Paleozoic serpentinite-enclosed chromitites from Tehuitzingo (Acatlán Complex, southern Mexico): a petrological and mineralogical study

Journal of South American Earth Sciences 16 (2004) 649–666 www.elsevier.com/locate/jsames

Paleozoic serpentinite-enclosed chromitites from Tehuitzingo (Acatla´n Complex, southern Mexico): a petrological and mineralogical study J.A. Proenzaa,*, F. Ortega-Gutie´rrezb, A. Camprubı´c, J. Tritllac, M. Elı´as-Herrerab, M. Reyes-Salasb a

Departament de Cristal·lografia, Mineralogia i Dipo`sits Minerals, Facultat de Geologia, Universitat de Barcelona, C/Martı´ i Franque`s s/n, 08028 Barcelona, Spain b Instituto de Geologı´a, Universidad Nacional Auto´noma de Me´xico, 04510 Me´xico DF, Mexico c Centro de Geociencias, Universidad Nacional Auto´noma de Me´xico, Carretera Qro.-S.L.P. km 15.5, Campus UNAM-Juriquilla, 76230 Santiago de Quere´taro, Qro, Mexico Received 31 May 2003; accepted 31 December 2003

Abstract The serpentinites and associated chromitite bodies in Tehuitzingo (Acatla´n Complex, southern Mexico) are in close relationship with eclogitic rocks enclosed within a metasedimentary sequence, suggesting that the serpentinites, chromitites and eclogitic rocks underwent a common metamorphic history. Primary chromites from the chromitite bodies at Tehuitzingo are of refractory-grade (Al-rich) and have a chemical composition similar to that expected to be found in an ophiolitic environment. The chromite grains in chromitites and serpentinites are systematically altered to ‘ferritchromite’. The alteration trend is usually characterized by a decrease in the Al, Mg and Cr contents coupled by an increase in Fe3þ and Fe2þ. The Tehutizingo chromitites have low Platinum Group Elements (PGE) contents, ranging from 102 to 303 ppb. The chondrite-normalized PGE patterns are characterized by an enrichment in the Ir-subgroup elements (IPGE ¼ Os, Ir, Ru) relative to the Pd-subgroup elements (PPGE ¼ Rh, Pt, Pd). In addition, all chromitite samples display a negative slope from Ru to Pd [(Os þ Ir þ Ru)/(Pt þ Pd) ¼ 4.78 2 14.13]. These patterns, coupled with absolute PGE abundances, are typical of ophiolitic chromitites elsewhere. Moreover, all the analyzed samples exhibit chondrite-normalized PGE patterns similar to those found for non-metamorphosed ophiolitic chromitites. Thus, the PGE distribution patterns found in the Tehuitzingo chromitites have not been significantly affected by any subsequent Paleozoic high-pressure (eclogite facies) metamorphic event. The chondrite-normalized PGE patterns of the enclosing serpentinites also indicate that the PGE distribution in the residual mantle peridotites exposed in Tehuitzingo was unaffected by high-pressure metamorphism, or subsequent hydrothermal alteration since the serpentinites show a similar pattern to that of partially serpentinized peridotites present in mantle sequences of non-metamorphosed ophiolites. Our main conclusion is that the chromitites and serpentinites from Tehuizingo experienced no significant redistribution (or concentration) of PGE during the serpentinization process or the high-pressure metamorphic path, or during subsequent alteration processes. If any PGE mobilization occurred, it was restricted to individual chromitite bodies without changing the bulk-rock PGE composition. Our data suggest that the Tehuitzingo serpentinites and associated chromitites are a fragment of oceanic lithosphere formed in an arc/backarc environment, and represent an ophiolitic mantle sequence from a supra-subduction zone, the chemical composition of which remained essentially unchanged during the alteration and metamorphic events that affected the Acatla´n Complex. q 2004 Elsevier Ltd. All rights reserved. Keywords: Tehuitzingo; Mexico; Acatla´n Complex; Serpentinites; Chromitites; Platinum-group elements; Eclogite-facies metamorphism

