Mineral assemblages of the Francisco I. Madero Zn–Cu–Pb–(Ag) deposit, Zacatecas, Mexico: Implications for ore deposit genesis

Ore Geology Reviews 35 (2009) 423–435

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Mineral assemblages of the Francisco I. Madero Zn–Cu–Pb–(Ag) deposit, Zacatecas, Mexico: Implications for ore deposit genesis Carles Canet a,⁎, Antoni Camprubí b, Eduardo González-Partida c, Carlos Linares a, Pura Alfonso d, Fernando Piñeiro-Fernández e, Rosa María Prol-Ledesma a a

Instituto de Geofísica, Universidad Nacional Autónoma de México, Ciudad Universitaria, Delegación Coyoacán, 04510 México D.F., Mexico Departamento de Geoquímica, Instituto de Geología, Universidad Nacional Autónoma de México, Ciudad Universitaria, Delegación Coyoacán, 04510 México D.F., Mexico Centro de Geociencias, Universidad Nacional Autónoma de México, Campus Juriquilla, Boulevard Juriquilla 3001, 76230 Santiago de Querétaro, Qro., Mexico d Departament d'Enginyeria Minera i Recursos Minerals, Universitat Politècnica de Catalunya, Avinguda Bases de Manresa 61-73, 08242 Manresa, Catalunya, Spain e Industrias Peñoles, S.A. de C.V., carretera a Francisco I. Madero n°1, Cieneguillas, 98170 Zacatecas, Mexico b c

a r t i c l e

i n f o

Article history: Received 31 May 2008 Received in revised form 12 February 2009 Accepted 13 February 2009 Available online 24 February 2009 Keywords: Skarn Massive sulfides Geothermobarometry Mineral chemistry Guerrero Terrane North American Cordillera

a b s t r a c t The Francisco I. Madero deposit, central Mexico, occurs in the Mesozoic Guerrero Terrane, which hosts many ore deposits, both Cretaceous (volcanogenic massive sulfides) and Tertiary (epithermal and skarn deposits). It is hosted by a 600 m-thick calcareous-pelitic unit, of Lower Cretaceous age, crosscut by porphyritic dikes that strike NW–SE. A thick felsic volcanic Tertiary sequence, consisting of andesites and rhyolitic ignimbrites, unconformably overlies the Cretaceous series. At the base, the mineralization consists of several mantos developed within calcareous beds. They are dominantly composed of sphalerite, pyrrhotite and pyrite with minor chalcopyrite, arsenopyrite and galena. At the top of the orebody, there are calcic skarns formed through prograde and retrograde stages. The resulting mineral assemblages are rich in manganoan hedenbergite (Hd75–28Di40–4Jh40–20), andraditic garnets (Adr100–62Grs38–0), epidote (Ep95–36Czo60–5Pie8–0), chamosite, calcite and quartz. The temperature of ore deposition, estimated by chlorite and arsenopyrite geothermometry, ranges from 243° to 277 °C and from 300° to 340 °C, respectively. The pressure estimated from sphalerite geobarometry averages 2.1 kbar. This value corresponds to a moderately deep skarn and agrees with the high Cu content of the deposit. Paragenesis, P–T conditions and geological characteristics are compatible with a distal, dike-related, Zn skarn deposit. Its style of mineralization is similar to that of many high-temperature carbonate replacement skarn deposits in the Southern Cordillera. © 2009 Elsevier B.V. All rights reserved.

1. Introduction The Francisco I. Madero (FIM) Zn–Cu–Pb–(Ag) deposit is located ~ 20 km west of the city of Zacatecas, in central Mexico (Fig. 1). It was explored between 1976 and 1994 by the Mexican Geological Survey (formerly Consejo de Recursos Minerales), using geophysical methods and drilling. The deposit was acquired by Servicios Industriales Peñoles in 1994, which performed 130,000 m of diamond drilling. Mining operations started in 2001 with a 7000 t/day processing plant. Historical production up to 2005 is 9.6 Mt of ore containing 4.74 MOz silver, 30.28 kt of lead and 309.7 kt of zinc (González and López-Soto, in press). The FIM deposit is an ore deposit whose genesis has been controversial since its discovery. The orebodies are stratabound, are hosted by Mesozoic marine sedimentary rocks deposited in a back-arc environment, and the only intrusive rocks observable near the ores are a few Tertiary, post-Laramide dikes. For that reason, syngenetic submarine exhalative models, either sedimentary exhalative (SEDEX) ⁎ Corresponding author. Tel.: +52 55 56224133. E-mail address: ccanet@geofisica.unam.mx (C. Canet). 0169-1368/$ – see front matter © 2009 Elsevier B.V. All rights reserved. doi:10.1016/j.oregeorev.2009.02.004

or volcanogenic (VMS), have been suggested (Gómez-Caballero, 1986; Clark, 1999; Góngora-Flemate, 2001; Miranda-Gasca, 2003; González and López-Soto, in press). On the other hand, the predominance of calc-silicate assemblages and replacement textures that developed selectively along limestone beds, and the absence of exhalites or underlying structures attributable to feeder zones, suggest a manto or distal skarn model (Caddey, 2003). In addition, lead isotope compositions of the FIM ores suggest that this deposit is Tertiary in age and, thus, epigenetic and related to the later continental arc magmatism (Mortensen et al., 2008). The aim of this study is to elucidate the nature and ore genesis of the FIM deposit through the detailed study of its paragenesis and mineral chemistry.

