MINERALOGY AND GEOCHEMISTRY OF TOURMALINE IN CONTRASTING HYDROTHERMAL SYSTEMS: COPIAPÓ AREA, NORTHERN CHILE. Ana C. Collins

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1 MINERALOGY AND GEOCHEMISTRY OF TOURMALINE IN CONTRASTING HYDROTHERMAL SYSTEMS: COPIAPÓ AREA, NORTHERN CHILE by Ana C. Collins A Prepublication Manuscript Submitted to the Faculty of the DEPARTMENT OF GEOSCIENCES In Partial Fulfillment of the Requirements for the Degree of MASTER OF SCIENCE In the Graduate College THE UNIVERSITY OF ARIZONA 2010

2 STATEMENT BY THE AUTHOR This thesis has been submitted in partial fulfillment of requirements for the Master of Science degree at the University of Arizona and is deposited in the Antevs Reading Room to be made available to borrowers, as are copies of regular theses and dissertations. Brief quotations from this manuscript are allowable without special permission, provided that accurate acknowledgment of the 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 Department of Geosciences when the proposed use of the material is in the interests of scholarship. In all other instances, however, permission must be obtained by the author. Ana Collins (author s signature) (date) APPROVAL BY RESEARCH COMMITTEE As members of the Research Committee, we recommend that this thesis be accepted as fulfilling the research requirement for the degree of Master of Science. _Dr. Mark D. Barton Major Advisor (type name) (signature) (date) _Dr. Eric Seedorff (type name) (signature) (date) _Dr. Robert Downs (type name) (signature) (date) 2

3 Abstract Tourmaline group minerals can be useful for petrogenetic studies due to their refractory nature, chemical and isotopic variability, and widespread occurrence in many geologic settings. Near Copiapó, Chile, tourmaline occurs in a wide range of igneousrelated hydrothermal systems with widely varying types of mineral assemblages and chemical compositions, making this area an exceptional locality for studying controls on tourmaline chemistry. Copiapó tourmalines cover the majority of known tourmaline compositions excluding those associated with Li-rich pegmatites. Tourmaline formed in multiple, complex stages and is commonly intermediate schorl-dravite with a general progression in later tourmaline generations towards more Fe-rich and Al-deficient compositions with a dominant substitution of Fe 3+ for Al. This compositional trend, along with the presence of several tourmaline generations, is consistent with time-varying, relatively oxidizing, saline, acidic, boron-bearing fluids and reflects a greater host rock influence with progressive hydrothermal alteration. Tourmalines from other saline environments, both mineralized and otherwise, show similar compositional trends, reflecting analogous tourmaline-forming fluid compositions. The correlation between iron enrichment and highly saline fluids may reflect progressively more effective leaching and transport of iron from the host rock with time. Boron isotope analyses of tourmaline indicate a mixed fluid source, reflective of both magmatic and evaporitic sources, and is consistent with previous fluid-related studies of mineralizing fluids associated with iron oxide-copper-gold (IOCG) mineralization in the Candelaria-Punta del Cobre district. The study of tourmaline in these settings has the potential to constrain the origin(s) of this puzzling style of mineralization and can yield insights on the diversity of conditions under which tourmaline forms. Introduction Tourmaline occurs in a variety of geological environments and is a common accessory mineral in granitic pegmatites, low- to high-grade metamorphic rocks, and clastic sedimentary rocks. However, hydrothermal environments comprise some of the most common and diverse occurrences. Tourmaline s complex composition reflects changes in its chemical and physical environment which, combined with its refractory 3

4 nature and wide range of stability, make it well-suited to explore the conditions under which it formed (Henry and Guidotti, 1985). Consequently, tourmaline has been the subject of many studies and is useful for investigating differences between contrasting hydrothermal systems. Tourmaline is a complex borosilicate mineral group that has a general structural formula of XY 3 Z 6 [T 6 O 18 ](BO 3 ) 3 V 3 W, where X = Na, Ca, K, and, Y = Li 1+, Mg 2+, Fe 2+, Mn 2+, Al 3+, and Ti 4+, Z = Al 3+, Mg 2+, Fe 3+, V 3+, and Cr 3+, T = Si, Al, and B, V = OH, O, and W = OH, O, and F (Dyar et al., 1998; Hawthorne and Henry, 1999). Usually tourmaline is considered in terms of its end members, of which there are fourteen IMArecognized species (Table 1; Hawthorne and Henry, 1999). Solid solution in tourmaline is ubiquitous as simple or coupled substitutions. Table 2 summarizes common exchange vectors in tourmaline. Al, Fe, Na, Ca, and Mg comprise some of the most important substituent elements. Li can be important in tourmalines from rare-metal granites and pegmatites; however, it is minor or absent in most other types of settings (e.g., Henry and Guidotti, 1985). Documenting and understanding variations in these constituents is central to interpreting the significance of tourmaline in hydrothermal systems. Tourmaline is commonly associated with and co-precipitated during the formation of numerous types of mineral deposits, including copper, silver-gold, tin(-tungsten), massive sulfide, and uranium deposits, and occurs as breccia cement and clasts, veins, alteration envelopes and assemblages, and other metasomatic bodies (Slack et al., 1984; Pirajno and Smithies, 1992; Slack, 1996; Xavier et al., 2008). Tourmaline is commonly the principal host of boron in these deposits, and its durability allows it to preserve a detailed record of its formation even when dispersed during weathering and erosion. Tourmaline chemistry reflects the diverse compositions of both host rock and hydrothermal fluids, as well as differences in temperature and pressure of formation. This compositional record provides insight into mineralizing conditions, fluid flow, and possible sources of constituents in hydrothermal systems (e.g., in magmatic-hydrothermal systems: Pirajno and Smithies, 1992; Mlynarczyk and Williams-Jones, 2006; Dini et al., 2008; those sourced from external fluids: Palmer and Slack, 1989; Peng and Palmer, 2002; Xavier et al., 2008). Major-element trends have been used as guides for exploration (i.e., Clarke et al., 1989). For instance, Fe/(Fe+Mg) ratios in tourmaline vary systematically in Sn and Sn-W 4

5 hydrothermal deposits, with ratios decreasing with increasing distance from the magmatic source of mineralizing fluids and increasing interaction with the host rock (Pirajno and Smithies, 1992). Boron isotopes in tourmaline have been used to fingerprint the source of mineralizing fluids and can provide new insight as to the metallogenesis of various hydrothermal deposits (Palmer and Slack, 1996; Xavier et al., 2008). The abundance of tourmaline can also serve as a prospecting guide for undiscovered borate bodies and stratabound mineral deposits (Peng and Palmer, 2002; Slack, 1982). Care must be taken when using tourmaline as an exploration tool, as the compositions can be strongly influenced by the composition of the host rock, and the final composition may reflect an amalgamation of multiple sources and chemical interactions. Relatively little work has been done on the mineralogy and stability of tourmaline in many high-temperature hydrothermal systems, and tourmaline petrology and geochemistry have only recently been considered in iron oxide-copper-gold (IOCG) deposits (i.e., Xavier et al., 2008). IOCG systems are characterized by voluminous magnetite and/or hematite, variable amounts of Cu- and Fe-sulfides, gold, and REE, and low Ti contents compared to most igneous rocks (Hitzman et al., 1992; Barton and Johnson, 1996; Williams et al., 2005). In contrast to porphyry-type systems, magmatic compositions play only a secondary role on IOCG alteration mineralogy and elemental abundances (Barton and Johnson, 1996). IOCG deposits are generated by hypersaline, variably CO 2 -bearing, Cl-rich, and S-poor fluids, are formed at shallow to mid-crustal levels, and are closely associated with variably intense and voluminous sodic(-calcic) and potassic alteration. The origins of the ore-forming fluids are unsettled: they might be exsolved from magmas (Marschik and Fontboté, 2001; Sillitoe, 2003; Pollard, 2006) or be derived from external, evaporitic brines (Barton and Johnson, 1996; Xavier et al., 2008). Thus, understanding tourmaline in these settings has the potential to elucidate and constrain the origin(s) of this puzzling style of mineralization and, in well-chosen cases, can yield an independent set of constraints on the diversity of conditions under which tourmaline forms. This study systematically looks at igneous-related tourmaline occurrences in the Copiapó region of northern Chile, a part of the Chilean Iron Belt and one of the world's classic areas for IOCG mineralization (Marschik and Fontboté, 2001; Sillitoe, 2003). This 5

