone of the Lucifer deposit is composed of a fine-grained sandstone sequence that overlies the limestone and contains thin manganese oxides lenses of about

40 cm thick interbedded within the fine-grained sequence (Freiberg 1983). These thin manganese oxide horizons (pyrolusite and cryptomelane) are crosscut by silica veins 2 to

5 mm thick. The Mn horizon zone is overlain by a 10 m thick horizon of chaotic breccia,

composed of angular manganese oxide fragments consisting of pyrolusite, cryptomelane,

and todorokite. The angular fragments (1 to 5 cm in diameter) represent around 30% of

the horizon and are supported in a Mn-oxide matrix (Fig. 2.4). This unit is overlain by a 5

m thick horizon with brecciated lens composed mostly of angular manganese oxide

(cryptomelane, pyrolusite, and todorokite) fragments, minor Fe-oxides (hematite and

goethite), and jasper fragments, all ranging from 5 to 15 cm in diameter and supported in

a Mn-oxide matrix. This lens constitutes the richest manganese ore in the Lucifer deposit,

with about 40% Mn. This horizon is overlain by other brecciated lenses, which are

massive bodies composed mainly of brecciated jasper (Fig. 2.4). The jasper bodies

contain some Fe oxides such as goethite and hematite, and are overlain by a breccia with

jasper fragments and minor Mn-oxide fragments, both ranging from 5 to 20 cm in

diameter, supported by a manganese oxide matrix (cryptomelane, pyrolusite, and

todorokite) with jasper. The distal facies of the main Mn ore bodies in Lucifer have minor

manganese mineralization within fine-grained sediments (Fig. 2.4). The incipient Mn

mineralization consists of thin manganese oxide horizons (10–15 cm thick), associated

with clay that shows load structures caused by the boulders and pebbles of the

conglomerate member.

2.4.4, Paragenesis of the Boléo district

The paragenetic sequence for the ore and gangue minerals in the Boléo district is shown in Figure 2.5. Framboidal pyrite is the first Fe-sulfide recorded prior the ore

mineralization, possibly during the early stages of diagenesis, and occurs along the

sedimentary lamination planes within the fine-grained sediments of mantos 2 and 3. The

framboids are generally less than 20 µm, and are usually coated with or totally replaced

by copper sulfides (Ochoa Landín 1998; Conly 2003).

Chalcocite is the first ore mineral found replacing the framboidal pyrite from

mantos 2 and 3, and also occurs as granular aggregates, isolated euhedral grains, and

felted masses of tabular chalcocite (Conly 2003). Detailed petrographic coupled with

XRD determined that the chalcocite is a complex mixture of chalcocite, digenite and

djurleite (Echávarri and Pérez-Segura 1975). The continuous copper sulfide

mineralization resulted in the coeval precipitation of covellite and the subsequent

replacement of chalcocite (Conly 2003). Afterward, bornite replaces both chalcocite and

covellite, and consists of fine-grained aggregates and euhedral crystals (Conly 2003). The

last copper sulfide mineralization is recorded by chalcopyrite intergrowths along covellite

and chalcocite (Conly 2003). The presence of cobalt-bearing sulfides was reported as intergrowths within chalcocite (Bailes et al., 2001), while Conly (2003) only identified carrollite and linnaeite by XRD. Sphalerite is the only zinc-bearing sulfide reported

petrographically and by XRD (Echávarri and Pérez-Segura 1975; Conly 2003), and

occurs as intergrowths within the sulfides apparently at the final stage of the ore-bearing

mineralization (Fig. 2.5).

Manganese oxide mineralization within the Boléo formed during the

hydrothermal, late diagenetic, and supergene stages (Conly 2003). The primary

manganese oxides within the mantos occur as interlamination within the claystone or as

disseminated grains associated to sulfide facies, as small nodules, dendrites within sediments directly over the ASL rocks, replacing the limestone member, or as disorganized veins from 1 to 5 cm thick cross-cut the mantos (Conly 2003, Del Rio Salas

et al., 2008).

The supergene oxide copper mineralization mainly consists of chrysocolla,

malachite, and azurite, and minor atacamite, melaconite, and cuprite (Touwaide 1930;

Wilson and Rocha 1955; Conly 2003). Also, secondary copper mineralization within the

Boléo district accounts for the Cu-bearing oxychlorides as the locality type for boleite

[Pb26Ag9Cu24Cl62(OH)48] (Mallard and Cumenge, 1981), pseudoboleite

[Pb31Cu24Cl62(OH)48] (Friedle 1906), and cumengite [Pb21Cu20Cl42(OH)40] (Mallard

1893). The secondary copper mineralization is associated to claystone facies in the manto zones following stratification planes, or as filling open spaces within the brecciated sediments in the upper manto zones (Fig. 2.5). The secondary manganese mineralization mainly consists of hollandite group minerals, and occurs as dendritic growths to massive replacement associated with other supergene oxide or silicate phases (Conly 2003).

Gangue minerals include clay minerals, carbonates, barite-celestite, gypsum- anhydrite, zeolites and silica (Conly 2003). The principal clay mineral consists of montmorillonite, although variable amounts of smectite and saponite are reported

(Wilson and Rocha 1955; Bailes et al., 2001; Conly 2003).

Calcium carbonates are the dominant carbonates within the mantos, although siderite, dolomite, Mn-calcite, and rhodochrosite are reported in lesser amounts (Conly

2003). Calcite is present from the early diagenetic stage trough the supergenic stages.

Carbonates occur primarily as micritic to sparry fillings of pore spaces within clay-rich manto and as cements within coarser grain units (Conly 2003). Copper carbonates occur as secondary mineralization filling open spaces within the claystones and siltstone above the mantos.

Sulfates are present through early diagenetic to the supergene stage (Fig. 2.5).

Acicular grains of barite and celestite are present in the matrix of the breccias and are commonly obscured by clays and Fe oxides (Conly 2003). Gypsum occurs as veins cross- cutting the mantos 4, 3, and 2. These veins are 3 mm thick and are usually perpendicular to the mantos. Disseminated gypsum occurs within the mantos but is difficult to distinguish from clays and carbonates (Conly 2003).

2.5, CONCEPCIÓN PENINSULA MN DEPOSITS

Several localities with manganese mineralization are along the Concepción

Peninsula to the southern Mencenares volcanic field (Fig. 2.1; Antúnez-Echegaray 1944;

Noble 1950; González-Reyna 1956). The manganese mineralization is hosted within the volcanic and volcanoclastic rocks, and constrained to NW-SE fault systems, similar to that in the Santa Rosalía region.

The Gavilán deposit is located at the northeastern tip of the Concepción

Peninsula, 15 km east of Mulegé (Fig. 2.1). This is the most important manganese oxide deposit in this region, with probable resources of approximately 200 000 tons, and grades up to 55 wt% Mn (González-Reyna 1956). The host rocks in Concepción Peninsula consist of the Upper Oligocene to Middle Miocene Comondú Group (Hausback 1984).

These volcanic rocks essentially belong to the same volcanic arc of the ASL rocks

(Sawlan and Smith 1984), but several name discrepancies exist for these Tertiary volcanic rocks (Umhoefer et al., 2001, and references therein). The nomenclature used for the volcanic rocks from the Concepción Peninsula is from Hausback (1984), since some of the formations that host the manganese mineralization are time constrained.

The manganese mineralization is hosted by andesitic-basaltic lavas of the Pilares

Formation from the Upper Oligocene to Middle Miocene Comondú Group (Camprubí et al., 2008), and consists of a series of NW–SE-oriented manganese oxide veins, which crosscut highly fractured volcanic flows of the Pilares Fromation (Wilson and Veytia

1949; Noble 1950; Camprubí et al., 2008). Two types of veins can be distinguished: (1) veins with massive or laminated manganese oxides, and (2) laminated veins with dolomite and quartz, that might contain minor manganese oxides (Camprubí et al., 2008).

Most veins are 1 to 10 cm thick, although along fault zones, some veins can reach more than 0.5 m in thickness. These veins are between 2 to 3 m apart (Wilson and Veytia

1949). Also, the manganese mineralization occurs as stockwork with veinlets 1-12 cm thick consisting of pyrolusite with minor coronadite and romanechite, along with dolomite, barite, and vanadinite (Camprubí et al., 2008). Finally, the manganese mineralization also occurs as breccia matrix; the fragments consist of basaltic and basaltic-andesite lavas with an average diameter around 10 cm. The matrix assemblage is composed by coronadite and minor pyrolusite, and dolomite (Camprubí et al., 2008).

Less important manganese deposits are present along the Concepción peninsula

(Guadalupe, Minitas, Pilares, Trinidad, Santa Teresa, Azteca mine, etc), and most of

them share the similar geological features; the manganese mineralization consists mostly of pyrolusite and romanechite, grades ranging between 4 and 42 wt% Mn, occurring generally in NW-SE veins systems, hosted in volcanic and volcanoclastic rocks from the

Comondú Group (Camprubí et al., 2008). Only two manganese deposits (Santa Rosa and

San Juanico mines) show slight differences. The mineralization at Santa Rosa mine located south of the Concepción peninsula occurs along N-S vertical veins hosted in

Pliocene alluvial conglomerates, although the mineralization is also replacing the matrix of the same conglomerates (Camprubí et al., 2008; Rodríguez Díaz 2009; Rodríguez Díaz et al., 2010). The manganese mineralogy consists of botroidal romanechite or coronadite, along with opal, and phanerocrystalline barite (Camprubí et al., 2008). San Juanico mine is located south of Concepción peninsula near Cerro Mencenares volcanic filed; the mineralization is related to fault zones within the Comondú Group volcanic rocks and

Pliocene limestones of the Infierno Formation. The mineralization is composed of pyrolusite and romanechite, along with Fe-oxides and quartz (Camprubí et al., 2008).

2.6, ANALYTICAL PROCEDURES

2.6.1, Rare earths and other trace elements

For the rare earth elements (REE) and other trace elements in the manganese ores, pure manganese samples were digested using HClO4 around 100°C for a few hours and then the solutions were evaporated to dryness. This step was repeated three times in order to ensure total digestion. Subsequently, the samples were treated with aqua regia

overnight and then the solutions were evaporated to dryness. The last step was repeated

one more time.

The Cu and Mn, REE, and the other trace elements concentrations were analyzed

in an Elan DRC-II ICP-MS system (Inductively Coupled Plasma-Mass Spectrometer) in the Arizona Laboratory for Emerging Contaminants (ALEC) at the University of

Arizona, and in a Perkin-Elmer ICP-OES model Optima 4200 DV inductively coupled plasma optical emission spectrometer in the Geochemistry Laboratory at the Geology

Department at the University of Sonora. Table 2.1 and 2.2 show the REE and other trace

element concentrations in the manganese oxides from the Boléo district mantos, Lucifer,

and Gavilán manganese deposits.

2.6.2, Sulfur and oxygen isotopes

The sulfur and oxygen isotopes in the sulfate samples were measured at the

Environment Isotope Laboratory at The University of Arizona. The sulfur isotopes were measured on a continuous-flow gas-ratio mass spectrometer (Finnigan Delta PlusXL). A range between 0.3 to 0.7 mg and ~1.0 mg of powder sample was used for sulfide and sulfate samples, loaded in a tin capsule along with V2O5 to buffer fO2. The samples were

combusted at 1,030ºC, using an elemental analyzer (Costech) coupled to the mass

spectrometer. Standardization is based on international standards NBS123 and OGS-1 for

the sulfides and sulfates respectively, and several other sulfide and sulfate in-house

standards that have been compared between laboratories. Calibration is linear in the range

-10 to +30 per mil. Analytical precision is ±0.15 per mil or better (1σ).

The gypsum samples were dissolved in 2N HCl and re-precipitated as BaSO4.

Approximately 0.3 mg of the powder was placed in tin capsules along with V2O5 to

18 buffer fO2. The δ O values in the sulfate samples were measured on a continuous-flow

gas-ratio mass spectrometer (Finnigan Delta PlusXL). The samples were combusted in

the presence of excess carbon at 1350°C using a ThermoQuest thermal combustion

elemental analyzer (TCEA) coupled to the mass spectrometer. Standarization is based on

the replicate analyses of the standard. Analytical precision obtained is ± 0.3 per mil or

better (1σ). The δ34S values for the sulfide samples are shown in Table 2.3, and the δ34S

and δ18O data for the sulfates are shown in Table 2.4.

2.6.3, Oxygen and carbon isotopes

Microdrilled sample powders from carbonates were collected for carbon and

oxygen analysis. Powdered samples between 20 and 150 μg were reacted with 100%

dehydrated phosphoric acid under vacuum at 70°C. The isotope ratio measurement is

calibrated on the basis of repeated measurements of NBS-19 and NBS-18, with a

precision of ±0.1‰ for δ18O and ±0.06‰ for δ13C (1σ). Samples were heated under

vacuum to 200°C prior to measurement. The δ18O and δ13C values for the carbonate

samples were measured using an automated carbonate preparation device (KIEL-III)

coupled to a gas-ratio mass spectrometer (Finnigan MAT 252). Table 2.5 shows the δ13C

and δ18O for the Boléo district.

2.6.4, Copper isotopes

Micro-drilled sample powders from the copper and manganese ore minerals were

collected for the analysis of the copper isotopes. Approximately 0.05 g of the sulfide and

oxide samples was digested in aqua regia overnight at 140ºC in Savillex teflon

containers. The samples were subsequently evaporated to dryness at 40ºC. The samples

were then treated with 8N HNO3 to ensure total digestion and evaporated to dryness

again at 40ºC. The Cu-silicate samples were digested in a mixture HF and HNO3 in a 5:1 ratio, and evaporated to dryness at 40ºC. The copper from the copper ore minerals are then separated using ion exchange chromatography (IEC) following the procedure of

Mathur et al., (2005) without the use of hydrogen peroxide. For the manganese oxide samples, the solutions were treated twice in the IEC in order to remove the manganese.

The copper solutions retrieved after the IEC produced yields greater than 95%. The copper isotope analyses of the purified copper solutions were performed on a GV

Instruments Multicollector Inductively Coupled Plasma Mass Spectrometer (MC-ICP-

MS) in the Geosciences Department at the University of Arizona. The copper isotope ratios are expressed as follows:

‰ 1 1000,

where the standard is SRM NIST 976 standard reference copper material (Shields et al.,

1964). Table 2.6 shows the copper isotope data. The analytical precision is better than

0.15 (2σ).

2.6.5, Lead and strontium isotopes

Whole rock powdered samples were digested using Savillex teflon containers in a mixture of HF and HNO3 in a 5:1 ratio. The samples were evaporated to dryness and treated with HClO4. After evaporation, concentrated HNO3 was added to the samples and evaporated again to dryness. The samples were subsequently treated with 8N HCl and evaporated to dryness. Some steps described above were repeated to ensure total digestion of the samples. Micro-drilled powders from copper and manganese ore samples were extracted using lead-free tungsten-carbide dental drill bit. The powders were subsequently digested using the same procedure described above.

Sr and Pb are then separated from the resulting solutions using a chromatographic method following the procedure described by Thibodeau et al. (2007). Lead isotope analyses were conducted on a GV Instruments Multicollector Inductively Coupled

Plasma Mass Spectrometer (MC-ICP-MS) in the Geosciences Department at the

University of Arizona according to the methods discussed in Thibodeau et al. (2007).

Analysis of NBS-981 standard produced the following results 206Pb/204Pb = 16.9405 (±

0.0029 2σ), 207Pb/204Pb = 15.4963 (± 0.0034 2σ), and 208Pb/204Pb = 36.7219 (± 0.0099

2σ).

The strontium solution separates were loaded on tantalum filaments with Ta gel to enhance ionization following the procedure in Chesley et al. (2002). The isotopic analyses were performed by Negative Thermal Ionization Mass Spectrometer (N-TIMS) on a VG 54 mass spectrometer in the Geosciences Department at the University of

Arizona, following the conditions used in Chesley et al. (2002). Analytical uncertainties

(2σ) are 87Sr/86Sr = 0.0011% or better.

2.7, RESULTS

2.7.1, Major and trace elements in manganese oxides

The major and trace element concentrations data is shown in Table 2.1. In

general, the manganese concentration in manto 4 in the Boléo district is greater than the

concentrations in mantos 3 and 2 (1.2 to 52.7 wt%), and produces an average

concentration of 25.9 wt%. The copper concentrations in the manganese oxides are lower

than mantos 3 and 2, and range between 0.2 to 8.4 wt%. Zinc concentrations range from

0.4 to 1.7 wt% and cobalt concentrations from 32 to 1600 ppm.

The manganese concentrations from manto 3 range from 0.8 to 19.7 wt%. The

copper concentrations in the manganese oxides are higher than the rest of the mantos and range from 0.03 to 68 wt%. The zinc and cobalt concentrations are higher than concentrations from manto 4, and range from 0.2 to 1.6 wt% and 393 to 4600 ppm respectively. Mantos 3A and 2 show lower manganese concentrations (~2.5 wt%), and exhibits lower copper content (5.5 and 13.4 wt% respectively), and higher cobalt and zinc concentrations than manto 3 and 4 (Table 2.1).

At Neptuno area, a single manganese oxide sample exhibit 58 wt% of manganese concentration. The copper concentration ranges from 0.05 to 2.7 wt% and produce an average of 0.69 wt%. The zinc concentrations are lower than in the Boléo district and

range between 0.05 to 0.4 wt%. Except for one sample, cobalt concentrations range from

36 to 560 ppm (Table 2.1).

The manganese oxides from Lucifer deposit show the highest manganese concentrations in the region, ranging between 18.8 and 61.8 wt%, with an average of

24.54 wt%. Cobalt concentrations are similar in the Neptuno area, and range between 20 to 460 ppm, whereas zinc concentrations are lower than in the previous deposits (0.01 to

0.25 wt%). A single manganese oxide sample vein from the Gavilán deposit shows manganese and copper values of 22.27 and 0.07 wt% respectively, and minimum zinc and cobalt (Table 2.1).

Table 2.2 shows the rare earth element concentrations for the manganese oxides from the Boléo district mantos, Lucifer, and Gavilán deposits. Table 2.2 also shows the normalized La/Sm, Gd/Yb, and La/Yb ratios from the manganese oxides from the different mantos and other manganese occurrences from the Boléo district. Figure 2.6 shows the REE patterns of the manganese samples of Baja California, normalized to the

North American shale composite (NASC) from Gromet et al., (1984).

The shaded area (Fig. 2.6a) represents the REE-normalized trend for the manganese oxides from the different mineralized mantos in the Boléo district. The normalized La/Sm and Gd/Yb ratios are variable, yielding an average of 2.9 and 1.1, respectively. In general the REE spectrum is relatively flat characterized with positive Eu anomaly, yielding an average (La/Yb)n ratio of 2.4.

In the Neptuno area, the normalized La/Sm and Gd/Yb ratios yielded an average of 1.9 and 2.0, respectively. The normalized La/Yb ratio for the manganese

mineralization is 4.1, and the normalized REE spectra show a slight positive Eu anomaly

(Fig. 2.6b). Only one sample (LF-46) is characterized by a high (La/Yb)n ratio of 12

(Table 2.1). Excluding this sample, the average normalized ratios for La/Sm, Gd/Yb, and

La/Yb is 2.0, 1.1, and 2.1, respectively.

The normalized La/Sm and Gd/Yb ratios from the manganese oxide ores from

Lucifer deposit are consistent, yielding an average of 1.0 and 1.3, respectively. The REE

trend is flat with no distinctive Eu anomaly (Fig. 2.6c), with an average (La/Yb)n ratio of

2.0. The manganese oxides from Lucifer deposit are characterized by low REE

concentrations with respect to NASC (Table 2.2).

The average normalized La/Sm and Gd/Yb ratios for the Gavilán manganese

deposit are 1.3 and 0.8 respectively. The REE spectrum for the manganese oxide veins

from the Gavilán deposit shows REE concentrations lower than NASC (Fig. 2.6d) with

an (La/Yb)n of 0.9. This manganese oxide has slightly negative Ce and Eu anomalies.

2.7.2, Sulfur and oxygen isotopes

The stable isotope data for the sulfide and sulfate samples from the Boléo district

is shown in Tables 2.3 and 2.4 respectively. The δ34S values for the sulfides from the

Boléo district range from –6.8 to –1.7‰ (Table 2.3). In addition, Tables 2.3 and 2.4 show the sulfur and oxygen isotope data in sulfates and also the sulfur isotope data from the sulfides documented by Ochoa Landín (1998) and Conly (2003) in the Boléo district.

The δ34S and δ18O values for the gypsum member range from +23.6 to +24.1‰

and +11.0 to +12.7‰ respectively. In contrast, the δ34S and δ18O values for the gypsum

samples within the mantos range from –33.5 to +14.8‰ and –0.8 to +8.3‰ respectively.

2.7.3, Oxygen and carbon isotopes

The carbon and oxygen isotope data determined in the carbonate phases from the

Boléo region and Gavilán are listed in Table 2.5. The δ13C and δ18O values for the

limestone member range from –5.7 to +2.2 and +21.6 to +30.8‰ respectively. The δ13C

and δ18O values for calcites from manto 4 range from –11.5 to +4.3 and +19.1 to +26.7‰

respectively. The δ13C values in the malachite from manto 3 range from –8.3 to +0.2‰,

whereas the δ18O values range from +17.7 to +31.6‰. The carbon and oxygen isotope

values for copper carbonates from manto 2 range from –4.7 to –0.7 and +24.1 to +27.3‰ respectively. Table 2.5 also includes the δ13C and δ18O values for carbonates from the

Boléo district documented by Ochoa Landín (1998) and Conly (2003). The δ13C and δ18O

data for Lucifer deposit range from –12.6 to –12.5 and +19.8 to +20.2‰ respectively,

whereas the δ13C and δ18O data for Gavilán deposit range from –2.1 to +2.9 and +28.0 to

+31.4‰ respectively.

2.7.4, Copper isotopes

The copper isotope data for the copper and manganese mineralization from the

Santa Rosalía region are shown in Table 2.6. The manganese mineralization from manto

4 exhibits δ65Cu values from –0.31 to –0.22‰. The δ65Cu values for the copper and

manganese mineralization from manto 3 range from –1.62 to –0.13‰ and –0.73 to –

0.17‰ respectively. The δ65Cu values for the copper mineralization in manto 2 range

from –1.39 to –0.43‰, whereas a single value of +0.16‰ was recorded in the manganese

oxides. Only one copper sample along the horizons within the gypsum member below the

clastic member in the Boléo Formation produced a δ65Cu value of –1.58‰. A manganese

oxide vein cross cutting the manto 3 produces a δ65Cu value of –0.73‰, whereas a

manganese oxide vein cross cutting the ASL rocks has a δ65Cu value of –0.20‰.

A single isotope copper data in manganese oxides from Neptuno area produces a

value of –0.31‰, whereas a range of –0.86 to –0.17‰ is produced for manganese oxides

from Lucifer deposit. Finally a single value of +0.48‰ is recorded for manganese oxides

from the Gavilán deposit (Table 2.6).

2.7.5, Pb and Sr isotopes

Lead and strontium isotopes for the copper and manganese mineralization from

the Boléo mantos, the marine members of the Boléo Formation, as well as the Cinta

Colorada unit, and the ASL rocks are shown in Table 2.7. This table also includes the

isotopic data for the Lucifer and Gavilán deposits. The lead isotope ratios for the

Peninsular batholith show constrained values of 206Pb/204Pb = 18.798 to 18.803,

207Pb/204Pb = 15.602 to 15.604, 208Pb/204Pb = 38.603 to 38.608. The lead isotope data for

the ASL rocks is 206Pb/204Pb = 18.679 to 18.776, 207Pb/204Pb = 15.595 to 15.608,

208Pb/204Pb = 38.495 to 38.559. The only Pb data for the Boléo Basalt unit is 206Pb/204Pb =

18.623, 207Pb/204Pb = 15.576, 208Pb/204Pb = 38.392. The Pb data for the Limestone and

Gypsum member are 206Pb/204Pb = 18.772, 207Pb/204Pb = 15.605, 208Pb/204Pb = 38.547 and

206Pb/204Pb = 18.720, 207Pb/204Pb = 15.591, 208Pb/204Pb = 38.484, respectively.

The lead isotope data for the manganese mineralization from manto 4 are

206Pb/204Pb = 18.726 to 18.833, 207Pb/204Pb = 15.589 to 15.595, 208Pb/204Pb = 38.471 to

38.499. The Pb data for the copper mineralization is 206Pb/204Pb = 18.727 to 18.866,

207Pb/204Pb = 15.583 to 15.607, 208Pb/204Pb = 38.440 to 38.534, whereas the manganese

oxides have 206Pb/204Pb = 18.708 to 18.800, 207Pb/204Pb = 15.588 to 15.598, 208Pb/204Pb =

38.452 to 38.511. The lead data in the manganese oxides from manto 3A is 206Pb/204Pb =

18.751, 207Pb/204Pb = 15.591 to 15.593, 208Pb/204Pb = 38.483 to 38.494. The copper

mineralization form manto 2 have 206Pb/204Pb = 18.496 to 18.876, 207Pb/204Pb = 15.590 to

15.611, 208Pb/204Pb = 38.481 to 38.535, whereas a single manganese oxide data has

206Pb/204Pb = 18.721, 207Pb/204Pb = 15.589, 208Pb/204Pb = 38.469.

The lead isotope data of manganese oxide from the Lucifer deposit exhibit

206Pb/204Pb = 18.743 to 18.788, 207Pb/204Pb = 15.597 to 15.606, 208Pb/204Pb = 38.512 to

38.590. The single lead data of the ASL rocks and a manganese oxide vein from the

Gavilán deposit have 206Pb/204Pb = 18.621, 207Pb/204Pb = 15.584, 208Pb/204Pb = 38.432,

and 206Pb/204Pb = 18.612, 207Pb/204Pb = 15.579, 208Pb/204Pb = 38.421, respectively.

The strontium isotopic composition of the Peninsular batholiths, the ASL, and the

ERV rocks show a constrained 87Sr/86Sr range values from 0.7035 to 0.7045 and agree

with previous Sr data reported by Conly 2003 (Table 2.7). The Sr isotope data for the

gypsum member range from 0.7082 and 0.7084, which agrees perfectly with the Sr

isotope composition of seawater at 7 Ma. The Sr isotope ratio for the limestone member

is 0.7062, lower than the seawater value at that time.

The Sr isotope data for the copper mineralization and the manganese oxides from

the Boléo district mantos range from 0.7062 to 0.7088 and 0.7043 to 0.7067 respectively

(Table 2.7).