2. Geological setting The FIM deposit is located in the Mesa Central physiographic province (Central Plateau or Mexican “Altiplano”) in central Mexico, which contains several economically and historically important mining districts, such as Zacatecas, Fresnillo, Guanajuato and Real de Catorce (Clark et al., 1982; Nieto-Samaniego et al., 2007). This deposit

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Fig. 1. Location and detailed geological map of the Francisco I. Madero deposit. The schematic stratigraphic section (right) is based in González and López-Soto (in press). Circles A, B and C indicate the location of the geological cross-sections shown in Fig. 2.

is hosted in the Guerrero Composite Terrane, which is a unit product of complex subduction-related processes influenced by major translation and rifting during the Mesozoic along the western margin of Mexico (Campa and Coney, 1983; Centeno-García et al., 2008). The Guerrero Composite Terrane consists mainly of metavolcanic-sedimentary sequences and is partially covered by the Tertiary felsic volcanics of the Sierra Madre Occidental. The base of the host sequence of the FIM deposit is a Mesozoic metapelitic unit composed of shales, meta-siltstones and -subarkoses (Fig. 1). This unit is probably Upper Jurassic to Lower Cretaceous in age (González and López-Soto, in press), although no radiometric or paleontological dating is available. It is conformably overlain by an up to 600 m thick Lower Cretaceous (González and López-Soto, in press) calcareous-pelitic series. These rocks are fine-grained and change progressively from black shales at the base to micritic limestones at

the top; the latter host the ores. Up to ~ 300 m thick volcanosedimentary sequence is deposited on the above rocks. It consists of basaltic pillow lavas and submarine tuffs interlayered with sandstones and shales. K–Ar dating of the volcanic rocks yielded an Upper Cretaceous age (94 Ma; González and López-Soto, in press). A small gabbroic stock, of probably Cretaceous age, intruded the Mesozoic sequence (Fig. 1). The Mesozoic sequence was deformed during the Laramide orogeny (Late Cretaceous to Paleocene) and metamorphosed under lower greenschist facies conditions (e.g., Salinas-Prieto et al., 2000). During the Late Cretaceous, a compression event produced the collision of several volcanic arcs, included those that formed the Guerrero Composite Terrane (Tardy et al., 1994; Centeno-García et al., 2008). In the FIM district, the Laramide compression led to the formation of large, open NW folds, with both limbs dipping about 20°.

Fig. 2. Schematic geological section of the Francisco I. Madero deposit (location shown on Fig. 1).

C. Canet et al. / Ore Geology Reviews 35 (2009) 423–435

Subsequently, an ENE–WSW post-Laramide extension event produced normal faults and developed horst and graben systems (González and López-Soto, in press). A N1000 m thick felsic volcanic Tertiary sequence associated with the Sierra Madre Occidental volcanic province unconformably overlies the Mesozoic series. It consists of andesites at the base and rhyolitic ignimbrites at the top, with interlayered beds of continental volcanic conglomerates. K–Ar dating of the earliest rocks of such sequence in the vicinities of the FIM district yielded an Eocene age (42 Ma; González and López-Soto, in press). A suite of meter-thick porphyritic dikes, associated with the Tertiary volcanics and ranging in composition from diorite to tonalite and granite, crosscuts the Mesozoic sequence. Such dikes mostly strike NW–SE, were emplaced along post-Laramide normal faults, and produced an intense marmorization of the Mesozoic limestone beds. Additionally, minor granitoid stocks crop out locally in the FIM district. The Guerrero Composite Terrane hosts a large number of Cretaceous and Tertiary ore deposits (Miranda-Gasca, 2000). Tertiary deposits include precious- and base-metal epithermal deposits,

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porphyry-copper deposits and skarns, whereas Cretaceous deposits are mostly of VMS type (Clark, 1999; Miranda-Gasca, 2000). 3. Structure of the orebodies The FIM deposit extends over an area of ~6 km2, at depths between 60 and 550 m below the present surface. It consists of several roughly stratiform orebodies hosted by a Mesozoic calcareous-pelitic series (Fig. 1). The thickness of the mineralization including non-economic portions changes abruptly from a few tens of meters up to ~ 170 m. 3 to 60 m thick individual ore lenses occur within calcareous beds. As a general feature, the mineralized unit is a large dome-shaped structure whose flanks are cut by NW normal faults (Fig. 2). Thus, the marginal portions of the mineralization were downthrown. Many of such faults host porphyritic dikes. Paterson (1995) found magnetic and gravimetric anomalies suggesting the presence of a deep magmatic intrusion. At its base, the mineralization consists of sulfide-rich (60 to 75 modal % sulfides) mantos, with major pyrite and pyrrhotite, and minor

Fig. 3. Photographs showing representative textures from the Francisco I. Madero deposit. (A) Photograph showing the irregular, replacive contact between a banded ore composed mainly of fine-grained pyrrhotite, pyrite and sphalerite, and a later coarse-grained massive ore mostly composed by pyrrhotite, sphalerite and galena. Slabbed core samples: (B) banded ore, with fine-grained pyrrhotite and pyrite; (C) massive ore, composed by coarse-grained sphalerite, galena, and pyrrhotite, accompanied by calcite and minor chlorite; (D) banded calc-silicate assemblage with calcic garnet (dark), and quartz with calcite and chlorite (light); (E) irregular calc-silicate assemblage with calcic clinopyroxene (brown), and chlorite with epidote (dark green), showing complex replacement textures; (F) banded chlorite-epidote assemblage, with coarse grained galena; (G) late quartz-calcite (white) brecciated vein, with fragments of black shale (dark) and banded chlorite-epidote with disseminated sulfides (grey). Abbreviations: Cal — calcite; Chl — chlorite; Ep — epidote; Gn — galena; Grt — garnet; Hd — clinopyroxenes (hedenbergite); Po — pyrrhotite; Py — pyrite; Q — quartz.

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sphalerite, galena and chalcopyrite, accompanied by quartz and pyrite. Locally, dm-thick Cu-(Ag-) and Pb- and Zn-rich lenses occur. Both banded and massive ores are observed (Fig. 3). Moderately silicified black shale beds with disseminated pyrite occur interlayered with the ores. The banded ore is mostly fine grained (up to 0.5 mm) and is very rich in pyrrhotite (~ 50 modal %) and pyrite (~20 modal %), with minor sphalerite and chalcopyrite (Fig. 3B). Towards the bottom of the banded ores, the pyrite content and grain size increase, and accessory arsenopyrite is present. The mm-scale layering of banded ores is defined by alternating sulfide- and chlorite-rich layers. Massive ores are coarse grained (up to 5 mm) and form irregular, replacive bodies composed of sphalerite and pyrrhotite (total ~ 50 modal %), with minor galena, chalcopyrite, calcite, chlorite, sericite, quartz and epidote (Fig. 3C). The upper part of the mineralized structure consists mainly of calcsilicate rich assemblages that strongly vary in texture, composition and grade. It contains, in order of abundance: (a) microcrystalline epidote-