1. Introduction * Corresponding author. E-mail address: [email protected] (J.A. Proenza). 0895-9811/$ - see front matter q 2004 Elsevier Ltd. All rights reserved. doi:10.1016/j.jsames.2003.12.003

Chromite composition is extensively used as a petrogenetic and geotectonic indicator (Irvine, 1967; Dick

650

J.A. Proenza et al. / Journal of South American Earth Sciences 16 (2004) 649–666

and Bullen, 1984; Arai, 1992). In the case of completely serpentinized ultramafic rocks, that preserve no primary silicate minerals, the composition of unaltered accessory chromite may provide useful petrogenetic information. Chromite is the only igneous mineral that is commonly preserved in serpentinite following metamorphism. The composition of chromite in chromitite bodies is also used as a petrogenetic and geotectonic indicator (e.g. Stowe, 1994), although, chemical modifications related to subsolidus reequilibration and metamorphic hydrothermal processes can significantly influence the primary high-T composition of Cr-spinel. However, Suita and Streider (1996) in a study of the metamorphic modifications of Cr-spinel (included chromite from massive stratiform and podiform chromitites) from different Brazilian mafic and ultramafic complexes, showed that, regardless of their metamorphic changes, the cores of Cr-spinel grains in massive chromitites preserved the chromite’s primary chemical composition. These Brazilian complexes were affected by varying degrees of metamorphism, ranging from greenschist to epidote-amphibolite, upper amphibolite and granulite facies. However,

there is very little information regarding changes in chromite composition during high-pressure metamorphism. Platinum-group elements (PGE) are also thought to be immobile during serpentinization processes (Groves and Keays, 1979; Prichard and Tarkian, 1988; Leblanc, 1991). However, significant redistribution of PGE (mainly Pdsubgroup elements or PPGE) during serpentinization or metamorphism has been reported in some ophiolites (e.g. Thalhammer et al., 1990; Malitch et al., 2002). Several lines of evidence also suggest PGE transport and deposition in hydrothermal solutions (e.g. Mountain and Wood, 1988); Au and Pt are more mobile than the other PGE during alteration processes, and Pd can be mobilized by hydrothermal fluids (Barnes et al., 1985). However, no extensive data are available concerning the mobilization of PGE from ophiolitic rocks during metamorphism. This paper focuses on the study of the petrological characteristics, mineral chemistry, and the distribution of PGE in chomitite bodies hosted in serpentinites at Tehuitzingo, southern Mexico (Fig. 1a). These chromitites and host serpentinites are associated with eclogitic rocks,

Fig. 1. (a) Location of the study area in southern Mexico. (b) Simplified geological map of the northern part of the Acatla´n Complex, showing the location of serpentinite in the Tehuitzingo area. Modified from Ortega-Gutie´rrez et al. (1999).

J.A. Proenza et al. / Journal of South American Earth Sciences 16 (2004) 649–666

providing an ideal setting for the study of possible changes in chromite composition, and the remobilization of PGE within them, during high-pressure metamorphism.

2. Geologic setting 2.1. The Acatla´n Complex The Acatla´n Complex is located in southern Mexico, and comprises a deformed assemblage of Paleozoic metasedimentary rocks, granitoid bodies, and metamorphosed mafic-ultramafic rocks (Fig. 1b). The stratigraphic and tectonostratigraphic units that compose this complex have been described in detail by Ortega-Gutie´rrez (1978, 1993); Ya´n˜ez et al. (1991); Weber et al. (1997); Ortega-Gutie´rrez et al. (1999); Malone et al. (2002), and Elı´as-Herrera and Ortega-Gutie´rrez (2002). According to Ortega-Gutie´rrez (1993), the Acatla´n Complex can be subdivided into two principal tectonic units or subgroups (Fig. 1b). The Petlalcingo subgroup (parautochthonous, lower plate) consists of a thick package of metasedimentary rocks (Cosoltepec, Chazumba and Magadalena formations). In addition, ocean-floor related rocks (pillow metabasalts