2.8, DISCUSSION

2.8.1, Mineralization and hydrothermal activity

The mineralization in the Boléo district has been documented in several studies, and most of them agree with the style of occurrence despite the discrepancies regarding the genetic model. Both the primary and secondary mineralization is constrained to the claystone facies. The primary mineralization consists of a mixture of ore-bearing sulfides

(Cu-Co-Zn), and associated manganese oxides. The secondary mineralization consists of

Cu-bearing carbonates, silicates, and oxychlorides, along with Mn and Fe oxides. The secondary mineralization is the result of meteoric water circulation, and coupled with the intense arid climatic conditions which enhance oxidation of the sulfide facies (Conly

2003). Field evidence around the district shows the NW-SE structures constituted by manganese oxides and copper silicates cross-cutting the ASL rocks; manganese oxide mineralization is also related also to the NW-SE structures cross cutting the ASL volcanic rocks in Lucifer area (Del Rio Salas et al., 2008a). The NW-SE structures served as conducts for the ascent of the mineralizing fluids were discharged into the Santa

Rosalía basin. The fact that the mineralizing fluids ascended through these structures is

inferred by the juxtaposition of high-grade Cu±Co zones and localized discordant to

stratabound zones of pervasive Mn-Fe-Si alteration (Conly et al., 2006). In addition,

Conly et al. (2006) pointed out the relationship between the hydrothermal activity and

tectonism in the Boléo district by the cyclical nature of the Boléo Formation clastic

sequence, and the similarities in the Cu/Zn and Co/Zn ratios of manto 1 within the south

sub-basin and manto 3 in the north sub-basin (Fig. 2.3).

The range of temperature of the hydrothermal activity responsible for the

mineralization at the Boléo district has been constrained by different methods. The formation of framboidal pyrite has been recognized as an indicative of relatively low temperature settings, usually below 200ºC (Scott et al., 2009). The framboidal textures in

pyrites from the different ore mantos constrain the formation low temperature conditions,

along with the lack of hydrothermal alteration above and below the mantos, support the

low temperature nature of the mineralization (Ochoa Landín 1998). Also, the presence of

orthorhombic chalcocite reported by Touwaide (1930), constrains the mineralization

temperature to below 91ºC (Posjnak et al., 1915). Later, Bailes et al. (2001) reported

chalcocite occurring as intergrowths with digenite and djurleite, whose assemblage

indicates temperatures between 70 and 93ºC (Vaughan and Craig 1978). In addition,

equilibration temperatures using the quartz-pyrolusite geothermometer from Zheng

(1991) produced a range of temperatures between 18 and 118ºC for the manganese oxide

mineralization in the different Boléo mantos (Conly 2003; Conly et al., 2006). Taken together all of these criteria suggest a conservative mineralization temperature range between 70 and 118ºC.

Late Tertiary hydrothermal activity along a 200 km segment of the eastern coast

of Baja California Sur is evidenced a series of mineralized localities (Fig. 2.1). North the

Tres Vírgenes volcanic field, evidence of hydrothermal activity is recorded by the copper

and manganese mineralization within the ASL rocks in the San Alberto prospect, and

copper mineralization at the Caracol alteration zone (Fig. 2.1). South the Boléo district,

the hydrothermalism is mostly represented by manganese deposits such as Mantitas,

Gavilán, La Trinidad, Pilares, Las Minitas, Santa Teresa, and Azteca (Bustamante García

1999; Camprubi et al., 2008).

The geothermal waters from springs and wells in and around the Tres Vírgenes

and La Reforma Caldera fields located north of the Boléo district (Fig. 2.1), are

characterized by temperatures from 21 to 98ºC (Portugal et al., 2000). South of the Boléo

district, the localities such as Saquicismunde, Los Volcánes, Piedras Rodadas, Agua

Caliente, El Imposible, and El Tejón (Fig. 2.l), are current examples of hydrothermal

activity that may correspond in type to that responsible for the mineralization at the Boléo

district. These geothermal emanations are quite geographically dispersed, and all of them

share similar geological features such as the structural northwest-southeast control, the

evidence of hydrothermal alteration, the occurrence within the Miocene volcanic or

volcanoclastic rocks, and the low temperatures range between 38 to 94ºC (Casarrubias

and Gómez López, 1994; Bustamante García 1999; Camprubí et al., 2008). Also

hydrothermal activity is present along the western coast of Concepción Bay, between the

Santispac and the Mapachitos areas (Fig. 2.l). Although some of these geothermal

emanations occur within the shallow submarine environment (<20 m depth), they share

similar features than those exposed above (Prol-Ledesma et al., 2004; Canet et al.,

2005a).

The hydrothermal activity and manganese mineralization appear to have migrated

southward along Baja California Sur, from the Lucifer deposit to the Cerro Mencenares

volcanic field (Fig. 2.l). In fact, such southerly migration is recorded at a smaller scale

within the Santa Rosalía basin, where the southward migration, due to the basin

subsidence, is recorded by the presence of the first two sedimentary cycles (4 and 3) in

the north sub-basin, whereas the last sedimentary cycles (2 to 0) are present in both north and south sub-basins (Conly et al., 2006). The migration of the hydrothermal activity can be explained as a response of the regional southward trend of Proto-Gulf extension

(Stock and Hodges 1989).

2.8.2, Mineralization age

The age of the Boléo Formation in the Santa Rosalía region has been previously constrained by Holt et al., (2000). The age was calculated using isotope data in conjunction with magnetostratigraphy, which is correlated with the geomagnetic polarity time scale. The most likely correlation yielded an age of 7.09–6.93 Ma for the base and

6.27– 6.14 Ma for the top of the Boléo Formation (Holt et al., 2000). Furthermore, the age of the manganese mineralization in the Boléo deposit has yielded 7.0±0.2 Ma (Conly

2003), which is stratigraphically and chronologically in agreement with the deposition age for the base of the Boléo Formation. Moreover, the manganese mineralization age is in agreement with the geochronologic data of the underlying volcanic rocks in the Santa

Rosalía region, which include the 24–13 Ma ASL, the 11–9 Ma Boléo basalts and Boléo basaltic andesites, and the 9–8 Ma El Morro tuff.

Moreover, the Gulf of California region has been subjected to intense volcanism and tectonic activity (Sawlan 1991). Around the Santa Rosalía region, Conly et al.,

(2005) reported K–Ar ages for the high-K andesites from Cerro San Lucas suite, which represents volcanic activity during the transition from arc- to rift magmatism between 9.5 and 7.7 Ma. Furthermore, Pallares et al., (2008) reported new geochronologic data supporting the continuation of volcanic activity until the Pleistocene, following the end of the calc-alkaline volcanism northern Santa Rosalía region.

Interestingly enough and as mentioned before, manganese mineralization is recorded along 200 km along the eastern coast of Baja California Sur. The southern manganese mineralization such as those in Santa Rosa and Juanico, is constrained to

Pliocene structures (González-Reyna, 1956; Terán Ortega and Ávalos Zermeño, 1993;

Umhoefer et al., 2002), therefore, the hydrothermal activity responsible for the manganese mineralization is younger than the geochronologic data reported in the Boléo district. In addition to this, volcanic activity such as the Pliocene Mencenares volcanic center is reported to be coeval to the manganese mineralization (Bigioggero et al., 1995), which suggests the migration of the mineralization activity south of the Boléo district, as evidenced also by the present day hydrothermal activity reported in Concepcion peninsula by Prol-Ledesma et al., (2004).

2.8.3, Geochemistry of manganese oxides from the Boléo mantos

Manganese discrimination diagrams have been proposed based on the cation-

adsorption capacity of manganese oxides, considering the concentrations of specific trace

elements present in manganese oxides. These discrimination diagrams have been used to

distinguish between a hydrothermal (continental or marine) or a hydrogenous origin. The

term hydrothermal refers to manganese oxides deposited directly from geothermal waters

around hot springs and pools in continental environments or sedimentary exhalative

manganese mineralization deposited in marine environments (Nicholson 1992 and

references therein). The term hydrogenous refers to deposits formed by slow precipitation

or adsorption of dissolved components from seawater (Bonatti et al., 1972; Crerar et al.,

1982; Nicholson 1992).

Elements such as Ba, Cu, Ni, Co, Pb, Sr, V, and Zn are frequently found in hydrothermal manganese-rich systems (Nicholson 1992). These elements are present in significant concentrations in the manganese ores from Santa Rosalía region and

Concepción peninsula (Table 2.1). Hydrothermal oxides have lower Co, Cu, Ni, and Zn concentrations, relative to hydrogenous deposits; hence, high cobalt concentrations are indicative of marine environments (hydrogenous) as pointed out in the discrimination diagram, which potentially could also distinguish between marine–freshwater vs. hydrothermal deposits with further development (Fig. 2.7; Nicholson 1992).

Previous studies in Lucifer deposit and the manganese oxides from Concepción

Peninsula have documented the hydrothermal nature of the manganese oxide deposits

(Canet et al., 2005a; Camprubí et al., 2008; Del Rio Salas et al., 2008a; Rodríguez Díaz

2009; Rodríguez Díaz et al., 2010). Figure 2.7 shows the manganese ore samples reported

and cited from Baja California Sur and those from the Boléo mantos are located within

the hydrothermal field. In general, the samples from Concepción Peninsula are characterized by the lowest Co and Ni concentrations, like those from the Lucifer deposit

(Fig. 2.7). The data from the Neptuno area are located above the Lucifer and the

Concepción Peninsula deposits, and are characterized by greater concentrations of Co and

Ni. The Boléo mantos samples are relatively richer in Co and Ni respect to the previous samples, and within these samples, it is possible to notice an enrichment from manto 4 to manto 2 (Fig. 2.7). The trace element concentrations in the manganese oxides show a clear hydrothermal origin for all manganese deposits in the Santa Rosalía region and the

Concepción Peninsula.

The shaded area (Fig. 2.6a) represents the North American shale composition

(NASC)-normalized trend for the manganese oxides from the Boléo mantos, and there is no particular difference within the REE enrichment within the mantos. The entire

spectrum is relatively flat and consists of a subtle enrichment of the light rare earth

element (LREE) over the heavy rare earth element (HREE), as pointed out with the

normalized La/Yb ratios, as well as the noticeable positive Eu anomaly. The REE-

normalized spectrum of the manganese mineralization from Neptuno area displays also

the positive Eu anomaly with relatively flat trends, except for sample LF-46, which is the

most enriched in the LREE relative to the NASC (Fig. 2.6b), as noticed by the high

normalized La/Yb ratio (Table 2.2).

The REE trend for the manganese oxides from Lucifer deposit is flat as noticed by

the average normalized La/Yb ratio of 2.0, whose signatures are characterized by the

absence of either positive or negative Ce or Eu anomalies (Fig. 2.6c). The REE trend

from the Gavilán deposit (Del Rio Salas et al., 2008a; Rodríguez Díaz 2009) show that

the REE concentrations are lower than NASC, and are relatively flat with an average

normalized La/Yb of 0.9, with slight negative Ce and Eu anomalies (Fig. 2.6d).

The REE data from Guadalupe manganese oxides deposits also show a depletion respect to NASC, with spectrums relatively flat and with slightly negative Ce and Eu anomalies (Fig. 2.6e; Rodríguez Díaz 2009), whereas those from Santa Rosa exhibit a negative Eu anomaly; (Fig. 2.6f). In general, the manganese oxides from Concepcion

Peninsula exhibit relatively flat patterns, characterized by slightly negative Ce and Eu

anomalies.

Table 2.8 shows the average of the total REE abundances of various manganese

oxide deposits of hydrothermal and hydrogenous nature. The total REE concentrations in

hydrothermal deposits range from 45 to 647 ppm, whereas those for hydrogenous

deposits range from 1,200 to 1,900 ppm. The average of the total REE for the Boléo

mantos ranges from 6 to 270 ppm and produce an average of 215 ppm. The average total

REE for the manganese oxides from Neptuno area is around 560 ppm; sample LF-46 is characterized by a total REE around 1,900 ppm, and excluding this particular sample, the

average of total REE is 220 ppm. The average total REE for Lucifer and the deposits in

Concepción Peninsula ranges from 25 to 46 ppm (Table 2.2 and 2.8).

Figure 2.6g shows the NASC-normalized REE patterns for manganese oxides from hydrogenous and hydrothermal deposits from Usui and Someya (1997). The pattern for the hydrogenous deposits is enriched relative to the NASC values, whereas the pattern for the hydrothermal deposits is either depleted or slightly enriched than NASC values.

This figure also shows the spectrum for the average modern and fossil hydrothermal deposits from Usui and Someya (1997). In addition, this figure includes the average

NASC-normalized REE spectra from the manganese deposits from the Boléo and adjacent areas, which is included in the hydrothermal field. The average REE patterns for the Lucifer, Santa Rosa, and Guadalupe deposits are the most depleted in REE and agree with the average modern hydrothermal deposits. The REE average signature from the

Gavilán deposit is located between the average modern and average fossil hydrothermal deposits. Finally, the REE signatures for the Boléo and Neptuno deposits are slightly more enriched in the REE average from fossil hydrothermal deposits, but still located in the hydrothermal field (Fig. 2.6g).

In general, NASC-normalized REE patterns for the manganese mineralization around the Boléo district and Concepción Peninsula are relatively flat as noticed from the consistent normalized La/Sm, Gd/Yb, and La/Yb ratios. The REE patterns of the Boléo district and Neptuno area show middle REE enrichment (Fig. 2.6a and b), which can be characteristic of hydrogenous manganese samples (Nath et al., 1992, 1997). Also, the

REE patterns exhibit very distinctive positive Eu anomaly, which is a characteristic feature of modern hydrothermal deposits in the ocean (Hodkinson et al., 1994), as opposed to the REE signature of seawater, which exhibits a negative Ce anomaly and a

slight enrichment of the HREE over the LREE (Douville et al., 1999). Although the

mineralization in these deposits is not properly from hydrothermal activity in a marine

environment, the Eu enrichment in the manganese ores can be explained by the presence

of plagioclase in the fine-grained sediments derived from the ASL rocks. Eu mobility

depends strongly on redox and temperature conditions (Michard et al., 1983), and Eu

enrichment involves hot and reduced fluids, whereas Eu depletion involves cold and oxidizing fluids (Parr 1992; Canet et al., 2005a). Since each manto experienced such

redox conditions, a positive Eu anomaly can be explained by the cyclical hydrothermal

activity related to the formation of each manto. Moreover, these samples are located

within the hydrothermal field (Fig. 2.7) and the hydrothermal nature is also confirmed by

the total REE in Table 2.8.

The NASC-normalized REE patterns for the manganese ores from the Lucifer and

the manganese deposits from Concepción Peninsula (Gavilán, Guadalupe, and Santa

Rosa) agree with the spectra of the average hydrothermal deposits (Fig. 2.6c; Usui and

Someya 1997); only the average of total REE for the Gavilán deposit exhibits a subtle

enrichment especially in the HREE relative to the average of modern hydrothermal

deposits (Fig. 2.6g). Hydrogenous manganese deposits are characterized by a positive Ce anomaly (Fleet 1983), as a result of the oxidation of Ce3+ to Ce4+, which forms highly insoluble CeO2 in seawater (Elderfield and Greaves 1981; Fleet 1983; Nath et al., 1997;

Canet et al., 2008). Conversely a negative Ce anomaly is characteristic of hydrothermal

Fe-Mn deposits. The slightly negative Eu could be explained by cold and oxidizing fluids

(Parr 1992), as previously reported for the recent Mn mineralization in Concepción Bay

(Canet et al., 2005a). Furthermore, the total REE content of these deposits confirms their

hydrothermal nature when compared to other hydrothermal deposits in Table 2.8.

In summary, trace element and REE geochemistry in the manganese oxides from the Boléo district and Concepción peninsula, demonstrates the hydrothermal nature of the ores (Figs. 2.6 and 2.7; Nicholson 1992). The REE enrichments for the Boléo mantos and

Neptuno mineralization can be explained as mixture of hydrothermal and hydrogenous sources or by supergene processes.

2.8.4, Sulfur and oxygen isotopes

Sulfur isotope data for the Boléo sulfides and sulfates have been documented in previous studies (Wilson and Rocha 1955; Ochoa Landín 1998; Conly 2003), and they

are shown in Tables 2.3 and 2.4 respectively. The sulfur isotope data from sulfides show

a smaller range between –33.6 to –1.7‰. The δ34S values for sulfides (pyrite-dominated)

from manto 3 show a range between –33.6 to –10.9‰ (Ochoa Landín 1998), whereas

Conly (2003) reported δ34S values for the Cu (±Co, Zn) sulfides from –13.7 to –1.8‰.

The new δ34S values reported in this study agrees with the range presented by Conly

(2003). Figure 2.8 shows a δ34S frequency diagram for the sulfides, sulfates from the

mantos, and the gypsum member of the Boléo Formation. The reported δ34S values for

the framboidal pyrite from the mantos are lighter than –15‰ (Ochoa Landín 1998),

suggesting that they formed as the result of bacterial sulfate reduction of seawater sulfate, whose SO4-H2S fractionation range from 40 to 55‰ (Ohmoto and Rye 1979). Conly et

al. (2006) suggested that the most likely source for the dissolved sulfate was the seawater

trapped in pore spaces within the uncompacted to partially compacted Boléo manto sediments, or possibly sulfate diffusion after sedimentation. The framboidal pyrite is the

earlier sulfide in the mantos 3 and 2.

The ore sulfides are characterized by heavier δ34S values (–8.0 to –1.8‰) and are consistent with a continued bacterial sulfate reduction under conditions closed to partially closed to sulfate (Conly et al., 2006). The higher δ34S values are explained by bacterial

sulfate reduction at higher temperatures, most likely between 80 to 110°C, which result

with a fractionation ~25‰, similar to the fractionation documented (20 to 36‰) for the

bacterial sulfate reduction at higher temperatures for sulfides from the Guaymas basin

(Conly et al., 2006). Variable mixing degrees between the pyrite with lower δ34S and the heavier sulfur associated to the ore bearing Cu-Co-Zn sulfides is clearly documented.

This is further supported by the correlation of the δ34S and the Cu/Fe ratios that

approximate the relative proportion of the Cu-sulfides to pyrite (Conly et al., 2006).

Figure 2.9 shows the sulfur and oxygen isotopes from the gypsum member and

sulfate samples from the Boléo mantos. The δ34S and δ18O values for the gypsum

member range from +23.6 to +24.1‰ and +11.0 to +12.7‰ respectively (Ortlieb and

Colleta 1984; Conly 2003; Conly et al., 2006; present study), and agree perfectly with

evaporite deposits precipitated from Miocene seawater (Claypool et al., 1980). Some

gypsum occurring as brecciate mound structures product of bedded sulfate tectonically

deformed by major basin structures, present a decrease in the δ18O values, which suggests

recrystallization due to interaction with circulating meteoric water (Conly et al., 2006).

The δ18O values of water samples from active shallow submarine vents in

Concepcion Bay range from –0.3 to –3.1‰, which suggests a maximum of 40% for the

thermal end-member and the rest is composed by seawater, at a discharge temperature of

50ºC (Pro-Ledesma et al., 2004). As mentioned before, the oxygen isotopic signature for

the recharge meteoric water in the Tres Vírgenes volcanic field is −9.7‰ (Portugal et al.,

2000), whereas the δ18O value for the shoreline waters in the Santa Rosalía region is

−9.0‰ (Conly 2003).

The later interaction of these fluids with the framboidal pyrite and the mineralized

manto sulfides promote the oxidation of the sulfides. Biological and abiological oxidation

of sulfides may produce very small negative sulfur isotope fractionation, but generally

oxidation products have very similar δ34S values to those of the source sulfide minerals

(Toran and Harris 1989; Gu 2005). The samples cross-cutting the mineralized mantos are characterized mainly by negative δ34S values and by δ18O values ranging from –0.8 to

+8.3‰ (Fig. 2.9). Only a couple of gypsum samples cross-cutting manto 3 are

characterized by high δ34S values (~15‰) which suggests the possibility of a mixture of

seawater sulfate with sulfate formed during the oxidation of the mantos sulfide (Fig. 2.9).

The range in the δ18O values in the gypsum could reflect a possible modification in the

δ18O caused by the interaction of meteoric water. Conly (2003) determined that the decrease in the δ18O, δ34S, and 87Sr/86Sr values in barites from manto 3 is due to the

mixing of warm hydrothermal fluids that leached Sr and low δ34S from basement and

volcanic rocks (samples with δ18O and δ34S similar to marine sulfate in Fig. 2.9). An

alternative explanation for the observed δ18O and δ34S values is that these sulfate samples

were precipitated from marine sulfate trapped within the pore sediments with minimum interaction with local meteoric water.

In summary, the most important source of sulfur in the system is undoubtedly the

Gypsum member (Conly et al., 2006). The first evidence of sulfide formation is recorded

by the framboidal pyrites via the bacterial sulfate reduction (Ochoa Landín 1998). Later,

the oxidizing mineralizing solutions that precipitate the ore-bearing sulfides coating the

framboidal pyrites; these ore-bearing sulfides are formed via the bacterial sulfate reduction at higher temperatures (Conly et al., 2006). After the formation of both the framboidal pyrites and the ore-bearing sulfides, both sulfides are oxidized during the interaction of a mix of meteoric water and seawater sulfate trapped within sediments.

These fluids precipitate the gypsum veinlets cross-cutting the mantos and are characterized with similar δ34S signatures from the sulfide mantos, and δ18O values located along the proposed mixing trend between meteoric water and seawater sulfate trapped within sediments.

2.8.5, Oxygen and carbon isotopes

Carbon and oxygen isotopes are important in the study of ore deposits because provide important information about the temperature and nature of the solutions involved in the precipitation of ore and gangue minerals. Even though the C and O isotope data from the carbonates do not reveal the nature of the early mineralizing fluids from the

Boléo, they can provide information about the nature of the fluids related to the secondary mineralization occurring as Cu-carbonates.

Figure 2.10 shows the δ13C and δ18O data for carbonates from the Boléo district,

Lucifer, and Gavilán deposits. The gray-shaded square represents the field for marine

carbonates precipitated in equilibrium with seawater followed Conly et al. (2006), where

the δ18O range was calculated using the calcite-water fractionation equation (O’Neil et

al., 1969) from 15 to 40ºC and δ13C data from marine mollusks (Keith and Weber 1964).

In general, the isotopic compositions of the carbonates from the Boléo district display a

variation trend from enriched δ13C and δ18O values towards depleted δ13C and δ18O values (Figure 2.10a-e). This variation is seen in carbonates from each stratigraphic level of the Boléo Formation (i.e. the limestone and the consecutive ascending mantos), and suggests a particular systematic at each stratigraphic level and therefore, a same systematic occurring at different time during the evolution of the Santa Rosalía basin.

The end-member with enriched δ13C and δ18O values represent more likely carbonates

precipitated directly from seawater, whereas the end-member with depleted δ13C and

δ18O values represent oxidation of organic matter.

The Tirabuzon limestone is at least 1 Ma younger than the Boléo Fm, the isotopic

data exhibits the lower δ13C and δ18O values and is perfectly located along the correlation

trend of the Limestone member of the Boléo Formation (Fig. 2.10a). Conly (2003) used

the isotopic data for the Tirabuzon limestone to constrain the isotopic composition of the

mixed fluvial-seawater shoreline waters in the Santa Rosalía region. The shoreline waters

18 13 are inferred to be brackish with δ O values around −9‰ and δ C HCO3 around −10‰ at

25ºC. In addition, Portugal et al. (2000) documented the isotopic composition of the

thermal waters from the Tres Vírgenes geothermal system located north of the Boléo

district. The intersection point between the local meteoric water line and the regression line from the geothermal waters from Tres Vírgenes volcanic field is δ18O = −9.7‰ and

δD = −67.3‰, and represents the initial isotopic signature of the recharge meteoric water

in the region (Portugal et al., 2000). In summary, the δ18O signature for the meteoric

water in Santa Rosalía region can range between −9 and −9.7‰. The range for the δ18O values for the meteoric water in the Santa Rosalía region is important since this type of water is a potentially fluid involved in the mineralization processes.

Figure 2.10a shows the C and O isotope data for the limestone member along with carbonaceous sediments filling the shells within the limestone, the manganiferous limestone zones, the scarce dolomite present within the limestone member, and finally the Tirabuzon limestone as a regional reference. The more positive values correspond to dolomite and manganiferous limestone samples, whereas the more negative values correspond to the Tirabuzon limestone. The possible oxygen isotopic composition for the fluids involved in the precipitation of the dolomite was calculated using the dolomite- water oxygen isotope fractionation from Land (1983). The highest temperature documented for the formation of the mineralization at the Boléo district is 118ºC (Conly et al., 2006), consequently the highest temperature considered in the modeling is 120ºC.

The lowest temperature considered is 10ºC. According to the model, the most likely temperature range is between 35 and 40ºC for δ18O values around zero, which is the

oxygen isotope composition of seawater. In contrast, for δ18O values lighter than −9‰,

the temperature for the formation of the limestone member should be lower than 10ºC.

Therefore, the isotope data indicate that the fluid involved in the formation of the

dolomites and limestone was seawater. For the case of the more negative values, the

range of temperature for the fluids involved in the precipitation of the Tirabuzon

limestone is 20 to 25ºC for δ18O values around −9‰, otherwise temperatures around

70ºC is needed if only seawater was involved. Therefore, according to the isotopic modeling, the trend of the limestone member suggests a re-equilibration of the C and O in response of the interaction with meteoric water with lighter δ13C and δ18O values (Fig.

2.10a).

Figures 2.10b-e shows the C and O isotopic data from manto 4, 3, 3A, and 2. The

more positive δ13C and δ18O values for the calcite from the mantos suggests an oxygen

isotope signature similar to the seawater composition, between –1 to –0.2‰ if the

temperature of the fluid involved in the precipitation was around 30ºC (O’Neil et al.,

1969). On the other hand, the depleted δ13C and δ18O values in the calcites mantos

suggest temperatures of formation between 70 to 75ºC if precipitated from seawater, whereas temperatures between 25 to 30ºC are needed if meteoric water was involved in

re-equilibration of the calcite mantos.

Even dough the reported evidence for the isotope disequilibrium of siderite with

co-existing calcite in the mantos (Conly et al., 2006), if considered the two possible fluid

sources and the temperature range exposed above, and with the application of the

siderite-water oxygen isotope fractionation from Carothers et al. (1988), a temperature

range between 50 to 70ºC is calculated if seawater was involved in the precipitation of

the siderite, otherwise a range between 15 to 25ºC is required if meteoric water was

involved during precipitation. The isotope signature in the siderite could be explained by

the oxidation of Fe-sulfides at relatively high temperature (50 to 70ºC), possibly in the

presence of a mixture of seawater and meteoric water.

The same systematic was applied to the secondary mineralization represented by

the copper carbonates (malachite and azurite) using the malachite-water oxygen isotope

fractionation (Melchiorre et al., 1999) and the azurite-water oxygen isotope fractionation

(Melchiorre et al., 2000). According to the model, a temperature between 15 and 20ºC is

required to precipitate malachite from meteoric water, whereas a temperature range between 60 to 70ºC is required if the malachite precipitated in equilibrium from seawater.

Finally, according the azurite-water oxygen isotope fractionation (Melchiorre et al.,

2000), suggests that a temperature range of 20 to 35ºC is needed if the meteoric water

was involved in the precipitation of secondary carbonates, whereas a range between 65

and 85ºC is calculated if the seawater was involved in the precipitation of the secondary

Cu-carbonates. The low temperature range is more reliable for the secondary copper

carbonates in the presence of meteoric water.