chlorite, (b) almost-monomineralic macrocrystalline calcic clinopyroxene, and (c) banded calcic garnet-rich assemblages (Fig. 3D–F). Calcite, quartz, adularia, sericite, Ca-amphibole, stilpnomelane, titanite, rutile and apatite occur in variable amounts. Ore minerals are, in order of abundance, sphalerite, galena, marcasite, pyrite, chalcopyrite, arsenopyrite and magnetite; they attain up to 25 modal %. Calc-silicate rich assemblages locally develop banded and fold-like patterns. Calc-silicate- and sulfide-rich units show complex contact relationships, which include either sharp-replacive or gradual contacts. Decimeter-thick crustiform-banded undeformed veins and seldom breccias crosscut the above assemblages (Fig. 3G). These occur around felsic porphyritic dikes, and are lined by essentially calcite and quartz. Calcite mostly occurs as bladed crystals and quartz as chalcedony. Such veins also contain fluorite, chlorite and dolomite, and are locally enriched in Au and Ag (González and López-Soto, in press). In addition, Yta et al. (2003) reported for these veins a complex assemblage of Ag-, Bi-, Pb- and Cu-tellurides, sulfides and sulfosalts, including hessite, matildite, tetradymite, bismuthinite, tsumoite, aikinite and wittichenite.

Fig. 4. Episodes and sequence of crystallization in the Francisco I. Madero deposit.

C. Canet et al. / Ore Geology Reviews 35 (2009) 423–435 Table 1 Chemical composition and structural formulas of selected silicates (anhydrous) from the Francisco I Madero deposit (electron-microprobe data).

SiO2 Al2O3 TiO2 CaO Na2O K2O BaO MnO Fe2O3 FeO MgO ZnO Cr2O3 Total Si Al Ti Ca Na K Ba Mn2+ Fe3+ Fe2+ Mg Zn Cr

wt.%

Apfu

Grt

Grt

Px

Px

Px

Kfs

Kfs

#1

#2

#3

#4

#5

#6

#7

38.78 1.19 0.00 32.35 0.00 0.19 n.a. 0.17 25.37 2.60 0.30 n.a. 0.00 100.95 3.201 0.116 0.000 2.860 0.000 0.020 – 0.012 1.576 0.179 0.037 – 0.000

37.29 3.92 0.04 34.94 0.00 0.01 n.a. 0.18 25.73 0.00 0.03 n.a. 0.00 101.64 3.028 0.375 0.003 3.039 0.000 0.001 – 0.012 1.572 0.000 0.004 – 0.000

49.65 0.27 0.00 22.56 0.09 n.a. n.a. 11.35 – 16.26 0.72 0.04 n.a. 100.95 2.008 0.013 0.000 0.978 0.007 – – 0.389 – 0.550 0.043 0.001 –

50.56 0.04 0.00 22.23 0.02 n.a. n.a. 9.03 – 17.28 1.94 0.04 n.a. 101.14 2.022 0.002 0.000 0.952 0.001 – – 0.306 – 0.578 0.116 0.001 –

47.56 0.28 0.05 25.32 0.03 n.a. n.a. 10.94 – 8.90 6.86 n.a. n.a. 99.95 1.905 0.013 0.002 1.087 0.002 – – 0.371 – 0.298 0.410 – –

64.62 18.82 0.02 0.01 0.25 15.39 0.72 0.00 – 0.07 0.10 n.a. n.a. 100.06 2.986 1.025 0.001 0.000 0.023 0.907 0.013 0.000 – 0.003 0.007 – –

63.95 19.19 0.04 0.03 0.39 16.07 0.56 0.00 – 0.13 0.00 n.a. n.a. 100.39 2.959 1.047 0.001 0.001 0.035 0.948 0.010 0.000 – 0.005 0.000 – –

Apfu: atoms per formula unit. Grt: Ca garnets (structural formula based on 12 O); andradite. Px: Ca clinopyroxenes (structural formula based on 6 O); manganohedenbergite. Kfs: K feldspars (structural formula based on 8 O); adularia. n.a.: not analyzed.

However, they do not contribute significantly to the overall economic resource of the FIM deposit. At the base of the orebody, cm-scale, irregular, metamorphic quartz veins, often with pyrite crystals but devoid of other metallic minerals, are hosted in the basal metapelites, which in the vicinities of orebodies are altered and contain chlorite and epidote.

4. Paragenetic sequence Textural patterns suggest an epigenetic sequential mineral deposition caused by a complex, multistage hydrothermal event. The sequence of crystallization of the FIM mineralization is shown in Fig. 4. Based on paragenetic relationships, three main stages can be inferred, from early to late: sedimentation and regional metamorphism, epigenetic hydrothermal stage, and supergene alteration. Replacement textures suggest that the mineralization formed for the most part at the expense of the carbonate beds of the Lower Cretaceous sequence, in an epigenetic, hydrothermal metasomatic stage. This stage formed a complex sequential replacive deposition of minerals, and is characterized by three substages: (a) prograde, (b) retrograde, and (c) vein filling. During the prograde substage Caclinopyroxene and garnet, along with minor magnetite and titanite, formed. During the retrograde substage hydrous silicates (mainly epidote and chlorite), sulfides and quartz formed. Such mineral assemblage occurs interstitially with respect to the clinopyroxene crystals, and pseudomorphose garnet. Although early banded ores and a later replacive massive ores can be macroscopically differentiated, microscopic and compositional differences between them are only slight. Intergrowths and textures with curvilinear equilibrium grain boundaries prevail in both cases, indicating a relevant overlap in the crystallization of sulfides.