651

and a sheeted dike complex) are found within the Cosoltepec Formation (Ortega-Gutie´rrez, 1993; Ortega-Gutie´rrez et al., 1999). The Acateco subgroup (allochthonous, upper plate) comprises by eclogitized mafic and ultramafic rocks interlayered with pelitic and siliceous metasedimentary rocks (Xayacatla´n Formation), that are tectonically overlain by high-pressure metagranitoids and migmatites (Esperanza Granitoids). The two tectonically superimposed units were covered by the volcano-sedimentary Tecomate Formation, and intruded by the La Noria and Totoltepec plutons (Ortega-Gutie´rrez, 1993; Ortega-Gutie´rrez et al., 1999). The Tecomate Formation is thought to be Devonian in age (Ortega-Gutie´rrez, 1993). However, unconfirmed recent studies published in abstract form or guide books (Keppie et al., 2003) suggest younger ages (Pennsylvanian– Permian) for the Tecomate Formation. Finally, the Acatla´n Complex was unconformably covered by shallow-marine rocks of Early Mississippian age (Fig. 1b; Ortega-Gutie´rrez, 1993). 2.2. The Tehuitzingo serpentinites Serpentinite bodies, containing chromitite lenses, are present in the Tehuitzingo area (Fig. 2; Ortega-Gutie´rrez,

Fig. 2. Geologic map and cross section of the Tehuitzingo ultramafic body, also showing location of some representative samples. Note that the ultramafic rocks (serpentinites) are associated with eclogitic rocks and embedded in a metasedimentary matrix, suggesting a common (eclogitic) metamorphic history for all three lithologies.

652

J.A. Proenza et al. / Journal of South American Earth Sciences 16 (2004) 649–666

1978, 1981; Carballido-Sa´nchez and Delgado-Argote, 1989; Gonza´lez-Mancera, 2001). These serpentinite bodies are part of the Xayacatla´n Formation (Fig. 1b). This unit comprises greenschists, gneisses, quartzites, amphibolites, metagabbros, mafic eclogitic and pelitic rocks, and serpentinites. These rocks preserve the mineral assemblages derived from eclogite-facies metamorphism (Ortega-Gutie´rrez, 1978, 1993), and have been interpreted to represent an obducted slice of oceanic lithosphere (‘Xayacatla´n ophiolite’), emplaced along the Iapetus suture during the Late Ordovician-Early Silurian collision between eastern Laurentia and Oaxaquia (Ortega-Gutie´rrez et al., 1995, 1999). The serpentinite bodies, up to 500 m thick, form lenses associated with elongate units of eclogitic metabasite and metapelitic rocks (Fig. 2). The serpentinites are commonly mylonitic, and frequently display a foliation defined by the preferential orientation of antigorite blades and spinels. The serpentinites are further characterized by having interpenetrative and interlocking (non-pseudomorphic) textures. The present mineralogical composition of these serpentinites consists on antigorite, with minor chrysotile, talc, calcite, dolomite, amphiboles, chlorite, epidote, brucite, quartz, spinel, stichtite, a Fe – Ni sulfide, and Fe –Ni alloys. The spinels found are magnetite, relict chromite inherited from the primary accessory spinel, and their alteration products (‘ferritchromite’ and newly formed magnetite). The grains of accessory chromite are systematically rimmed by ‘ferritchromite’, a more porous phase with higher reflectance. In addition, the ‘ferritchromite’ zone is usually overgrown by an outer rim of magnetite. The outer magnetite rims are narrow, frequently incomplete and have sharp contacts with the ‘ferritchromite’ zones. The zone boundaries are distinct in reflected light and in backscattered-electron images. The serpentinite bodies can be interpreted as a completely serpentinized tectonite harzburgite, where the dominant occurrence of antigorite and non-pseudomorphic textures is congruent with the geological history suggested for the ‘Xayacatla´n ophiolite’ (Ortega-Gutie´rrez, 1978, 1981). In a subduction context, characterized by high pressures and relatively low temperatures, antigorite is the most stable mineral (Ulmer and Trommsdorff, 1995; see P – T conditions for the MgO – SiO2 – H2O system by Berman et al., 1986). Like other serpentinite bodies present in the Acatla´n Complex, the studied serpentinites and their associated chromitites are closely associated with eclogitic rocks and embedded in a metasedimentary matrix, suggesting that the entire assemblage shared a common metamorphic history, like that as described elsewhere (Bebout and Barton, 1989; Auzende et al., 2002). Across the entire Acatla´n Complex eclogite facies rocks of mafic, granitic and pelitic composition are intimately associated with serpentinites, strongly indicating a common high pressures metamorphic event affecting the crustal