Figure 2.10f shows the C and O data from calcites from the Gavilán and Lucifer

deposits. The calcites from the Gavilán deposit are located within the seawater field, and

according to the oxygen isotope fractionation (O’Neil et al., 1969), the temperature of the

seawater involved in the precipitation of the calcite was around 20ºC. In contrast, the

calcites from Lucifer deposit have lower δ13C and δ18O values, and most likely the

calcites precipitated mostly in the presence of meteoric water around 20 to 25ºC.

Carbonate species in fresh water and carbonate minerals in fresh water

environments tend to be more negative and variable with δ13C values between −2 to

- −10‰, because both CO2 and HCO3 produced by the oxidation of organic carbon and

atmospheric CO2 are important constituents in fresh waters (Ohmoto and Rye 1979).

Conly (2003) documented C and O isotopic signatures from the mantos at the Boléo

district and interestingly enough, the more negative δ13C data correlate positively with the

highest total organic carbon in the samples. The isotopic data exposed above agrees with

the geological observations by Touwaide (1930) and Wilson and Rocha (1955), who documented carbonized plants and petrified wood partially replaced by chalcocite from the mantos 4 and 3, and particularly in the facies A, they defined a proportion of organic carbon (Ochoa Landín 1998). In effect, the CO2 generated during decomposition of organic matter have been an important source of organic carbonate in some mineral deposits. Such feature is common in carbonates from sediment-hosted sulfide deposits,

characterized by a large variation in the δ13C values, with a clear predominance of

negative values (Ohmoto 1986).

2.8.6, Pb and Sr isotopes

Some Pb and Sr isotopic data for the peninsular batholith and the volcanic rocks

in Santa Rosalía region, the Boléo conglomerate, and the sulfide and manganese oxide

mineralization have been reported previously by Conly (2003). Osmium isotope data

from the manganese oxide mineralization in the Santa Rosalía region are reported here in

order to constrain the metal sources involved in the manganese mineralization (Chapter

3). The present study contributes with more precise Pb and Sr data for the copper and

manganese oxide mineralization in the different mantos in the Boléo district as well as for

the mineralized zones around the Santa Rosalía district (Table 2.7).

Figure 2.11 shows that the lead isotope data from the ASL rocks and the

Peninsular batholith in Santa Rosalía region exhibit a constrained 206Pb/204Pb ratio ranging from 18.67 to 18.80, whereas the single values for the EBR basalt in Santa

Rosalía and the ASL rocks from the Gavilán deposit exhibit a lower 206Pb/204Pb ratio of

18.62; the previous lead isotope data published by Conly (2003) for the ASL rocks

cluster around this smaller value. Important to note is that the fine-grained sediments, the

limestone and gypsum members, as well as the Cinta Colorada unit are characterized by

206Pb/204Pb ratios within the range for the new ASL rock data reported in the present

study. The new lead data reported for the ASL and the EBR volcanic rocks also show

relatively uniform 207Pb/204Pb ratios ranging from 15.57 to 15.60, and are located within

the field of lower continental crust and above the Pacific MORB field (Fig. 2.11).

Detailed geochemical and petrological characterization of the volcanic rocks from the

Santa Rosalía area suggests a mantle signature with a metasomatic overprint of slab-

derived aqueous fluids and silicic melts (Conly et al., 2005). Figure 2.12 shows in detail

the close relation between the ASL rocks, the Cinta Colorada, the gypsum and limestone

members of the Boléo Formation. The lead isotopes from the ASL and the Boléo Basalt

rocks agree with a mantle source metasomatized by slab-derived material. Also, the Pb isotopes suggest that the lead source in the Boléo district rocks (the Cinta Colorada, the gypsum, limestone, and clastic members of the Boléo Formation) is most likely the ASL rocks in conjunction with the Peninsular batholith.

In addition, the 87Sr/86Sr ratios for the ASL, the Peninsular batholith, and the

Cinta Colorada rocks show a limited range from 0.7040 to 0.7045, similar to the modern

mantle values. The lowest 87Sr/86Sr ratio value (0.7035) belongs to the Boléo basalt (rift-

related volcanic rock), a value which is closer to modern mantle values. The 87Sr/86Sr ratio for the Limestone is slightly greater than that of the volcanic rocks (0.7062), whereas the value for the gypsum member is closer to the seawater value at 7 Ma. Even though seawater was involved during the precipitation of these members, the difference in the Sr isotopic signature between these two units indicates that the Limestone member precipitated from brackish water (Conly et al., 2005), whereas the Gypsum member precipitated from 7 Ma seawater.

2.8.7, Source of metals

The use of the lead and strontium isotopic methods has been commonly applied to constrain the metal sources involved in the formation of ore deposits. Limited Pb and Sr data from the rocks units and mineralization from the Boléo district was published previously (Conly et al., 2006). New Pb and Sr isotope data from the copper and manganese oxide mineralization from the Boléo mantos along with most of the rock units within the Boléo Formation are shown in Figure 2.13.The lithological units included within the Boléo formation (limestone, gypsum, cinta colorada and the clastic member) are located inside or slightly below the ASL volcanic rocks field, suggesting that the Pb in those lithological units is mainly derived from the ASL volcanic rocks.

Figure 2.13 shows that the Pb data from the copper and manganese mineralization are located inside and below the ASL rocks field, where the 207Pb/204Pb ratios are constrained to a limited range between 15.583 and 15.608, and between 38.440 to 38.590 on the 208Pb/204Pb ratio (Fig. 2.13), vaguely deviated from the ASL field and farther than the Peninsular granites, which suggest that the principal source of metals in the different mantos in the Boléo district and the manganese mineralization in the Lucifer deposit is mainly the ASL volcanic rocks.

Despite the scarcity of isotopic data for the rift-related suite rocks, the single data for the Boléo basalt rock reported in this study, and complementing with the data for the

Cerro San Lucas volcanic rocks, whose 206Pb/204Pb ratios range from 18.43 to 18.61

(Conly 2003), the rift-related suite rocks are characterized by less radiogenic lead, which suggests that the volcanism related to the rifting of the Gulf of California is not contributing to the metal budget in the mineralized areas reported in the present study.

On the base of a single isotope measurement in the Gavilán deposit, the ASL rocks in that area are less radiogenic that those from the Santa Rosalía region (Fig. 2.13), suggesting slight isotopic differences for the sources involved in the volcanic activity of the ASL rocks trend along Baja California (Fig. 2.1). An alternative explanation for the less radiogenic nature of the ASL rocks in the Gavilán deposit is the possibility of the continuation of the isotope trend of the ASL rocks from the Boléo district. The manganese oxide sample from the Gavilán deposit exhibits the same isotopic signature as the ASL rocks, suggesting again that the source of metals in the mineralized zones southern the Boléo district is the ASL rocks.

Figure 2.14 shows the strontium and lead isotope compositions of the volcanic

rocks and mineralized samples in the district. In general the 87Sr/86Sr ratios for the ASL rocks and the Peninsular batholith have a constrained range between 0.7040 to 0.7045 and form a restricted field in the 206Pb/204Pb vs. 87Sr/86Sr plot, as oppose to the rift-related

rocks, which are characterized by lower 87Sr/86Sr ratios (0.7035 to 0.7040) and less

radiogenic lead. The most elevated 87Sr/86Sr ratios characterize the gypsum member of

the Boléo Formation and clearly represent the Sr isotope composition of seawater at the

time of deposition (~7 Ma).

The copper and manganese oxide mineralized samples are clearly located on a

mixing trend between the high 87Sr/86Sr gypsum end-member and the low 87Sr/86Sr ASL and Peninsular batholith end-member (Fig. 2.14). The limestone member is the only rock unit that is located in this trend, most likely as a result of re-equilibration with meteoric water or mixing between seawater and meteoric fluids at the time of deposition. The elevated 87Sr/86Sr ratios in the mineralized samples indicate the incorporation of more

radiogenic strontium in the mineralizing fluids, and could suggests an important

involvement of seawater fluids during the formation of the mantos. But the most probable explanation for the 87Sr/86Sr ratios in the mineralized samples is the interaction of the

ascending fluids with the gypsum member located below the clastic sequence that hosts

the mineralized mantos. Figure 2.14 clearly shows that the rift-related volcanism is not

responsible for the metal sources for the copper and manganese mineralization in the

Boléo district, and the ASL rocks undoubtedly act as the most important source, with a

minor role of the Peninsular batholith as shown in Figure 2.13.

2.8.8, Copper isotopes in nature

The first copper isotope measurements from copper mineralization in ore deposits

and organics were made by Walker et al. (1958) and Shields et al. (1965), who reported a

range from −1 to +8‰ with a considerable error. Due to the improvement of mass

spectrometry, in the last decade several authors developed new and more precise

analytical techniques for the stable isotopes of the transition metals group (i.e. Fe, Cu,

Zn, Mo), and new variations have been reported for different ore-forming environments

(Maréchal et al., 1999; Larson et al., 2003; Siebert et al., 2003; Albarède 2004; Anbar

2004; Graham et al., 2004; Johnson et al., 2004; Rouxel et al., 2004; Mathur et al., 2005;

Asael et al., 2007; Markl et al., 2006; Mathur et al., 2009).

Figure 2.15 shows the copper isotope variations for continental and marine environments. Botfield (1999) originally defined a continental igneous range for δ65Cu

between −0.5 and +0.5‰. Li et al. (2009) analyzed the δ65Cu values for batholiths from

the Lachlan Fold Belt and reported a range from −0.46‰ to 1.51‰. The mean δ65Cu

values for the I-type and S-type granites for the Lachlan Fold Belt are 0.03 ± 0.15‰ and

−0.03 ± 0.42‰ respectively. The higher δ65Cu values were assumed to be affected by

later hydrothermal activity.

A narrow δ65Cu range is also documented for the majority of high temperature ore

deposits, although some deposits can exhibit a consistent enrichment or decrease in 65Cu.

Larson et al. (2003) reported a narrow range for chalcopyrites related to mafic intrusions

(+0.25 to +0.16‰). Primary high temperature copper sulfides from porphyry copper and

skarn deposits exhibit δ65Cu values from −0.83 to +1.34‰ (Maréchal et al., 1999; Larson

et al., 2003; Graham et al., 2004; Mathur et al., 2005), whereas δ65Cu values for the low

temperature secondary copper mineralization from porphyry copper deposits shows a

wider range from −16.06 to +9.98‰ (Mathur et al., 2009).

The reported copper isotope data from active and fossil submarine hydrothermal

systems show subtle differences but are consistent within each locality. The δ65Cu values from black smoker sulfides, massive sulfides and their alteration products from three sea- floor hydrothermal systems from the Mid-Atlantic Ridge (Lucky Strike, Logatchev, and

Rainbow) produce a range for the active and inactive hydrothermal fields from −0.34 to

+3.14‰ and from −1.52 to +3.31‰ respectively (Rouxel et al., 2004). Chalcopyrites from active hydrothermal vents from the East Pacific Rise show a δ65Cu range between

+0.34 to +1.15‰, whereas the δ65Cu values for the inactive vents range from −0.48 to

−0.18‰ (Zhu et al., 2000). Chalcopyrites from the inactive vents from the Galapagos Rift

produce also a similar range from −0.45 to −0.24‰ (Zhu et al., 2000).

Markl et al. (2006) reported copper isotope data from hydrothermal vein-type deposits from the Schwarzwald mining district in southwest Germany. Primary copper sulfides deposited at temperatures between 120 and 200°C gave a narrow cluster of δ65Cu

values around 0 ± 0.5‰, although some relicts of primary ore with evidence of oxidation

show copper isotope depletion down to −2.9‰. The δ65Cu values for secondary supergene copper mineralization are −1.55 to +2.41‰ (Fig. 2.15). Also, Haest et al.

(2009) reported the copper isotope variations for the Cu-Ag deposit of Dikulushi in the

Democratic Republic of Congo. Chalcopyrite and chalcocite show variable low Cu

isotope compositions (0.0 to –2.3‰ δ65Cu). The supergene mineralization (malachite,

azurite, and chrysocolla) are enriched in 65Cu (+1.37 to +2.65‰ δ65Cu).

Copper isotope data in sediment-hosted stratiform copper deposits were first

documented by Jiang et al. (2002), who reported a δ65Cu range between −3.7 and +0.3‰

for chalcopyrites and tetrahedrites from the Jinman deposit in China, which is a low-

temperature sandstone-hosted vein copper deposit. Recently, Asael et al. (2007, 2009)

published copper isotope data from the Timma Valley and Kupferschiefer copper districts. Primary copper sulfides from the Kupferschiefer district show δ65Cu values between −2.73 to +0.65‰. The copper sulfides from the Timma Valley district exhibit copper isotope values between −3.78 and −1.24‰, whereas secondary copper mineralization show a smaller range between −1.73 and −0.09‰ (Fig. 2.15).

Copper isotope data in seawater, river, and estuarine systems were published by

Bermin et al. (2006) and Vance et al. (2008). The δ65Cu values for dissolved copper in

rivers range between +0.02 to +1.45‰, and an average of δ65Cu = +0.68‰ is estimated

to be delivered to the oceans (Vance et al., 2008). The Cu-isotope data for dissolved

copper from two estuaries in SE England (Itchen and Beaulieu) range between +0.42 and

+0.94‰. Only the particulate copper phase from the Itchen estuary was reported and

presents a range from −0.24 to −1.02‰, which is lighter than the dissolved copper in that estuary. The δ65Cu values of seawater are higher than in the previous systems, and range

between +0.75 and +1.44‰ (Fig. 2.15); the enrichment in 65Cu in the dissolved copper in

the different waters is attributed to the Cu-bonding to organic complexes (Vance et al.,

2008).

2.8.9, Copper isotope fractionation

Copper isotope fractionation has been studied for abiotic and biotic processes,

although most of the studies have been focused on the abiotic processes. Zhu et al. (2002)

published experimental fractionation data for biotic and abiotic processes. The

65 65 65 fractionation factor Δ Cu (ΔCu(II)–Cu(I) = δ Cu(II) – δ Cu(I)) for the reduction and precipitation from Cu(II) solution to Cu(I) iodide solution at 20°C resulted in 4.03‰.

Lower factors were documented during the anoxic precipitation of covellite from aqueous copper sulfate solution, with Δ65Cu values of 3.47 and 2.72‰, at 2 and 40°C respectively

(Ehrlich et al., 2004). Asael et al. (2006) obtained similar fractionation factors for Cu

sulfides (chalcopyrite, covellite, chalcocite) precipitated by the reaction of pyrite and pyrrohotite with CuSO4 solution under anoxic conditions. The fractionation determined

using pyrrohotite as reactant at 40 and 100°C is 3.032 and 2.67‰ respectively, and using

pyrite is 2.66 and 2.69‰ for the same temperatures. Similar data was documented by

Mathur et al. (2005), who reported fractionation data from oxidative dissolution

experiments. Leached fluids released during abiotic oxidation of both chalcopyrite and

chalcocite were enriched in the δ65Cu= +1.90 and +5.34‰, respectively, than the starting

material δ65Cu= +0.58 and +2.60‰, respectively.

On the other hand, in processes with no redox change, fractionation is small.

Fractionation between 0.20 and 0.38‰ was documented when malachite precipitated

from Cu(II) solutions at 30°C, and from 0.17 to 0.31‰ at 50°C (Marechal and Sheppard

2002). Similar values were obtained for the precipitation of Cu(OH)2 at 20°C, and produced a mean value of 0.27‰ (Ehrlich et al., 2004).

A few experimental data show that lower fractionation factors apply for biotic processes. Zhu et al. (2002) published Δ65Cu values from −0.98 to −1.71‰ for the copper

uptake from a Cu(II) solution to azurin and yeast proteins, which demonstrate that these

particular proteins preferentially incorporate 63Cu. Mathur et al. (2005) demonstrated that

the δ65Cu value of aqueous copper from the dissolution of chalcopyrite and chalcocite

inoculated with Thiobacillus ferrooxidans was similar to that of the starting material,

suggesting the uptake of the 65Cu by the bacteria cells, as evidenced by the formation of

Cu-Fe oxide minerals surrounding the cell.

2.8.10, Copper isotope data for the Boléo mineralization

Copper isotopes were measured in the Cu-mineralization and in the Cu-rich

manganese oxide minerals. The pristine copper isotopic composition for the ore-solutions

in the Boléo district is impossible to know because of the scarcity and the micron-scale size of the primary sulfides from the Boléo mantos. However, based on the published copper isotope data for primary mineralization in several ore deposits, a δ65Cu value

~0‰ can be assumed for the sulfide mineralization for the Boléo district. The subsequent

oxidation of the primary sulfide ores from each manto could produce relicts slightly

depleted in the δ65Cu as reported previously in different mineral deposits (Zhu et al.,

2000; Rouxel et al., 2004; Markl et al., 2006). Experimental copper fractionation data

reported suggest that the remaining solutions after the oxidation would have higher δ65Cu values (Mathur et al., 2005).

Figure 2.16 models the copper isotope fractionation during the formation of the

Boléo mineralized mantos. Firstly, the mineralizing fluids ascended along the NW-SE faults system cross-cutting the ASL rocks, and encountered the biogenic pyrite reduced horizons within the fine facies at the beginning of the first sedimentary cycle of the Boléo

Formation (Fig. 2.16a). The mineralizing solutions replaced the biogenic pyrite and the copper isotope signature of the copper ores remained most likely around 0‰. At this stage the manganese oxides did not precipitate because of the unsuitable reducing conditions, therefore the manganese most likely remained in solution.

As mentioned before, the Boléo Formation records the continental-marine transition documented by the fluvial-marine prograding fan-delta system, characterized by a lacustrine environment in the fine-grained facies of the clastic sequence (Wilson and

Rocha 1955; Ochoa Landín 1998; Conly et al., 2006). Although the beginning of each sedimentary cycle was characterized by a lacustrine environment with low energy, followed by the input of meteoric water probably mixed with seawater, and served to increase the oxidation state and the pH conditions, and therefore oxidized the ore-bearing sulfides and allowed the precipitation of the manganese oxides.

Figure 2.16b depicts the proposed systematic for the copper fractionation in the

Boléo mineralizing mantos. Initially, the interaction of the probable mixture of meteoric and seawater with the mineralized manto promotes the oxidation of the sulfide ores. The copper-bearing fluids resulted from the oxidation remained in solution, and the copper

probably kept bonded to organic complexes, most likely enriched in the 65Cu as

documented in dissolved copper in estuarine environments (Vance et al., 2008); probably

a considerable amount of the heavy copper was flushed out the Santa Rosalía basin

through the pre-initial stages of the Gulf of California (Fig. 2.16a). The residual copper

sulfides and copper oxides are characterized by low δ65Cu values as seen in the different mantos (Table 2.6). Once the oxidizing conditions were appropriate, the manganese oxides precipitated with slightly enriched δ65Cu values, and the reason is that some of the

65Cu was still in solution (Fig. 2.16). The negative δ65Cu values from the chrysocolla and copper carbonates observed in the Boléo can be explained by the fractionation for the low temperature Cu(II)-minerals precipitated from the solutions, which could be insignificant

(0.2 to 0.4‰) as reported by the fractionation experimental data at temperatures lower than 50°C (Marechal and Sheppard 2002; Ehrlich et al., 2004).

The proposed copper isotope fractionation is supposed to occur in each manto, during the redox changes within the lacustrine conditions at the beginning of each sedimentary cycle. The copper fractionation in each manto was concomitant to each sedimentary cycle, and therefore was independently and subsequent to the fractionation in the manto located below in the stratigraphic column (Fig. 2.2). In contrast, if the hydrothermal copper-bearing fluids with δ65Cu ~0‰ interacted the manto formed before

(i.e. mineralizing fluids from manto 3 passing through the previously formed manto 4),

the copper fractionation during this process could be negligible, because of the lack of the

chemical traps for the ore-forming fluids, and the upwelling driven solutions caused by

subsidence of the active Santa Rosalía basin. An alternative explanation for the range of

fractionation for the copper isotopes in the Cu and Mn mineralization (~1.7 and ~1.4‰

respectively), could also be in terms of the smaller fractionation as the result of larger

mass transfer during the precipitation of copper, in conjunction with the biotic

fractionation as suggested in the experimental copper fractionation by Ehrlich et al.

(2004).

Figure 2.17 shows the histograms for the copper isotope data in the Cu and Mn

mineralization in the mantos. The copper isotope data for the manganese mineralization

from the different mantos vary around 0‰ (Fig. 2.17a-c). The limited available copper

isotope data from Cu-sulfides from manto 3 has two peaks. The lower peak is slightly

negative and close to the 0‰, whereas the higher peak has more negative δ65Cu (Fig.

2.17b). The copper isotope data for the secondary copper mineralization has also two

peaks, and are next to the Cu-sulfides picks, which suggests that the secondary copper

mineralization formed after the oxidation of the primary mineralization.

The single δ65Cu value for the manganese oxides from the Neptuno area agrees

with the range for the δ65Cu values for the Lucifer deposit (Fig. 2.15). According to the

geological observations, the manganese oxides from these localities are constrained to the first sedimentary cycle of the Boléo Formation, and the copper mineralization either as sulfide or oxide is essentially insignificant in Lucifer deposit and scarce at the Neptuno area. Although the copper content is much higher in the Boléo mantos, a similar copper fractionation systematic could explain the copper data for these deposits. The proposed fractionation in Lucifer deposit can be explained with hydrothermal solutions entering the

Lucifer sub-basin with a δ65Cu ~0‰ (Fig. 2.17d). The precipitated manganese oxides are

slightly enriched in the 63Cu, and the remaining solution is suggested to be enriched in the

heavier Cu. Ultimately, because of the continued contribution of the clastic material in

the sub-basin, the heavy copper is most likely flush out the basin.

The copper isotope data for the manganese oxide vein from the Gavilán deposit is

the highest δ65Cu value reported in this study (+0.48‰).Although the scarcity of the data

from the Gavilán deposit, the high copper isotope value can be explained in terms of the

geological context. As exposed above, the Gavilán deposit consists of manganese oxide

veins within the ASL volcanic rocks. The δ65Cu value agrees perfectly within the range for the primary hydrothermal mineralization and the igneous rocks, and an undistinguishable or no isotope fractionation is suggested during the deposition of the manganese oxide mineralization in this deposit.

The only copper isotope data from manganese oxides published in the literature is from the surface layers of ferro-manganese nodules from the Central Pacific core RC 17-

203, whose δ65Cu values range from 0.05 to 0.6‰ with a mean value of 0.31 ± 0.23‰, slightly differing from the related basaltic rocks (Albarède 2004).

2.8.11, Metal budget

Studies from the sediment-hosted stratiform copper deposits from the White Pine

and Kupferschiefer districts provide data about the volumes required to form an ore fluid

(Hitzman 2000). Approximately a volume of 8 × 1011 m3 of the Copper Harbor

Conglomerate is estimated to contribute the total metals at the White Pine deposit considering a basin of 2 km × 10 km × 40 km (Hitzman et al., 2005). For the case of the

Polish Kupferschiefer district, a sub-basin with dimensions of 900 m × 40 km × 160 km is needed to produce a volume of 5.8 × 1012 m3 of Rotliegendes source rocks.

The geological and geochemical features for the case of the copper mineralization

in the Boléo district, suggest the leaching of the ASL rocks and possibly the peninsular batholith, followed by the ascent of the mineralizing fluids, as oppose to the previous examples exposed above, which indicate that the metals are leached from the sedimentary units that hosts the mineralization or spatially related within the sedimentary pile.

Considering only the thickness (800 m), abundant distribution (see Figure 2.1), and the copper content in the ASL rocks (40-80 ppm), approximately a volume of at least 3 ×

1010 m3 is needed to be leached in order to contribute with the copper budget in the Boléo

mantos. The geological context suggests the possibility that part of the Santa Rosalía

basin and consequently part of the Boléo mineralization, have been washed away to the

Gulf of California during the tectonic evolution of the gulf, and therefore greater volume

of rocks is needed to contribute to the metal budget.

2.8.12, Origin of the Cu-Co-Zn and Mn mineralization in Santa Rosalía region

Field evidence for hydrothermal activity is recorded along mineralized NW-SE

structures around the Santa Rosalía region, as well as the manganese deposits mentioned

above; also hydrothermal activity is inferred by the juxtaposition of the high-grade

Co±Cu zones along the faults (Conly et al., 2006). Geochemistry of the manganese

mineralization from the different mantos at the Boléo and the manganese deposits

supports the hydrothermal origin for the mineralizing fluids and records the exhalative-

intraformational nature of the mineralizing system (Del Rio Salas et al., 2008). Even

though the ascending nature of the mineralizing fluids is clearly documented, downward

infiltration for the metal-bearing saline brine is inferred by several geological

observations, where the most significant is the stratigraphic distribution of the metals

within the mantos (Conly et al., 2006).

Figure 2.18a shows the formation of the first sedimentary cycle of the Boléo

Formation, which consists of fine-grained sediments (clays and silts). The first evidence

of mineralization is the high Mn and Fe oxides present along manto 4 which is hosted

within the fine-grained sediments and sometimes replacing the basal limestone overlying

the ASL volcanic rocks. The Pb isotope data in the Cu and Mn mineralization from the

Boléo mantos demonstrate that the sources of metals are the ASL rocks and the

peninsular batholith. Cation geothermometers applied in the vents from Concepcion Bay

constrain the reservoir temperature around 200°C (Pro-Ledesma et al., 2004). Deeper

fluids interacting with the ASL rocks below the Santa Rosalía basin had similar

temperatures or greater. This could explain the leaching of the metals from the ASL and

peninsular batholith, as well as the assumption of the δ65Cu for the mineralizing fluids.

The ascent of the mineralizing fluids is evidenced along the NW-SE structures

and clearly served as feeders zones. The mineralizing fluids encounter the fine-grained sediments in a low energy environment, and the formation of the mineralized manto 4 started. As the tectonic activity of the Santa Rosalía basin continued, a high energy environment allowed the deposition of the conglomerate member of the first sedimentary cycle. The abrupt input of continental material disrupted the hydrothermal activity and

probably the mineralizing fluids were washed out the basin through the pre-gulf of

California. Later, a calm period is evidenced again by the fine-grained sediments of the second sedimentary cycle during the evolution of the tectonic activity (Fig. 2.18b). The continue subsidence of the Santa Rosalía basin allowed the reactivation of the fault system and the ascent of the mineralizing fluids encounter the fine-grained sediments and started the formation of manto 3.

The first evidence of sulfide formation is recorded by the framboidal pyrites via the bacterial sulfate reduction (Ochoa Landín 1998). Later, the oxidizing mineralizing solutions that precipitate the ore-bearing sulfides coating the framboidal pyrites; these ore-bearing sulfides are formed via the bacterial sulfate reduction at higher temperatures

(Conly et al., 2006). After the formation of both the framboidal pyrites and the ore- bearing sulfides, both sulfides are oxidized during the interaction of a mix of meteoric water and seawater sulfate trapped within sediments. These fluids precipitate the gypsum veinlets cross-cutting the mantos and are characterized with similar δ34S signatures from

the sulfide mantos, and δ18O values located along the proposed mixing trend between

meteoric water and seawater sulfate trapped within sediments.