6 area is also the locus for widespread, though economically unimportant, porphyry copper mineralization (Maksaev et al., 2007). The first part of this project evaluates the petrographic, chemical, and isotopic characteristics of tourmaline in the diverse hydrothermal environments present near Copiapó, of which the IOCG systems in the Candelaria-Punta del Cobre district are pre-eminent examples. We then use this record to interpret the chemistry of the various hydrothermal fluids, the role of host rocks, and the identity of potential fluid sources. This involves geochemical and simple thermodynamic interpretations of the chemical and isotopic compositions in conjunction with ancillary data from other studies. Finally, we compare the Copiapó tourmaline patterns to those of tourmalines from other hydrothermal systems with the goal of gaining better insight into the controls on tourmaline composition in various hydrothermal settings and its use for interpreting their origins. Tourmaline-bearing hydrothermal systems near Copiapó Geologic context Chilean Coastal Batholith and related hydrothermal systems: Northern Chile contains a spatial and temporal progression of sub-parallel belts of iron, copper, and gold-rich hydrothermal systems (Fig. 1). Within the westernmost portion of these belts, the Late Jurassic to Early Cretaceous Coastal Batholith and related volcanic rocks of northern Chile host numerous hydrothermal systems, which fall mainly along the Chilean Iron Belt and can be divided into deposits affiliated with the IOCG class (Marschik and Fontboté, 2001; Sillitoe, 2003) or deposits of the porphyry copper family (e.g., Maksaev et al., 2007). Although the abundance of tourmaline within the Chilean Iron Belt and in younger porphyry copper deposits has long been recognized (e.g., Sillitoe and Sawkins, 1971), little detailed work has been done with regard to tourmaline geochemistry and its genetic implications in this region or elsewhere. The Chilean Iron Belt occurs principally along the long-lived Atacama fault system and its various splays. Hydrothermal alteration of sodic(-calcic), potassic, and hydrolytic types is nearly ubiquitous (Barton et al., 2005: unpublished mapping). The porphyry copper hydrothermal systems are present in association with the most felsic intrusive centers (tonalite to granodiorites), whereas IOCG deposits form independently of 6

7 magmatic composition and occur mainly in intrusive and volcanic rocks with intermediate compositions (e.g., diorites, tonalites, andesites) and to a lesser extent in the broadly coeval clastic to carbonate sedimentary rocks of the Early Cretaceous Chañarcillo Group (Marschik and Fontboté, 2001; Williams et al., 2005). Deposits range from magnetite-rich accumulations mined solely for iron with but minor amounts of copper mineralization to those with variable but typically abundant hematite (or magnetite replacing hematite) that commonly have more abundant copper, locally present in economic quantities (Sillitoe, 2003; Williams et al., 2005). Candelaria-Punta del Cobre district: The Candelaria-Punta del Cobre district is located in the Chilean Coastal Cordillera about 20 km south of Copiapó and contains a number of deposits that are classic examples of the IOCG family (Fig. 1; Williams et al., 2005)). The Candelaria deposit has proven reserves of 368 Mt at 0.55 percent Cu, 0.11 percent Au, and 1.97 percent Ag and probable reserves of 23 Mt at 0.54 percent Cu, 0.11 percent Au, and 1.91 percent Ag (Freeport-McMoran, 2009). The combined deposits from the Punta del Cobre district contain reserves of >120 Mt at 1.5 percent Cu, g/t Au, and 2 8 g/t Ag (Marschik and Fontboté, 2001). Smaller, vein-hosted deposits occur elsewhere in the region, the largest of which contains upwards of 10 Mt of ore with grades of >1.5 percent Cu. The Candelaria and Punta del Cobre deposits are hosted by volcanic and volcaniclastic rocks of the Punta del Cobre Formation, that underlie evaporite-bearing, carbonate-dominated sedimentary rocks of the Chañarcillo Group (Marschik and Fontboté, 2001). The deposits contain chalcopyrite, magnetite, hematite, and pyrite as the principal ore minerals, with minor sphalerite and, at Candelaria, metamorphic pyrrhotite. The andesites of the Punta del Cobre Formation are intensely potassically altered over a large region, including the areas that host the IOCG mineralization; locally, the upper parts of the Punta del Cobre Formation contain high metal grades, but ore-grade mineralization is nearly absent in the overlying Chañarcillo Group (Ryan et al., 1995). To the west of Candelaria and Punta del Cobre, the Copiapó batholith intrudes the supracrustal rocks and consists of calc-alkaline plutons that range from diorite through monzodiorite, quartz monzonite and tonalite, to granodiorite. These plutons range in age from 119 to 95 Ma and have O, Os, Pb, Nd, and Sr isotopic signatures indicative of 7

8 mixed mantle and crustal sources (Mathur et al., 2002; Marschik and Söllner, 2006; M.D. Barton, unpublished data). Although the plutons contain voluminous sodic(-calcic) alteration and minor IOCG vein deposits (Barton et al., 2005; Kreiner and Barton, 2009), none of these intrusions appears to be associated with the IOCG mineralization in the older rocks (e.g., not found either through mapping or exploratory drilling), nor is there any recognizable zonation of alteration or mineralization away from the batholith. In contrast, local porphyry-style mineralization is linked to particular granodioritic intrusions (Ryan et al., 1995; Barton et al., 2005). Conversely, the batholith is clearly responsible for a well-developed, 1-3 km wide metamorphic aureole that contains widespread copper-poor skarn alteration in the carbonate rocks and has deformed and overprinted the ores in the adjacent Candelaria deposit, but not in the more distal orebodies of the Punta del Cobre district (Fig. 1). Mineralization involved multiple magmatic, metamorphic, and metasomatic events, with a diverse set of structural controls and incontrovertible evidence from crosscutting relationships for discrete episodes of mineralization. Ar-Ar geochronology yields ages on silicate alteration minerals that can be broken into two groups: an older group of about 114 to 116 Ma and a younger group of 110 to 112 Ma with broadly similar Re-Os ages (115 Ma) on molybdenite from Candelaria (Mathur et al., 2002). Although Marschik and Fontboté (2001) argue that copper mineralization displays a close spatial and temporal correlation with post-magnetite, calcic amphibole alteration, magnetite is abundant and widespread throughout the deposit, and there appears to be no direct correlation between copper grades and the extent of magnetite mineralization (Ryan et al., 1995). Moreover, geologic evidence, including crosscutting relationships in dated rocks from the Punta del Cobre side of the district, indicates that at least some and perhaps much of the mineralization is of Punta del Cobre age (ca. 130 Ma) and predates the batholith (Pop et al., 2000; M.D. Barton, unpublished data). Most workers in the district have inferred a magmatic source of hydrothermal fluids based on the isotopic similarity of the ores and igneous rocks and sulfide sulfur isotopic compositions that are near zero per mil (e.g., Mathur et al., 2002). Even in these proposed magmatic systems, however, the participation of non-magmatic fluids in mineralization 8

9 has not been ruled out (Marschik and Söllner, 2006), and recent work has demonstrated a significant component of external, non-magmatic fluids in the ores and related alteration (Barton et al., 2005; Chiaradia et al., 2006). Tourmaline occurrences in the Copiapó region Petrographic characterization of tourmaline in over 100 polished thin sections (out of more than 1000 available sections) has been carried out as part of a multi-disciplinary geologic and geochemical study of the region (M.D. Barton and others, in progress). The optical properties, mineral assemblages, and timing relationships were determined using standard transmitted and reflected light techniques. Tourmaline is by far the most common boron mineral in the Copiapó area, although datolite and dumortierite also occur locally (Ryan et al., 1995; Kreiner and Barton, 2009). Tourmaline occurrences are diverse and include: occurrences with pluton- and volcanic-hosted IOCG mineralization of multiple styles and with several alteration types, high-temperature veins associated with biotite(-hornblende) quartz monzodiorites, locally in the contact metamorphosed rocks, and in many areas with intensely sodicly-altered, metal-depleted rocks. With the exception of tourmaline associated with quartz-feldspar assemblages in some of the granitoids, nearly all tourmaline is quite fine-grained (<100 microns), rather massive, and, therefore, subtle in appearance. This may account for the sparse attention that it has received in earlier studies. Petrographic studies, summarized in Table 3, show that tourmaline formed during multiple stages of hydrothermal activity, as indicated by crosscutting, overgrowth, and replacement relationships. Appendix 3 contains petrologic descriptions of each sample that was analyzed chemically for major and minor elements. Tourmaline is scarce in the Candelaria deposit, perhaps because of highly alkaline conditions which may have produced some of the late datolite (Ryan et al., 1995). Conversely, tourmaline is abundant in the little metamorphosed Punta del Cobre district. Tourmaline is also found in iron-dominated IOCG occurrences, such as the Cerro Iman and Cerro Negro Norte deposits. Tourmaline occurs as euhedral columnar, granular, massive, and acicular grains (Fig. 2B) in veins (as the dominant or an accessory mineral), breccia cement, brecciated 9

10 clasts (Fig. 2A), and as part of alteration assemblages (Fig. 2H) in advanced argillic, potassic, sericitic, sodic, and sodic-calcic types of alteration assemblages (Fig. 2). Characteristics of Tourmalines in the Copiapó Area Tourmaline from multiple localities within the Copiapó area were analyzed via electron microprobe and laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS) methods to characterize the chemical and boron isotopic compositions of tourmaline and associated minerals (see footnotes of Tables 4 and 6 for operating conditions and Appendix 1 for further details). These results are combined with petrographic data to document correlations between tourmaline occurrences, textures and optical properties, and to provide the basis for interpreting the significance of tourmaline in the district. Tourmaline compositions Tourmaline from Copiapó varies widely in chemical composition, reflecting systematic changes associated with types of hydrothermal mineral assemblages and geologic settings. Overall, tourmalines show the greatest variation in Al, Mg, Ca, Na, Fe +2 and Fe +3 concentrations and contain insignificant K, Mn, or Cl. Representative analyses of tourmaline from each sample can be found in Table 4 (complete results are provided in the Appendix). Structural formulae were calculated on the basis of 15 cations (T + Y + Z = 15). Light-element (H, Li, B, O) contents were not determined, but boron was assumed to be stoichiometric (Grice and Ercit, 1993), and hydrogen and oxygen were adjusted to meet charge balance constraints. Fe +3 /Fe total calculations were derived from working curves using the methods outlined by Fialin et al. (2004), which is outlined in detail in Appendix 1. Analyses of Copiapó tourmalines largely fall within the alkali compositional group defined by Hawthorne and Henry (1999) (Fig. 3). The composition of the mineral assemblage exerts the primary control on tourmaline compositions that plot outside of the alkali compositional field. Alkali-deficient tourmaline from the advanced argillic type of hydrothermal mineral assemblage contain little Ca and plot along the base of the ternary diagram along the [Na(Mg,Fe)]( Al) -1 exchange vector (refer to Table 2 for a list of common exchange vectors). Tourmalines from calcic assemblages show a core-to-rim 10