and mantle units; unfortunately, the ultramafic rocks were 100% serpentinized and expected high pressures mineral indicators such as garnet relics (garnet peridotite) are not found. However, minimum pressures of about 9 kbar (at presumed peak temperatures of 600 8C) implied by the stability of antigorite (e.g. Berman et al., 1986) instead of the assemblage forsterite þ talc or equivalent pseudomorphs throughout the Acatla´n serpentinites, locate the Tehuitzingo ultramafic body indeed close to the eclogite field. Moreover, the anomalous development of Cr-rich silicates (Cr-epidote, Cr-actinolite, Cr-chlorite, and fuchsite) associated with high pressure phases in interlayered eclogitic rocks, indicate coeval eclogite facies interaction between these rocks and the chromite-bearing ultramafic units. 2.3. The chromitite bodies The chromitite bodies at Tehuitzingo were mined in the 1950s and 1960s. However, accurate data about the mining history of the area and its ore production are unavailable. From old mining works, we believe the chromitite bodies were small and variable in size, the deposits rarely extending more than a few tens of meters in length, and less than 2 m in thickness.

3. Analytical methods Ten serpentinite and twenty-one chromitite samples were collected from the serpentinite bodies found in the Tehuitzingo area. The majority of the collected chromitite samples came from blocks accumulated in dumps close to the old workings. Polished and thin sections of chromitite and thin sections of host serpentinites were studied in detail by optical methods prior to the analysis of mineral phases. The analyses of chromite and silicates were done with a four-channel CAMECA S £ 50 electron microprobe at the Serveis Cientificote`cnics of the Universitat de Barcelona. The analytical conditions were 20 kV accelerating voltage, 20 nA beam current, 2 mm beam diameter, and counting time of 10 s per element. Calibrations were performed using natural and synthetic standards: chromite (Cr, Al, Fe), periclase (Mg), rhodonite (Mn), rutile (Ti), NiO (Ni) and metallic V. The chemical data for Cr-spinels were stoichiometrically recalculated in order to distinguish FeO from Fe2O3 according to the procedure described by Carmichael (1967). Selected analytic results for chromite cores from chromitite and serpentinite samples are listed in Table 1. Seven chromitite and three serpentinite samples were analyzed for PGE and Au in the Genalysis Laboratory Services Pty. Ltd. at Maddington (Western Australia). These elements were analyzed by ICP-MS after nickel sulfide fire essay collection. Detection limits were 1 ppb for Rh, and 2 ppb for Os, Ir, Ru, Pt, Pd, and 5 ppb for Au.