After the deposition of the fine-grained sediments of sedimentary cycle 2, a

change in the tectonic activity caused the immense continental input into the sedimentary

cycle 2 and caused the disruption of the mineralizing fluids probably in a similar

systematic to the sedimentary cycle 1. The formation of the mantos in the following

sedimentary cycles was most likely similar to the formation of the previous mantos.

The tectonic activity of the Santa Rosalía basin recorded in the sedimentary cycles

within the Boléo Formation allowed the cyclical conditions for the formation of each

mineralized manto. However, a similar geochemical systematic occurred during the formation of each manto traced by the geochemistry of the ore- and gangue-minerals.

The late diagenetic nature of the mineralization is evidenced by tectonic activity reflected in brecciation of the mantos 4, 3, 3A, and 2 (facies A). The brecciation is more conspicuous in manto 3 and 3A, and the sulfide mineralization is found in the clasts, mainly following the laminar planes, and less frequently in the matrix, indicating that the brecciation was prior to sulfide mineralization (Ochoa Landín 1998).

The current hydrothermal fluids in the intertidal hot springs and submarine hydrothermal vents reported in Concepción Bay support the low temperatures nature and the continuation of the hydrothermal activity and migration southern the Santa Rosalía region. The difference between the hydrothermal activity related to the formation of the

Boléo mineralization and the current hydrothermalism in Concepcion Bay, is the presence of the basins and sub-basins acting as physical traps during the early opening of the Gulf of California.

2.9, CONCLUSIONS

The geological context, the hydrothermal activity, and the Cu-Co-Zn mineralization in the Boléo district, along with the manganese deposits along the eastern coast of Baja California, confirm the strong relationship between the tectonic evolution of the Gulf of California and the ore-forming processes along the eastern coast.

The trace element concentrations in the manganese oxides from the Boléo mantos, along with the manganese mineralization from Lucifer, Neptuno, and the manganese deposits from Concepción Peninsula, demonstrate the hydrothermal origin and the exhalative nature for all manganese deposits reported in the present study. The REE geochemistry in the manganese oxides also support the hydrothermal nature and exclude the hydrogenous nature for the manganese oxides in the Santa Rosalía region.

The stable isotope geochemistry of the ore- and gangue-minerals from the localities contributes with the understanding of the formation of the Boléo Formation and the hosted mineralization. Sulfur isotopes indicate that the Gypsum member and the marine sulfate in pore sediments are the most important sources of sulfur for the mineralization processes in the Boléo district (Conly et al., 2006), and the recycling of the sulfur in the system is evidenced along the different stratigraphic levels and mineral stages. The carbon and oxygen isotope geochemistry indicates the existence of two end- members (seawater and meteoric water-organic material) for the source of carbonates in the district. Also, the C and O isotopes indicate that a similar isotope systematic is occurring at different stratigraphic levels within the Boléo Formation (Limestone member and the subsequent mantos above), and evidence of mixing between the end-members is recorded during the formation of each manto. However, the fluids involved during the formation of the secondary copper carbonate mineralization consist mostly of the local meteoric water. The C and O isotopic data recorded in the carbonates in the Gavilán manganese deposit indicate that the fluids involved in their formation are similar to carbonates precipitated from Miocene seawater at low temperatures, whereas the

carbonates from Lucifer manganese deposit also precipitated at lower temperatures but

fluids involved correspond most likely to meteoric water.

The model proposed for the copper isotope fractionation at the reducing

conditions during the formation of the Boléo mineralized mantos, consists of initial δ65Cu

values for the ore-sulfides around 0‰. The consecutive oxidation of the primary Cu-

sulfides produced fluids enriched in 65Cu, similar to the systematic documented in dissolved copper in estuarine environments (Vance et al., 2008), and the fluids enriched in the 65Cu most likely were flushed out from the Santa Rosalía basin system. The

residual Cu-sulfides are characterized by low δ65Cu values as seen in previous

experimental copper isotope fractionation. The formation of the manganese oxide mineralization is characterized by high δ65Cu values, and is explained by the slightly enriched δ65Cu values of the oxidizing ore fluids still in solution. The proposed copper

isotope fractionation is supposed to occur in each manto, during the redox changes within

the lacustrine conditions at the beginning of each sedimentary cycle.

Lead isotopes demonstrate that the sources for the ore metals in the Santa Rosalía

region are the ASL rocks and the Peninsular batholith, and that the early-rift volcanic

rocks, whose magmatism and volcanism are nearly coeval to the mineralization age, do

not contribute with metals for the ore-forming processes, although could contribute with

thermal energy to promote the hydrothermal activity. Lead and strontium isotopes show

that strontium isotopes in the Cu and Mn mineralization from the Boléo district and

Lucifer deposit are the result of mixing between two possible end-members: 1) ASL

rocks and the Peninsular batholith, and 2) the Gypsum member. The strontium isotope

signature in the mineralization from the Boléo district is the result of the interaction of the hydrothermal fluids with the gypsum member and marine sulfate in pore sediments, whereas in Lucifer slightly change in the 87Sr/86Sr ratio is recorded, which suggests almost negligible interaction of the hydrothermal fluids.

The metallogenic features in the Boléo mineralization agree with those established for the SSC deposits. The geological and geochemical data contributed in the present study and the previous studies, supports the sources and fluids involved in the formation of the mineralization in the Boléo district.

)

Figure 2.4, (a) Cross-section and (b) schematic stratigraphic column of the Lucifer manganese oxide deposit, northern the Boléo district. 99

100

Figure 2.5, Paragenetic sequence for the Boléo district (after Conly 2003)

101

Figure 2.6, NASC-normalized REE patterns for the manganese oxide mineralization from the Boléo region and the Mn deposits from Concepción Peninsula. The REE data of Mn deposits from Concepción Peninsula taken from Rodríguez Díaz (2009) and Rodríguez Díaz et al. (2010). REE data from hydrothermal and hydrogenous fields, and average fossil and modern deposits from Usui and Someya (1997).

102

Figure 2.7, Trace element discrimination diagram for manganese oxides deposits between supergene (or hydrogenous) and hydrothermal deposits (Nicholson 1992).

103

Figure 2.8, Histogram showing the sulfur isotope data for the sulfide and sulfate samples from the Boléo district (Ortlieb and Colleta 1984; Ochoa Landín 1998; Conly et al., 2006; present study).

104

Figure 2.9, Oxygen and sulfur isotope data of sulfates from the Boléo district. Samples located within the dotted field correspond to gypsum veinlets cross-cutting the indicated mantos; sulfate data outside dotted field from manto 3 and 4 taken from Conly et al. (2006). Gray square represents evaporite deposits precipitated from Miocene seawater (Claypool et al., 1980; Ortlieb and Colleta 1984).

105

106

Figure 2.11, 207Pb/204Pb vs. 206Pb/204Pb diagram after Rollinson (1993). The mantle reservoirs of Zindler and Hart (1986) are as follows: DM - depleted mantle; BSE - bulk silicate earth; EMI and EMII - enriched mantle. EMII coincides with the field of oceanic pelagic sediments; PREMA - prevalent mantle composition; MORB - mid-ocean ridge basalts. Note that the Miocene volcanic rocks from the Boléo district plots within the lower continental crust.

107

Figure 2.12, Lead isotope diagram showing in detail the Miocene volcanic rocks from the Santa Rosalía region. NHRL - Northern Hemisphere Reference Line; MORB gray field.

108

Figure 2.13, Lead isotope diagram showing the isotope data for the copper and manganese mineralization from the different mineralized mantos and the manganese deposits around Santa Rosalía. Also is shown the lead data fields for the Miocene volcanic rocks from the Boléo district and the peninsular batholith. NHRL - Northern Hemisphere Reference Line.

109

Figure 2.14, 206Pb/204Pb vs. 87Sr/86Sr diagram for the rocks and Cu and Mn mineralization from the Boléo district.

110

Supergene mineralization PCD (12)

Sulfides PCD (8,9,10,11) Cpy related to mafic intrusions (7) Continental igneous rocks (5,6) Schwarzwald Cu district (4) (s) (p) (s) Timma Cu (p) deposit (3) Kupferschiefer (3)

Jinman Cu deposit (2)

Calculated total Cu Estuary (1) Particulate Cu Estuary (1) Dissolved Cu Estuary (1) Dissolved Cu rivers Average river (1) Seawater (1) Gavilán (Mn) Lucifer* (Mn) Neptuno* (Mn) Boléo* (Mn) Boléo* (Cu)

‐5 ‐4 ‐3 ‐2 ‐10123 65 δ Cu (‰) Figure 2.15, Copper isotope variations for continental and marine environments. (*) Present study; (1) Vance et al., 2008; (2) Jiang et al., 2002; (3) Asael et al., 2009; (4) Markl et al., 2006; (5) Botfield 1999; (6) Li et al., 2009; (7 ) Larson et al., 2003; (8) Maréchal et al., 1999; (9) Larson et al., 2003; (10) Graham et al., 2004; (11) Mathur et al., 2005; (12) Mathur et al., 2009.

111

Figure 2.16, Schematic model for the copper isotope fractionation in the mineralized mantos from the Boléo district. a) Mineralizing fluids ascended along the fault system and encountered the biogenic pyrite reduced horizons within the fine facies at the beginning of the sedimentary cycle of the Boléo Formation. b) Proposed systematic for the copper isotope fractionation. (1) Seawater and meteoric water; (2) Oxidation of the sulfide ores producing fluids with higher δ65Cu; (3) Mineralization relicts with lower δ65Cu; (4) Continental flow of meteoric water through the pre-Gulf of California.

Table 2.1, Trace element concentrations in the manganese oxide ores from the Boléo District, Lucifer and manganese deposits from Concepción peninsula. Concentrations are expressed in ppm, otherwise indicated. Sample Comment Cu (wt%) Co Zn Mn (wt%) Fe (wt%) Ni V As Pb Sr Ba Cr Mn/Fe Co/Zn Ref. Boléo district BO 3507 Mn-oxide (º) 5.33 4,898 3,362 1.34 0.28 106 nd nd 106 na na na 4.72 1.46 2 BO 1007 Manto 2 13.42 12,479 38,847 2.41 1.51 3,856 125 50 719 na na na 1.60 0.32 2 95-231b Manto 3A 5.51 17,407 11,743 2.62 0.89 415 11 604 151 na na na 2.95 1.48 2 95-231e Manto 3 21.73 1,705 15,833 19.70 2.07 nd 49 nd 73 na na na 9.52 0.11 2 BO 0107 Manto 3 0.72 393 2,606 0.84 0.58 4 nd nd 519 na na na 1.46 0.15 2 BO 0407 Manto 3 4.19 3,218 4,800 0.69 0.40 64 11 32 329 na na na 1.74 0.67 2 BO 1907 Manto 3 68.17 nd 8,091 9.34 3.29 nd nd nd 4,162 na na na 2.84 - 2 BO 2107 Manto 3 14.83 4,682 12,864 2.09 1.32 nd 651 399 529 na na na 1.58 0.36 2 BO 2207 Manto 3 0.03 nd 1,735 5.52 0.73 nd nd nd 0 na na na 7.52 - 2 BO 1507 Manto 4 6.63 1,640 10,250 8.19 2.00 nd nd nd 82 na na na 4.11 0.16 2 BO 2907 Manto 4 5.29 699 11,561 1.25 0.42 nd 350 78 1,516 na na na 2.99 0.06 2 RE 9902a Manto 4 0.19 1,608 17,025 52.66 3.90 56 207 na na 2,408 7,875 6 13.51 0.09 1 RE 10002 Manto 4 8.38 32 3,625 41.47 2.02 58 1,228 na na 13,700 49,750 245 20.48 0.01 1 BH-1 Mn-vein (*) 13.80 nd nd 8.61 0.01 nd 101 nd 101 na na na 856.00 - 2 Avg. 0.44

Neptuno area RE 9702 Manto 4 2.70 36 1,580 na na 34 185 na na 2,640 9,140 230 - 0.02 1 RE 9802a Manto 4 0.23 1,608 3,875 58.47 1.87 49 nd na na 2,395 5,300 8 31.19 0.41 1 LF-45 Manto 4 0.10 150 500 na na 10 20 na 40 1,190 1,475 na - 0.30 1 LF-46 Manto 4 0.05 195 858 na na 15 10 na 45 296 331 na - 0.23 1 LF-47 Manto 4 0.39 562 2,030 na na 15 135 na 3,470 900 1,345 na - 0.28 1 Avg. 0.25

Lucifer Deposit RE 4102 Manto 4 na 22 78.90 na na 72 154 na na na na na - 0.28 1 RE 4302 (j) Manto 4 0.20 30 2,408 3.72 11.80 9 644 na na 91 127 nd 0.32 - 1 RE 4402 Manto 4 0.00 226 1,440 37.75 6.35 27 639 na na 2,480 5,460 nd 5.94 0.16 1 RE 4502a Manto 4 0.32 359 2,480 na na 32 967 na na 2,530 6,060 nd - 0.14 1 RE 4602 Manto 4 0.00 350 2,043 18.76 3.39 20 480 na na 2,378 4,550 35 5.53 0.17 1 RE 5102 (j) Manto 4 na 2,975 242 0.67 15.95 1 497 na na 78 62 nd 0.04 - 1 114 RE 5302 Manto 4 0.32 345 793 na na 59 658 na na 5,690 9,350 209 - 0.44 1

Table 2.1, (Continued) Sample Comment Cu (wt%) Co Zn Mn (wt%) Fe (wt%) Ni V As Pb Sr Ba Cr Mn/Fe Co/Zn Ref. RE 5402 Manto 4 0.00 463 1,735 61.80 0.29 49 790 na na 4,150 3,975 19 212.92 0.27 1 RE 5502 Manto 4 0.10 112 464 na na 39 522 na na 3,230 3,870 11 - 0.24 1 LF-24 Manto 4 0.20 404 1,140 na na 5 745 na 2,570 4,890 9,500 na - 0.35 1 LF-25 Manto 4 0.12 43 1,330 na na 5 550 na 1,105 487 1,355 na - 0.00 1 LF-26 Manto 4 0.14 231 1,240 na na 5 590 na 1,770 3,830 5,180 na - 0.19 1 Avg. 0.22

Gavilán deposit, Concepción peninsula - RE 8602 Mn-vein 0.07 6 903 22.27 1.20 19 372 na na 440 1,583 114 18.56 0.01 1 Ga1-1A Mn-vein 0.05 496 419 63.08 0.32 50 684 399 1002 1290 7520 13 197.13 1.18 3 Gav-3A Mn-vein 0.02 212 1970 64.74 0.26 45 320 144 4467 1140 1150 5 249.00 0.11 3 Gav-4A Mn-vein 0.12 445 1619 56.02 0.22 21 933 230 430 5190 38100 8 254.64 0.27 3 GavF-4A Mn-vein 0.05 149 1175 36.84 2.39 20 876 225 990 5940 28600 13 15.41 0.13 3 Gav-PA Mn-vein 0.06 41 4223 56.62 0.62 40 779 361 950 2490 10900 8 91.32 0.01 3 Avg. 0.28

Guadalupe, Concepción peninsula Man-4 Mn-Qz vein 0.003 1 80 32.95 3.83 20 24 61 5 923 64000 <20 8.60 0.01 3 Man-5a Mn-Qz breccia 0.022 14 140 32.48 1.2 20 118 519 103 1380 85500 <20 27.07 0.10 3 Man-5b Mn-calcite vein 0.03 14 70 37.81 0.14 <20 76 44 127 4240 28300 <20 270.07 0.20 3 Avg. 0.10

Santa Rosa, Concepción peninsula SR-b Breccia 0.062 55 280 18.26 2.92 40 778 1380 245 1020 99000 50 6.25 0.20 3,4 SR-c Breccia 0.114 77 340 38.7 0.37 30 1410 1340 49 796 124000 <20 104.59 0.23 3,4 SR-h Breccia 0.133 57 330 33.15 0.47 20 1310 937 82 898 137000 50 70.53 0.17 3,4 SR-Rod Breccia 0.003 27 80 34.73 0.79 <20 23 30 6 1020 65700 20 43.96 0.34 3,4 SR-y-2 Breccia 0.163 91 340 17.05 0.22 30 1110 1960 58 666 117000 <20 77.50 0.27 3,4 Avg. 0.24

Notes: (na) Not analized; (nd) Below detection limit ; (-) not calculated; (*) Manganese oxide vein within the ASL rocks; (º) Manganese oxide horizon in sediments from the base of Gloria Formation ; (j) Jasperoid; Sources (1) Del Rio Salas et al 2008, (2) Present study, (3) Rodríguez Díaz 2009, (4) Rodríguez Díaz et al 2010. 115

Table 2.2, Rare earth element concentration in ppm from manganese oxides from the Boleo district and adjacent areas in Baja California Sur, México.

Sample Comment La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Σ REE (La/Sm)n (Gd/Yb)n (La/Yb)n Ref. Boléo district BO 3507 Mn-oxideº 63.4 99.9 10.1 34.5 6.1 3.6 7.9 0.4 6.8 0.6 5.0 0.3 5.0 0.3 243.8 2.0 0.9 1.3 2 BO 1007 Manto 2 16.4 16.5 1.2 11.9 0.9 0.7 1.0 0.1 0.8 0.2 0.5 0.1 0.4 0.1 50.7 3.6 1.5 4.1 2 95-231b Manto 3A 25.9 75.4 4.9 49.6 1.8 23.3 1.9 0.3 1.8 0.4 1.2 0.2 1.3 0.2 188.2 2.8 0.8 2.0 2 95-231e Manto 3 35.2 70.4 4.0 35.1 2.6 9.2 3.0 0.4 2.4 0.5 1.8 0.2 1.4 0.3 166.6 2.6 1.2 2.5 2 BO 0107 Manto 3 11.5 26.5 3.2 12.7 2.7 0.4 2.8 0.2 2.3 0.2 0.5 0.1 0.6 0.1 63.6 0.8 2.8 2.1 2 BO 0407 Manto 3 81.4 230.2 15.0 49.7 7.6 7.6 7.8 0.5 7.8 1.7 5.8 0.3 6.4 0.3 422.1 2.1 0.7 1.3 2 BO 1907 Manto 3 9.6 22.9 2.7 11.2 2.0 0.5 2.2 0.3 1.8 0.4 1.1 0.2 1.1 0.2 56.1 0.9 1.2 0.9 2 BO 2107 Manto 3 231.8 302.0 17.0 72.2 9.1 22.0 10.0 0.6 10.1 0.8 8.3 0.4 8.7 0.5 693.4 4.9 0.7 2.7 2 BO 2207 Manto 3 1.2 2.0 0.2 1.3 0.2 0.3 0.2 0.0 0.2 0.0 0.1 0.0 0.1 0.0 5.9 1.3 1.2 1.2 2 BO 1507 Manto 4 104.7 126.4 12.9 44.4 4.7 3.5 5.6 0.9 5.4 1.2 3.6 0.5 3.1 0.5 317.5 4.2 1.0 3.4 2 BO 2907 Manto 4 8.0 9.7 1.1 10.1 0.2 4.0 0.3 0.0 0.2 0.0 0.1 0.0 0.1 0.0 34.0 7.0 1.4 7.5 2 RE 9902a Manto 4 90.1 113.3 4.2 21.7 4.3 0.1 4.8 0.7 5.6 1.3 4.8 0.5 4.3 0.6 256.1 4.1 0.6 2.1 1 RE 10002 Manto 4 147.1 267.3 6.7 35.0 11.4 nd 6.8 1.1 9.4 1.9 7.1 0.9 7.9 1.0 503.6 2.5 0.5 1.9 1 BH-1 Mn-vein* 0.5 2.9 0.1 1.1 0.1 0.6 0.1 0.0 0.1 0.0 0.1 0.0 0.1 0.0 5.6 1.6 0.4 0.5 2 Average 214.8 2.9 1.1 2.4 Neptuno area LF-45 Manto 4 9.3 12.5 1.9 5.5 1.0 0.6 1.1 0.2 1.0 0.2 0.7 0.1 0.7 0.1 34.9 1.8 0.9 1.3 1 LF-46 Manto 4 424.0 996.0 90.3 285.0 46.1 20.7 36.1 4.4 18.2 2.6 6.4 0.6 3.5 0.4 1934.3 1.8 5.8 12.1 1 LF-47 Manto 4 41.7 122.0 10.1 30.0 5.5 2.6 3.9 0.6 3.5 0.6 1.6 0.3 1.9 0.3 224.6 1.5 1.2 2.2 1 RE 9702 Manto 4 28.9 83.2 6.1 21.2 3.4 8.4 3.1 0.5 3.1 0.6 1.8 0.1 1.6 0.2 162.0 1.6 1.1 1.8 1 RE 9802a Manto 4 122.3 233.0 13.8 51.8 7.9 3.8 7.6 1.0 7.2 1.4 4.8 0.5 3.8 0.5 459.5 3.0 1.1 3.3 1 Average 563.1 1.9 2.0 4.1 Lucifer deposit LF-24 Manto 4 8.6 10.5 1.3 4.5 0.8 0.2 0.8 0.1 0.5 0.1 0.3 0.1 0.3 0.1 28.2 2.1 1.5 2.9 1 LF-25 Manto 4 2.7 7.0 0.6 2.5 0.5 0.1 0.8 0.1 0.6 0.1 0.4 0.1 0.4 0.1 16.0 1.0 1.1 0.7 1 LF-26 Manto 4 5.1 7.5 0.7 2.5 0.5 0.1 0.6 0.1 0.4 0.1 0.2 0.1 0.2 0.1 18.2 2.0 1.7 2.6 1 LF-27 Manto 4 0.5 1.0 0.2 0.5 0.1 0.1 0.1 0.1 0.1 0.1 0.1 0.1 0.1 0.1 3.2 1.0 0.6 0.5 1 RE 4302 Manto 4 2.0 7.4 0.4 1.9 0.4 0.1 0.6 0.1 0.5 0.1 0.4 nd 0.3 0.1 14.5 0.9 1.1 0.6 1 RE 4402 Manto 4 4.0 8.3 0.5 2.8 1.1 0.2 0.5 0.0 0.4 0.1 0.3 nd 0.2 0.0 18.5 0.7 1.5 2.1 1

RE 4602 Manto 4 20.0 22.7 2.8 12.3 2.3 0.2 2.0 0.3 1.6 0.3 1.0 nd 0.8 0.1 66.5 1.7 1.4 2.5 1 116 RE 4702 Manto 4 9.7 18.3 3.6 15.0 2.8 0.7 2.4 0.4 2.0 0.4 1.3 0.1 1.2 0.2 57.9 0.7 1.2 0.8 1

Table 2.2, (Continued)

Sample Comment La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Σ REE (La/Sm)n (Gd/Yb)n (La/Yb)n Ref. RE 4802a Manto 4 9.2 15.5 0.7 3.6 1.8 nd 0.6 0.1 0.5 0.1 0.4 nd 0.3 0.0 32.8 1.0 1.3 3.4 1 RE 5202 Manto 4 2.7 3.4 0.4 2.1 1.6 nd 0.2 nd 0.2 0.0 0.1 0.1 0.1 nd 11.0 0.3 1.7 3.4 1 RE 5402 Manto 4 1.8 3.7 0.2 1.0 0.7 nd 0.2 0.0 0.1 0.0 0.0 nd 0.1 0.0 7.9 0.5 1.6 2.9 1 RE 5902 Manto 4 8.1 7.1 0.8 4.6 1.8 nd 1.0 0.1 0.9 0.1 0.6 0.0 0.5 0.0 25.7 0.9 1.2 1.7 1 Average 25.0 1.0 1.3 2.0

Gavilán deposit RE 8602 Mn-vein 9.8 8.1 1.1 5.3 1.5 0.1 1.6 0.2 1.9 0.4 1.4 0.1 1.1 0.1 32.9 1.3 0.8 0.9 1 Ga1-1A Mn-vein 7.0 8.3 1.3 5.9 1.4 0.4 1.9 0.3 2.0 0.4 1.5 0.2 1.4 0.2 32.3 1.0 0.8 0.5 3 Gav-3A Mn-vein 5.0 5.6 1.1 5.4 1.4 0.5 1.8 0.3 2.0 0.4 1.3 0.2 1.1 0.2 26.3 0.7 0.9 0.5 3 Gav-4A Mn-vein 22.7 7.7 2.8 12.6 2.8 0.5 3.9 0.6 4.0 0.8 2.5 0.4 2.2 0.3 63.8 1.6 1.0 1.0 3 GavF-4A Mn-vein 14.1 31.0 2.6 11.8 2.9 0.4 3.4 0.7 4.2 0.9 2.9 0.4 2.6 0.4 78.3 0.9 0.7 0.5 3 Gav-PA Mn-vein 11.5 5.1 1.8 9.4 2.5 0.6 3.7 0.6 3.6 0.8 2.5 0.3 1.9 0.3 44.6 0.9 1.1 0.6 3 Average 46.3 1.1 0.9 0.7

Guadalupe Man-4 Mn-Qz vein 19.2 20.6 2.4 6.5 1.2 <0.05 1.5 0.2 0.9 0.1 0.3 <0.05 0.1 <0.04 53.0 3.1 8.5 19.2 3 Man-5a Mn-Qz bx 15.4 16.1 2.3 7.5 1.5 0.1 1.6 0.2 1.2 0.2 0.6 <0.05 0.4 0.1 47.2 2.0 2.3 3.9 3 Man-5b Mn-cal vein 1.9 5.5 0.4 1.5 0.3 <0.05 0.5 <0.1 0.2 <0.1 <0.1 <0.05 <0.1 <0.04 10.3 1.2 - - 3 Average 36.8 2.1 5.4 11.5

Santa Rosa SR-b Breccia 14.6 30.4 3.0 10.6 2.0 0.0 1.8 0.3 1.7 0.3 0.8 1.8 0.3 1.7 69.4 1.4 3.4 4.9 3,4 SR-c Breccia 7.1 5.8 1.13 3.3 0.6 0.0 0.7 <0.1 0.5 <0.1 0.3 0.7 <0.1 0.5 20.7 2.3 - - 3,4 SR-h Breccia 6.2 7.3 1.02 2.7 0.4 0.0 0.4 <0.1 0.4 <0.1 0.2 0.4 <0.1 0.4 19.5 3.0 - - 3,4 SR-Rod Breccia 3.6 8.7 1 3.5 0.8 0.0 0.7 0.1 0.7 0.1 0.3 0.7 0.1 0.7 21.0 0.9 4.0 3.6 3,4 SR-y-2 Breccia 11.2 12.7 2.33 7.5 1.4 0.1 1.5 0.2 1.2 0.2 0.6 <0.05 0.6 0.1 39.6 1.5 1.4 1.9 3,4 Average 34.0 1.8 2.9 3.4

Notes: (na) not analized; (nd) below detection limit; (-) not calculated; (*) manganese oxide vein within the ASL rocks; (º) manganese oxide horizon in sediments from the base of the Gloria Fm; (bx) breccia; (cal) calcite; Sources (1) Del Rio Salas et al 2008, (2) Present study, (3) Rodríguez Díaz 2009, (4) Rodríguez Díaz et al 2010.