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The latest hydrothermal substage produced veins, breccias and void-lining quartz and carbonates, accompanied by minor chlorite and fluorite. Finally, a supergene assemblage, consisting of goethite and carbonates, is locally developed. 5. Analytical conditions Seventy-three samples were taken from drill cores and 15 were collected from the mine. The mineral associations have been studied in 30 polished thin sections using a standard petrographic microscope. Bulk mineralogy was confirmed by X-ray diffraction (XRD) analyses, performed on 14 samples, using a Philips 1400 diffractometer equipped with a Cu anode tube X-ray source and directing the collimated Cu Kα1,2 radiation (λ = 0.15405 nm) towards a randomly oriented sample. X-radiation was generated at 40 kV and 20 mA. Scans were recorded from 4° to 70° (2θ) with a step-scan of 0.02° and 2 s/ step. Images and quantitative analyses were obtained from 16 selected thin sections, with an electron probe microanalyzer (EPMA) JEOL JXA8900XR. The samples were examined in backscattered electron (BSE) mode. Wavelength dispersive spectrometry (WDS) analyses were carried out using following conditions: (a) for silicates: 20 keV, 20 nA, beam diameter of 1 μm, and a counting time of 30 s; the used standards were biotite (FeKα, SiKα, MnKα, KKα, TiKα, MgKα) and pyrope (MgKα), chlorite (AlKα) and albite (AlKα), sanidine (BaLα, KKα), plagioclase (CaKα), kaersutite (NaKα, CaKα, SiKα), diopside (CrKα), obsidian (FKα, ClKα), sphalerite (ZnKα) and ilmenite Table 2 Chemical composition and structural formulas of selected silicates (hydrated) from the Francisco I Madero deposit (electron-microprobe data). Chl

Chl

Stl

Stl

Ms

Ep

Ep

#1

#2

#3

#4

#5

#6

#7

#8

30.98 17.82 0.03 0.10 0.00 0.00 n.a. 0.28 – 14.26 – 24.65 0.03 0.00 0.87 11.80 100.82 3.043 2.063 0.002 0.011 0.000 0.000 – 0.024 – 1.171 – 3.609 0.002 0.000 0.271 7.729

39.45 9.52 0.14 0.00 0.03 2.80 0.00 4.21 – 30.32 – 7.26 0.12 0.03 c1.12 5.12 100.13 9.206 2.619 0.024 0.000 0.013 0.834 0.000 0.832 – 5.918 – 2.527 0.021 0.012 0.827 2.766

38.39 10.54 0.15 0.00 0.02 3.70 0.00 3.70 – 29.87 – 7.33 0.13 0.01 0.98 5.18 100.01 8.987 2.908 0.027 0.000 0.010 1.105 0.000 0.734 – 5.848 – 2.559 0.022 0.004 0.722 2.873

49.04 35.29 0.26 0.00 0.30 8.11 0.26 0.00 – 1.27 – 1.07 0.03 0.00 0.07 4.58 100.28 3.190 2.706 0.013 0.000 0.038 0.673 0.007 0.000 – 0.069 – 0.104 0.001 0.000 0.015 1.985

38.51 25.15 0.12 23.25 n.a. n.a. n.a. – 0.10 – 11.76 0.00 0.02 0.00 0.48 1.70 100.65 3.027 2.331 0.007 1.958 – – – – 0.006 – 0.696 0.000 0.001 0.001 0.118 0.890

38.79 27.71 0.02 22.83 n.a. n.a. n.a. – 0.69 – 8.21 0.01 0.03 0.02 0.29 1.80 100.13 3.017 2.540 0.001 1.903 – – – – 0.041 – 0.480 0.002 0.002 0.003 0.071 0.932

37.13 26.80 0.06 20.46 n.a. n.a. n.a. – 0.77 – 11.72 1.65 0.01 0.01 0.54 1.59 100.26 3.033 2.580 0.004 1.791 – – – – 0.048 – 0.721 0.200 0.001 0.002 0.140 0.869

SiO2 wt.% 25.32 Al2O3 20.60 TiO2 0.00 CaO 0.00 Na2O 0.00 K2O 0.00 BaO n.a. MnO 0.07 Mn2O3 – FeO 40.10 Fe2O3 – MgO 3.51 ZnO 0.00 Cl− 0.02 − F 0.82 H2O 10.39 Total 100.83 Si Apfu 2.815 Al 2.700 Ti 0.000 Ca 0.000 Na 0.000 K 0.000 Ba – Mn2+ 0.007 Mn3+ – Fe2+ 3.730 Fe3+ – Mg 0.582 Zn 0.000 Cl− 0.003 F− 0.289 OH−⁎ 7.708

Ep

Apfu: atoms per formula unit. Chl: chlorite group (structural formula based on 18 O,OH); analysis #1 chamosite, #2 clinochlore. Stl: stilpnomelane (structural formula based on 32 O,OH; 2.4 H2O). Ms: muscovite (structural formula based on 12 O,OH). Ep: epidote group (structural formula based on 14 O,OH); epidote. ⁎Only the hydroxyl group (combined water). n.a.: not analyzed.

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Table 3 Chemical composition and structural formulas of selected sulfides from the Francisco I. Madero deposit (electron-microprobe data).

S As Zn Co Ni Fe Sb Cu Cd Ag Sn Pb Bi Te Total S As Zn Co Ni Fe Sb Cu Cd Ag Sn Pb Bi Te

wt.%

atom %

Sp

Sp

Sp

Py

Po

Apy

#1

#2

#3

#4

#5

#6

34.20 0.01 54.68 0.00 0.00 10.32 0.00 n.a. n.a. n.a. n.a. n.a. n.a. n.a. 99.21 51.095 0.004 40.045 0.000 0.000 8.856 0.000 – – – – – – –

34.12 0.02 51.68 0.02 0.01 13.59 0.02 n.a. n.a. n.a. n.a. n.a. n.a. n.a. 99.45 50.713 0.015 37.648 0.016 0.005 11.596 0.007 – – – – – – –

33.77 0.00 54.35 0.02 0.00 10.53 0.00 0.00 0.09 n.a. n.a. n.a. n.a. n.a. 98.76 50.784 0.000 40.067 0.016 0.000 9.096 0.001 0.000 0.036 – – – – –

53.56 0.05 0.00 0.05 0.00 47.11 0.00 n.a. n.a. 0.00 0.00 n.a. n.a. n.a. 100.80 66.409 0.026 0.000 0.030 0.000 33.535 0.000 – – 0.000 0.000 – – –

39.15 0.01 0.03 0.05 0.00 61.16 0.00 0.03 0.00 n.a. n.a. n.a. n.a. n.a. 100.43 52.671 0.008 0.018 0.039 0.000 47.243 0.000 0.022 0.000 – – – – –

23.51 41.76 0.00 0.05 0.00 32.01 1.62 0.00 n.a. 0.02 0.00 0.05 0.00 0.00 99.01 39.033 29.674 0.000 0.049 0.000 30.516 0.708 0.000 – 0.007 0.000 0.013 0.000 0.000

Apy, arsenopyrite; Po, Pyrrhotite; Py, pyrite; Sp, Sphalerite. n.a.: not analyzed.