J.A. Proenza et al. / Journal of South American Earth Sciences 16 (2004) 649–666

653

Table 1 Representative electron microprobe analyses of chromite from the chromitites and serpentinites of the Tehuitzingo area 1

2

3

4

5

6

7

8

9

10

11

12

13

14

0.20 0.16 26.03 0.08 43.88 2.45 16.64 0.16 10.65 0.16 0.03 100.42

0.11 0.12 25.78 0.08 44.00 3.03 17.60 0.16 9.02 0.09 0.11 100.10

0.09 0.15 25.71 0.13 43.64 2.90 16.33 0.15 11.06 0.13 0.00 100.30

0.14 0.10 24.32 0.08 44.83 3.02 15.68 0.36 11.55 0.13 0.05 100.26

0.12 0.12 24.50 0.08 45.29 2.47 15.74 0.26 11.70 0.18 0.08 100.54

0.06 0.12 24.64 0.10 44.90 2.88 16.01 0.18 11.36 0.11 0.03 100.40

0.18 0.03 23.47 0.12 45.11 3.87 15.70 0.15 11.52 0.21 0.06 100.42

0.12 0.05 26.03 0.09 40.81 4.63 14.19 0.52 13.99 0.01 0.46 100.90

0.19 0.13 26.05 0.13 40.07 3.58 11.90 0.43 17.51 0.05 0.32 100.37

0.08 0.13 26.00 0.12 40.31 3.65 11.89 0.44 17.62 0.00 0.28 100.50

0.12 0.16 26.54 0.20 39.87 3.25 11.78 0.44 17.85 0.10 0.23 100.54

0.13 0.19 26.43 0.19 39.81 3.42 11.93 0.36 17.70 0.03 0.30 100.49

0.09 0.14 26.03 0.04 39.96 4.01 12.23 0.41 17.07 0.02 0.19 100.17

0.15 0.13 25.84 0.15 39.94 3.88 11.57 0.33 18.12 0.00 0.24 100.35

Si Ti Al V Cr Fe3þ Mg Mn Fe2þ Ni Zn

0.05 0.03 7.24 0.02 8.19 0.44 5.86 0.03 2.10 0.03 0.00

0.03 0.02 7.16 0.02 8.20 0.54 6.18 0.03 1.78 0.02 0.02

0.02 0.03 7.19 0.02 8.19 0.52 5.78 0.03 2.20 0.02 0.00

0.03 0.02 6.87 0.02 8.49 0.55 5.60 0.07 2.31 0.02 0.01

0.03 0.02 6.90 0.02 8.55 0.44 5.60 0.05 2.34 0.04 0.02

0.02 0.02 6.93 0.02 8.47 0.52 5.69 0.04 2.27 0.02 0.01

0.04 0.01 6.64 0.02 8.56 0.70 5.62 0.03 2.31 0.04 0.01

0.03 0.01 7.34 0.02 7.72 0.83 5.06 0.11 2.80 0.00 24.01

0.05 0.02 7.48 0.03 7.72 0.66 4.32 0.09 3.57 0.01 0.06

0.02 0.02 7.46 0.02 7.76 0.67 4.31 0.09 3.59 0.00 0.05

0.03 0.03 7.60 0.04 7.66 0.60 4.27 0.09 3.63 0.02 0.05

0.03 0.04 7.57 0.04 7.65 0.63 4.32 0.08 3.60 0.01 0.05

0.02 0.03 7.47 0.01 7.70 0.74 4.44 0.09 3.48 0.00 0.03

0.04 0.02 7.44 0.03 7.72 0.71 4.21 0.07 3.70 0.00 0.04

Cr# Mg# Fe3þ#

0.53 0.74 0.03

0.53 0.78 0.03

0.53 0.72 0.03

0.55 0.71 0.03

0.55 0.71 0.03

0.55 0.72 0.03

0.56 0.71 0.04

0.51 0.64 0.05

0.51 0.55 0.04

0.51 0.55 0.04

0.50 0.54 0.04

0.50 0.55 0.04

0.51 0.56 0.05

0.51 0.53 0.04

SiO2 TiO2 Al2O3 V2O3 Cr2O3 Fe2O3 MgO MnO FeO NiO ZnO Total

1–7: chromite in chromitite. 8 –14: chromite in serpentinite. Cations calculated on the basis of 32 oxygens.

Au is below detection limits in all the analyzed samples, except one (CR-05). Analytical results are listed in Table 5.