117

118

Table 2.3, Sulfur isotope data for the sulfide phases in the Cu-Co-Zn Boléo District δ34S Sample Comment (‰) References 06-964 Manto 2 -1.7 1 95-3339 Manto 3 -5.3 1 95-255C Manto 3 -6.8 1 272-180.40 Manto -6.1 2 272-181.25 Manto -3.7 2 281-145.04 Manto -3.8 2 370-58.00 Manto -5.6 2 370-58.00c Manto -1.8 2 391-160.30 Manto -8.0 2 391-161.45 Manto -5.4 2 466-94.80 Manto -7.6 2 466-95.50 Manto -9.6 2 527-74.28 Manto -10.7 2 527-75.60 Manto -4.5 2 527-75.70 Manto -3.3 2 527-79.20 Manto -13.7 2 95-50 Manto 3 -10.9 3 95-50 Manto 3 -33.6 3 94-26 Manto 3 -29.6 3 94-26 (r) Manto 3 -33.4 3 94-26 Manto 3 -17.6 3 94-32 Manto 3 -22.1 3 95-5 Manto 3 -20.5 3 94-44 Manto 3 -20.5 3 95-10 Manto 3 -27.7 3 94-87 Manto 3 -27.7 3

Sources: (1) Present study, (2) Conly et al 2006, (3) Ochoa Landín 1998.

119

Table 2.4, Sulfur and oxygen isotope data for sulfate phases in the Cu-Co-Zn Boléo District, Baja California Sur 34 18 δ S δ O(SO4) Sample Comment (‰) (‰) Ref. 515e Manto 2 -19.2 -0.5 1 515c Manto 2 -18.6 0.9 1 515d Manto 2 -17.4 -0.8 1 BO05-07 Manto 3 14.8 6.8 1 840 CRb Manto 3 -20.8 6.6 1 BOb06-07 Manto 3 -21.4 4.1 1 BOa06-07 Manto 3 -18.4 8.3 1 840 Cra Manto 3 -24.8 2.6 1 BO16-07 Manto 3 13.3 7.0 1 BO32-07 Manto 3 -33.5 3.7 1 BO14-07 Manto 4 -16.9 5.2 1 96-TR072-1.50 Manto sulfate 18.9 10.7 2 397-198.90 Manto sulfate 17.5 11.2 2 527-83.25 Manto sulfate 21.1 12.3 2 429-48.65 Manto sulfate -5.1 14.5 2 BO25-07 Mn Basal gypsum 23.6 12.7 1 BO25-07 Cu Basal gypsum 24.1 11.0 1 BO34-07 Basal gypsum 24.0 11.9 1 454-36.85 Basal gypsum 21.7 6.5 2 746-108.54 Basal gypsum 22.3 13.1 2 CPP6-188.54 Basal gypsum 22.1 12.8 2 96-0827-026 Basal gypsum 22.6 10.8 2 96-0924-136 Basal gypsum 21.6 6.9 2 00-0503-217 Basal gypsum 21.6 8.4 2 00-0503-218 Basal gypsum 21.9 8.7 2 00-0503-220 Basal gypsum 21.7 8.8 2 Gypsum Basal gypsum 22.0 13.0 3

Sources: (1) Present study, (2) Conly et al 2006, (3) Ortlieb and Colleta 1984.

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Table 2.5, Carbon and oxygen isotope data for the Cu-Co-Zn Boléo District and adjacent areas δ13C δ18O Sample (‰) (‰) Comment Locality References 96-0810-011 -6.0 20.8 cal from BL Boléo 2 96-0827-040 4.3 28.7 dol from BL Boléo 2 96-0827-040 4.0 28.3 Repeat Boléo 2 96-0827-040 0.0 24.6 cal from BL Boléo 2 96-0827-040 0.6 25.3 Repeat Boléo 2 96-0910-099 4.4 29.3 dol from BL Boléo 2 96-0910-099 -0.6 26.2 cal from BL Boléo 2 96-0921-126 -7.7 19.5 cal from TL Boléo 2 BO 1207a -0.2 26.5 cal from BL Boléo 1 BO 1207e -1.8 25.1 cal from BL Boléo 1 BO 1207c -1.7 25.2 BL fsl Boléo 1 BO 1207d -3.1 24.1 BL fsl Boléo 1 BO 2307 2.3 30.8 BL Boléo 1 BO 2707 -0.8 28.9 BL Boléo 1 BO 1207b -5.7 21.6 fsed Boléo 1 BO 1307 -1.2 24.7 carb-seds from BL Boléo 1 429-53.70 -9.6 18.6 cal from cgl, manto 4 Boléo 2 705-101.00 -4.4 24.4 dol from cgl, manto 3 Boléo 2 036-82.00 -9.7 18.7 cal from cgl, manto 3 Boléo 2 97-0220-140 -2.4 20.4 cal from Tcbx Boléo 2 451-86.04 1.4 25.6 cal, manto 0 Boléo 2 455-90.97 1.5 25.6 Repeat Boléo 2 455-90.97 -6.1 22.5 cal, manto 0 Boléo 2 520-112.25 -1.3 25.0 cal, manto 1 Boléo 2 032-21.16 -6.6 23.9 dol, manto 2 Boléo 2 294-222.20 -8.5 24.7 sd, manto 2 Boléo 2 SM3-1 -0.4 28.5 cal, manto 2 Boléo 3 94-70 -9.7 22.4 cal, manto 2 Boléo 3 94-70 -11.1 21.1 cal, manto 2 Boléo 3 SR-PRC4-2 -10.9 21.4 cal, manto 2 Boléo 3 515e -4.7 26.5 az, manto 2 Boléo 1 515e -2.8 26.9 az, manto 2 Boléo 1 515c -3.5 24.1 az, manto 2 Boléo 1 515d -0.7 27.3 az, manto 2 Boléo 1 345-42.50a -6.2 25.1 sd, manto 3 Boléo 2 345-42.50a -11.6 19.0 cal, manto 3 Boléo 2 345-42.50b -6.8 23.4 sd, manto 3 Boléo 2 397-198.90a -1.9 24.2 cal, manto 3 Boléo 2 397-198.90a -2.7 24.3 sd, manto 3 Boléo 2 397-198.90b -1.4 24.7 cal, manto 3 Boléo 2

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Table 2.5, (Continued) δ13C δ18O Sample (‰) (‰) Comment Locality References 397-202.11 -2.9 24.5 sd, manto 3 Boléo 2 397-202.11 -2.6 23.9 cal, manto 3 Boléo 2 709-52.15 -7.7 23.2 cal, manto 3 Boléo 2 527-74.28 -9.2 26.1 sd, manto 3 Boléo 2 527-74.28 -7.9 22.1 cal, manto 3 Boléo 2 94-79 -3.9 22.0 cal, manto 3 Boléo 3 95-141 -7.1 21.7 cal, manto 3 Boléo 3 95-308 -1.1 31.6 cal, manto 3 Boléo 3 95-344 -4.0 20.9 cal, manto 3 Boléo 3 96-644 -7.1 21.4 cal, manto 3 Boléo 3 96-588 -8.3 19.9 cal, manto 3 Boléo 3 96-149C -2.1 25.6 cal, manto 3 Boléo 3 840 Cra-Ia -3.5 26.4 az, manto 3 Boléo 1 840 Cra-Ib -3.1 26.9 az, manto 3 Boléo 1 95-149 -2.0 25.7 mal, manto 3 Boléo 1 96-149c -2.8 25.1 mal, manto 3 Boléo 1 96-149c -2.0 26.0 mal, manto 3 Boléo 1 96-682 0.2 27.4 carb-seds, manto 3 Boléo 1 96-682 -3.1 25.2 carb-seds, manto 3 Boléo 1 96-682 -3.4 23.0 carb-seds, manto 3 Boléo 1 96-824 -6.4 17.7 carb-seds, manto 3 Boléo 1 96-542 6.7 31.1 cal, manto 3A Boléo 3 96-479a -6.7 22.4 cal, manto 3A Boléo 3 96-479b -4.9 22.5 cal, manto 3A Boléo 3 96-642 -1.8 24.9 cal, manto 3A Boléo 3 95-100 -7.3 31.7 cal, manto 4 Boléo 3 429-50.70 -3.1 21.5 cal, manto 4 Boléo 2 429-50.70 -3.2 21.8 Repeat Boléo 2 429-50.70 0.0 24.4 sd, manto 4 Boléo 2 429-50.70 -0.4 24.3 Repeat Boléo 2 429-52.95 -3.8 23.6 cal, manto 4 Boléo 2 BO 1407 4.3 26.7 cal, manto 4 Boléo 1 95-231f-Ia -11.5 19.1 cal, manto 4 Boléo 1 95-231f-Ib -10.9 19.2 cal, manto 4 Boléo 1 R4 -9.9 22.9 cal, manto 4 Boléo 3 BH-1 -4.1 -6.5 cal, Mn vein Boléo 1 RE8802a -0.7 28.0 cal vein within ASL Gavilán 1 RE8802b -2.1 29.1 cal vein within ASL Gavilán 1 RE8802c 2.9 31.4 cal vein within ASL Gavilán 1 RE8802d 1.9 30.3 cal vein within ASL Gavilán 1

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Table 2.5, (Continued) δ13C δ18O Sample (‰) (‰) Comment Locality References RE8802d 1.8 30.4 cal vein within ASL Gavilán 1 RE8802e -0.6 29.2 cal vein within ASL Gavilán 1 RE8802e -2.0 28.6 cal vein within ASL Gavilán 1 RE8802e -1.0 28.9 cal vein within ASL Gavilán 1 JLC-1 -12.5 20.2 cal with jasper Lucifer 1 JLC-1 rep -12.6 19.8 cal with jasper Lucifer 1 JLC-2 -12.5 20.2 cal with jasper Lucifer 1

Notes: (BL) Boléo limestone member; (TL) Tirabuzon limestone; (cgl#) Boléo conglomerate from manto; (fsl) fossiliferous; (fsed) sediments filling fossils from Boleo limestone; (Tcbx) carbonate cement of brecciated arc-to-rift volcanic rocks; (carb-seds) carbonaceous sediments; (ASL) Andesite of Sierra Santa Lucia; (cal) calcite; (dol) dolomite; (az) azurite; (mal) malachite; (sd) siderite; Sources: (1) Present study, (2) Conly et al 2006, (3) Ochoa Landín 1998.

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Table 2.6, Copper isotope data from the Boléo district, Lucifer, and Gavilán deposit in Concepción peninsula. δ65Cu Sample (‰) Comment Locality IEC 287b -1.39 ccl, Cu Manto 2 Boléo yes 06-964 -1.01 ccl, Cu Manto 2 Boléo yes 06-964rep -1.03 ccl, Cu Manto 2 Boléo yes 06-964rep -1.05 ccl, Cu Manto 2 Boléo yes BO 0907 -0.43 ccl, Cu Manto 2 Boléo yes BO 0907rep -0.40 ccl, Cu Manto 2 Boléo yes BO 1007 0.16 MnOx, Mn Manto 2 Boléo yes 95-255c -0.32 ccl, Cu Manto 3 Boléo yes 95-333a -0.13 pcs, Cu Manto 3 Boléo yes BO 0407 -0.47 ccl, Cu Manto 3 Boléo no BO 1907 -0.19 ccl, Cu Manto 3 Boléo yes BO 2007 -0.20 ccl, Cu Manto 3 Boléo yes BO 2107 -0.22 ccl, Cu Manto 3 Boléo yes 95-231d -0.54 mal, Cu Manto 3 Boléo no 95-255c -1.33 ccl, Cu Manto 3 Boléo yes 95-255crep -1.39 ccl, Cu Manto 3 Boléo yes 95-333b1 -1.17 pcs, Cu Manto 3 Boléo yes 95-333b2 -1.20 pcs, Cu Manto 3 Boléo yes MBLU -1.57 bol, Cu Manto 3 Boléo yes MBRA -1.62 bol, Cu Manto 3 Boléo yes BO 1907a -0.47 MnOx, Mn Manto 3 Boléo yes BO 2107 -0.17 MnOx, Mn Manto 3 Boléo yes BO 1507 -0.22 MnOx, Mn Manto 4 Boléo yes RE 9902 -0.31 MnOx, Mn Manto 4 Boléo yes CuBolYes -1.58 bol in gy Boléo yes BO 2207 -0.73 MnOx vein, Mn Manto 3 Boléo yes BH-1 -0.20 ccl in MnOx vein Boléo yes RE 4802a -0.85 MnOx, Mn Manto 4 Lucifer yes RE 4802b -0.86 MnOx, Mn Manto 4 Lucifer yes RE 4502 -0.17 MnOx, Mn Manto 4 Lucifer yes RE 9802 -0.31 MnOx, Mn Manto 4 Neptuno yes RE 8602 0.48 MnOx vein Gavilán yes

Notes: (ccl) chrysocolla; (bol) boleite; (mal) malachite; (gy) gypsum; (pcs) primary copper sulfides; (MnOx) manganese oxides; (IEC) ion exchange chromatography.

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Table 2.7, Lead and strontium isotope data of the Boléo district and adjacent deposits Sample Description Locality 87Sr/86Sr 206Pb/204Pb 207Pb/204Pb 208Pb/204Pb BO3307 Peninsular batholith Boléo 0.7043 18.803 15.602 38.603 BO3307(r) Peninsular batholith Boléo 0.7044 18.798 15.604 38.608 RE0102 Andesitic flow (ASL) Boléo 0.7045 18.752 15.608 38.559 RE0102(r) Andesitic flow (ASL) Boléo 0.7045 18.747 15.600 38.534 RE3602 Basaltic andesite flow (ASL) Boléo 0.7041 18.748 15.600 38.506 RE3802 Basaltic flow (ASL) Boléo 0.7041 18.776 15.606 38.545 RE4002 Basaltic andesite flow (ASL) Boléo 0.7041 18.748 15.601 38.506 BO2607 Basaltic andesite flow (ASL) Boléo 0.7041 18.679 15.595 38.495 RE11002 Boléo basalt (ERV) Boléo 0.7036 18.623 15.576 38.392 RE8302 Basaltic andesite flow (ASL) Gavilán 0.7038 18.621 15.584 38.432 BO2307 Limestone member Boléo 0.7062 18.772 15.605 38.547 BO2507 Gypsum member Boléo 0.7085 18.720 15.591 38.484 BO3607 Cinta Colorada member Boléo 0.7042 18.735 15.595 38.517 BO 2007 Boléo sediments Boléo - 18.761 15.596 38.499 95-231b Boléo sediments Boléo 0.7055 18.770 15.597 38.500 BO35-07 base de la Fm Gloria Boléo 0.7071 18.727 15.585 38.461 06-964 Cu Manto 2 Boléo - 06-964 Cu Manto 2 Boléo 0.7067 18.910 15.619 38.540 287b Cu Manto 2 Boléo - 18.724 15.590 38.481 287b Cu Manto 2 Boléo 0.7067 18.672 15.569 38.468 BO0907 Cu Manto 2 Boléo - 18.720 15.593 38.481 BO0907(r) Cu Manto 2 Boléo 0.7088 BO0907 Cu Manto 2 Boléo 0.7070 18.712 15.588 38.477 BO1007 Mn Manto 2 Boléo - 18.721 15.589 38.469 95-231b Mn Manto 3A Boléo - 18.751 15.593 38.494 95-231b(r) Mn Manto 3A Boléo 0.7063 18.751 15.591 38.483 06-1018c Cu Manto 3 Boléo - 06-1018 Cu Manto 3 Boléo 0.7068 18.657 15.573 38.469 95-231d Cu Manto 3 Boléo 0.7066 18.747 15.598 38.507 95-255c Cu Manto 3 Boléo - 18.727 15.583 38.440 95-333a Cu Manto 3 Boléo - 18.866 15.607 38.500 BO0507 Cu Manto 3 Boléo 0.7062 18.747 15.603 38.534 BO2007 Cu Manto 3 Boléo 0.7066 18.742 15.599 38.512 BO2107 Cu Manto 3 Boléo - BO2107 Cu Manto 3 Boléo - BO2107 Cu Manto 3 Boléo 0.7065 BoleoML Cu Manto 3, Boleita Boléo - 18.734 15.594 38.513 95-231e Mn Manto 3 Boléo 0.7065 18.720 15.593 38.481 BO0107 Mn Manto 3 Boléo 0.7043 18.727 15.598 38.497 BO0407 Mn Manto 3 Boléo - 18.762 15.597 38.511 BO1907 Mn Manto 3 Boléo - 18.708 15.588 38.452 BO2107 Mn Manto 3 Boléo - 18.740 15.591 38.473 BO2207 Mn Manto 3 Boléo - 18.800 15.594 38.487 BO0407 Mn Manto 3 Boléo 0.7054 18.7580 15.5994 38.5157 CBY Cu in Gypsum member Boléo 0.7057 18.721 15.591 38.496

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Table 2.7, (Continued) Sample Description Locality 87Sr/86Sr 206Pb/204Pb 207Pb/204Pb 208Pb/204Pb RE9902 Mn Manto 4 Boléo 0.7068 18.731 15.592 38.478 RE10002 Mn Manto 4 Boléo - 18.833 15.595 38.499 BO1507 Mn Manto 4 Boléo 0.7067 18.726 15.589 38.471 BO2907 Mn Manto 4 Boléo 0.7064 18.742 15.596 38.518 RE9802 Mn Manto 4, Neptuno Neptuno 0.7067 18.741 15.593 38.493 BO2507 Mn in Gypsum member Boléo 0.7083 18.719 15.591 38.480 RE4502 Mn Manto 4 Lucifer 0.7052 18.743 15.597 38.512 RE4802 Mn Manto 4 Lucifer - 18.788 15.606 38.590 RE5402 Mn Manto 4 Lucifer 0.7052 18.759 15.601 38.546 RE5302 Mn Manto 4 Lucifer 0.7050 18.761 15.600 38.546 RE4402 Mn Manto 4 Lucifer 0.7050 18.788 15.604 38.588 RE8602 Mn oxide Gavilán 0.7052 18.612 15.579 38.421 BH-1 Mn oxide Boléo 0.7079 18.784 15.594 38.480

Note: (r) repeated sample

Table 2.8, Rare earth element average concentrations from hydrothermal and hydrogenous manganese oxide deposits including those from Baja California. Locality La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Σ REE Deposit type Ref. Indian Ocean 153.4 714.1 34.8 142.0 33.8 8.9 43.7 5.7 30.6 5.6 15.5 2.3 15.0 2.2 1207.6 Hydrogenous 1 Johnston Island 260.0 1201.0 44.8 192.7 35.4 10.1 48.5 6.8 43.7 9.9 28.9 4.3 27.9 3.6 1917.7 Hydrogenous 2 Pitcairn Island hotspot 276.8 741.7 50.8 214.4 42.7 10.2 49.9 7.2 44.7 9.2 26.4 3.9 26.0 4.0 1508.1 Hydrogenous 3 Pacific seamounts 202.2 1104.6 106.2 162.4 41.6 9.9 26.0 7.5 57.8 6.6 31.9 4.3 17.7 3.3 1782.0 Hydrogenous 4 Marginal seamounts 228.2 740.6 51.4 259.4 45.6 10.4 57.0 6.9 46.3 8.9 24.7 3.5 19.4 3.2 1505.4 Hydrogenous 4 Marginal abyssal plain 228.4 918.4 47.9 220.8 47.5 11.4 53.0 7.3 40.5 7.3 20.2 2.9 17.9 2.8 1626.2 Hydrogenous 4 Indian Ocean 183.2 163.8 - 164.4 35.3 9.1 37.2 - 34.7 - - - 17.6 2.6 647.9 Hydrothermal 5 Pitcairn Island hotspot 33.8 48.3 9.9 39.1 8.4 2.5 7.5 1.1 6.4 1.1 3.2 0.4 2.6 0.4 164.9 Hydrothermal 3 Modern hydrothermal 18.9 16.3 - 7.2 1.0 0.3 - 0.3 - - - - 0.8 0.1 44.8 Hydrothermal 4 Fossil hydrothermal 27.7 25.8 - 24.4 4.2 1.1 - 0.8 - - - - 2.8 0.3 87.1 Hydrothermal 4 Hokkaido Japan 30.5 140.3 - 43.3 12.6 2.4 - 2.2 - - - - 2.3 1.6 235.0 Hydrothermal 6 Lucifer deposit 6.2 9.4 1.0 4.4 1.2 0.2 0.8 0.1 0.7 0.1 0.4 0.1 0.4 0.1 25.2 Hydrothermal 7 Neptuno area 125.2 289.3 24.4 78.7 12.8 7.2 10.3 1.3 6.6 1.1 3.1 0.3 2.3 0.3 563.1 Hydrothermal 7 Gavilán deposit 11.7 11.0 1.8 8.4 2.1 0.4 2.7 0.5 2.9 0.6 2.0 0.3 1.7 0.3 46.3 Hydrothermal 7,8 Guadalupe 12.2 14.1 1.7 5.2 1.0 0.1 1.2 0.2 0.8 0.2 0.5 - 0.3 0.1 37.3 Hydrothermal 8 Santa Rosa 8.5 13.0 1.7 5.5 1.0 0.0 1.0 0.2 0.9 0.2 0.4 0.9 0.3 0.7 34.0 Hydrothermal 8,9 Boléo district - all mantos 59.1 97.5 5.9 27.9 3.8 5.8 3.9 0.4 3.9 0.7 2.9 0.3 2.9 0.3 215.2 Hydrothermal 10 Boléo district - manto 2 16.4 16.5 1.2 11.9 0.9 0.7 1.0 0.1 0.8 0.2 0.5 0.1 0.4 0.1 50.7 Hydrothermal 10 Boléo district - manto 3A 25.9 75.4 4.9 49.6 1.8 23.3 1.9 0.3 1.8 0.4 1.2 0.2 1.3 0.2 188.2 Hydrothermal 10 Boléo district - manto 3 61.8 109.0 7.0 30.4 4.0 6.7 4.3 0.3 4.1 0.6 2.9 0.2 3.0 0.2 234.6 Hydrothermal 10 Boléo district - manto 4 87.5 129.1 6.2 27.8 5.2 2.5 4.4 0.7 5.1 1.1 3.9 0.5 3.8 0.5 278.4 Hydrothermal 10 Boléo district - Gloria Fm 63.4 99.9 10.1 34.5 6.1 3.6 7.9 0.4 6.8 0.6 5.0 0.3 5.0 0.3 243.8 Hydrothermal 10

References (1) Nath et al. 1992; (2) Wiltshire et al. 1999; (3) Glasby et al. 1997; (4) Usui and Someya 1997; (5) Nath et al. 1997; (6) Miura and Hariya 1997; (7) Del Rio Salas et al. 2008; (8) Rodríguez Díaz 2009; (9) Rodríguez Díaz et al 2010; (10) Present study. 126

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CHAPTER 3: GEOLOGY, GEOCHEMISTRY, AND U-PB GEOCHRONOLOGY OF THE MARIQUITA PORPHYRY COPPER AND LUCY CU-MO DEPOSITS, CANANEA DISTRICT, MÉXICO

3.1, ABSTRACT

The Mariquita porphyry Cu and Lucy Cu-Mo deposits are located in the western

section of the Cananea porphyry copper district. Four hydrothermal stages are found in

Mariquita: Stage I is composed of quartz-pyrite-biotite-magnetite; Stage II corresponds to orthoclase-quartz; Stage III consists of unidirectional veinlets of quartz-pyrite- chalcopyrite-magnetite and minor molybdenite; and Stage IV alunite veinlets. Lucy deposit is characterized by quartz-molybdenite-chalcopyrite.

The temperatures and emplacement depths of the mineralization in Mariquita is 1 to 1.2 km and 430 to 380ºC, respectively. In contrast, the mineralization in Lucy deposit shows deeper emplacement depths (~3km) and higher mineralization temperatures (550-

500 ºC). Sulfur isotopes indicate that the source of sulfur for both deposits is clearly magmatic. The δ34S values (1 to 2.5‰) from alunites from stage IV in Mariquita are the

result of oxidation of previous hydrothermal stages. The isotope composition of the ore

fluids involved during the mineralization of hydrothermal stages I to III in Mariquita

determine a magmatic origin, whereas the stage IV consists of the mixing between fluids

of the magmatic and meteoric water components. The isotope data of the ore fluids from

Lucy deposit show a magmatic origin.

The new U-Pb zircon data for the hosting rock in Lucy deposit produces an age of

63.8 ± 1.1 Ma. The new U-Pb ages in zircons reported for two mineralizing quartz

feldspathic porphyries in Mariquita yield crystallization ages of 60.4 ± 1.1 Ma and 62.7 ±

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1.3 Ma. The magmatic-mineralizing pulses reported for the western portion of Cananea

district increase the potential for the presence of undiscovered mineralized bodies either

emplaced within the Cuitaca granodiorite (e.g. Lucy), or within the Laramide volcanic

rocks. The geology and geochronology reported here contribute with the understanding of

the western and northwestern section of the Cananea district.

3.2, INTRODUCTION

In the North American southwest region, the states of Arizona and New Mexico

are characterized by a rich endowment of porphyry copper deposits (PCDs). This region has been target of numerous geologic studies that collectively have enhanced the understanding of this metallogenic province (Titley 2001). The PCDs of Sonora essentially represent a southern extension of the porphyry copper province of southwest

North American (Titley 1982). Most of the magmatic-hydrothermal activity of the PCDs of this province occurred during the Laramide orogeny. Despite their economic importance, the PCDs of northwestern Mexico have received less attention, although some important pioneer works have contributed significantly (e.g. Valentine 1936;

Wantke 1925; Velasco1966; Echavarri 1971; Berchenbriter 1976; Bushnell 1980;

Meinert 1982).

The Cananea mining district is the largest copper producer in Mexico, and among the largest and most productive districts in the world. The district is a good example of a

cluster of PCDs, breccia pipe and skarn mineralization formed during a limited time

range, and has an estimated ~7,500 million metric tons of ore, with grades ranging from

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0.35 to 1.7 wt percent of Cu (Valencia-Moreno 2007). Despite the economic importance of the Cananea district, the geology and the mineralization features of the western and northwestern portions of the district have been barely documented in the literature.

Given the significant contributions of the Mariquita porphyry copper and Lucy

Cu-Mo deposits to the economic value of the whole district, there is a need for a better of

these deposits. The present study contributes with new geological, geochemical, and

geochronological data to constrain the genesis of the copper mineralization and provides

more information from the exploration perspective to the western sector of the Cananea

porphyry copper district.