(TiKα, MnKα); (b) for carbonates: 15 KeV, 10 nA, beam diameter of 1 μm, and a counting time of 30 s; the used standards were kaersutite (KKα, NaKa, MgKa), dolomite (MnKa, SrLa, FeKa), calcite (CaKa), sphalerite (ZnKα) and plagioclase (BaLα); and (c) for sulfides: 20 KeV, 20 nA, beam diameter of 1 μm, and a counting time of 30 s; the used standards were skutterudite (NiKα, AsKα, FeKα, CoKα), sphalerite (ZnKα, SnLα), cuprite (CuKα), marcasite (SKα), stibnite (SbLα), Ag (AgLα) and crocoite (PbMα). The number of analyzed points is: chlorite, 75; muscovite, 16; stilpnomelane, 12; epidote, 87; Capyroxene, 24; Ca-garnet, 19; feldspar, 6; sphalerite, 104; arsenopyrite, 7; pyrrhotite, 15; carbonates, 56. 6. Mineral chemistry, geothermobarometry and paragenesis Ore-bearing mineral assemblages are calc-silicate rich and show replacement textures that suggest sequential mineral deposition. Details on mineral chemistry and textural features of the major mineral phases are provided in Tables 1–3, and in Figs. 5 and 6, respectively. 6.1. Calcic clinopyroxenes Calcic clinopyroxenes occur mostly in almost-monomineralic layers, as euhedral, randomly oriented prismatic crystals, a few tens of µm to 5 mm long. Quartz, calcite and minor epidote and sulfides (sphalerite, galena and pyrite) are interstitial to clinopyroxene crystals (Fig. 5C). Reddish weathering patinas of iron oxyhydroxides occur along cleavage and cracks in clinopyroxene crystals. These belong to the diopside-hedenbergite series (Hd75–28Di40–4Jh40–20; Table 1). As a distinctive feature, most of the analyzed clinopyroxenes have high contents of Mn (ranging from 5.76 to 11.35 wt.% MnO) and can be classified as manganoan hedenbergite (Fig. 7; Morimoto et al., 1988).

The highest Mn values were obtained from microcrystalline clinopyroxenes occurring in upper part of the mineralized structure, nearby the calc-silicate-limestone contact. The Zn content is up to 0.4 wt.% ZnO. 6.2. Calcic garnets Calcic garnets develop in roughly banded aggregates, associated with quartz, calcite, chlorite, muscovite, pyrite, and minor epidote, hedenbergite, chalcopyrite, titanite, rutile (altered to leucoxene) and apatite. Garnets constitute ~ 15 modal % of the total assemblage, and occur as up to 2 mm subhedral crystals (Fig. 6D). They are partially replaced by an assemblage of quartz, chlorite and sulfides, and are crosscut by calcite veinlets. They belong to the grossular-andradite series (Adr100–62Grs38–0; Fig. 8), and their Mn content is very low (b0.5 wt.% MnO; Table 1). Their chemical composition is very variable, as reflected in their concentric zoning, where crystal rims are richer than cores in the grossular end-member and in Mn (Fig. 9A). 6.3. Epidote Epidote is a major component of the mineralization as it occurs in all the calc-silicate assemblages, in the sulfide mantos and in the footwall altered metasedimentary rocks. Its main occurrence is as microcrystalline epidote–chlorite assemblages (up to 60 modal % of epidote) with quartz and calcite, and minor amphiboles of the actinolite–tremolite series, calcic garnet and clinopyroxene, titanite, stilpnomelane and sulfides (galena, sphalerite, pyrite, marcasite and chalcopyrite). In such associations epidote crystals are 20 to 400 µm long, randomly oriented subhedral prisms (Fig. 6C). In the sulfide mantos, epidote occurs as up to 2.5 mm euhedral prismatic crystals that are embedded into a sulfide groundmass. Such crystals show concentric zonation, reflecting changes in chemical composition (Fig. 9B). At a deposit scale the epidote end-member (Ep95–36Czo60–5Pie8–0) is dominant, and the Mn content is very low, between 0.10 and 1.46 wt.% Mn2O3 (Fig. 10; Table 2). The highest Mn contents in epidote correspond to crystals in the orebodies, whereas the Mn content in epidote from the basal metasediments is almost negligible (Fig. 10). 6.4. Sheet silicates Chlorite is very abundant, both in the mineralization and the footwall altered metapelites. In the calc-silicate rich units, it occurs as up to 500 μm-wide platelets that can form spherulitic aggregates, commonly in association with quartz and epidote (Fig. 6B). Chlorite may form pseudomorphs, probably after clinopyroxene. In the sulfiderich mantos, chlorite usually occurs disseminated within sulfide groundmasses of pyrrhotite, sphalerite and galena. In the crustiformbanded veins and breccias chlorite is scarce and forms aggregates of fine-grained crystals (up to 10 μm across). EMP analyses of chlorite crystals indicate that they are trioctahedral (Type “I” of Zane and Weiss, 1998) and belong to the clinochlorechamosite series (Fig. 11; Table 2). Chlorites from the orebodies are richer in the Fe2+ end-member (chamosite), whereas those from the footwall metapelites plot close to the Mg end-member (clinochlore). In both cases, the Mn content is almost negligible (Table 2). Octahedral vacancies per formula unit (□) based on Σ(O,OH) = 18 vary between 0.03 and 0.40, showing a progressive variation with depth. Thus, the lowest values correspond to the footwall metapelites (clinochlore), and the highest to the orebody (chamosite). Two generations of muscovite were identified. The first generation is formed by diagenetic to low-grade metamorphic processes, and is an essential mineral of the host metapelites. The second generation, of hydrothermal origin, occurs associated to chlorite in subordinate amounts. The Ba content in muscovite is negligible (Table 2); Cr was not detected.

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Fig. 5. Photomicrographs (plane polarized light) of mineral assemblages from the Francisco I. Madero deposit. Intergrowths and curvilinear equilibrium grain boundaries between sulfides (reflected light): (A) galena, pyrite and sphalerite, and (B) sphalerite, pyrrhotite and chalcopyrite; note that sphalerite lacks chalcopyrite blebs. (C) Replacement of calcic clinopyroxenes by an assemblage of bladed calcite, quartz and minor chlorite; (D) late vein filling of calcite and chalcedony (transmitted light, crossed nicols). Abbreviations: Cal — calcite; Chd — chalcedony quartz; Ccp — chalcopyrite; Gn — galena; Po — pyrrhotite; Px — calcic clinopyroxenes; Py — pyrite; Q — quartz; Sp — sphalerite.