11 progression from moderately sodic to calcic end members along the NaAl[Ca(Mg,Fe)] -1 exchange vector in response to increasing Ca concentration with progressive evolution of the hydrothermal system. Figure 4 shows discrimination diagrams for common tourmaline end members. Tourmaline compositions trend from alkali-free (magnesiofoitite-foitite) to dravitic compositions (Fig. 4A) or from schorl-dravite to calcic (uvite-feruvite) end members (Fig. 4). These latter compositions reflect the overall assemblage, commonly towards non-acidic compositions, with higher Fe, Mg, Na, and/or Ca. For the most part, early tourmalines trend from intermediate schorl-dravite compositions to another tourmaline compositional end member. Overall, the majority of tourmaline corresponds to the dravite-schorl solid-solution series (X Mg = ) with a dominant trend towards povondraitic compositions. Copiapó tourmalines cover the majority of known tourmaline compositions in the Al-Fe- Mg and Ca-Fe-Mg compositional diagrams of Henry and Guidotti (1985), occupying all fields except that associated with Li-rich pegmatites (Figs. 5 and 6), denoting a wide array of possible exchange vectors. Within the Al-Fe-Mg diagram, the greater proportion of tourmaline generations fall within field 6, representing Fe 3+ -hydrothermally altered quartz-tourmaline, calc-silicate, and metapelitic rocks. According to Henry et al. (1999), tourmaline samples that fall below the schorl (buergerite)-dravite join are Al-deficient (less than 6 apfu Al) and either Fe- or Ca-Mg-rich. In the absence of a substantial calcic component, which would result in the coupled substitution of Ca and Mg or Fe 2+ for Na and Al, it can be inferred that Fe 3+ is a major substituent of Al in the Z-site to maintain charge balance (Fig. 7A). The implications of this will be covered in further detail in the discussion. The elements that display the greatest variation are plotted in Figure 7 to show potential exchange vectors that may play a role in the chemical complexity of Copiapó tourmaline. For the most part, tourmaline compositions (with the exception of those from potassic assemblages) do not fall along the schorl-dravite solid-solution line (Fig. 7B). Al-deficient tourmaline tends to plot along the CaMg(NaAl) -1 or the FeAl -1 exchange vector, whereas tourmalines with slightly more than 5.5 apfu Al appear to lie along the Al(NaMg) -1 exchange vector, as indicated by the negative slope in Figure 7C and the 11

12 increasing influence of X-site vacancies in tourmalines with greater than 5.5 apfu Al (Fig. 8). Tourmalines from sodicly and some sericitically altered rocks appear to be largely dominated by the FeAl -1 exchange vector (Fig. 7C). Magmatic tourmalines, as well as tourmalines from sodic-calcic, potassic, and some sericitic assemblages, display a dominant negative correlation between Al and Ca, whereas sodic and advanced argillic tourmalines show no correlation (Fig. 7E). Overall, tourmalines trend from greater than 6 apfu Al to less than 6 apfu Al contents. Petrography and assemblage data Petrographic studies suggest that tourmaline was deposited in several stages of hydrothermal alteration and is found in association with many diverse types of mineral assemblages (see Table 3). Elongate tourmaline crystals show evidence of crack-seal mechanisms for their formation, indicative of multistage development of tourmaline at the vein scale. Tourmaline compositions reflect both the host rock and fluid source characteristics and are strongly controlled by the mineral assemblage. In the following sections, tourmaline-bearing mineral assemblages are discussed especially in terms of their influence and correlation to tourmaline chemical variations. Advanced argillic assemblages: Advanced argillic alteration assemblages are stable under extremely acidic conditions (i.e., low ph or low a K+ /a H+, broadly coincident with the stability of kaolinite, pyrophyllite, or andalusite in originally K-feldspar-bearing rocks) (Seedorff et al., 2005). Tourmaline-bearing advanced argillic assemblages in the Copiapó area are both hydrous and anhydrous. Hydrous tourmaline-bearing samples contain dumortierite, pyrophyllite, hematite, calcite, andalusite, and quartz. Tourmaline is light tan to light blue-brown with Mg concentrations increasing from core to rim. The latest generation of tourmaline growth, occurring as comb-like needles on preexisting tourmaline, is more aluminous than the earlier phases (Fig. 9). Overall, these tourmalines have high F contents compared to other assemblages from the Copiapó region. It is likely that this assemblage represents a metamorphic overprint by the batholith or preexisting advanced argillic assemblages in the volcanic rocks. The two chemically and morphologically distinct generations of tourmaline represent discrete periods of growth possibly reflective of formation before and following the emplacement of the intrusion. 12

13 Early Fe-rich cores, possibly accompanied by other Al-rich phases such as albite, were replaced by dumortierite and pyrophyllite, a low temperature, Al-rich, Na-poor assemblages with a fine-grained, calcitic groundmass. Dumortierite and pyrophyllite are then replaced by Mg-rich tourmaline, which is common for metamorphic environments (Henry and Dutrow, 1996), and calcite is recrystallized to form equant interlocking grains. Hematite probably formed early and was replaced by magnetite during the metamorphic overprint. Late hematite after magnetite ("martite") then formed during some later event. Higher temperature, largely anhydrous, advanced argillic assemblages contain finegrained, acicular tourmaline in association with andalusite, quartz, hematite, and minor sericite and lazulite (Fig. 2B). The presence of andalusite in these assemblages may suggest a contact metamorphic overprint. Tourmalines from these assemblages have a narrow range of X Mg values ( ), are highly aluminous, alkali-deficient (X Na <0.7 apfu), and plot along the dravite-magnesiofoitite, Al(NaMg) -1, exchange vector (Fig. 7C). Sericitic-chloritic assemblages: Sericitic-chloritic assemblages are stable is an acidic environment, though at somewhat higher ph or higher a K+ /a H+ conditions than those characteristic of advanced argillic alteration (Seedorff et al., 2005). Tourmalines present in within sericitic or chloritic assemblages range from well formed to poorly formed with corroded grain boundaries. The compositions of tourmaline in these assemblages vary greatly, trending from foitite to schorl-dravite and commonly plot along the exchange vector Na(Mg, Fe)( Al) -1. Tourmalines are progressively enriched in Fe, possibly indicative of an increase of host rock control on tourmaline composition. Periods of distinct growth and dissolution are particularly noticeable in tourmalines from sample C2J-124, in which both early and late tourmaline have embayed cores and inclusion edges. Any zones that are present are commonly discordant. C2J-124 appears to have at least four generations of tourmaline, with an early aluminous, vacancy-rich, light blue-gray tourmaline generation followed by brown, Mg-Fe-rich anhedral tourmaline. Well-formed tourmalines in sericitic assemblages are found in the vicinity of the San Gregorio pluton. Overall, albite phenocrysts in these rocks are intact and only display sericitic alteration in the presence of tourmaline. In this case, tourmalines trends from 13

14 schorl to feruvite, which is possibly correlated to the breakdown of magmatic pyroxene or amphibole. A later brown, Mg-rich schorl generation, intergrown with titanite, contains among the highest Ti concentrations in the dataset. In other samples from this area, the relationship to sericitic alteration is unclear (i.e., intact tourmaline vein with scant amounts of partially sericitized plagioclase). Sodic assemblages: Sodic assemblages form as a result of Na-enrichment of host rocks by highly saline fluids (Seedorff et al., 2005). Albite is the most common mineral associated with tourmaline in these assemblages and is the only other sodic phase. Hematite and Fe-sulfides or sulfates are the next most abundant association. Sodiclyaltered rocks commonly have volcaniclastic or volcanic protoliths and are variably overprinted by breccias associated with hydrothermal alteration. Tourmaline is common, typically occurring as part of the breccia cement or in veinlets. These tourmaline are largely dravitic and commonly trend from Al-saturated to slightly Al-deficient compositions. Tourmalines that occur with specular hematite are Al-deficient and Ferich, commonly falling within the schorl-buergerite compositional field, and are intricately zoned, which may be indicative of a fracture-fill style of formation. Early phases are granular, Mg-rich, and contain higher F (up to 0.19 apfu) and Ti than later generations, which become progressively more Fe-rich through time. In some cases, there is no compositional continuity between different tourmaline generations, indicating distinct tourmaline-forming conditions. Moreover, the grain boundaries of the earliest generations are not embayed, showing that they were not destabilized by the later mineralizing fluid. Sodic-calcic assemblages: Sodic-calcic alteration is associated with the addition of Na ± Ca by hot (>350 C), saline fluids resulting in higher temperature assemblages than those associated with sodic alteration (Seedorff et al., 2005). Tourmalines from sodiccalcic assemblages are commonly dravitic to uvitic in composition and operate along the exchange vector CaMg(NaAl) -1. These tourmalines also tend to have higher Ti values. Epidote, titanite, actinolite, and calcite are common minerals that occur with tourmaline in these assemblages. Tourmalines are also found in association with actinolite as part of a calcic overprint on earlier, potassically altered assemblages as evidenced by abundant biotite with chlorite in mafic sites. Similarly, calcic overprints on sodic assemblages are 14