3.3. PREVIOUS STUDIES OF THE CANANEA DISTRICT

The Cananea district is known by the world class Cananea PCD, although there is

the presence of other smaller PCD’s, in addition other ore deposits like breccia pipes,

skarns, and manto deposits (Fig. 3.1). Emmons (1910) and Valentine (1936) originally

established the basic geology of the district. Later, several authors complemented and

documented important geological issues (Mulchay and Velasco 1954; Velasco 1966;

Ochoa Landín and Echavarri 1978; Wodzicki 1995, Wodzicki 2001; Cox et al 2006). The

different aspects concerning the mineralization styles in the Cananea PCD have been

studied (Weed 1902; Austin 1903; Lee 1912; Virtue 1996). The breccia pipes, skarn, and

manto deposits have been studied in detail by Perry (1933), Perry (1961), Bushnell

(1980), Meinert (1982), and Bushnell (1988). Also, several studies have been performed

in Milpillas PCD (Carreón-Pallares 2002; de la Garza et al 2003; Valencia et al 2006;

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Noguez-Alcántara et al., 2007; Noguez-Alcántara 2008), in Mariquita PCD (Woodburne

2000; Del Rio Salas et al., 2006; Zúñiga Hernández 2006), Lucy (Del Rio Salas et al.,

2006), and El Alacrán PCD (Amaya-Martínez 1970; Dean 1975; Arellano 2004).

Various authors have documented geochronological data concerning the

mineralization and the magmatic activity along the Cananea district (Varela 1972;

Anderson and Silver 1977; Meinert 1982; Damon and Mauger 1966; Damon et al., 1983;

McCandless et al 1993; Wodzicki 2001; Barra et al., 2005; Cox et al., 2006; Del Rio

Salas et al., 2006; Valencia et al., 2006).

3.4, REGIONAL GEOLOGIC SETTING

3.4.1, Cananea district stratigraphy

The oldest unit exposed in the district is the 1,440 ± 15 Ma Cananea granite (Fig.

3.2, Anderson and Silver 1977), which includes the Precambrian basement in northeastern Sonora represented by the 1.7 Ga Pinal schist (Silver et al., 1977; Anderson and Silver 1979; Anderson and Schmidt 1983). Valentine (1936) described the Cananea granite in two facies: (1) a coarse granitoid to pegmatitic rock composed of orthoclase, oligoclase, quartz, and smaller amounts of hornblende, magnetite, and apatite; and (2), the most abundant type, a granophyric granitoid with phenocrysts of quartz and a microgranitoid matrix composed of orthoclase, microcline, quartz, and oligoclase.

The Cananea granite is unconformably overlain by a Paleozoic sedimentary sequence that includes the Bolsa (Cambrian), Abrigo (Cambrian), Martín (Devonian), and Escabrosa (Mississippian) Formations, and part of the Permian Naco Group

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(Mulchay and Velasco 1954; Velasco 1966; Meinert 1982). Notwithstanding the intense

faulting, metasomatism, and hydrothermal alteration in the Paleozoic sequence, Mulchay

and Velasco (1954) suggested a correlation between the Paleozoic sedimentary sequence

at Cananea and similar sedimentary rocks in southeast Arizona. The Paleozoic

sedimentary sequence in Cananea is economically important because it hosts the Zn-Pb-

Cu skarn mineralization described by Meinert (1982).

The Proterozoic and Paleozoic rocks are unconformably overlain by a pile of

Mesozoic to Early Tertiary volcanic rocks (Valentine 1936). The Mesozoic rocks include

the Triassic-Jurassic and the Laramide magmatic arcs. The oldest rocks in the volcanic pile are the volcanic rocks of the Elenita Formation, composed of rhyolitic to andesitic tuffs and lavas with interbedded sandstone and quartzite. The Elenita Formation outcrops in the west and the southwest portions of the Cananea district (Fig. 3.2), and a thickness of 1,800 m is estimated (Valentine 1936). This formation is similar to the Late Triassic-

Mid Jurassic Wrightson Formation in southern Arizona described by Drewes (1971) and

Riggs and Blakey (1993). The Henrietta Formation is overlies the Elenita Formation (Fig

3.3; Valentine 1936), and is composed by medium to high-K, calc-alkaline, dacitic to rhyolitic tuffs and flows (Wodzicki 1995). The Henrietta Formation occurs in a northwest trending belt across the center of the Cananea district (Fig 3.2), and generally dip E-NE except in the western part, where dips are W-NW (Ochoa Landín et al., 2007), and a thickness of 1,700 m is estimated (Valentine 1936). An Ar-Ar age in hornblende from a volcanic flow of the Henrietta Formation produced a minimum age of 94 Ma (Wodzicki

1995). This formation is important in the district because hosts part of the copper

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mineralization of the Cananea ore body (Velasco 1966). The intrusive counterpart of the

Jurassic rocks within the Cananea district is the 175 Ma Torre syenite, which intrudes

both the Elenita and Henrietta Formations (Wodzicki 2001; Noguez-Alcántara 2008).

The oldest Laramide rocks correspond to the Mariquita diabase, which consists of a high-K basaltic-andesite flows and intrusive bodies, and is characterized by a porphyritic “turkey tracks texture” (Wodzicki 2001). The Mariquita diabase occurs as volcanic flows shallowly dipping to the East, and make up the upper 400 m of the Sierra

Mariquita located East and North of the Mariquita and Maria deposits respectively (Fig.

3.2). Between Sierra Mariquita and Cananea the Mariquita diabase occurs as dikes and

stocks intruding the dacitic tuffs of the Henrietta Formation, and also as a thick flow that overlies the Henrietta Formation and grades upward into the overlying Mesa Formation

(Wodzicki 2001).

The Laramide Mesa Formation represents most of the Cretaceous volcanic

activity in the district (Valentine 1936). Compositions vary from bottom to top, from

trachy-basaltic, basaltic-andesite, andesitic, dacitic, to trachy-andesitic. Tuffs,

agglomerates, lahars, and flows of andesitic composition are present (Valentine 1936;

Wodzicki 2001). The Mesa Formation crops out in the eastern portion of the district (Fig.

3.2) and a thickness of 1,500 m is estimated (Valentine 1936). These rocks are important

within the district because they host the disseminated copper mineralization. A flow

within this formation has been dated 69 ± 0.2 Ma using the 40Ar/39Ar method in biotite

(Wodzicki 1995), although a span of 72 to 68 Ma has been documented with the same

dating method around the Cananea district (Cox et al., 2006; Noguez-Alcántara 2008).

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The Mesa formation overlies the Elenita and Henrietta Formations, and is intruded by the

Tinaja-Cuitaca granodiorite, and the Mariquita Formation of Laramide age, and by

younger intrusive bodies.

The earliest Laramide intrusive unit is the Tinaja-Cuitaca batholith (Fig. 3.2), which occurs as two spatially distinct, composite equigranular intrusive bodies named the

Tinaja diorite and the Cuitaca granodiorite (Valentine 1936). The Tinaja diorite intrudes the Henrietta and Elenita Formations in the western portion of the Cananea mine. The composition varies from gabbro to monzonite to quartz monzonite, but the predominant compositional phase is the monzodioritic (Wodzicki 1995). Previous studies in the district support the idea that the Tinaja and Cuitaca intrusions belong to the same batholith (Valentine 1936; Meinert 1982, Bushnell 1988). Isotopic data support the idea of a genetically related polyphase batholithic body (Wodzicki 1995). The Cuitaca granodiorite is a large batholithic body with a northwest-southeast major axis and is many kilometers in length (Valentine 1936). It is 64 ± 3 Ma (Anderson and Silver 1977), and intrudes the Elenita, Henrietta, and Mariquita Formations. The composition ranges from monzonitic to granodioritic to granitic, but the main compositional phase is granodioritic (Wodzicki 1995).

The Tinaja-Cuitaca batholith is intruded by numerous near-vertical mafic dikes oriented NW 60-80 and NE 40 (Valentine, 1936). These intrusions are dominated by the

Campana dikes and are dated at 58.4 ± 0.6 Ma (Carreón-Pallares 2002). The Henrietta and Mesa Formations are locally cross-cut by similar dikes. These mafic dikes are not

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cross-cut by younger quartz feldspar porphyries, and apparently were emplaced close to the time of solidification of the Cuitaca intrusive body (Wodzicki 1995).

Several monzonitic and quartz monzonitic mineralized porphyry plugs are present along the Cananea district. The oldest mineralizing porphyry documented within the district is located in Milpillas PCD, which yielded a U-Pb age in zircons of 63.9 ± 1.3 Ma

(Valencia et al., 2006). Younger mineralizing quartz-monzonitic and granodioritic porphyries are present along the Cananea mine, Maria, La Colorada, and Alacrán deposits, whose mineralization events yield Re-Os ages from 59 to 60 Ma (Barra et al.,

2005).

3.4.2, Structural geology

The structural framework within the Cananea district is complex, and is characterized by important pre- and post-mineralization faulting episodes. Two major pre-mineralization structural events are observed in the district: early set of northwest and northeast trending faults, which are locally followed by the intrusion of quartz-feldspar porphyries and basaltic dikes (Fig. 3.2), and a later set of NNE trending normal faults that tilt the entire section (Wodzicki 1995).

The pre-mineralization faulting consist of a northwest-southeast trend that can be separated into NW-SE 60-80º and NW-SE 40-50º trends (Valentine 1936), and of a NE-

SW 40º trend (Meinert 1982). These structures cross cut the entire pre-Laramide section, and control the emplacement of the mineralizing quartz-feldspathic porphyries (Valentine

1936; Wodzicki 2001).

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The post-mineralization faulting at the eastern portion of the district consists of

steeply NE-SW 5-10º normal faults, and tilt the entire rock section around 15º to the east

(Wodzicki 1995). The western section of the Cananea district is severely affected by the post-mineralizing tectonics, and is responsible for the development of the Cuitaca half- graben, the sedimentary cover, and erosion of older rock units. This post-mineralizing faulting is most likely the direct result of the Basin and Range tectonic event (Fig. 3.2).

3.5, GEOLOGY OF MARIQUITA DEPOSIT

3.5.1, Geology

The Mariquita PCD is located 13 Km northwest of the town of Cananea and is

one of the porphyry copper deposits forming the northwestern trending of the Cananea

district (Fig. 3.2). The deposit contains 100 million metric tons at 0.3 percent copper

(Consejo de Recursos Minerales, 1994). The oldest rocks in the area are the volcanic lava

flows of the Henrietta Formation. These rocks outcrop in the southern and eastern

portions of the Mariquita area and the thickness of this volcanic sequence can exceed 200

m (Fig. 3.3b and 3.4). The Henrietta Formation exhibits both structural and intrusive

contact with the Cuitaca granodiorite, and structural contact with the Mesa Formation

(Fig. 3.4). Moderate to intense quartz-sericite alteration characterizes these rocks, the

texture and composition is locally completely obscured by the intense pervasive

alteration. Nevertheless, the stratigraphic position can be constrained with certainty by

the presence of the Mariquita formation, which directly overlies the Henrietta Formation.

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The Mariquita Formation crops out in the central and northern parts of the

Mariquita area (Fig. 3.4). The maximum thickness in this area is 40 m. The Mariquita

Formation is easily distinguished in the field by the plagioclase phenocrysts immersed in an aphanitic matrix, and particularly by the “turkey track” texture. This unit serves as a stratigraphic index because is located between the Henrietta and Mesa Formations, as previously pointed out by Valentine (1936) for the eastern and southeastern sections of

the Cananea district. The Mariquita diabase is covered by gravels along the western side

of the area, and it is intruded by the Cuitaca granodiorite on the eastern side (Fig. 3.4).

The Mesa Formation is overlies the Mariquita Formation, and crops out in the

north-central portion of Mariquita area (Fig. 3.4). In the Mariquita area, the Mesa

Formation is composed mostly of pyroclastic flows and minor lava flows, both ranging

from andesitic to dacitic composition. The thickness of these rocks can reach 200 m.

These rocks show a moderate to intensive quartz-sericite alteration, locally obscuring the

original texture and composition.

The Cuitaca granodiorite crops out at the eastern side of the Mariquita area, and is

in structural contact with the Mesa Formation, along a north-south fault locally referred

as the eastern fault (Fig. 3.4). The Cuitaca granodiorite also intrudes the Mariquita and

Mesa Formations at the northeast, central, and south-central portions of the Mariquita

area (Fig. 3.4). The granodiorite is coarse-grained, and is mostly fresh, although it shows

local moderate to intense quartz-sericitic alteration which masks the original texture. The

importance of this intrusive body within the district is that it hosts Cu-Mo breccia or

pegmatitic bodies such as the Maria, La Colorada, and Lucy deposits (Fig. 3.2).

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All the volcanic rocks mentioned above are intruded by quartz-feldspathic porphyry stocks with a NW-SE trend (Fig. 3.4). They are composed of quartz and feldspar phenocrysts, the latter mostly altered to sericite and minor quartz, embedded in a fine matrix of quartz-sericite or quartz>>sericite. These porphyry stocks are similar to those present in the Cananea district, and are responsible for the mineralization in the

Mariquita PCD.

A mafic dike similar to those reported by Valentine (1936) crops out at the northern section of the Mariquita PCD. The thickness of the mafic dike ranges from 1 to

3 m, and it occurs mostly as a steeply dipping tabular body in a striking NE-SW 35-40º.

The Mesa Formation and apparently the quartz-feldspathic porphyry are cross-cut by this mafic dike. The relationship with the rhyolite unit is not clear. This mafic dike exhibits moderate propylitic alteration.

A rhyolite porphyry crops out locally in the northern area of the Mariquita PCD

(Fig. 3.4). This intrusive unit is spatially related to the NE-SW fault system, which apparently controlled the intrusion. Similar rhyolite porphyry plugs crop out further to the northwest, between the Mariquita PCD and the Lucy deposit (Fig. 3.2).

Tertiary gravel and Quaternary alluvium deposits are the youngest units in the area, and cover the western portion of the Cananea district (Fig. 3.2). The clastic material fills the Cuitaca half-graben, a basin bounded by north-south structures. The gravel deposits, known as the Sonora Group (Grijalva-Noriega and Roldán-Quintana 1998), are composed mostly of moderately consolidated to semi-consolidated clasts derived from the volcanic formations described above, and the thickness of the clastic fill can exceed

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the 600 m (Woodburne 2000). The gravel deposits overlie the Mesozoic rocks unconformably, and they have been subdivided into lower and upper basin fill based on a gradational depositional contact between the two units (Woodburne 2000).

3.5.2, Structure

Figure 3.4b shows the spatial and chronological relationship between the fault systems recorded in the Mariquita area. The Mariquita area is limited by a north-south fault system locally known as the western and eastern boundary faults (Fig. 3.4). Three dominant fault systems are recorded in the Mariquita area; the oldest structures are steeply dipping NE-SW 40-60º faults that are cross cut by NE-SW 20º faults, dipping 25-

60º SE (Fig. 3.4b). The youngest structures recorded in the area intersect the previous two fault systems, and are near N-S trending faults. This fault system defines the western scarps of the Elenita and Mariquita Mountains, as wells as the limits of the Cuitaca half- graben (Fig. 3.2).

3.5.3, Alteration and mineralization

Several aspects of the hydrothermal alteration in the Cananea district has been documented decades ago (Valentine 1936; Perry 1961; Velasco 1966; Meinert 1982;

Bushnell 1988; Wodzicki 1995; Virtue 1996; Noguez-Alcántara 2008). As most of the porphyry copper deposits within the North American southwest and in the Cananea and the Nacozari districts, the most abundant hydrothermal alteration consists of phyllic alteration superimposed on a previous potassic and propylitic alteration.

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Potassic alteration

The first phase of hypogene alteration in the Mariquita area is the potassic

alteration, which is present mostly in the northern section (Fig. 3.5). The mineral

assemblages are quartz-orthoclase-biotite and biotite-magnetite. The Mariquita diabase is

characterized by a selectively pervasive biotitic alteration, along with magnetite. In thin

section, fine-grained secondary biotite is distributed within the matrix and replaces the

original hornblende micro-phenocrysts. The mineral assemblage in the veinlets that cross cut the Mariquita diabase consists of biotite-orthoclase-magnetite, an association that shows evidence of a superimposed quartz-sericite alteration event. In addition, the volcanic rocks from the Mesa Formation show evidence of early potassic alteration represented by the mineral assemblage quartz-biotite-pyrite, along with an intense silicification (Fig. 3.5). Samples retrieved from core-drill in the Mariquita area show the mineral association composed by quartz-pyrite-chalcopyrite-enargite-anhydrite, and quartz-pyrite-chalcopyrite-molybdenite. In general, evidence of potassic alteration is often superimposed by phyllic alteration in the Mariquita area.

Propylitic alteration

The distal zones of the Mariquita area are characterized by propylitic alteration represented by the mineral association of chlorite-epidote, most commonly present in the

Cuitaca granodiorite and locally intensely developed in the Mesa and Mariquita

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Formations in the north and northeastern portions of the deposit (Woodburne 2000;

Aponte-Barrera 2006).

Phyllic alteration

The dominant hydrothermal event is evidenced by widely distributed phyllic

alteration represented by the mineral assemblage quartz-sericite (Fig. 3.5). In the northern and central portions of the Mariquita area, intense to moderate quartz-sericite alteration is

characteristic near and in the quartz-feldspathic porphyry stocks. In thin section, the

quartz-feldspathic porphyry stocks show strong quartz-sericite alteration that obscures the

original texture, except for quartz “eye” crystals. In addition, primary textures in the

rocks of the Mesa Formation are locally completely obscured by pervasive quartz-sericite alteration, and is cross-cut by veinlets of quartz-sericite-pyrite.

In the central portion of the Mariquita area, the rocks from the Henrietta,

Mariquita, and Mesa Formations are mostly dominated by moderate phyllic alteration, along with disseminated sulfides, also cross-cut by quartz-sericite-pyrite veinlets.

Locally, there is a strong phyllic alteration along the structures and near the quartz- feldspathic porphyry stocks (Fig. 3.5). The open spaces of the breccias associated to the

NE-SW faults are filled by quartz-pyrite-chalcopyrite, commonly rimmed or replaced by

chalcocite. In thin section, the intensity and the types of alteration within these structures

can vary (quartz>sericite, sericite, >>kaolinite;

quartz<>kaolinite).

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Argillic alteration

A later strong argillic alteration represented mostly by kaolinite is associated with

NE-SW faults that brecciate the rocks of the Mariquita and Mesa Formations. In the NE-

SW faults-related breccias, there is a clear argillic alteration imposed over the previous strong phyllic alteration (quartz-sericite-sulfides assemblage described above).

The southern portion of the Mariquita area is mostly characterized by a moderate phyllic alteration in the Henrietta Formation rocks, along with some argillic and silicification zones in both the Henrietta Formation and the Cuitaca granodiorite, spatially associated to NE-SW structures (Fig. 3.5). In addition, along the intrusive contact between the Henrietta Formation and the Cuitaca granodiorite, there is a moderate to strong silicification with disseminated pyrite (1-3%).

Sulfide mineralization

The economic mineralization consists of an enriched chalcocite blanket spatially oriented WNW-ESE and slightly tilted to the southwest, with copper grades ranging between 0.4 and 0.6%. The average thickness of the chalcocite manto is about 100 m and a similar thickness is estimated for the upper oxide zone of hematite-limonite. In the northwestern portion of the deposit the chalcocite zone thins to less than 30 m, and the oxide zone, jarosite-limonite in this area exceeds 200 m thickness, and contains tenorite, neotocite and malachite (Aponte-Barrera 2006). The enriched zone is displaced by the last faulting stage.

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At least four hydrothermal events were identified in the Mariquita Pit, where

temporal and spatial relationships can be constrained. These are in order from earliest to

latest: (Stage I) thin veinlets composed of quartz-pyrite-biotite-magnetite, with thickness

ranging from 1 to 2 mm; (Stage II) irregular veinlets of orthoclase-quartz, with an

average thickness of 1 cm, clearly cross cutting the event 1; (Stage III) unidirectional

veinlets of quartz-pyrite-chalcopyrite-magnetite and minor molybdenite, with an average

thickness of 3 cm; and (Stage IV) alunite veinlets, with thickness varying from 1 to 8

mm, which clearly cross-cut all the previous vein types. The first two stages correlate to

the potassic alteration. The third stage corresponds to the phyllic alteration. Stage IV veins correspond to alunite veinlets that cross cut the earlier stages. A final event, characterized as supergene in nature, consists in the development of argillic alteration, which is clearly restricted to faults and breccias, and is always superimposed on all the previous types of alteration describe above, and also post-dates the enrichment of the chalcocite blanket.

Moreover, petrographic and mineragraphic studies in La Verde located in the northern section of Mariquita area (Fig. 3.4), indicate that the hypogene mineralization consists of disseminated pyrite-chalcopyrite and veinlets composed of quartz-pyrite.

Weak supergene enrichment is evidenced by the replacement of chalcopyrite by chalcocite in the disseminated mineralization and by chalcocite coating the pyrite in the quartz-pyrite veins. Copper oxide mineralization is more abundant in this area, and consists mostly of chrysocolla and neotocite.

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3.6, GEOLOGY OF LUCY DEPOSIT

The Lucy deposit is located at the northwestern end of the linear trend of the porphyry copper deposits in the Cananea district, 5 Km northwest of the Mariquita PCD

(Fig. 3.2). The geology in the Lucy area is simple: the Henrietta and Mesa Formations are

intruded by the Cuitaca granodiorite. The next geologic episode consists in the intrusion

of rhyolitic dikes within the Henrietta formation. Finally, gravel deposits of the Sonora

Group were deposited in the Cuitaca basin. The Lucy Cu-Mo deposit is hosted within the

Cuitaca granodiorite, and originally was manifested on the surface as irregular veinlets

composed of quartz-sericite-sulfides, along with the notable presence of tourmaline and

Fe-oxides in the shallower zones.

Figure 3.6 is a schematic cross-section of the Lucy deposit and shows the

different hydrothermal alteration zones. The first alteration event is evidenced by patches

of orthoclase, over which is imprinted an intense phyllic alteration represented by the

assemblage of quartz-sericite-pyrite. Propylitic alteration is present at the distal zones of

the deposit, where chlorite replaces the hornblende and biotite of the Cuitaca

granodiorite.

Most of the mineralization in the Lucy Cu-Mo deposit occurs as ellipsoidal bodies

of breccia. The breccia is located along north-south structures, with open-spaces filled by

quartz-molybdenite-chalcopyrite-tourmaline. Drilling campaigns in this deposit indicate

no continuation of deeper mineralized zones. Near horizontal fractures filled by quartz-

orthoclase-chalcopyrite-molybdenite are common in the shallower zones of the Lucy

deposit. Secondary copper minerals (malachite) coat disseminated pyrite crystals in the

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Cuitaca granodiorite. Petrographic studies show that secondary copper enrichment occurs as chalcocite coating pyrite. Although the supergene enrichment is weak, the primary grades make this deposit economic with 0.8% Cu and 0.1% Mo.

3.7, ANALYTICAL PROCEDURES

3.7.1, Stable isotopes

The oxygen isotopes in the silicates and magnetite samples were analyzed in the

Stable Isotope Laboratory at The Southern Methodist University, and the sulfur, oxygen, and hydrogen isotopes in the sulfides, sulfates, and mica samples were measured at the

Environmental Isotope Laboratory at The University of Arizona.

The δ18O values from the silicates and magnetite samples were analyzed

following the methods of Clayton and Mayeda (1963) and Borthwick and Harmon

(1982). Approximately 10 mg of sample was placed in a nickel reaction vessel with

approximately 170 torr of BrF5 and reacted at 400°C for at least 14 hours. The liberated

O2 was then converted to CO2 by heating it with a graphite rod. The CO2 was collected

and analyzed on a Finnigan MAT 251 mass spectrometer. The data are reported with

respect to Vienna Standard Mean Ocean Water. Replicate analyses of the standard NBS-

28 give an average value of 9.64‰ with a standard deviation of 0.12‰. The data are

reported in Table 3.1.

For the oxygen isotopes in the sulfate samples, approximately 0.3 mg of powder

sulfate was placed in silver capsules. The δ18O values in the sulfates samples were

measured on a continuous-flow gas-ratio mass spectrometer (Finnigan Delta PlusXL).

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The samples were combusted at 1350°C using a ThermoQuest thermal combustion

elemental analyzer (TCEA) coupled to the mass spectrometer. The isotope data for the

sulfates is shown in Table 3.1. Analyses are reported with analytical precision of ± 0.3‰

or better (1σ).

The δD values in the mica and alunite samples were measured also in a

continuous-flow gas-ratio mass spectrometer (Finnigan Delta PlusXL) H2 gas was

generated at 1400°C using the Thermal Combustion Elemental Analyzer (TCEA) coupled to the mass spectrometer (Table 3.1). Standardization is based on NBS-30 and

IAEA-CH-7. Precision (1s) is better than ± 2.5‰ based on repeated internal standards.

Isotopic data are reported in standard δD notation relative to Vienna SMOW.

The sulfur isotopes in the sulfide and sulfate samples were measured on a continuous-flow gas-ratio mass spectrometer (Finnigan Delta PlusXL). Depending on the mineral species, 0.3 to 1.0 mg of powdered sample was, loaded in a tin capsule along with V2O5 as an oxygen buffer. The samples were combusted at 1030ºC, using an

elemental analyzer (Costech) coupled to the mass spectrometer. Standardization is based

on international standards NBS-123 and OGS-1 for the sulfides and sulfates respectively,

and several other sulfide and sulfate in-house standards that have been compared between

laboratories. Calibration is linear in the range -10 to +30‰. Analytical precision is

±0.15‰ or better (1σ). The sulfur isotope data are reported in Table 3.2.

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3.7.2, U-Pb method

For the U-Pb dating method in magmatic zircons, around 1 kg of the intrusive

rocks were crushed and milled. Heavy mineral concentrates smaller than 350 µm were

separated using the Wilfley Table. The zircons were concentrated using di-iodomethane

heavy liquid and magnetic techniques. Later the zircons were handpicked under a binocular microscope, and were mounted in an epoxy resin and polished. Around 30 zircons from each sample were analyzed by laser ablation multicollector inductively

coupled plasma mass spectrometry (LA-MC-ICPMS) at the Arizona LaserChron Center.

The analyses involve ablation of zircon with a New Wave/Lambda Physik DUV193

Excimer laser (operating at a wavelength of 193 nm) using a spot diameter of 35 microns.

The ablated material is carried with helium gas into the plasma source of a GV

Instruments Isoprobe, which is equipped with a flight tube of sufficient width for simultaneous measurements of U, Th, and Pb isotopes. All measurements are made in static mode, using Faraday detectors for 238U and 232Th, an ion-counting channel for

204Pb, and either faraday collectors or ion counting channels for 208-206Pb. Ion yields

are ~1 mv per ppm. Each analysis consists of one 20-second integration on peaks with the

laser off (for backgrounds), 20 one-second integrations with the laser firing, and a 30-

second delay to purge the previous sample and prepare for the next analysis. The ablation

pit is ~15 microns in depth.