Stilpnomelane, with an approximate structural formula K(Fe2+,Mg, Mn2+)9(Si,Al)12 (O,OH)32·nH2O, occurs as small platelets, up to 75 μm across, and is sparsely distributed in the chlorite-rich assemblages (Table 2; Fig. 6B).

6.5. Sulfides Pyrrhotite is the most common sulfide in the ore mantos, but is scarce in the calc-silicate rich assemblages. It forms mosaic aggregates with curvilinear equilibrium grain boundaries, in association with sphalerite and minor chalcopyrite. These aggregates contain minor interstitial, fine-grained (up to 30 μm) pyrite. The grain size of pyrrhotite crystals is up to 2 mm. The Fe contents in pyrrhotite crystals range from 46.9 to 47.8 at.% (Table 3). The average Fe per formula unit (based on 1 S) is 0.897, and thus its composition approaches Fe9 S10. According to Lusk et al. (1993) and Lynch and Mengel (1995), such composition is distinctive of the 5H polytype (hexagonal) pyrrhotite. Pyrite is widespread in the mineralized zone, in a large variety of textures and grain sizes. At the bottom of the orebody, the pyrite content is up to 75 modal %, occurring as coarse grained aggregates of euhedral crystals (up to 1 cm in diameter) with interstitial calcite, chlorite and quartz. These crystals contain numerous inclusions of pyrrhotite and sphalerite, and are accompanied by minor arsenopyrite (Fig. 6A). In contrast, in the banded ore pyrite occurs mostly as anhedral grains, up to 2 mm across, forming intergrowths and curvilinear equilibrium grain boundaries with sphalerite and galena (Fig. 5A). In the calc-silicate rich assemblages, pyrite occurs interstitially and as pseudomorphs after garnet (Fig. 6D). Arsenopyrite occurs in subordinate amounts in the coarse-grained pyrite basal aggregates and in the garnet-rich rocks. It forms up to 100 μm long euhedral, diamond-shaped, occasionally twinned crystals

(Fig. 6A). The arsenopyrite shows relatively low arsenic contents, from 29.0 to 29.7 at.%; the Co content is b0.1 at.% (Table 3). Sphalerite is the second metallic mineral in abundance (after pyrrhotite) in the deposit, and occurs all through the mineralized zone. It develops anhedral grains, up to few mm across, with curvilinear equilibrium grain boundaries, forming mosaic aggregates and intergrowths with galena, pyrrhotite, pyrite and seldom chalcopyrite (Fig. 5A, B). In general, sphalerite grains lack “chalcopyrite disease”. The analyzed Fe contents in sphalerite show a wide range of variation, between 15.3 and 24.6 mol% FeS (Fig. 12; Table 3). This variation meets up with the different paragenetic relationships. 6.6. Carbonates Carbonate minerals occur: (a) in the host sedimentary rocks; (b) interstitially in the calc-silicate rich assemblages; and (c) in late veins and breccias. In the calc-silicate assemblages, manganoan calcite (Ca0.91–0.82Mn0.09–0.16CO3) is abundant, and develops (i) anhedral, poikilitic patches up to 5 mm across, (ii) bladed crystals up to 1 mm in diameter, and (iii) microcrystalline pseudomorphs, in association with chlorite, probably after clinopyroxene. In the late veins, carbonates range from pure calcite to manganoan calcite, from 0.6 to 16.8 mol % MnCO3, and form mosaic textures associated with chalcedony (Fig. 5D). In addition, mm-sized euhedral dolomite crystals occur as late breccia cement. 6.7. Quartz In the sulfide mantos quartz generally forms up to 1 mm long subhedral to euhedral crystals, embedded in a sulfide groundmass. Such crystals are rich in inclusions of epidote and calcic amphiboles. In the calc-silicate assemblages quartz crystals occur in mosaic

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Fig. 6. SEM-BSE images of mineral assemblages in the Francisco I. Madero deposit. (A) euhedral arsenopyrite and pyrite crystals, and a late marcasite–siderite assemblage; (B) stilpnomelane associated to quartz, chlorite and epidote; (C) euhedral titanite crystals associated to epidote, with interstitial quartz, calcite and sphalerite; (D) pseudomorphic replacement of andradite by a sulfide assemblage (pyrite and chalcopyrite), calcite and quartz. Abbreviations: Adr — andradite; Apy — arsenopyrite; Cal — calcite; Ccp — chalcopyrite; Chl — chlorite; Ep — epidote; Gn — galena; Mrc — marcasite; Ms — muscovite; Po — pyrrhotite; Py — pyrite; Q — quartz; Sd — siderite; Sp — sphalerite; Stp — stilpnomelane; Ttn — titanite.

Fig. 7. Composition of calcic clinopyroxenes (manganohedenbergite) from the Francisco I. Madero deposit plotted on a johannsenite (Jh) — diopside (Di) — hedenbergite (Hd) diagram. Filled diamonds indicate microcrystalline (up to 50 μm) manganohedenbergite disseminated within an epidote-quartz assemblage; open diamonds indicate coarse-grained (up to 1 cm) almostmonomineralic manganohedenbergite. Dashed areas indicate the pyroxene composition range for Zn skarn deposits summarized by Meinert et al. (2005).