4. Textural characteristics and mineralogy of chromitites Primary chromitite textures are predominantly massive, locally grading from densely to thinly disseminated. Also, banded chromitites, with alternating chromite-rich and serpentine-rich layers, are present (Fig. 3). Late shearing and faulting obliterated these textures giving rise to mylonitic, cataclastic, and brecciated textures. Massive chromitites (. 80% chromite) are mainly composed by coarse aggregates of chromite, comprising subhedral to anhedral crystals up to 1.2 cm in size. However, in some cases, the chromite grains show elliptical shapes resembling typical nodular chromitites. Individual chromite grains show effects of corrosion or reaction in addition to mechanical disruption, as observed in many massive chromitites from various ophiolitic complexes (e.g. Robinson et al., 1997). Disseminated chromitites (30 – 80% chromite) have smaller and more regular chromite grains, and usually grade towards more massive textures with increasing chromite contents.

No traces of primary silicate minerals were preserved in the matrix of the chromitites. Instead, intergranular minerals are mainly chlorite and minor serpentine. In massive chromitites chlorite is the only silicate present. Traces of disseminated magnetite, Fe – Ni sulfide and Fe– Ni alloys are also found in the chromitite matrix. Chlorite and serpentine also occur as inclusions in chromite grains. The chromite grains in chromitites systematically exhibit rims of ‘ferritchromite’, which are highly porous and contain numerous chlorite inclusions. Massive chromitites are usually less affected by the ferritchromitization process than the disseminated chromitites. However, sample M-106, a massive chromitite, shows a high degree of alteration to ‘ferritchromite’.

5. Mineral chemistry 5.1. Chromitite According to Rollinson (1995); Suita and Streider (1996), use of chromite composition as a petrogenetic and geotectonic indicator requires a rigorous petrography study prior to any electron microprobe analyses. The massive chromitite samples analyzed in this work normally preserve unaltered chromite in the grain cores. Only those analyses

654

J.A. Proenza et al. / Journal of South American Earth Sciences 16 (2004) 649–666

sitional changes that occurred during metamorphism and/or hydrothermal alteration. 5.1.1. Chromite The chemical composition of chromite cores are plotted in Fig. 4a and b for Cr# [Cr/(Cr þ Al)] vs. Mg# [Mg/(Mg þ Fe)], #Cr vs. TiO2, respectively. These results show that with the sole exception of sample CR15, the primary chromite composition is uniform. Cr# varies from 0.53 to 0.57 (corresponding to Cr2O3 contents between 43.27 and 46.05 wt%), Mg# ranges from 0.69 to 0.78, and the TiO2 content is generally very low (TiO2 , 0.18 wt%). Fe3þ# [Fe3þ/(Fe3þ þ Cr þ Al)] is lower than 0.04, corresponding to Fe2O3 contents between 2.04 and 4.16 wt%, and MnO (, 0.30 wt%), V 2 O3 (, 0.19 wt%), ZnO (, 0.12 wt%), and NiO (, 0.23 wt %) are very low (Table 1). CR-15 shows the highest Cr2O3 values of any of the massive chroimitite. Cr# ranges between 0.64 and 0.65 (corresponding to Cr2O3 contents between 50.5 and 51.9 wt%), Mg# ranges between 0.59 and 0.67, and the TiO2 content is low (TiO2 , 0.14 wt%) (Fig. 4a and b). Fe3þ# is lower than 0.02, corresponding to Fe2O3 contents between 1.2 and 2.0 wt%, and like the other chromitite samples, CR-15 shows very low MnO (, 0.4 wt%), V2O3 (, 0.21 wt%), ZnO (, 0.1 wt%), and NiO (, 0.1 wt%). All chromites from the chromitites at Tehuitzingo plot within the podiform (ophiolitic) chromitite field (Fig. 4a), with compositions located mainly in the Al-rich part (#Cr , 0.6) of this field. This is similar to other Al-rich ophiolitic chromitites, such as those in Coto, Philippines (Leblanc and Violette, 1983), Sartohay, China (Zhou et al., 2001), Tari-Misaka, Japan (Arai and Yurimoto,