15 also present, with early cores of unzoned, dravitic tourmaline slightly replaced by albite and later rimmed by calcic, euhedral tourmaline. The overall increase in Fe, Mg, and Ca moving from core to rim, which is evident in all samples, implies that reacting fluids became increasingly enriched in these elements during tourmaline formation. Tourmalines occur as veins and isolated clusters. Vein tourmaline occurs in sodiclyaltered rocks with albite or in calcicly-altered rocks with actinolite and/or epidote. Compositionally, vein tourmaline trends from sodic to calcic compositions and generally has sharp boundaries with included minerals. In some cases late generations of tourmaline may have reaction rims (Fig.10) and commonly have dravitic overgrowths. Tourmaline clusters are generally well formed, but may have irregular edges in the presence of epidote. Tourmalines from the aureole of the San Gregorio pluton are the most oxidized and Fe-rich samples in the dataset. These tourmalines plot along two different exchange vectors: CaFe(NaAl) -1 (schorl-feruvite substitution) and FeAl -1 (schorl-povondraite substitution). This is readily evident in Figures 7A and 7E, where both Ca and Fe concentrations show a negative correlation with Al content. Accordingly, tourmaline compositions progress from the compositional ranges normally associated with Li-rich granites in the earliest generations to that of Fe-rich hydrothermally altered rocks in the latest generation, as shown in Figure 5. Potassic assemblages: Potassic assemblages are stable in highly alkaline environments with high a K+ /a H+ ratios. Potassic alteration commonly occurs at high temperatures and is linked to magmatic-dominated fluids (Seedorff et al., 2005). The majority of drill core samples from Santos, believed to represent the intermediate levels of the hydrothermal system, are characterized by potassic alteration (biotite ± K-feldspar) with actinolite, epidote, and albite present locally (Marschik and Fontboté, 2001). Allanite can also be locally abundant. Tourmalines from this area either occur as veins, breccia clasts and cement, part of the alteration envelope, in altered clasts, or columnar crystals commonly in larger biotite grains. Breccia clasts are indicative of a history of repeated brittle deformation as well as distinct phases of tourmalinization. These Santos samples are predominantly dravitic and either trend towards more Fe-rich or Mg-rich compositions. For example, tourmalines from the deepest sample (DDH ) in the 15

16 dataset tend to plot mostly in the Li-poor granite field with trends towards Alundersaturated and intermediate schorl-dravite compositions from core to rim (Fig. 5). However, this chemical progression is not systematic, given that elemental abundances fluctuate at the micron scale (Fig. 11). At higher levels in the hydrothermal system, Santos tourmalines trend from the Al-saturated and Al-undersaturated metapelitic fields towards more Fe-rich and Al-deficient compositions (Fig. 5). Although tourmalines from different depths appear to have different geneses, the compositions of their later generations overlap, suggesting a second-order control on tourmaline chemistry other than depth (e.g., interaction with wall rocks). Potassic alteration postdates sericitic alteration in some areas, as evidenced by late tourmaline-biotite veins in a sericitic matrix (i.e., sample C2B-576c). Whereas Santos tourmalines are distinctly chemically zoned, these tourmalines are light blue in transmitted light, unzoned, and largely fall within field 4 of Figure 5 associated with Alsaturated phases. Optical characteristics: The petrographic characteristics of tourmaline vary considerably, and pleochroic color and intensity do not always correlate with chemical composition. In general, tourmaline is weakly to strongly pleochroic. The pleochroic characteristics of tourmalines from each sample are summarized in Table 5. In the case of pleochroic tourmaline, strong optical absorption and pleochroism suggest that the tourmaline may contain significant amounts of Fe with mixed valences (2+ and 3+), resulting in intervalence charge transfer (Mattson and Rossman, 1987). On the other hand, tourmaline that contains monovalent iron should exhibit weak absorption. Nonpleochroic tourmaline tends to also be complexly zoned and dark in color. An interesting characteristic of some of the vein tourmaline associated with the San Gregorio pluton is their intense pleochroism and optical characteristics. Sample C2B-655 contains tourmalines that are nearly optically opaque in thin section, similar to the ferridravite (now recognized as povondraite) species identified by Walenta and Dunn (1979). These are also some of the most Fe-rich tourmalines, suggesting that optical opacity may be related to high Fe concentrations relative to Al. Although Fe 2+ => Ti 4+ charge-transfer has been cited as the cause for blue color in kyanite and dumortierite (Parkin et al., 1977; Platonov et al., 2000), blue tourmaline from 16

17 Copiapó samples have among the lowest Ti concentrations and the highest Fe values in the dataset, suggesting that Fe is the primary chromophore. Brown colors appear to correspond to Mg and Ti contents, although Ti content in the Copiapó tourmaline is highly variable. Taylor and Slack (1984) interpreted the blue to black colors associated with schorl to be dominantly influenced by Fe 2+ => Fe 3+ and O 2- => Fe 3+ charge-transfer processes, and the brown hues of dravite by uv-centered O 2- => Fe 2+ and Fe 2+ => Ti 4+ processes. Copiapó tourmalines are generally optically zoned, sometimes spectacularly, varying from concentric to highly irregular. The color change can be either gradational or display sharp optical discontinuities. Cores range from irregular (common) to euhedral with euhedral to corroded rims. Unzoned tourmalines are brown or blue. The complexity of zoning evident in thin section and backscattered-electron imaging (Fig. 12) is reflected in the compositional variation of tourmalines from even one thin section. Summary Tourmalines formed in distinct episodes of mineralization in diverse mineral assemblages. Copiapó tourmaline is commonly pleochroic and optically and chemically zoned. The chemical variation evident in tourmaline from this area is a function of fluid composition and type of hydrothermal mineral assemblage. Tourmaline compositions operate on a variety of exchange vectors, with Fe, Mg, Al, Na, and Ca contents showing the greatest variation. Most notably, tourmalines from all types of assemblages (except advanced argillic) show a negative correlation between Fe and Al (Fig. 7A). Copiapó area tourmalines largely have sodic compositions. However, advanced argillic assemblages contain the most aluminous and alkali-deficient tourmalines in the dataset. Calcic tourmalines (operating along the Ca(Mg,Fe)(NaAl) -1 exchange vector) are found in sodic-calcic, potassic, and some sericitic types of hydrothermal mineral assemblages. Magmatic tourmalines, however, also show a dominant negative correlation between Al and Ca. Tourmalines within sericitic types of hydrothermal mineral assemblages consist of two distinct populations: one that is aluminous and trends toward more Fe-rich compositions and another that is more calcic as the result of the breakdown of magmatic mafic minerals (i.e., pyroxene or amphibole). Sodic alteration assemblages 17

18 contain tourmalines that show the same trend as some sericitic assemblages, but display more compositional scatter that is dominated by substitution along the FeAl -1 exchange vector. Tourmalines present in potassic types of alteration assemblages are the only tourmalines to plot along the FeMg -1 exchange vector and, along with sodic-calcic tourmalines, also trend towards more Fe-rich compositions. Boron isotopic compositions Boron isotopic analyses were accomplished via LA-ICP-MS techniques at the University of Arizona. Isotope ratios for selected coarse-grained samples and analytical methods are presented in Table 6. Given the fine-grained nature of most Copiapó area tourmaline, these results are quite selective. δ 11 B values range from per mil, with over half of the analyses falling below -1 per mil (Fig. 13). The most positive values come from vein tourmaline in sample C2B-352e, which is considered the clearest and simplest example of magmatic tourmaline in the vicinity of the San Gregorio pluton. Thus, this would indicate a mixed boron source or equilibration of magmatic boron with the host rock, thus driving δ 11 B values to more positive values. Smith and Yardley (1996) found that isotopically lighter tourmaline was associated with Fe enrichment and Li enrichment in the Cornwall district. However, there is no observed correlation between elemental abundances and isotopic composition in the Copiapó tourmaline, nor is there any correlation between δ 11 B values and tourmaline morphology or sample locality. Overall, there appears to be variations in isotope compositions where no chemical zoning is evident. The variation from magmatic to more positive values is indicative of the influx of a source with a positive signature. Potential sources will be covered in further detail in the discussion. Controls on Tourmaline Compositions Tourmaline compositions reflect the host rock and hydrothermal fluid compositions with progressive evolution of the hydrothermal system, as well as differences in temperature and pressure of formation. This compositional record provides insight into mineralizing conditions, fluid flow, and possible sources of constituents in hydrothermal systems. Although tourmaline has a wide range of stability, it is strongly controlled by the composition of the mineralizing fluid and associated mineral assemblage. 18