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Fig. 8. Composition of calcic garnets from the Francisco I. Madero deposit (pyralspite-grossular-andradite diagram). Dashed areas indicate the garnet composition range for Zn skarn deposits summarized by Meinert et al. (2005).

aggregates associated with chlorite, calcite and minor muscovite. Quartz also occurs in the late veins and breccias forming (a) early euhedral crystals normal to the wall rock, and (b) late chalcedony aggregates of up to 300 μm long fibrous crystals (Fig. 5D). 6.8. Geothermobarometry A wide range of composition and a non-stoichiometric behavior make the minerals of the chlorite group attractive for geothermometry (de Caritat et al., 1993). Many studies are based on an empirical observation stating that, in geothermal systems, SiIV and, hence, □VI decrease with increasing depth. Thus, equations relating temperature and chlorite composition have been deduced from microprobe chlorite analysis and microthermometric data from fluid inclusions (Cathelineau and Nieva, 1985; Kranidiotis and MacLean, 1987; Cathelineau, 1988). However, besides temperature, many other parameters such as fO2, pH and the Fe/(Fe + Mg) ratio in the host rock could affect chlorite composition (de Caritat et al., 1993). Although the empirical chlorite geothermometry must be interpreted with caution, it can be used satisfactorily in geologic settings comparable to those where the calibration was performed and for temperatures below 350 °C, since they consider, not explicitly, many thermodynamic variables (de Caritat et al., 1993). Using the empirical □-T plot for chlorites of Cathelineau and Nieva (1985), we obtained temperatures between 270° and 277 °C for chlorites at the footwall metapelites (hydrothermally altered) and the bottom of orebodies, and between 243° and 262 °C at the central and upper parts of orebodies. Arsenopyrite displays a significant solid solution range in the Fe– As–S system that is sensitive to temperature, which allows it to be used as a geothermometer (Kretschmar and Scott, 1976; Sharp et al., 1985). The estimated pressure effect on the composition of arsenopyrite in equilibrium with pyrite is irrelevant (Kretschmar and Scott, 1976). Arsenopyrite with b0.5 at.% Co can be used for geothermometry (Sharp et al., 1985). However, arsenic contents below 30.0 at.% As do not allow to use the fS2 vs. T equilibria established by Kretschmar and Scott (1976), which has been applied in the study of several

metamorphosed ore deposits (e.g., Lynch and Mengel, 1995; Lentz, 2002). In this study, petrographic and scanning electron microscope observations suggest that arsenopyrite formed in the retrograde hydrothermal substage of the mineralizing event, together with pyrite (Figs. 4 and 6). Therefore, a coprecipitation under equilibrium conditions can be assumed. Thus, we obtained, from the arsenopyrite–pyrite stability field of the pseudobinary T–X diagram of Kretschmar and Scott (1976), temperatures that range between 300° and 340 °C (Fig. 13). The Fe content in sphalerite is largely affected by pressure and, consequently, it can be used for geobarometry (Scott and Barnes, 1971; Hutchison and Scott, 1981). The sphalerite geobarometer can be applied with confidence in several skarn deposits, in view of their sulfide assemblages and range of temperature and pressure conditions (Shimizu and Shimazaki, 1981). To estimate the pressure from XFeS it is required equilibrium of sphalerite with pyrite and hexagonal pyrrhotite in order to buffer the sulfur fugacity. XFeS in sphalerite crystallized along the pyrrhotite–pyrite equilibrium ranges between 10 and 20% (Scott and Barnes, 1971). Sphalerite rich in chalcopyrite blebs is not suitable for geobarometry calculations (Hutchison and Scott, 1981). In this work, sphalerite crystals with XFeS N 20% are mostly in paragenesis with marcasite, which may be secondary; these were thus rejected for geobarometry purposes. On the other hand, sphalerite grains with XFeS b20% mostly coexist with pyrite and minor pyrrhotite (without marcasite; Fig. 12). Thus, they are suitable for pressure calculations. The estimated pressures from sphalerite within such assemblages range from 0.8 to 4.2 kbar, with average of 2.1 kbar, corresponding to a depth of ~8 km.

7. Discussion The ore-bearing rocks of the FIM deposit are composed largely by calc-silicates (manganoan hedenbergite, andraditic garnet, epidote, titanite), which are accompanied by other silicates (quartz, chlorite), carbonates (manganoan calcite), oxides (magnetite), and sulfides

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Fig. 9. Zoning patterns showing variations in Fe2O3, Al2O3 and Mn2O3 contents (electron microprobe data) for (A) calcic garnet, and (B) epidote crystals. SEM-BSE images (right) indicate the position of electron microprobe profiles.

(sphalerite, pyrrhotite, pyrite, marcasite, chalcopyrite, galena and arsenopyrite). According to Einaudi et al. (1981), Delgado et al. (1997) and Meinert et al. (2005) such mineralogy and the association of orebodies with limestones allow us to categorize the FIM deposit as a skarn. Skarn deposits form by metasomatism, through a succession of replacive mineralizing events triggered by a magmatic-hydrothermal system, generally including a prograde (high temperature) stage, and a subsequent retrograde (lower temperature) stage (e.g., Einaudi et al., 1981). In the FIM deposit, the succession of prograde and retrograde stages resulted in a large exoskarn characterized by overprinted and recrystallized paragenesis (Fig. 4). The prograde stage caused the extensive occurrence of manganoan hedenbergite and minor andraditic garnet in the FIM deposit. In calcic skarns of the Zn–Pb type (Einaudi et al., 1981), calcic clinopyroxenes predominate over garnets. Besides, the highest pyroxene to garnet ratios are found in skarns that are distal to the causal magmatic sources (Meinert et al., 2005). The high manganese contents in pyroxenes (Mn/ Fe ratio: 0.3–1.2) of the FIM deposit agree with the johannsenite-rich

composition of pyroxenes from most Zn–Pb skarns deposits (Einaudi et al.,1981; Nakano,1998). Nakano (1998) found in Japanese skarns, that the Mn/Fe ratio is N0.2 in Zn–Pb deposits, ranges between 0.2 and 0.1 in W deposits, and is b0.1 in Cu–Fe deposits. According to the same author, an association of high Mn/Fe pyroxenes and andraditic garnet indicates oxidizing conditions. Such conditions agree with the early occurrence of magnetite at the FIM deposit. Zoning in garnet crystals reflects oscillations in the composition of metasomatic fluids (Fig. 9A; Jamtveit et al., 1993). The FIM retrograde skarn assemblage is largely made up of epidote, chlorite (chamosite) and quartz, and is accompanied by an extensive deposition of base metal sulfides. Fe-rich members of the epidote-clinozoisite series are relevant minerals in retrograde assemblages of many Zn–Pb skarn deposits (Einaudi et al., 1981), and their composition generally depends on the composition of the protolith and the fO2 (Delgado et al., 1997). In addition, epidote and chlorite (clinochlore) occur in the footwall metapelites due to the hydrothermal alteration by metasomatic fluids. Thus, these rocks can be considered periskarns (as defined by Zharikov, 1970).

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Fig. 10. Composition of epidote from the Francisco I. Madero deposit. Filled diamonds indicate crystals from the orebody; open squares indicate crystals from the altered metapelites of the basal series. Abbreviations: Pie — piemontite; Czo — clinozoisite; Ep — epidote.