For each analysis, the errors in determining 206Pb/238U and 206Pb/204Pb result in a

measurement error of ~1% (at 2-sigma level) in the 206Pb/238U age. The errors in

measurement of 206Pb/207Pb and 206Pb/204Pb also result in ~1% (2-sigma) uncertainty in

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age for grains that are >1.0 Ga, but are substantially larger for younger grains because of

the low intensity of the 207Pb signal. For most analyses, the cross-over in precision of

206Pb/238U and 206Pb/207Pb ages occurs at ~1.0 Ga.

Common Pb correction is accomplished by using the measured 204Pb and assuming an initial Pb composition from Stacey and Kramers (1975) (with uncertainties of 1.0 for 206Pb/204Pb and 0.3 for 207Pb/204Pb). Our measurement of 204Pb is unaffected by

the presence of 204Hg because backgrounds are measured on peaks (thereby subtracting any background 204Hg and 204Pb), and because very little Hg is present in the argon gas.

Inter-element fractionation of Pb/U is generally ~20%, whereas fractionation of

Pb isotopes is generally <2%. In-run analysis of fragments of a large zircon crystal

(generally every fifth measurement) with known age of 564 ± 4 Ma (2-sigma error) is

used to correct for this fractionation. The uncertainty resulting from the calibration

correction is generally ~1% (2-sigma) for both 206Pb/207Pb and 206Pb/238U ages.

The analytical data are reported in Tables 3.3-3.5. Uncertainties shown in these tables are at the 1-sigma level, and include only measurement errors.

The reported ages are determined from the weighted mean (Ludwig 2003) of the

206Pb/238U or 206Pb/207Pb ages of the concordant and overlapping analyses. Analyses that

are statistically excluded from the main cluster are shown in blue on these figures. Two uncertainties are reported on these plots. The smaller uncertainty (labeled “mean”) is based on the scatter and precision of the set of 206Pb/238U or 206Pb/207Pb ages, weighted according to their measurement errors (shown at 1-sigma). The larger uncertainty

(labeled “age”), which is the reported uncertainty of the age, is determined as the

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quadratic sum of the weighted mean error plus the total systematic error for the set of

analyses. The systematic error, which includes contributions from the standard

calibration, age of the calibration standard, composition of common Pb, and U decay

constants, is generally ~1-2% (2-sigma).

3.8, RESULTS

3.8.1, Oxygen and hydrogen isotopes

The oxygen and hydrogen isotope data from the phyllosilicate (biotite and

sericite) and alunite samples are included in Table 3.1. The δ18O values in the biotite

from stage I range from +4.1 to +4.7‰, whereas the δD values range from −74 to −82‰.

The δ18O values of the coexisting water were calculated using the fractionation equation

for the hydroxyl-bearing silicates (Zheng 1993) and generate a range from +6.7 to +7‰.

The δD values of coexisting water were calculated using the hydrogen fractionation

factor from Suzuoki and Epstein (1976), and have a range between −29 and −37‰. The isotope data from the sericites of stage III range from +4.1 to +10.7‰ for the δ18O

values, and from −32 to −77‰ for the δD values. Coexisting water ranges from +4 to

+9.9‰ in δ18O, and 0 to −36‰ in δD (Table 3.1).

The δ18O and δD values of alunites from stage IV range between −59 to −84‰ and between +2.9 to +8.6‰ for the δ18O and δD respectively. The δ18O and δD values of

the water coexisting with alunite were made using the fractionation factors of Stoffregen

et al. (1994), and the δ18O values range between −3.2 to 2.5‰, and the δD values range between −56 to −87‰.

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The δ18O values in anhydrite range from +12.2 to +13.7‰, and the calculated

δ18O of the fluids coexisting with anhydrite from stage III range between +7.5 to +9.0‰

(Chiba et al., 1981). The δ18O values in the quartz samples range from +9.7 to +10.1‰.

The δ18O values of the fluids coexisting with the precipitated quartz range between +4.1

to +6.4‰ using the Quartz-H2O fractionation factor (Clayton et al., 1972).

The δ18O and δD values of the sericite from Lucy are +10.9 and −54‰

respectively. The δ18O and δD values of the fluids involved in the precipitation of that sericite are +11.9 and −42‰ respectively (Suzuoki and Epstein 1976; Zheng 1993). The only δ18O value from quartz in Lucy is +11.8‰, and the calculated isotope composition

for the fluids involved during the precipitation is +10.4‰ (Clayton et al., 1972).

3.8.2, Sulfur isotopes

The sulfur isotope data from the sulfides and sulfates samples from Mariquita en

Lucy deposits are shown in Table 3.2. The sulfides analyzed consist of pyrite,

chalcopyrite, molybdenite, tennantite, and bornite. The δ34S values in the sulfides and

sulfate samples range from −4.6 to 3.8‰ and 1.0 to 24.8‰ respectively. Figure 3.7 shows the frequency diagrams for the sulfides from the Mariquita and Lucy deposits.

3.8.3, U-Pb ages

The U-Pb zircon ages for Mariquita and Lucy deposits are shown in Tables 3.3-

3.5 and Figures 3.8-3.9. The mineralizing porphyry 604 from Mariquita PCD produces a

206Pb/238U weighted average age of 60.4 ± 1.1 Ma (n = 21, MSWD = 1.3), with one

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inherited zircon of Middle Jurassic age (~170 Ma). The mineralizing porphyry 104 from

Mariquita yields a 206Pb/238U weighted average age of 62.7 ± 1.3 Ma (n = 30, MSWD =

1.4), also with one inherited zircon of Middle Jurassic age (~180 Ma).

The U-Pb zircon age for the granodiorite that is hosting rock in Lucy deposit

produces a 206Pb/238U weighted average age of 63.8 ± 1.1 Ma (n = 21, MSWD = 3.2).

3.9, DISCUSSION

The geology of the Mariquita PCD is similar to the geology of the Cananea district, although some notable variations are present. The most striking differences are the absence of the 1,440 ± 15 Ma Cananea granite and the Paleozoic sedimentary sequence, and the presence of the Tertiary gravel deposits of the Sonora Group. Another remarkable feature is the 35 to 45º west-northwest tilting of the lithological units as oppose of the shallower east-southeast tilting in the Cananea mine (Fig. 3.2).

The Tertiary extensional tectonics is responsible for the development of the nearly north-south trending of the Cuitaca half-graben at the western section of the Cananea district (Fig. 3.2). This half-graben suffered various subsidence stages as response of the extensional tectonic regime during Tertiary times. The first subsidence episode is evidenced by the lowermost, well-consolidated, and west-tilted conglomerate unit of the

Baucarit Formation. The continuous subsidence is evidenced by the younger overlain and non-consolidated gravels belonging to the Sonora Group (Grijalva-Noriega and Roldán-

Quintana 1998), which show slightly unconformities, slightly tilting, and a series of faults

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cross-cutting both, the lower and upper sedimentary fill of the basin, according to the proposed subdivision by Woodburne (2001).

In Mariquita area, the volcanic rocks of the Mesa Formation whose thickness range from 150 to 250 m and is intruded by the Cuitaca granodiorite, both units covered by the sediment fill in the Cuitaca half-graben. Further north, between the Mariquita and

Milpillas PCDs, and also northern of the Milpillas PCD, there is a notable increase in the thickness of the sediment fill as well as the volcanic pile, which proof major subsidence rates and faulting at the northern portion of the Cananea district.

The north-south post-mineralization faulting at the eastern section of the Cananea range exhibit less displacement and slight inclination to the east, although the faulting apparently controlled the oxidation and supergene enrichment in the Cananea deposit.

The uplifting in the Cananea district was definitely very important for the development of the 500 m chalcocite blanket seen in the Cananea Mine (Wodzicki 2001). In contrast, the deposits located at the western portion of the district (i.e. Mariquita, Milpillas, Lucy, El

Toro, etc) were severely affected by the Tertiary extensional tectonics. In the Mariquita area, the fault systems that cross cut the deposit are both syn- and post-mineralization

(Fig. 3.4), although major displacements are recorded by the post-mineralization Tertiary faulting (Woodburne 2000). The extensional tectonics reflected by the formation of the

Cuitaca half-graben in Mariquita and adjacent areas, most likely produced erosion, dismemberment, and displacement of the mineralized bodies, which consequently were positioned at deeper levels, or partially covered by the gravel deposits, or eroded away by fluvial processes. On the other hand, the Tertiary tectonics un-roofed and uplifted deeper

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deposits (i.e. Lucy, La Milpa, El Toro), which are located on the western horst of the

Cuitaca half-graben (Fig. 3.4). The Lucy deposit does not have a leach capping or enrichment blanket, possibly because of the deeper level of exposure of the mineralized

Cuitaca granodiorite at Lucy, or on the other hand, by intensive erosion, if in fact an enriched capping was originally developed.

3.9.1, Alteration and mineralization

In Mariquita, the high-grade mineralization is commonly found along the contact between the mineralizing porphyritic stocks and the host rocks from the Henrietta,

Mariquita, and Mesa Formations, and along the brecciated zones and the fault systems

(Fig. 3.4). The most significant difference between the Cananea PCD and the Mariquita porphyry copper deposit is the absence of the 500 m chalcocite blanket. In Mariquita the chalcocite blanket is near horizontal distributed with dimensions 1500 × 170 m with an average thickness of 60 m (Aponte-Barrera 2009), and is covered by the gravel deposits from the Sonora Group (Fig. 3.10). Preliminary themobarometric data reveal that the emplacement depths of the Mariquita mineralization range from 1 to 1.2 km. The systematics of the hypogene mineralization and alteration in Mariquita deposit is similar to that observed in Milpillas, Cananea Mine, Alacrán deposits (Virtue 1996; Arellano

2004; Noguez Alcantara 2008).

In contrast, the Cu-Mo Lucy deposit does not have an enrichment blanket, and from the geological and structural framework exposed above, along with the thermobarometric data from fluid inclusions studies, Lucy represents a deeper

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mineralized body formed at ~3 km (Ochoa Landín et al., 2007; González-Partida et al.,

2009). Considering these facts, a combination of two possible options arises if an enriched capping was developed at any point: 1) intense erosion, and 2) segmented by the intense Tertiary tectonics characteristic of the western portion of Cananea district.

3.9.2, Supergene events

After the Laramide faulting and uplift of the Cananea district, an important erosion episode started after 54 Ma, which removed the overlain Late Cretaceous and

Early Tertiary rocks, allowing the unroofing of the Cananea PCD (Virtue 1996). The erosion episode started to be disrupted at 35 Ma by the deposition of the Sierra Madre

Occidental ignimbrites flows, which are dated at ~ 25 Ma in northern Milpillas area, which suggests a possible supergene enrichment prior 25 Ma in the district (Noguez-

Alcántara 2008).

Subsequent to the cessation of the volcanism at 25 Ma, is precisely the time when the extensional Tertiary tectonics in Sonora started, therefore the unroofing of the mineralized rocks continued, and along the climate change to arid conditions in the region, allowed the proper conditions for the development of the supergene enrichment in the district. A supergene episode in the district is evidenced in the Milpillas area, and consists of a horizon composed by red hematite along with jarosite, kaolinite, and alunite.

K-Ar geochronologic data in the supergene alunite produce an age interval between 19 to

17 Ma (Noguez-Alcántara 2008).

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Recent Ar-Ar geochronology of supergene alunite from Mariquita produces an

age ~9 Ma (Perez-Segura personal communication), which suggests so far the youngest

supergene process in the district. The available geochronological data for the Cananea

district suggests that the supergenic processes probably started since Eocene time and

continued episodically prior to the early opening of the Gulf of California (~7 Ma).

3.9.3, Sulfur isotopes

The sulfides from Mariquita mostly exhibit a narrow range of δ34S between +0.3

to +3.8‰. The uniform δ34S values suggest a homogeneous hydrothermal system, and

most likely, the sulfur isotope data indicate sulfur related to magmatic sources (Fig. 3.7).

Only one pyrite sample exhibits a lower δ34S value of −3.5‰. The scarce δ34S data for

sulfides from Lucy deposit show two groups; the δ34S values of the molybdenite samples range from +1.1 to +3.3‰ and agree perfectly with magmatic sources. The second group consists of pyrite samples whose δ34S values range −4 to −4.6‰.

The Δ34S value of the anhydrite-sulfide pairs from the early mineralization stages

is about 20‰, which is consistent with isotopic equilibrium at temperatures around 380ºC

(Ohmoto and Rye 1979), indicating no external incorporation of sulfate into the system,

and essentially the sulfate and sulfide minerals are precipitating in equilibrium within a

magmatic sulfur source for the earlier hypogene mineralization stage. This contrasts with

the Δ34S anhydrite-sulfide pairs for La Caridad in Nacozari, where isotope disequilibrium is documented for the earlier mineralization stages (Valencia et al 2008).

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The δ34S values of sulfates of the later hydrothermal activity (stage IV) are lower

(1 to 2.5‰) and contrast with the values of the earlier stages (~24‰). The isotope compositions of the later sulfates (alunites) suggest mixing with external meteoric water, which promotes oxidation of sulfides from previous stages I and III. Biological and abiological oxidation of sulfides may produce very small negative sulfur isotope fractionation, but generally oxidation products have very similar δ34S values to those of the source sulfide minerals (Toran and Harris 1989; Gu 2005). This explains the similitude of the δ34S values range for the sulfides and alunites from Mariquita (Fig. 3.7).

3.9.4, Isotope geothermometry

Geothermometric calculations in sulfides pairs from Ohmoto and Rye (1979),

34 where the Δ Spy-cpy values range between 0.4 to 1‰, produce temperatures from 388 to

844oC (Table 3.2). The lowest temperature agrees with the obtained temperatures from preliminary fluid inclusion studies. The calculated higher temperatures are excessively high and most likely reflect isotope disequilibrium. More geothermometric data using the sulfide pairs molybdenite-chalcopyrite and molybdenite-pyrite calibrated by Suvorova

(1974) generate temperature ranges between 530 to 570oC and 560 to 620oC, respectively. These temperatures are also higher than those obtained from the preliminary fluid inclusion analysis, also reflecting isotope disequilibrium.

Evidence of more isotope disequilibrium is also documented with the oxygen isotope systematic. Thermometry data in the quartz-magnetite pair from stage I generate

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excessively high temperatures of 746 and 759oC using the geothermometers of Clayton

and Keiffer (1991) and Chiba et al. (1989), respectively.

3.9.5, Ore fluids

The nature and origin of the fluids involved during the formation of ore deposits has been matter of interest to constrain genetic models in the study of ore deposits. In the

case of porphyry copper deposits, the involvement magmatic vs. meteoric-formation-

seawater waters involved in the mineralization has been subject of considerable debate.

Initially, studies of the ore fluids in porphyry copper systems concluded that magmatic

water was dominantly involved in the mineralization, and that later external heated

waters influenced the later alteration products (Taylor 1974; Sheppard 1969, 1971).

Later, Henley and McNabb (1978), based on isotope and fluid inclusion data documented

that at some stage of the hydrothermal system, the interaction of meteoric ground waters

with saline fluids evolved from a magmatic system. Bowman et al. (1987) demonstrated

that fluids in the potassic core at Bingham had isotope composition of magmatic waters,

and that fluids in the outer transitional and propylitic zones were progressively enriched

in δD and depleted in δ18O, leading to the conclusion that the external hydrothermal system was dominated by formation waters. Another example of non-magmatic fluids is documented in the isotope data from sericites and kaolinites, which indicate a seawater component (Chivas et al., 1984).

Few fluid inclusion studies showing the thermobarometric features and nature of the ore fluids of some deposits are available for the Cananea district (Wodzicki 1995;

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Esquivias-Flores 1998; Arellano 2004; González-Partida et al., 2009). This study presents

some preliminary data that help to constrain the nature of fluids involved during and post- mineralization stages in Mariquita deposit (Table 3.1).

Figure 3.11 shows the calculated isotope composition of the water in equilibrium with the minerals from the different hydrothermal stages from Mariquita, and a single data point for the high temperature hydrothermal minerals from Lucy deposit. In this

figure is also shown the range of the δ18O reported for the magmatic water in the Cananea

district (8 ± 1‰), based on the equilibrium with magmatic quartz from the quartz

feldspar porphyries (Wodzicki 2001). Along the global meteoric water line (GMWL) is

plotted the isotope composition of the winter and summer precipitation from Sierra Vista

Arizona (Coes and Pool, 2007), located 60 km north Cananea, and with an altitude

similar to that of Cananea (~1400 m).

Figure 3.11 also shows the calculated isotope compositions of the water in

equilibrium with quartz from stages I and III, based on preliminary fluid inclusions

temperatures and using the fractionation factors reported in Table 3.1. The isotope

composition of the water in equilibrium with the hydrothermal biotite from stage I is

located just above the magmatic water box of Taylor (1974), and plots close to the waters

from arc and crustal melts, and high temperature vapors related to convergent arc

volcanoes (Giggenbach 1992; Taylor 1992). The sericites related to stage III are

scattered, essentially located around the field of high temperature vapors from convergent

arc volcanoes (Giggenbach 1992). In addition, high temperature sericite from Lucy plots

near the water related to arcs and crustal melts.

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The fluids involved during the precipitation of the last hydrothermal stage (IV) in

Mariquita are calculated at 300 and 200ºC. These temperatures are not based on fluid

inclusions data, but instead, are assumed based on the lower temperature nature of the

latest hydrothermal stage. For 300ºC, the coexisting waters are closer to magmatic water,

whereas for 200ºC, they are closer to the meteoric water line (Fig. 3.11).

Figure 3.11 also shows the summer and winter precipitation data from Sierra

Vista region in Arizona. These isotope compositions are assumed similar to those from

the Cananea region since both are nearby and share similar altitude features (~1400 m).

Additionally, the porphyry copper deposits in this province were emplaced in a mountain

chain in which the altitude may have been comparable to the present day in the region.

Therefore, the isotope compositions of the meteoric water of winter and summer seasons shown in Figure 3.11 are assumed for the Cananea region at the time of the PCD formation.

Snowmelt recharge is significantly important as a contributor to the watershed in a mountainous system, particularly in the southwestern USA (Flerchinger et al., 1992;

Earman et al 2006). Winter precipitation can be more important than summer, since slow melting contributes more water to the aquifer mantle. Figure 3.11 shows the mixing between vapors related to arc volcanoes and the winter precipitation. Considering the exposed above, external waters with isotope compositions similar to those of winter season are more likely to be involved during the last hydrothermal stage in Mariquita.

From the isotope composition of the fluids related to mineralization in Mariquita, the mineral assemblages from mineralization stages I to III are clearly ore fluids of

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magmatic nature. Stage IV appears to be the result of the mixing between magmatic fluids and meteoric water. Figure 3.11 shows the range of the oxygen composition of the fluids in equilibrium with quartz from hydrothermal stages I, III, and from Lucy deposit, between the hydrothermal range.

3.9.6, Magmatic-hydrothermal geochronology

The first attempts to constrain the mineralization age in the Cananea district were done by Damon and Mauger (1966) and Varela (1972), who reported the first K-Ar ages in the district in K-rich minerals like phlogopite and sericite, both directly or indirectly to the mineralization or the mineralizing magmatism. The first mineralization age range reported in the district ranges between 56.7 and 59.9 Ma (Varela 1972; Damon et al.,

1983). The K-Ar geochronologic method has a lower closure temperature, and does not constrain the magmatic-hydrothermal event, and most likely represent either cooling ages or a disturbance in the isotopic clock due to a thermal geological event. The K-Ar geochronological method has shown this erroneous feature, and it has been demonstrated for some of the Mexican PCD’s (i.e. El Arco PCD in Baja California, K-Ar of ~98–106

Ma (Barthelmy, 1975) versus U-Pb and Re-Os ages ~164 Ma (Valencia et al., 2006);

Cumobabi in Sonora, K-Ar ~55 to 63 Ma (Scherkenbach et al., 1985) versus Re-Os age

~59 Ma (Barra et al., 2005).

The magmatic activity in the Mariquita area is recorded by different intrusion episodes. The most voluminous magmatism is represented by the Cuitaca granodiorite, and is present along the Cananea lineament. The Cuitaca batholith is presumably the

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precursor for the mineralization in the district (Valentine 1936; Noguez-Alcántara 2008).

The new U-Pb zircon data for the Cuitaca granodiorite, which hosts the Cu-Mo

mineralization in Lucy deposit, produces an age of 63.8 ± 1.1 Ma (Table 3.5). This

crystallization age agrees perfectly with the 64 ± 3 Ma reported initially in the Cuitaca town (Anderson and Silver 1977), located 20 km west of Cananea town. In addition, a similar age has been reported for the quartz monzonite porphyry that hosts the mineralization at the Milpillas PCD (Valencia et al., 2006).

The geographic position of Lucy deposit and the fact that the mineralization is

hosted within the ~64 Ma granodiorite, could suggest similar mineralization age to that from the Milpillas PCD (~64 Ma). The reported new Re-Os age for the molybdenite

mineralization from the Cu-Mo Lucy deposit is 1 Ma younger than the hosting rock, and no similar molybdenite mineralization age has been reported previously in the Cananea

district (Chapter 5).

The new U-Pb ages in zircons reported for two mineralizing porphyries in the

Mariquita PCD produce ages of 60.9 ± 1.2 Ma and 62.7 ± 1.3 Ma (Tables 3.3 and 3.4).

These two ages agree with those from the Cananea mine porphyries (Chapter 4), which

confirm a coeval magmatic activity across the district. These stocks are coeval with the

occurrence of the mineralizing quartz feldspathic stocks reported in the Cananea Mine

(see Chapter 5). In the Mariquita area, these quartz feldspathic stocks are mineralizing the

Cretacic Henrietta and the Laramide volcanic Formations, as seen in the rest of the

PCD’s in the Cananea district.

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The new Re-Os ages for the molybdenite mineralization in Mariquita PCD are

59.2 ± 0.3 and 59.3 ± 0.3 Ma (Chapter 5). This age range is consistent with the younger

mineralizing porphyry in the area, and suggests a constrained spatial and temporal

relationship between the porphyry and the primary sulfides mineralization in the

Mariquita area.

Even though field evidence and drilling exploration campaigns suggest the

absence of a porphyritic lithic unit in Lucy area, the available geochronological data suggest the possible relationship between the hydrothermal mineralization at Lucy coeval

to a magmatic-hydrothermal activity similar to the oldest porphyry in Mariquita PCD.

3.10, CONCLUSIONS

Despite the Mariquita and Lucy deposit belong to the western section of the

Cananea district, they are characterized by distinctive geological features that differ

slightly from the geology of the entire district.

The available thermobarometric data for the Mariquita deposit indicate

emplacement depths from 1 to 1.2 km with mineralization temperatures from 430 to

380ºC, similar to the P-T conditions reported in the Cananea Mine and Alacrán deposits

(Virtue 1996; Arellano 2004). In contrast, Lucy deposit shows deeper emplacement

depths (~3km) and higher mineralization temperatures (550-500 ºC) (Gonzalez-Partida et

al., 2009).

The sulfur isotope data demonstrate that the source of the sulfur is clearly

magmatic, and no external sulfur is later incorporated into the Mariquita system. The δ34S

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values of the last stage (alunite) are the result of oxidation of previous sulfides. In

addition, the isotope composition of the ore fluids involved during the mineralization of hydrothermal stages I to III determine essentially a magmatic origin, whereas the last stage (IV) consists of the mixing between fluids of the magmatic component and

meteoric water. The isotope data from Lucy deposit show that the source of sulfur and

the nature of the ore fluids are magmatic in essence.

The magmatic-mineralizing pulses reported here create the perfect scenario and

increase the potential for the presence of undiscovered mineralized bodies either

emplaced within the Cuitaca granodiorite (e.g. Lucy), or within the Laramide volcanic

rocks. Whatever the case might be, there is the potential of undiscovered mineralizing

pulses that aggregate economic value to the western and northwestern section of the

Cananea district.

In addition, the development of the Cuitaca graben by the Tertiary tectonics

creates suitable conditions for the generation of multiple supergene enrichment episodes

due to the constant subsidence of the Cuitaca basin, although the mineralized bodies

would be placed at deeper levels (i.e. Milpillas PCD). In addition, the Cuitaca graben

could promote to the formation of exotic copper mineralization in the basin fill (i.e.

Pilar), or the enrichment of a previous mineralized zone by the input of copper-bearing

solutions.

The understanding of the structural framework at the western portion of the

Cananea district plays an important role because Tertiary tectonics also potentially

contributes to the dismemberment and displacement of ore bodies. This can shed new

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light on the discovery and prospecting of new mineralized bodies that can highlight the attractiveness for the mineral exploration at the western and northwestern section of the district.

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Figure 3.1, Regional map showing the porphyry copper deposit belt northwestern Mexico and southeastern Arizona.

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Figure 3.2, Geologic map of the Cananea district modified after Wodzicki (1995) and Noguez-Alcántara 2008.

166

167

Figure 3.4, a) Geologic map of Mariquita PCD area; b) structural framework of Mariquita area showing the faulting stages.

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Figure 3.5, Map showing the different alteration zones from Mariquita area.

169

Figure 3.6, Schematic cross section of Lucy deposit showing the different alteration zones.

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Figure 3.7, Histogram showing the sulfur isotope data from Mariquita (gray columns) and Lucy (black bars) deposits.

171

Figure 3.8, U-Pb zircon ages from the mineralizing porphyritic units in the Mariquita PCD.

172

Figure 3.9, U-Pb zircon age from the Cuitaca granodiorite that hosts the Cu-Mo mineralization at Lucy deposit.