19 Tourmaline stability Stability of tourmaline end-members: Tourmaline is stable over a broad range of pressure and temperature conditions (Henry and Dutrow, 1996); however, the composition of tourmaline solid solutions should be governed by the coexisting mineral assemblages and fluid compositions. Here we use chemographic techniques combined with available thermodynamic data to evaluate the effect of independent compositional variables on tourmaline compositions Thermodynamic data for minerals and aqueous species from the SUPCRT92 database (Johnson et al., 1992) were used to create a chemical potential diagrams for several projections within the system Na 2 O B 2 O 3 Al 2 O 3 MgO CaO SiO 2 HCl H 2 O at 400 C and 500 bars. Tourmaline, quartz, and an aluminum mineral are considered to be saturated throughout each activity diagram. The fugacity of oxygen is defined by magnetite-hematite. Comparing modeled and measured tourmaline compositions provides information on the conditions prevailing in its host environment. Tourmaline stabilities for magnesiofoitite dravite and dravite uvite end members are plotted in Figure 14. In highly acidic environments, magnesiofoitite predominates over dravite. Dravite is stable in environments with high Na and Mg activities, whereas uvite stability appears to be constrained to lower Mg and higher Ca activities. Stability of tourmaline: Tourmaline has a wide range of stability from low to high temperatures and pressures (<150 C to >700 C and 1 bar to >10 kbars) but is strongly influenced by the composition of the fluid phase and mineral assemblage (Henry and Dutrow, 1996). The solubility of aluminosilicate phases and components (e.g., Al) in borate fluids increases with increasing fluid alkalinity and may be indicative of changing speciation mechanisms in solution as a function of ph (Morgan and London, 1989). Thus, minerals containing relatively large amounts of alkalis that react with water to produce alkaline solutions inhibit tourmaline growth, as tourmaline formation is favored in strongly to weakly acidic fluids (Frondel and Collette, 1957; Morgan and London, 1989). Moreover, the minimum aqueous boron necessary to stabilize tourmaline increases with increasing temperature and ph. Above ph 6.5, no level of boron can stabilize 19

20 tourmaline and another boron-bearing phase forms. This would explain the absence of authigenic tourmaline in most evaporitic deposits (Henry and Dutrow, 1996). The activity of Al 2 O 3 or equivalent aqueous species also contributes to tourmaline stability and formation, which are favored in acidic fluids with high availability of Al species. However, since Al transport is facilitated by alkali borate species, such as Na 2 B 4 O 7, a mixture of acidic and alkaline boron compounds is essential to provide the necessary Al for tourmaline-forming reactions (Morgan and London, 1989). Oxygen fugacity (fo 2 ) is another important constraint on tourmaline stability (Morgan and London, 1989). Irregular distributions of tourmaline could reflect limited boron and water availability, variable fo 2, and restricted mobility of mafic cations (e.g., Fe and Mg) necessary for tourmaline formation (Gawęda et al., 2002). Stability of tourmaline relative to other boron-bearing minerals: Tourmaline is the most common borosilicate in geologic environments. In highly alkaline and/or silica- or aluminum-undersaturated conditions, however, tourmaline growth is inhibited and other borosilicates form instead (Grew, 1996). High Ca concentrations, or high Ca/Al ratios, may not be favorable for the formation of tourmaline, and boron may be distributed in other species, such as danburite, serendibite, and axinite, to name a few (Frondel and Collette, 1957). Like tourmaline, danburite (CaB 2 Si 2 O 8 ) is preferentially stable in acidic fluids (Morgan and London, 1989). In highly acidic environments, dumortierite occurs over tourmaline due to the lack of alkalis, Fe, and Mg necessary for tourmaline formation (Taner and Martin, 1993). However, water is a key component for dumortierite formation (Werding and Schreyer, 1990). Tourmaline that does occur with dumortierite tends to be alkali-deficient, with some proportion of Na, and Al-rich, commonly with tetrahedral Al (Foit et al., 1989). Other borosilicates, such as serendibite, commonly form in calcic, silica-undersaturated conditions, and tourmaline that forms in these settings are typically uvitic (Grew, 1996). Major-element variation Copiapó tourmalines are compositionally complex and formed in multiple generations. Most of the compositional variability observed in these tourmalines involves Al and Fe with lesser involvement of Mg and the alkalis. This could reflect the changing 20

21 nature of the mineralizing fluid as it interacts with the host rock, becoming increasingly oxidized and Fe-rich. Zoning: Tourmaline chemical composition readily responds to changes in its chemical environment. According to several authors (e.g., Taylor and Slack, 1984; Dutrow et al., 1999; Henry and Dutrow, 1996), a large, external influx of boron is commonly accompanied by the generation of considerable amounts of unzoned, weakly zoned, oscillatory zoned, and/or complexly zoned tourmaline. The fine, oscillating chemical zoning of a wide range of cations of varying charge and density recorded in tourmaline suggests that cation diffusion in tourmaline is slow (Palmer et al., 1992). Zoning in Copiapó tourmaline ranges from concentric to discordant. The complexity of zoning present in most tourmaline is indicative of open-system behavior, with possible fluctuations in chemistry, pressure, and temperature. This is demonstrated by periods of distinct growth followed by dissolution and replacement by tourmaline with a different composition, as well as fine-scale zonation, which is indicative of rapid growth in a changing chemical environment (London and Manning, 1995). Dissolution can occur as a result of changing conditions and can be recognized by discordant zoning patterns with truncated compositional zones. Alkali variation: Alkali-deficient compositions can be obtained through the exchange vector Al[Na(Mg,Fe)] -1. These tourmalines are associated with alkali-deficient, highly acid, low-temperature environments (Rosenberg and Foit, 1979). Early, Al-rich cores reflect growth in Al-rich, highly oxidized, but alkali-poor environments. The formation of strongly Na-deficient tourmaline requires, for any given temperature, a very low concentration of Na in the fluid, and it is unlikely that this Na-poor phase is in equilibrium with other sodic phases such as albite. Metamorphic and synthesized tourmalines have shown an increase in Na with increasing temperature (Henry and Dutrow, 1996; von Goerne et al., 2001). Although such generalizations can be made, it is essential to compare tourmaline composition to the mineral assemblage, as Na content depends on several parameters (von Goerne et al., 2001). Overall, later generations of tourmaline have higher alkali (Na and Ca) contents that are commonly reflective of the alteration mineral assemblages. The alkali-defect exchange vector, Al[Na(Mg,Fe)] -1, accounts for the chemical difference observed in most cores. 21

22 Fe variation: Tourmaline compositions from various localities generally trend toward or fall within the Fe-rich hydrothermally altered rocks domain in the Al-Fe-Mg diagram of Henry and Guidotti (1985), suggesting an important host rock control of volcaniclastic and andesitic rocks. Tourmalines associated with metapelites and metavolcanic terranes tend to have intermediate schorl-dravite compositions, similar to the early generations of Copiapó tourmaline. However, there is no consistent correlation between the Fe/(Fe+Mg) ratio of tourmaline and a particular mineral assemblage (Power, 1968; Taylor and Slack, 1984). A remarkable feature noted in the majority of tourmalines from the Copiapó region is the inverse relationship between total Fe and Al. In other studies, the transition from Alrich to Fe-rich compositions has been linked to increasing distance from a magmatic source coincident with decreasing temperature and increasing differentiation of late magmatic fluids (Caverretta and Puxeddu, 1990) or from the breakdown of another Febearing phase. Whitney et al. (1995) also found that Fe contents tend to increase with increasing salinity in sulfur-free systems. Decreasing Al content in tourmaline can be linked to feruvite-uvite substitution, CaMg(NaAl) -1, and/or povondraite substitution, FeAl -1 (Bačík et al., 2008). However, subsequent decreases below 5.0 apfu Al necessitate the substitution of Fe 3+ for Al. There is a clear FeAl -1 substitution trend in Copiapó tourmalines that is indicative of a povondraitic component. An increase in the Fe 3+ /Fe total ratio may reflect the primary oxygen fugacity of the tourmaline-forming fluid, decreased Al activity, and/or higher overall Fe either from the fluid or increased fluid interaction with the host rock (i.e., Fe/(Fe+Mg) ratio of tourmaline reflects that of the bulk rock with progressive alteration; London and Manning, 1995; Gawęda et al., 2002). Galbraith et al. (2009) noted that the greater the fluid to rock ratio, the greater the evidence for hybridization of the source fluid and country rock compositions. As Fe is effectively transported as chloride complexes in hot, relatively acidic and saline solutions, this would also suggest that later evaporitic fluids were more effective at leaching Fe from the host rock. 22