The temperature of formation of the retrograde skarn, estimated through the arsenopyrite and chlorite geothermometers, ranges from 300° to 340 °C, and from 243° to 277 °C, respectively. These temperatures may reflect a cooling path in the retrograde skarn stage, as arsenopyrite predates chlorite (Fig. 4). The arsenopyrite temperature range could represent the temperature for ore deposition. This temperature range agrees with that of sulfide-rich skarn deposits (250° to N500 °C; Megaw, 1998), although it is notably lower

than most data after fluid inclusion microthermometry in the prograde, pre-sulfide skarn stages (Einaudi et al., 1981, and references therein). The average pressure estimated from the sphalerite geobarometry is 2.1 kbar. Pressure estimates for skarn deposits range from 0.3 to 3 kbar, although pressure for most of Zn-rich skarn deposits range from 0.5 to 1.8 kbar (Einaudi et al., 1981, and references therein). According to Shimazaki (1975), the depth of emplacement of skarn

Fig. 11. Classification of chlorites from the studied deposit. Chlorite diagram after Zane and Weiss (1998). Filled diamonds indicate crystals from the orebody; open diamonds indicate crystals from the altered metapelites of the basal series.

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deposits determine their base metal contents. Thus, shallow skarns (P b 1 kbar) are typically Zn- and Pb-rich, whereas deep skarns (P N1 kbar) are Cu-rich. In view of that, the high pressure values estimated for the FIM deposit agree with the high copper contents that it bears. Considering the relationship and nature of causative magmatism, as well as the morphology of the orebodies, the FIM deposit is a distal, dike-related skarn deposit (Einaudi et al., 1981). According to Meinert et al. (2005) most Zn-rich skarns occur distally with respect to associated igneous rocks. In Japan, near 70% of the Zn–Pb skarn deposits occur in close proximity to dikes, but distal to the unobserved plutons (Nakano et al., 1990). According to Einaudi et al. (1981), these dikes act as pathways for metasomatic fluids that come from a deeper, cogenetic, plutonic body. Such conditions are in accordance with those observable at the FIM district. Megaw (1998) defined Carbonate Replacement Deposits (CRD), based on the abundant base-metal distal skarn deposits of the Southern Cordillera (northwestern Mexico and southwestern USA), as “epigenetic, intrusion-related, high-temperature (N250 °C), sulfidedominant Pb–Zn–Ag–Cu–Au-rich deposits that typically form lenses or elongate to elongate-tabular bodies referred to as mantos or chimneys”. The FIM deposit fulfills all geological and mineralogical characteristics of this particular kind of skarn deposit. These include: (a) relatively large deposits (N50 Mt for the largest CRD deposits); (b) dome-like, fault-controlled deposits; (c) mineralization associated with dikes; (d) lack of exposed plutons or stocks; (e) the occurrence of volcanic capping rocks contemporaneous with the intrusions; (f) orebodies mostly shaped as lenses or mantos; (g) textural evidence for replacement; (h) sequential formation of the deposits, including prograde and retrograde mineralization stages; (i) occurrence of late veins (typically rich in quartz and fluorite) adjacent to the dikes; and (j) the ore (sphalerite, galena, pyrrhotite, pyrite, chalcopyrite and arsenopyrite) and (k) gangue (dominantly pyroxene and garnet) mineralogy. The adscription of the FIM deposit to CRD-style skarn deposits significantly enlarges the prospective area for these in Mexico towards the south, thus increasing the regional potential for base metal exploration. Most of the CRD-style skarn deposits, as well as the FIM deposit, are blind or poorly exposed (Megaw, 1998). Therefore, the exploration

Fig. 13. Arsenopyrite geothermometry based on the pseudobinary T–X diagram of Kretschmar and Scott (1976). Abbreviations: Apy — arsenopyrite; As — native arsenic; L — liquid; Lo — löllingite; Po — pyrrhotite; Py — pyrite.

for such deposits is complex and should be supported by geophysical techniques (Megaw, 1998). 8. Conclusions The petrography and mineral chemistry of the mineral associations allowed us to properly characterize the deposit model of a controversial, and metallogenetically and economically important deposit. The FIM deposit is a distal, dike-related, Zn skarn akin to the Carbonate Replacement Deposits, a style of skarn mineralization that is relatively common in the Southern Cordillera of North America. Prograde and retrograde mineralizing stages produced overprinted and recrystallized paragenesis. The prograde stage caused the extensive occurrence of manganoan hedenbergite and minor andraditic garnet, whereas the retrograde assemblage consists essentially of epidote, chamosite and quartz, and is accompanied by an extensive deposition of base metal (Zn–Cu–Pb) sulfides, pyrrhotite, pyrite and minor arsenopyrite. Temperature of formation of the retrograde skarn, estimated from arsenopyrite and chlorite geothermometry, ranges from 300° to 340 °C, and from 243° to 277 °C, respectively. The average pressure estimated from the sphalerite geobarometry is 2.1 kbar, corresponding to a moderately deep skarn, and is in agreement with the high Cu contents in this deposit. The classification of the FIM deposit as a CRDstyle skarn may increase the expectative for regional exploration of base metal deposits. Acknowledgements

Fig. 12. Histogram showing mol% FeS in sphalerite (electron microprobe data) from the Francisco I. Madero deposit. Abbreviations: Cal — calcite; Chl — chlorite; Dol — dolomite; Ep — epidote; Gn — galena; Gt — garnet; Mrc — marcasite; Po — pyrrhotite; Py — pyrite; Q — quartz; Sp — sphalerite.

This project was sponsored by Industrias Peñoles, SA de CV. We are grateful to the IGCP Project 502 (2004–2008). The SEM-EDS and WDS analyses and BSE images were obtained at the Laboratorio Universitario de Petrología (LUP), and petrography studies were performed in the Laboratorio de Yacimientos Minerales, both located in the Instituto de Geofísica, Universidad Nacional Autónoma de México (UNAM). Rufino Lozano Santa Cruz and Teresa Pi Puig are thanked for the XRD analyses that were carried out in the Instituto de Geología (UNAM). Javier García-Fons, Lepoldo González (Industrias Peñoles, SA de CV) and J. Richard Kyle are thanked for their assistance and comments during

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