173

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Table 3.1, Oxygen and hydrogen stable isotope data for the Mariquita PCD and Cu-Mo Lucy deposit. FI (Stage) Sample Location Mineral T (°C) δD (‰) δ18O (‰) δD (‰) δ18O (‰) Assamblage H2O H2O E1 Mariquita Bt (I) Qz-Py-Bt 438 -74 4.4 -29 6.7 EV1 Mariquita Bt (I) Qz-Py-Bt 438 -74 4.5 -29 6.8 E2 Mariquita Bt (I) Qz-Py-Bt 438 -74 4.4 -29 6.7 MA027 Mariquita Bt (I) Cpy-Mt-Qz-Bt 438 -82 4.7 -37 7.0 MA041 Mariquita Mt (I) Qz-Mt-Py 436 - 3.8 - 11.4 MA041A Mariquita Qz (I) Qz-Mt-Py 436 - 9.7 - 6.4 MA002 Mariquita Qz (II) Qz-Kfs - - 20.0 - - MA002(r) Mariquita Qz (II) Qz-Kfs - - 19.7 - - (IIIc) Qz-Ser-Py- MA016A Mariquita Qz Cpy 388 - 9.8 - 5.5 (IIIc) Qz-Ser-Py- -46 MA016 Mariquita Ser Cpy 388 -77 10.7 10.6 (IIIc) Qz-Ser-Py- -23 MA017 Mariquita Ser Cpy 388 -54 4.4 4.3 (IIIc) Qz-Ser-Py- -23 MA017(r) Mariquita Ser Cpy 388 -54 4.1 4.0 (IIIc) Qz-Ser-Py- -5 MA 604 Mariquita Ser Cpy 388 -37 8.1 8.0 (IIIc) Qz-Ser-Py- 0 MA 604(r) Mariquita Ser Cpy 388 -32 8.1 8.0 (IIIa) Qz-Mo-Py- Mo-73 (1b) Mariquita Qz Cpy 388 - 10.1 - 5.8 BDO-08 269.4(a) Mariquita Anh (IIIb) Py-Tn-Anh - 12.2 - 7.4 BDO-08 269.4(b) Mariquita Anh (IIIb) Py-Tn-Anh - 13.7 - 8.9 LVD 63-340m Mariquita Alu (IV) Alu -84 7.6 -90(a); -87(b) -3.0(a); 1.1(b) MA020 Mariquita Alu (IV) Alu -59 9.7 -65(a); -62(b) -0.9(a); 3.2(b)

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Table 3.1, Continued. FI (Stage) Sample Location Mineral T (°C) δD (‰) δ18O (‰) δD (‰) δ18O (‰) Assamblage H2O H2O MA020(r) Mariquita Alu (IV) Alu -60 9.7 -66(a); -63(b) -0.9(a); 3.2(b) MA023 Mariquita Alu (IV) Alu -53 6.7 -59(a); -56(b) -3.9(a); 0.2(b) BL-1 Lucy Ser 500§ -54 10.9 223 -18.4 BL-2 Lucy Qz 500§ - 11.8 - -30.1

Notes: (Qz) quartz; (Py) pyrite; (Cpy) chalcopyrite; (Bt) biotite; (Mt) magnetite; (Ser) sericite; (Alu) alunite (Kfs) K-feldspar; (Mo) molybdenite; (r) repeated analysis; FI - Fluid inclusion. Calculated fluid isotope composition of Qz (Clayton et al., 1972), Bt and Ser (Suzuoki and Epstein, 1976), Alu (Stoffregen et al., 1994), and Mt (Zheng and Simon 1991). Isotope data calculated at (a) 205 ºC and (b) 300 ºC; (§) temperature data from González-Partida et al., 2009.

176

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Table 3.2, Oxygen and sulfur stable isotopes and fluid inclusion data for the Mariquita and Lucy deposits. FI δ34S MP T T Sample Location Mineral (Stage) Assamblage (‰) (°C) (°C) MA 041 (3) Mariquita Py (I) Qz-Mt-Py 436 0.7 Mo-73 (1a) Mariquita Cpy (IIIa) Qz-Mo-Py-Cpy 388 0.7 Mo-73 (1c) Mariquita Mo (IIIa) Qz-Mo-Py-Cpy 388 2.1 Mo-73 (1d) Mariquita Mo (IIIa) Qz-Mo-Py-Cpy 388 1.8 Mo-73 (2a) Mariquita Cpy (IIIa) Qz-Mo-Py-Cpy 388 0.3 388* Mo-73 (3a) Mariquita Py (IIIa) Qz-Mo-Py-Cpy 388 1.3 388* BDO-08 269.4 Mariquita Tn (IIIb) Py-Tn-Anh 3.0 1465 N (2) Mariquita Cpy (IIId) Qz-Ser-Py-Cpy 2.4 1465 W (1) Mariquita Py (IIId) Qz-Ser-Py-Cpy 3.1 474* 1465 W (2) Mariquita Cpy (IIId) Qz-Ser-Py-Cpy 2.2 474* 1480 NW (1) Mariquita Cpy (IIIc) Qz-Cpy-Bn 2.3 1480 NW (2) Mariquita bor (IIIc) Qz-Cpy-Bn 2.5 MA 040 (2) Mariquita Py (I) Qz-Mt-Py 0.9 MA026 (1) Mariquita Py (IIId) Qz-Ser-Py-Cpy 0.9 MA027 (2) Mariquita Py (IIId) Qz-Ser-Py-Cpy 1.3 Mo-71 (1) Mariquita Py (IIIa) Qz-Mo-Py-Cpy 2.4 844* Mo-71 (2) Mariquita Cpy (IIIa) Qz-Mo-Py-Cpy 2.0 844* Mo-74 (1a) Mariquita Mo (IIIa) Qz-Mo-Py-Cpy 2.2 PQR 3 (1) Mariquita Mo (IIIa) Qz-Mo-Py-Cpy 403 1.2 PQR 3 (2) Mariquita Py (IIIa) Qz-Mo-Py-Cpy 403 -3.5 Sulfuros (1a) Mariquita Py (I) Qz-Mt-Py 1.3 Sulfuros (2) Mariquita Py (I) Qz-Mt-Py 0.7 Sulfuros 1(1) Mariquita Cpy (IIId) Qz-Ser-Py-Cpy 0.6 BDO-08 269.4 Mariquita Py (IIIb) Py-Tn-Anh 3.8 298* BDO-08 269.4 Mariquita Py (IIIb) Py-Tn-Anh 3.7 283* BDO-08 269.4(a) Mariquita Anh (IIIb) Py-Tn-Anh 24.1 298* BDO-08 269.4(b) Mariquita Anh (IIIb) Py-Tn-Anh 24.8 283* 1465 N (1) Mariquita Alu (IV) Alu 1.0 1465 N(1) (r) Mariquita Alu (IV) Alu 1.0 LVD 63-340 m Mariquita Alu (IV) Alu 1.7 MA020 Mariquita Alu (IV) Alu 2.2 MA020(r) Mariquita Alu (IV) Alu 2.2 MA023 Mariquita Alu (IV) Alu 2.5 PQR-4 Mariquita Alu (IV) Alu 2.3

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Table 3.2, Continued. FI δ34S MP (Stage) Sample Location Mineral T (°C) (‰) T (°C) Assamblage Lucy 07 Lucy Mo Qz-Py-Mo 3.3 MA 049 (1) Lucy Py Qz-Py-Mo -4.6 MA 050 (1) Lucy Py Qz-Py-Mo -4.0 MA 049 (2) Lucy Mo Qz-Py-Mo 1.1

Notes: (Qz) quartz; (Py) pyrite; (Cpy) chalcopyrite; (Bt) biotite; (Mt) magnetite; (Ser) sericite; (Alu) alunite (Kfs) K-feldspar; (Mo) molybdenite; (Tn) tennantite; (Anh) anhydrite; (Bn) bornite. FI - Fluid inclusion; MP - Mineral pair; (*) Sulfur isotope thermometers from Ohmoto and Rye (1979).

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Table 3.3, U-Pb geochronologic analyses of the mineralizing porphyry 104 from Mariquita PCD U Analysis (ppm) U/Th 206Pb/204Pb 206Pb*/238U± (%) 206Pb*/238U* ± (Ma) ratio age 1 726 2.2 3532 0.0096 2.3 61.8 1.4 2 851 1.5 4028 0.0099 3.5 63.5 2.2 3 1259 1.9 5124 0.0096 2.9 61.3 1.8 4 719 2.4 3668 0.0098 1.1 62.9 0.7 5 788 2.5 4318 0.0098 1.6 62.6 1.0 6 734 1.9 3914 0.0098 1.8 62.8 1.1 7 810 1.5 3202 0.0100 1.8 63.8 1.2 8 1296 1.8 7442 0.0096 4.5 61.8 2.8 9 1396 1.7 7256 0.0097 3.1 62.1 1.9 10 864 1.8 3388 0.0099 1.6 63.6 1.0 11 948 2.2 5670 0.0099 2.6 63.6 1.7 12 717 2.3 3318 0.0100 1.2 63.9 0.7 13 1243 2.0 6062 0.0098 2.5 62.7 1.6 14 775 3.1 3978 0.0099 3.8 63.4 2.4 15 662 2.0 4388 0.0097 1.4 62.3 0.8 16 616 3.3 3402 0.0096 2.0 61.7 1.2 17 1186 2.1 12512 0.0097 3.8 62.0 2.3 18 367 2.5 2296 0.0098 1.3 62.7 0.8 19 201 3.2 1378 0.0101 1.7 65.1 1.1 20 364 3.6 2732 0.0098 1.8 62.7 1.1 21 1051 2.4 5256 0.0095 3.4 61.3 2.1 22 564 2.3 2976 0.0097 1.4 62.3 0.9 23 980 1.9 4868 0.0096 3.4 61.3 2.1 24 1625 1.2 6320 0.0099 2.5 63.3 1.6 25 581 2.7 3388 0.0096 2.2 61.9 1.4 26 1017 2.1 6914 0.0099 2.0 63.5 1.2 27 979 2.2 4544 0.0097 0.9 62.5 0.6 28 586 1.7 3220 0.0097 1.8 62.4 1.1 29 1021 1.8 4636 0.0098 1.6 62.7 1.0 30 1087 2.0 4616 0.0097 3.0 62.5 1.9

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Table 3.4, U-Pb geochronologic analyses of the mineralizing porphyry 604 from Mariquita PCD. U Analysis (ppm) U/Th 206Pb/204Pb 206Pb*/238U± (%) 206Pb*/238U* ± (Ma) ratio age 1 526 2.7 20814 0.0096 2.2 61.5 1.3 2 857 2.7 4486 0.0096 2.7 61.5 1.6 3 857 2.2 4434 0.0095 3.6 61.0 2.2 4 456 4.0 3536 0.0093 3.8 59.6 2.3 5 752 4.6 6464 0.0096 2.2 61.6 1.4 6 514 2.8 2994 0.0095 2.1 60.7 1.3 7 418 2.2 6640 0.0095 1.9 61.1 1.1 8 583 4.3 4808 0.0094 1.3 60.5 0.8 9 694 2.6 6124 0.0095 2.4 60.9 1.5 10 512 2.5 2954 0.0096 2.9 61.6 1.8 11 509 3.4 3962 0.0093 1.9 59.7 1.1 12 454 1.8 2118 0.0092 3.0 58.9 1.8 13 1129 2.4 6450 0.0093 2.8 59.8 1.7 14 719 2.2 4582 0.0094 1.0 60.2 0.6 15 773 2.4 6100 0.0094 2.5 60.3 1.5 16 908 1.4 5224 0.0093 1.4 59.9 0.8 17 1096 1.7 3564 0.0093 2.8 59.8 1.6 18 785 2.8 5228 0.0095 2.3 61.2 1.4 19 566 2.6 3210 0.0097 1.3 62.3 0.8 20 975 0.9 2654 0.0092 2.8 58.8 1.6 21 977 2.1 5878 0.0095 2.1 60.7 1.3

181

Table 3.5, U-Pb geochronologic analyses of the Cuitaca granodiorite from Lucy deposit U Analysis (ppm) U/Th 206Pb/204Pb 206Pb*/238U± (%) 206Pb*/238U* ± (Ma) ratio age 1 757 1.6 4404 0.0100 2.2 63.9 1.4 2 729 2.5 3266 0.0100 1.0 64.1 0.6 3 522 1.3 2904 0.0098 1.0 63.0 0.6 4 459 1.6 2516 0.0099 1.6 63.7 1.0 5 1341 3.2 7642 0.0099 1.8 63.6 1.1 6 363 2.1 2104 0.0101 1.0 64.8 0.6 7 752 2.2 3906 0.0096 1.3 61.6 0.8 8 876 2.1 3722 0.0098 1.0 63.0 0.6 9 1089 2.3 6064 0.0098 1.0 63.0 0.6 10 708 2.0 4606 0.0098 1.0 63.2 0.6 11 715 2.1 3514 0.0099 1.0 63.2 0.6 12 647 2.8 3512 0.0101 1.0 64.7 0.6 13 501 1.6 2174 0.0096 1.0 61.9 0.6 14 737 2.8 3956 0.0098 1.0 63.0 0.6 15 757 1.8 4482 0.0100 1.0 64.0 0.6 16 868 2.5 4472 0.0099 1.0 63.4 0.6 17 783 2.6 4462 0.0101 1.0 64.9 0.6 18 953 1.4 4500 0.0098 1.3 63.1 0.8 19 1065 2.2 6166 0.0100 1.0 64.3 0.6 20 831 1.5 4286 0.0100 1.1 64.4 0.7 21 339 1.7 1992 0.0101 1.7 64.5 1.1 22 773 1.4 2594 0.0103 1.1 66.3 0.7 23 868 2.0 1546 0.0090 1.0 57.7 0.6 24 748 2.4 3072 0.0101 1.0 64.9 0.6 25 640 1.7 3426 0.0105 1.0 67.3 0.7 26 650 2.4 1654 0.0097 1.9 62.5 1.2 27 517 2.2 3082 0.0096 2.4 61.6 1.4 28 1961 2.7 5390 0.0102 1.2 65.2 0.8 29 904 2.0 2672 0.0103 2.3 66.0 1.5 30 737 1.9 3010 0.0102 1.0 65.5 0.7 31 603 1.4 2750 0.0099 1.0 63.6 0.6 32 656 1.8 1338 0.0101 2.4 64.5 1.5

182

CHAPTER 4: GEOCHRONOLOGY OF THE PORPHYRY COPPER AND RELATED DEPOSITS IN THE CANANEA DISTRICT, NORTHWESTERN MEXICO

4.1, ABSTRACT

The Cananea mining district is among two of the most important sources and

producers of copper in Mexico, and represents the southern continuation of the porphyry

copper deposit province of the North American southwest. The district is mainly

characterized by the presence of the world-class porphyry copper deposit of the Cananea mine, along with other porphyry copper deposits (e.g. Mariquita, Milpillas, Alacran) and other mineralization styles (skarn, manto, and breccia pipe deposits).

With the new and previous geochronologic data using the Re-Os in molybdenites demonstrate so far five well-constrained mineralization events in the Cananea district.

The order of the mineralizing events from older to younger is: the Pilar (73.9 ± 0.3 Ma),

Milpillas (63.1 ± 0.4 Ma), Lucy (61.8 ± 0.3 Ma), Maria and Alacran (60.9 ± 0.3 to 60.4 ±

0.3 Ma), and Cananea mine and Mariquita (59.3 ± 0.3 Ma). The main mineralization in

the district is constrained over a short period of time (~4 Ma).

The U-Pb geochronological ages suggest a continued magmatic activity period of

~6 Ma for the mineralizing porphyries. Considering the new U-Pb zircon ages of the host

rock from the Pilar deposit, a period of ~20 Ma is suggested for the entire magmatic

activity in the Cananea district. The new data suggest a periodicity of the magmatic-

hydrothermal events responsible for the endowment and formation of the porphyry

copper deposits in Cananea district. These magmatic-hydrothermal pulses have been

recognized in the porphyry copper deposits from Arizona. 183

4.2, INTRODUCTION

The Cananea mining district, located in northern Sonora, is among two of the most important sources and producers of copper in Mexico (Fig. 4.1). The Cananea and

Nacozari districts in northwestern Mexico, belong to the southern continuation of the porphyry copper deposit (PCD) province of the North American southwest (Titley 1982), also known as the “great cluster” (Keith and Swan 1996). The Cananea district lies along a ~350 km northwest-trending regional line defined as the Cananea Lineament (Hollister

1978), from the Silver Bell PCD at the northwestern end in Arizona, through La Caridad

PCD at the southeastern end in Sonora (Fig. 4.2).

The Cananea district is mainly characterized by the presence of the famous world- class porphyry copper mineralization of the Cananea mine deposit. But this district also includes other smaller and important PCD’s along with other copper mineralization styles like those of skarn, manto, and breccia pipe deposits. At a district scale, these mineralized occurrences are also located along a NW-SE belt, and some of the outstanding mineral deposits in the district include those from the Cananea mine, Maria, Mariquita, Milpillas,

Alacrán, Puertecitos, Capote Basin, Elisa, Lucy, Pilar, El Toro, etc.

The Cananea region has been target of several geologic and geochemical studies, and certainly, the majority of these studies have been focused on the principal ore bodies

(e.g. Meinert 1982; Bushnell 1988; Wodzicki 1995; etc). Despite the economic importance of the Cananea district, few geochronologic data have indirectly constrained the ages of the mineralization. Recently, geochronologic data have been documented for 184

some of the mineralization and associated magmatism in the district (Barra et al 2005;

Valencia et al 2006; Chapter 4).

The age determination is fundamental for the understanding and evolution of any

geological processes. In particular, the determination of the mineralization age in the

study of ore deposits is essential for the understanding of a single deposit or at a mining

district scale. This study uses the U-Pb and Re-Os geochronological methods in zircons

and molybdenites respectively to constrain the duration and evolution of the

mineralization, and the associated magmatism from some of the smaller deposits of the

Cananea district.

4.3, CANANEA DISTRICT

The Cananea mining district is located just inside the western edge of the

Precambrian North American craton, within the Basin and Range extensional province,

and forms part of the PCD province of southwestern North American and western

Mexico (Titley 1982). Cananea is the most important mining district in Mexico and is among the most important copper producers in the world. Consequently, this mining district has been the subject of several geological studies, with emphasis on economic viability, geology, and mineralizing processes.

The Cananea district is known by the world class Cananea PCD, although there is the presence of other smaller PCD’s, in addition other ore deposits like breccia pipes, skarns, and manto deposits (Fig. 4.1). Emmons (1910) and Valentine (1936) originally established the basic geology of the district. Later, several authors complemented and 185

documented important geological issues (Mulchay and Velasco 1954; Velasco 1966;

Ochoa Landín and Echavarri 1978; Wodzicki 1995, Wodzicki 2001; Cox et al 2006). The different aspects concerning the mineralization styles in the Cananea PCD have been studied (Weed 1902; Austin 1903; Lee 1912; Virtue 1996). The breccia pipes, skarn, and manto deposits have been studied in detail by Perry (1933), Perry (1961), Bushnell

(1980), Meinert (1982), and Bushnell (1988). Also, several studies have been performed in Milpillas PCD (Carreón-Pallares 2002; de la Garza et al 2003; Valencia et al 2006;

Noguez-Alcántara et al., 2007; Noguez-Alcántara 2008), in Mariquita PCD (Woodburne

2000; Del Rio Salas et al., 2006; Zúñiga Hernández 2006), Lucy (Del Rio Salas et al.,

2006), and El Alacrán PCD (Amaya-Martínez 1970; Dean 1975; Arellano 2004).

Geochronological data concerning the mineralization and the magmatism across the Cananea district have been documented by various authors (Varela 1972; Anderson and Silver 1977; Meinert 1982; Damon and Mauger 1966; Damon et al. 1983;

McCandless and Ruiz 1993; Wodzicki 2001; Barra et al. 2005; Cox et al. 2006; Del Rio

Salas et al. 2006; Valencia et al 2006).

4.3.1, Cananea district geology

The oldest unit exposed in the district is the 1,440 ± 15 Ma Cananea granite (Fig.

4.3, Anderson and Silver 1977), which intrudes the Precambrian basement, the 1.7 Ga

Pinal schist, in northeastern Sonora (Silver et al., 1977; Anderson and Silver 1979;

Anderson and Schmidt 1983). Valentine (1936) described the Cananea granite as comprising two facies: (1) a coarse granitoid to pegmatitic rock composed of orthoclase, 186

oligoclase, quartz, and smaller amounts of hornblende, magnetite, and apatite; and (2),

the most abundant type, a granophyric granitoid with phenocrysts of quartz and a

microgranitoid matrix composed of orthoclase, microcline, quartz, and oligoclase.

The Cananea granite is unconformably overlain by a Paleozoic sedimentary

sequence that includes the Bolsa (Cambrian), Abrigo (Cambrian), Martín (Devonian),

and Escabrosa (Mississippian) Formations, and part of the Permian Naco Group

(Mulchay and Velasco 1954; Velasco 1966; Meinert 1982). Notwithstanding the intense

faulting, metasomatism, and hydrothermal alteration in the Paleozoic sequence, Mulchay

and Velasco (1954) suggested a correlation between the Paleozoic sedimentary sequence

at Cananea and similar sedimentary rocks in southeast Arizona. The Paleozoic

sedimentary sequence in Cananea is economically important because it hosts the Zn-Pb-

Cu skarn mineralization described by Meinert (1982).

The Proterozoic and Paleozoic rocks are unconformably overlain by a pile of

Mesozoic to Early Tertiary volcanic rocks (Valentine 1936). The Mesozoic rocks include

the Triassic-Jurassic and the Laramide magmatic arcs. The oldest rocks in the volcanic pile are the volcanic rocks of the Elenita Formation, composed of rhyolitic to andesitic tuffs and lavas with interbedded sandstone and quartzite. The Elenita Formation outcrops in the west and the southwest portions of the Cananea district (Fig. 4.3), and a thickness of 1,800 m is estimated (Valentine 1936). This formation is similar to the Late Triassic-

Mid Jurassic Wrightson Formation in southern Arizona described by Drewes (1971) and

Riggs and Blakey (1993). The Henrietta Formation is overlies the Elenita Formation (Fig

4.4; Valentine 1936), and is composed by medium to high-K, calc-alkaline, dacitic to 187

rhyolitic tuffs and flows (Wodzicki 1995). The Henrietta Formation occurs in a northwest

trending belt across the center of the Cananea district (Fig. 4.3), and generally dip E-NE except in the western part, where dips are W-NW (Ochoa Landín et al., 2007), and a thickness of 1,700 m is estimated (Valentine 1936). An Ar-Ar age in hornblende from a volcanic flow of the Henrietta Formation produced a minimum age of 94 Ma (Wodzicki

1995). This formation is important in the district because it hosts part of the copper mineralization of the Cananea ore body (Velasco 1966). The intrusive counterpart of the

Jurassic rocks within the Cananea district is the 175 Ma Torre syenite, which intrudes both the Elenita and Henrietta Formations (Wodzicki 2001; Noguez-Alcántara 2008).

The oldest Laramide rock is the Mariquita diabase, which consists of a high-K basaltic-andesite flows and intrusive bodies, and is characterized by a porphyritic “turkey tracks texture” (Wodzicki 2001). The Mariquita diabase occurs as volcanic flows shallowly dipping to the east, and makes up the upper 400 m of the Sierra Mariquita located east and north of the Mariquita and Maria deposits respectively (Fig. 4.3).

Between Sierra Mariquita and Cananea the Mariquita diabase occurs as dikes and stocks intruding the dacitic tuffs of the Henrietta Formation, and also as a thick flow that overlies the Henrietta Formation and grades upward into the overlying Mesa Formation

(Wodzicki 2001).

The Laramide Mesa Formation represents most of the Cretaceous volcanic activity in the district (Valentine 1936). From bottom to top, the compositions vary from trachy-basaltic, basaltic-andesite, andesitic, dacitic, to trachy-andesitic. Tuffs, agglomerates, lahars, and flows of andesitic composition are present (Valentine 1936; 188

Wodzicki 2001). The Mesa Formation crops out in the eastern portion of the district (Fig.

4.3) and a thickness of 1,500 m is estimated (Valentine 1936). These rocks are important

within the district because they host the disseminated copper mineralization. A flow

within this formation has been dated 69 ± 0.2 Ma using the 40Ar/39Ar method in biotite

(Wodzicki 1995), although a span of 72 to 68 Ma has been documented with the same

dating method around the Cananea district (Cox et al., 2006; Noguez-Alcántara 2008).

The Mesa formation overlies the Elenita and Henrietta Formations, and is intruded by the

Tinaja-Cuitaca granodiorite, and the Mariquita Formation of Laramide age, and by

younger intrusive bodies.

The earliest Laramide intrusive unit is the Tinaja-Cuitaca batholith (Fig. 4.3), which occurs as two spatially distinct, composite equigranular intrusive bodies named the

1977), and intrudes the Elenita, Henrietta, and Mariquita Formations. The composition 189

ranges from monzonitic to granodioritic to granitic, but the main compositional phase is

granodioritic (Wodzicki 1995).

The Tinaja-Cuitaca batholith is intruded by numerous near-vertical mafic dikes

oriented NW 60-80 and NE 40 (Valentine, 1936). These intrusions are dominated by the

Campana dikes and are dated at 58.4 ± 0.6 Ma (Carreón-Pallares 2002). The Henrietta and Mesa Formations are locally cross-cut by similar dikes. These mafic dikes are not cross-cut by younger quartz feldspar porphyries, and apparently were emplaced close to the time of solidification of the Cuitaca intrusive body (Wodzicki 1995).

(Valencia et al., 2006). Younger mineralizing quartz-monzonitic and granodioritic porphyries are present in the Cananea mine and the Maria, La Colorada, and Alacrán deposits, whose mineralization events yield Re-Os ages from 59 to 60 Ma (Barra et al.,

2005). Table 4.1 shows a summary of the general geologic features of the porphyry copper deposits from the Cananea district.

4.4, ANALYTICAL PROCEDURES

4.4.1, Re-Os method

The rhenium and osmium isotopes in the sulfides were analyzed following the procedure described in Barra et al. (2003). Approximately 0.05 to 0.1 g of each molybdenite sample was handpicked and loaded in a Carius tube. Spikes of 185Re and 190

190Os were added, along with 16 ml of a 3:1 mixture of HNO3 (16 N) and HCl (10 N), following the procedure described by Shirey and Walker (1995). About 2 to 3 mL of hydrogen peroxide (30%) was added to ensure complete oxidation of the sample and spike equilibration. The tube was heated to 240ºC overnight, and the solution later treated in a two-stage distillation process for osmium separation (Nagler and Frei 1997).

Osmium was further purified using a microdistillation technique, similar to that of Birck et al. (1997), and loaded on platinum filaments with Ba(OH)2 to enhance ionization. After osmium separation, the remaining acid solution was dried and later dissolved in 0.1 N

HNO3. Rhenium was extracted and purified through a two-stage column using AG1-X8

(100–200 mesh) resin and loaded on platinum filaments with Ba(SO)4.

The Re–Os analyses were performed by negative thermal ion mass spectrometry

(NTIMS) (Creaser et al. 1991) on a VG 54 mass spectrometer in the Geosciences

Department at the University of Arizona. Molybdenite Re-Os ages were calculated using an 187Re decay constant of 1.666 × 10–11 a–1 (Smoliar et al., 1996). Uncertainties were calculated using error propagation, taking into consideration errors from spike calibration, the uncertainty in the rhenium decay constant (0.31%), and analytical errors.

All rhenium and osmium in molybdenite samples were measured with Faraday collectors.

Blank corrections were insignificant for molybdenite.

Long-term instrument reproducibility is monitored using in-house standard solutions. The 187Os/188Os ratio of our standard was 0.148817 ± 0.000036 (1SD, n=25), whereas the 185Re/187Re ratio of our in-house liquid Re standard was determined to be 191

0.59542 ± 0.00036 (1SD, n=21). Osmium and rhenium blanks were less than 1.2 pg and

less than 15 pg, respectively. The osmium blank 187Os/188Os was ~0.181.

4.4.2, U-Pb method

For the U-Pb dating method in magmatic zircons, around 1 kg of the intrusive rocks were crushed and milled. Heavy mineral concentrates smaller than 350 µm were separated using the Wilfley Table. The zircons were concentrated using di-iodomethane heavy liquid and magnetic techniques. Later the zircons were handpicked under a binocular microscope, and were mounted in an epoxy resin and polished. Around 30 zircons from each sample were analyzed by laser ablation multicollector inductively