23 Geologic controls The abundance of tourmaline in the Copiapó area is indicative of acidic, boronbearing hydrothermal fluids. The diversity of tourmaline compositions is a function of both host rock composition, reflected in later tourmaline generations, and the composition of the original fluid, which appears to have higher Al concentrations and lower overall alkali and FeMg contents. It also appears that these hydrothermal fluids became increasingly oxidized, as reflected by the increase in Fe 3+ in later tourmaline generations. The lack of significant amounts of other borosilicates indicates acidic fluids with high Al activities and sufficient ferromagnesian components to create tourmaline. The extensive sodic(-calcic) and potassic alteration in the Copiapó area suggests that alkalis are not a limiting factor; therefore, conditions for tourmaline formation would necessitate a driver towards more acidic compositions. An evaporitic source would provide the chloride ligands necessary for metal transport as well as the Na evident in hydrothermal alteration, while driving the fluids towards oxidized, relatively S-poor conditions (Barton and Johnson, 1996). The δ 11 B values in tourmaline, as with other compositional studies of fluid compositions in the Candelaria-Punta del Cobre deposits, suggest a mixed fluid source, suggesting both magmatic and sedimentary influences. Further evidence of highly-saline, oxidized fluids: Several interesting minerals indicative of highly saline solutions are scattered throughout the tourmaline-bearing Copiapó samples. Tourmalines from samples C2J-124 and C7B-003a, associated with sericitic and advanced argillic assemblages, respectively, contain inclusions of hypogene anhydrite 10s of microns across. These anhydrite crystals are not present within the matrix or in any other minerals. Sample C2B from the San Gregorio pluton contains chloro-potassichastingsite, a relatively uncommon Cl-rich amphibole (in this example containing over 5 wt percent Cl) that is indicative of alkali-chloride metasomatism (Mazdab, 2003). This mineral appears to be selectively replacing an early Al-, Fe-rich, blue tourmaline and may be coincident in time with the formation of Mg-, Ti-rich tourmaline within the same sample. In the Santos mine, marialitic scapolite is also present in association with tourmaline at depth, although it appears to postdate it. This scapolite end-member forms at temperatures around 400 C and is only stable in fluids that contain greater than 40 mol percent NaCl (Vanko and Bishop, 1982). 23

24 Comparison with other tourmaline-bearing hydrothermal systems Copiapó tourmaline compositions are compared to other occurrences from IOCG and other economic deposits, as well as to sediment-dominated systems, to establish any common crystal-chemical characteristics. Tourmaline compositions that plot along the oxydravite (a hypothetical sodic variation of magnesiofoitite)-povondraite solidsolution line defined by Henry et al. (2008) are considered a consequence of tourmaline formation in an oxidizing environment, possibly in the presence of high salinity aqueous fluids with reduced H 2 O activities (Henry et al., 2008). Povondraite is indicative of high oxygen fugacity during crystallization (Žáček et al., 1998), and the correlation to highly saline fluids may reflect more effective leaching and transport of elements from host lithologies (Henry et al., 2008). Magmatic-hydrothermal systems (Cu-Mo, Sn-W systems): Tourmaline compositions in porphyry and other magmatic-hydrothermal ore deposits may contain substantial Fe 3+ - rich and uvitic components (Fig. 15A). The existence of Fe 3+ -rich hydrothermal tourmalines in the majority of granitoid-related Cu Mo Au deposits suggests formation under relatively oxidizing conditions, in contrast to tourmalines from Sn W deposits which formed under more reducing conditions and lack significant ferric iron (Slack, 1996). In most porphyry Cu-Mo systems, tourmaline compositions range from Mg-rich to Fe-rich, and mineralization is associated with calc-alkaline to moderately felsic rocks. In rare cases, vein tourmaline associated with Sn mineralization appears to follow the povondraite-dravite trend as well. These tourmalines are found in mafic to ultramafic host rocks and thus have higher Fe contents (Mao, 1995). IOCG and similar systems: Tourmalines from the Larderello geothermal field in Italy are similarly Fe-rich (Cavaretta and Puxxedu, 1990). As expected, IOCG-related tourmaline show the same enrichment in Fe as Copiapó tourmaline (Fig. 15B). Out of all the IOCG deposits in the Carajás district, the Igarapé Bahia tourmalines appear to be the most chemically similar to the Copiapó tourmaline. IOCG mineralization is best developed in the breccias that lie between the mafic metavolcanic and metapyroclastic/metasedimentary units. Tourmalines from this locality are commonly Ferich dravites (Galarza et al., 2008). Boron isotopic work carried out by Xavier et al. (2008) points towards an evaporitic origin for these tourmalines, as reflected by high δ 11 B 24

25 values. Tourmaline boron isotope results from the Yerington and Humboldt IOCG deposits in Nevada also show evidence of external, non-magmatic fluid sources for mineralization, with average values ranging from -0.1 to +6.1 per mil (F.K. Mazdab, M.D. Barton, and R.L. Hervig, unpublished data) (Fig. 13) Metamorphic hydrothermal deposits: Tourmalines from various mesothermal Auquartz deposits display intermediate schorl-dravite to dravitic compositions (Fig. 15C) and appear to be geochemically similar to non-hydrothermal, metamorphic tourmaline. The composition of these tourmalines reflects both the chemistry of the mineralizing fluid and the composition of the host rock. Metamorphic hydrothermal tourmaline is most similar to early, Al-saturated generations of Copiapó tourmaline and do not show the same oyxdravite -povondraite trend. Stratabound deposits: Stratabound tourmalines, associated with massive sulfide and other deposits, are commonly dravitic in composition (Slack, 1982) (Fig. 15D). The compositional variation in these tourmalines is linked to the proportions of hydrothermal fluids and seawater, the water/rock ratio of the system, and the composition of the host rock (Slack, 1996). In non-massive sulfide deposits, Mg-rich tourmalines are believed to have formed through the circulation of a seawater component or a Mg-rich evaporitic brine. Tourmalines from all of these environments commonly plot along the schorldravite solid-solution and only rarely show the Fe-enrichment evident in Copiapó tourmaline. Sediment-dominated systems: Similarly Fe-rich tourmalines are also found in the Challenger Dome in the Gulf of Mexico (Henry et al., 1999) and in metamorphosed evaporites (Žáček et al., 1998; Henry et al., 2008) (Fig. 15E). Tourmalines associated with meta-evaporites are commonly sodic, magnesian, moderately to highly depleted in Al, and enriched in Fe 3+. These tourmalines typically fall along the oxydravite - povondraite join. Peng and Palmer (2002) found that although tourmalines may not be directly related to meta-evaporites, they may still exhibit an isotopic and fluid influence from meta-evaporitic sources. Tourmaline compositions that deviate from this trend likely reflect other superimposed reactions and a probable influx of reactive fluids. Tourmalines associated with the Barberton Greenstone in South Africa are very Al-enriched, even though they 25

26 are found in association with volcanics and evaporites, suggesting that the mineralizing fluid was far more aluminous than other evaporite-related areas (Byerly and Palmer, 1991). Notably, the Mg-rich tourmalines from the Fowler talc belt, New York (Ayuso and Brown, 1984), display similar characteristics to the distinctive early Mg-rich, Al-deficient generation in a sodic alteration assemblage observed in the Al-Fe-Mg ternary diagram and fall in field 8 of Figure 5, which is associated with metacarbonate or metapyroxenoid hosts. Ayuso and Brown (1984) hold that these tourmalines formed in specialized environments associated with evaporites, consequently having a profound effect on the bulk composition. Moreover, tourmalines from this locality also have heavy δ 11 B values (+13 ) Origin of boron isotope systematics Boron is highly mobile and easily leached and redistributed, especially under alkaline ph conditions (Dutrow et al., 1999). Tourmaline clearly shows a close relationship with certain distinct types of geologic settings, including quartz monzodiorites, IOCG deposits, and sodicly altered rocks but is not present or rare in other types of rocks and mineral deposits (e.g., porphyries, diorites, and some vein type IOCG occurrences). Although boron may be present in the latter settings, other conditions such as high alkalinity may have inhibited tourmaline formation (Morgan and London, 1989). In general, boron isotopic data show no correlation between mineral formation temperature, major-element composition, mineral assemblage, or age of deposit (Palmer and Slack, 1989; Swihart and Moore, 1989). Moreover, the mechanical and chemical stability of tourmaline prevents significant boron isotope fractionation, preserving its boron isotopic signature through time (Slack et al., 1989). However, boron isotope systematics are sensitive to ph, as it determines the relative abundance of B(OH) 3 and B(OH) 4 complexes in solution, and to the lithologic setting (Palmer and Slack, 1989). The dominant control on δ 11 B values in tourmalines is the composition of the boron source (Slack et al., 1989). As 10 B is preferentially incorporated into the solid phase during fluid-solid interactions, tourmaline δ 11 B values are systematically lower than the liquid from which it forms and are minimum estimates of the actual fluid composition 26

27 (Slack et al., 1989). Fractionation experiments predict a 5 per mil to 8 per mil difference between the mineralizing fluid and tourmaline (Palmer et al., 1992). As previously mentioned, the results of the boron isotopic work are biased towards coarse-grained tourmaline because of analytical constraints and thus may not reflect the complete picture of boron isotope variation in this district. The range of δ 11 B values of Copiapó tourmaline from -7.5 to +4.2 per mil overlaps a number of geological environments including granites, volcanics, metasediments, and non-marine evaporites. It is highly possible that the boron involved in tourmaline formation came from a mixture of these sources and maybe others that have more positive or negative δ 11 B values (see Fig. 13). Lighter δ 11 B values can be attained through mixing with a lighter source, such as granitic and volcanic host rocks, or through the loss of 11 B due to vapor phase separation (Smith and Yardley, 1996), whereas heavier values commonly necessitate a marine influence. Notably, all of the potential boron sources are present in the Copiapó region, underscoring the fact that data from even the best isotopic tracers require careful interpretation and commonly have non-unique interpretations. Fluid inclusion studies on quartz and ore minerals also support a mixed source for the origin of mineralizing fluids. The initial Sr isotope compositions of the fluid inclusions found in Candelaria are notably more radiogenic than the magmatic host rock, thus providing evidence for a non-magmatic source of Sr (Barton et al., 2005; Chiaradia et al., 2006). Cl isotopic work further indicates that the ore-forming fluid is a mixture of magmatic mantle-derived fluids (low radiogenic Sr and 37 Cl-rich) with a radiogenic Srrich and 37 Cl-poor crustal source such as a basinal brine (Chiaradia et al., 2006). Chiaradia et al. (2006) suggest that the magmatic fluids mixed with basinal brines or leached evaporites after exsolution from the magma, necessitating a mixed source model for the development of these deposits. These results could also be applied to the tourmaline boron isotope results, where the range from magmatic composition to more positive values may coincide with the influx of a non-marine brine. Conclusions Tourmalines provide a distinctive record of compositional variation in hydrothermal systems due to its wide range of stability and chemical variability. Tourmaline chemistry 27

28 can be used to provide a clearer understanding of ore-forming processes, related depositional environments, and the location of prospective exploration targets. Compositions of tourmalines from Copiapó reflect formation in oxidizing, acidic, highly saline environments and compositional trends toward Fe-rich compositions that reflect the compositions of the host lithologies. These results are consistent with relatively oxidized fluids that varied in composition through time, suggesting that tourmaline compositions distinguish different hydrothermal environments, even in the same area. The trend of tourmalines from Copiapó towards povondraitic compositions and their similarity to tourmalines from evaporitic sources attest to the highly saline and oxidizing nature of the mineralizing fluid. As reflected in δ 11 B values in Copiapó tourmaline, the boron necessary for tourmaline formation appears to have a mixed signature, with both evaporitic and magmatic fluid sources, an interpretation that is consistent with previous fluid inclusion studies in this area. This study is part of an ongoing study of IOCG mineralization in the Copiapó area and additional research may further constrain the enigmatic origin(s) of these deposits. Acknowledgments This study and associated field work has been financially supported by National Science Foundation (NSF) grant EAR , an NSF Graduate Student Research Fellowship, Science Foundation Arizona, and Freeport-McMoRan Inc. We would like to thank Ken Domanik for technical assistance with electron microprobe analyses and Mark Baker for his help creating an analytical method for boron isotope work on the isoprobe. Thanks also go to Doug Kreiner for providing some of the samples for this study, to David Cooke for tourmaline compositional data from the Río Blanco Cu-Mo deposit in Chile, and to Eric Seedorff and Bob Downs for their reviews of this manuscript. References Cited Arévalo, C. (1999). The Coastal Cordillera/Precordillera boundary in the Tierra Amarilla Area (27 20'-27 40'S/70 05'-70 20'W), Northern Chile, and the structural setting of the Candelaria Cu-Au ore deposit. Unpublished PhD, Kingston University, Kingston-upon-Thames, UK. 204 p. Ayuso, R. A., and Brown, C. E. (1984). Manganese-rich red tourmaline from the Fowler talc belt, New York. The Canadian Mineralogist, 2,

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37 Figure Captions Fig. 1. Regional geologic map showing the volcanic and sedimentary sequences intruded to the west by the Cretaceous Copiapó batholith (from Barton et al., 2005, modified from Arévalo, 1999). Fig. 2. Photomicrographs in plane polarized light of various tourmaline textures and mineral assemblages from the Copiapó region. A. DDH (Santos). Tourmaline clast with magnetite in a chlorite matrix surrounded by quartz. Tourmaline is intermediate schorl-dravite in composition. B. C6B-160b (Vinita Azul). Advanced argillic assemblage with acicular, foititic tourmaline, quartz, hematite, and andalusite. See text for further discussion. C. DDH (Santos). Potassic assemblage with dravitic tourmaline, biotite, magnetite, and quartz. D. C2B (San Gregorio S Granate). Fe-rich, blue tourmaline cluster with chloro-potassic-hastingsite replacing the bottom part of the tourmaline cluster. The host is coarse-grained albite. E. C2J-124 (Cerro Buitre Radio Tower). Sericitic assemblage with tourmaline, quartz, sericite, and zircon. Early violet gray, Al-rich tourmaline generation is rimmed by light brown tourmaline that is Na- and Mg-rich. F. C3B (Ojancos Viejo) Sodic assemblage with strongly zoned tourmaline, albite, limonite after pyrite, and fine-grained, indeterminable Al ± Fe sulfate. G. C6B-101b (Falla Ojancos Transito San Francisco). Granular, dravitic tourmaline intergrown with albite. H. C2J-334 (Cerro Buitre Radio Tower). Calcic assemblage with uvitic tourmaline and actinolite overprinting earlier potassic alteration (bioitite). Abbreviations represent: Act = actinolite, Alb = albite, And = andalusite, Bt = biotite, CPH = chloro-potassic-hastingsite, Hem = hematite, Mt = magnetite, Qtz = Quartz, SS = sheet silicate, Tour = tourmaline. Fig. 3. Alkali classification diagram after Hawthorne and Henry (1999). Analyses of tourmalines from various localities within the Copiapó district are plotted based on alteration assemblage type. The majority of tourmaline compositions fall within the alkali compositional group. Tourmalines from advanced argillic assemblages trend towards vacancy-rich compositions, whereas tourmalines of magmatic origin or in calcic assemblages trend towards more Ca-rich compositions. Fig. 4. Discrimination diagrams for naming tourmaline species, with analyses of tourmalines from the Copiapó district plotted according to type of alteration assemblage. A. Sodic and vacancy-rich end members. This discrimination diagram shows solid solutions between Na- and vacancy-rich end members with varying Mg/(Mg+Fe). B. Sodic and calcic end members. This discrimination diagram shows solid solutions between Na- and Ca-rich end members with varying Mg/(Mg+Fe). Fig. 5. Al-Fe-Mg compositional diagram of Henry and Guidotti (1985) with Copiapó tourmaline analyses plotted based on alteration assemblage type. Common end members are plotted for reference. Each field delineates various tourmaline-bearing rock types. Copiapó tourmalines cover the range of known tourmaline compositions outside of pegmatitic environments. The numbers correspond to distinct rock types: (1) Li-rich granitoid pegmatites and aplites; (2) Li-poor granitoids and their associated pegmatites and aplites; (3) Fe 3+ -rich quartz-tourmaline rocks (hydrothermally altered granites); (4) 37

38 Metapelites and metapsammites coexisting with an Al-saturating phase; (5) Metapelites and metapsammites not coexisting with an Al-saturating phase; (6) Fe 3+ -rich quartztourmaline rocks, calc-silicate rocks, and metapelites; (7) Low-Ca metaultramafics and Cr-, V-rich metasediments; and (8) Metacarbonates and meta-pyroxenites. Fig. 6. Ca-Fe-Mg compositional diagram of Henry and Guidotti (1985) with Copiapó tourmaline analyses. Common end members are plotted for reference. Each field delineates distinct tourmaline-bearing rock types. The majority of tourmaline analyses falls within field 4. The numbered fields correspond to: (1) Li-rich granitoid pegmatites and aplites; (2) Li-poor granitoids and their associated pegmatites and aplites; (3) Ca-rich metapelites, metapsammites, and calc-silicate rocks; (4) Ca-poor metapelites, metapsammites, and quartz-tourmaline rocks; (5) metacarbonates; (6) metaultramafics. Fig. 7. Potential exchange vectors. These figures denote the complexity of tourmaline compositional trends in this area. As discussed in the text, Copiapó tourmalines are largely Al-deficient, indicating formation in Al-undersaturated conditions. A. Al vs. Fe tot. Note the negative correlation between Fe and Al. B. Mg vs. Fe graph. The bold line on the graph indicates the FeMg -1 exchange vector. As is apparent from this graph, the majority of tourmaline compositions do not fall along the schorl-dravite solid-solution line. Only tourmalines from potassic assemblages show any correlation. C. Al vs. Na graph. Tourmalines from sodic-calcic and some sericitic assemblages plot along the NaAl(CaMg) -1 exchange vector, whereas tourmaline with greater than 6 apfu Al plot along the NaMg( Al) -1 exchange vector. Tourmalines from sodicly and some sericitically altered rocks appear to be largely dominated by the FeAl -1 exchange vector. D. Al+Ca vs Fe graph after Henry et al. (2008). This graph reinforces the negative correlation between Al and Fe along the exchange vector FeAl -1. E. Al vs Ca graph. There is a dominant negative correlation between Al and Ca for magmatic tourmaline, as well as for tourmaline from sodic-calcic, potassic, and some sericitic assemblages. Sodic and advanced argillic tourmalines show no correlation. F. X-site-vacancy vs. Ca graph. There is little correlation between X-site vacancies and overall Ca content. Fig. 8. Al vs X-site vacancy graph showing that X-site vacancies only begin to take a role in tourmaline chemical variations at Al contents greater than 5.5 apfu. Fig. 9. Photomicrograph in plane-polarized light of various stages of tourmaline growth in hydrous advanced argillic assemblage (sample C7B-003a from Jesus Maria). The tourmaline core tends to be Mg-rich compared to the later comb-like overgrowth that is more aluminous. Mineral abbreviations: Cal = calcite, Mar = hematite after magnetite ( martite ), Tour = tourmaline. Fig. 10. Backscattered electron image of a complexly zoned tourmaline sample from San Gregorio (sample C2B-655). This tourmaline shows an early, embayed and slightly fractured, Fe-rich generation overgrown by an aluminous rim that is later replaced by a complexly zoned reaction rim with numerous inclusions. Mineral abbreviations: Cal = calcite, Mt = magnetite, Tour = tourmaline. 38

39 Fig. 11. Backscattered electron image and X-ray maps for Mg, Na, Fe, Ca, and Ti of tourmaline in Santos sample DDH , showing micron-scale chemical variations and coalescence phenomena (subrounded cores overgrown by sub- to euhedral rims) (Pesquera et al., 2005), most evident in the Ca X-ray map. Numbers on the lower left hand corner of each X-ray map represents the approximate range of values for each element in the tourmaline cluster. Fig. 12. Backscattered electron image and X-ray maps of a complexly zoned, blue tourmaline (sample C2B from San Gregorio). Numbers on the lower left hand corner of each X-ray map represents the approximate range of values for each element in the tourmaline cluster. In general, early, Mg-rich tourmalines are overgrown by more aluminous, Ca-deficient, and slightly more Fe-rich compositions, which is consistent with the later superposition of sericitic alteration on an earlier igneous (roughly potassic) assemblage. Fig. 13. Histograms of δ 11 B values of tourmaline from Copiapó and other mineralized locations along with δ 11 B values of different geological environments. A. Histogram with the δ 11 B values of select samples of Copiapó tourmaline. Error bar indicates an uncertainty of ± 2 per mil. B. Histogram of unpublished boron isotope data of F. Mazdab, M. Barton, and R. Hervig. This dataset contains boron isotope data from IOCG deposits (Cloncurry district, Cortez Mountains, El Romeral, Humboldt, Yerington), sodicly altered rocks (Ajo, Superior-Engels, Granite Mountain), and porphyry-style occurrences (i.e., Cananea, Copper Creek, Nacozari de Garcia, Zaaiplaats), as well as miscellaneous tourmaline occurrences in kyanite ore (Cargo Muchacho-Vitrefax Hill) and sedimentary rocks (Belt Supergroup, Boyer Ranch). The standard deviation is about ±1.1 per mil. C. Compilation of δ 11 B values for different geological environments modified from Palmer and Swihart (1996). The gray box indicates the range of values covered by Copiapó tourmaline. Fig. 14. Mineral phase diagrams showing the topology of magnesiofoitite-dravite and dravite-uvite solid solution reactions at 400 ºC and 500 bars. These diagrams assume quartz and water saturation, with tourmaline stable throughout. Al is treated as an immobile element. Thermodynamic data was acquired using the SUPCRT92 program of Johnson et al. (1992). A representative reaction between tourmaline end members in the albite field is presented as an example of the effect of increasing log(k) (i.e., activity ratios of key tourmaline end members). The blue line represents the topology of pure end-member reactions between tourmaline end members. The green line represents an increase in log(k) by one log unit, and the orange line represents a decrease by one log unit. A. Magnesiofoitite to dravite reaction. This reaction is not dependent on the fugacity of oxygen. Na content increases to the right and Mg increases to the top. The stability fields move up and to the right with increasing dravite activity and down and to the left with increasing magnesiofoitite coincident with increasing acidity. B. Dravite to uvite reaction. The line on this diagram represents the topology of pure end-member reactions between dravite and uvite. Log(Ca/H 2 ) is held constant at 2. The stability fields move up and to the left with increasing dravite activity and down and to the left with increasing 39

40 uvite activity. Mineral abbreviations: Alb = albite, Dr = dravite, MgF = magnesiofoitite, Qtz = quartz, Uv = uvite. Fig. 15. Al-Fe 75 -Mg 75 diagrams after Henry et al. (2008) comparing Copiapó tourmaline to tourmaline from other hydrothermal systems, both mineralized and otherwise. Tourmalines from other IOCG, Cu-Mo, and sediment-dominated systems appear to show similar chemical variations to Copiapó tourmaline. A. Magmatic hydrothermal systems. References: Au (green): Koval et al. (1991), Golani et al. (2002); Cu-Mo (purple): Koval et al. (1991), Lynch and Ortega (1997), Frikken (2003), Yavuz et al. (1999); Pb-Zn-Cu (pink): Yavuz et al. (2008); Sn (blue): Mao (1995); Sn-Cu (red): London and Manning (1995), Mlynarczyk and Williams-Jones (2006); Sn-W (light blue): Pirajno and Smithies (1992); W (yellow): Clarke et al. (1989); Hydrothermal breccias and veins (peach): Dini et al. (2008), Demirel et al. (2009); Yavuz et al. (2008). B. Tourmaline from IOCG and related systems. Purple: Cavarretta and Puxeddu (1990); Yellow: Frietsch et al (1997); Orange: Xavier et al. (2008). C. Metamorphic tourmaline associated with Au-quartz veins. Brown: Garda et al. (2009); Blue: King and Kerrich (1989). D. Stratabound tourmaline occurrences. Green: Jiang et al. (1995); Red: Jiang et al. (1997b); Light blue: Mao (1995); Yellow: Plimer (1986); Orange: Plimer and Lees (1988); Purple: Raith (1988); Pink: Slack and Coad (1989); Blue: Taylor and Slack (1984). Abbreviations: MSD = massive sulfide deposits (includes both SEDEX and VMS deposits), SEDEX = sedimentary exhalative, VMS = volcanogenic massive sulfide. E. Sedimentary (evaporite-related) tourmaline occurrences. Dark red: Ayuso and Brown (1984). This tourmaline is found in a Fowler talc belt in New York and is unusually Mn rich. Light blue: Bačík et al. (2008). Tourmalines from this study are reworked tourmalinites now in conglomerates with carbonates and clastic sedimentary rocks with subordinate volcanics. Late tourmaline contains ferric iron and is Al-deficient. Purple: Henry et al. (1999). This tourmaline was found in the cap rock of a salt dome and is believed to have formed diagenetically within the bedded salt sequence that ultimately formed diapirs. Red: Henry et al. (2008). These tourmalines are found in meta-evaporite, non-marine sequences in Namibia and are moderately to highly deficient in Al and enriched in ferric iron. Green: Jiang et al. (1997a). Tourmalines from this study are found in meta-evaporite sequences in association with borate deposits. There is evidence for both non-marine and marine influences on the chemistry of the Houxianyu borate deposit, China. Blue: Peng and Palmer (2002). Tourmaline is found in meta-evaporite sequences enclosed by volcanic tuffs. Orange: Žáček et al. (1998). Tourmaline occurs in the meta-evaporite Locotol Breccia. Along with other silicate minerals, tourmaline formed along the reaction rim between dikes of highly alkaline volcanic rocks and B-rich evaporites. F. Copiapó tourmaline. Note how tourmaline from this area covers a wide range of compositions and displays the same compositional trends as tourmalines from highly saline environments. Tables Table 1. Tourmaline end members after Hawthorne and Henry (1999) Table 2. Exchange vectors after Henry and Guidotti (1985) and Dutrow and Henry (2000) Table 3. Tourmaline-bearing mineral assemblages 40

41 Table 4. Representative tourmaline analyses of Copiapó samples Table 5. Pleochroic characteristics of and zoning in Copiapó tourmaline Table 6. Boron isotope results of Copiapó tourmaline 41

42 Figure 1

43 Figure 2 A Tour B Tour Mt Chl Hem And Qtz Qtz 100 μm 100 μm C D Qtz Tour Mt Alb Bt 100 μm Tour CPH 100 μm E F Tour Lim Tour Alb Tour Qtz 100 μm AFS 100 μm G H Bt Alb Alb Tour Act Tour 100 μm 100 μm

44 Figure 3

45 Figure 4A

46 Figure 4B

47 Figure 5

48 Figure 6

49 Figure 7A

50 Figure 7B

51 Figure 7C

52 Figure 7D

53 Figure 7E

54 Figure 7F

55 Figure 8

56 Figure 9 Tour Cal Mar Tour 100 μm

57 Figure 10 Tour Mt Tour 100 μm Tour Cal

58 Figure 11 BSE Mg 20 μm Na Fe wt% High wt% wt% Low Ca Ti wt% wt%

59 Figure 12 BSE Al High 200 μm Mg Fe wt% Low Na wt% Ca wt% wt% wt%

60 Figure 13 A B marine brines non-marine brine C seawater marine evaporites non-marine evaporites limestones and dolomites magmatic fresh island arc volcanics granite and rhyolite fresh mantle-derived rocks δ 11 B