Fig. 3. Principal textures in the chromitites from the Tehuitzingo serpentinites: (a) massive, (b) densely disseminated, and (c) banded.

performed in these unaltered cores have been considered in the interpretation of the primary chromite, and used for petrogenetic indication. However, the zones of altered chromite were also analyzed to investigate the compo-

Fig. 4. (a) #Cr [Cr/(Cr þ Al)] versus #Mg [Mg/(Mg þ Fe)], and b) #Cr versus TiO2 content for chromite in Tehuitzingo chromitites and their hosting serpentinites. The podiform and stratiform fields are from to Irvine (1967), and Leblanc and Nicolas (1992). Boninitic and MORB fields were defined by Arai (1992).

J.A. Proenza et al. / Journal of South American Earth Sciences 16 (2004) 649–666

1994), and Moa-Baracoa, Cuba (Proenza et al., 1999). Moreover, the TiO2 values are similar to those accepted as most common in ophiolitic chromitites (usually , 0.25 wt%; Leblanc and Nicolas, 1992). In the Cr# vs. TiO2 plot (Fig. 4b), the Tehuitzingo chromite compositions plot outside the fields defined by boninites and mid-oceanic ridge basalts (MORB) (Arai, 1992). Instead, the compositions fall between these fields and within the compositional range defined by the Al-rich chromites from the Sagua de Ta´namo district in eastern Cuba (Proenza et al., 1999).

655

below 0.03, making the chlorites clinochlore after the classification of Hey (1954) (Fig. 5a). In this case, however the clinochlore is Cr-rich (chromian clinochlore), and contains up to 5.08 wt% of Cr2O3 (Fig. 5b, Table 3). 5.2. Serpentinite Accessory chromites in serpentinites are generally altered into ‘ferritchromite’ and magnetite, although relicts of primary chromites are found in some serpentinites (e.g. sample CR-3). The analyses of altered accessory chromites are not included in the interpretation of primary accessory chromite composition.

5.1.2. Serpentine and chlorite Serpentine is only found in the matrix of disseminated chromitites, where it exhibits SiO2 contents from 41.91 to 42.34 wt%, and FeO from 1.99 to 2.41 wt%. Serpentine Al2O3 (up to 2.39 wt%) and Cr2O3 contents (up to 2.17 wt%) are relatively high (Table 2). The electron microprobe analyses of serpentine from chromitites probably can be influenced by their inhomogeneity and by limited resolution of the electron microprobe. Single analyses represent, in most case, the bulk composition of two phases (serpentine and chlorite). Chlorite from the matrix of the chromitites shows SiO2 contents between 27.97 and 33.66 wt%, and low Fe contents (, 2.24 wt% of FeO). The Fe/(Fe þ Mg) ratio is normally

5.2.1. Accessory chromite Primary chromites from the serpentinites have similar Cr# (, 0.50), but lower Mg# (0.53 –0.67) than those found in the chromitites (Fig. 4a). The TiO2 contents range from 0.05 to 0.16 wt% (Fig. 4b). The accessory chromites also show a systematical enrichment in Fe2þ and depletion in Mg compared to the associated chromitites. These compositions match other results, which indicate that the Mg# of chromite is higher in the chromitites than in the host peridotites or serpentinites (Leblanc and Nicolas, 1992). Such variation in Mg# depends mainly on the chromite/Mg-silicate ratio, since it

Table 2 Representative electron microprobe analyses of serpentinite from the chromitites and serpentinites of the Tehuitzingo area

SiO2 TiO2 Al2O3 V2O3 Cr2O3 MgO CaO MnO FeO NiO Na2O K2O H2O Total Si Ti Al V Cr Mg Ca Mn Fe Ni Na K

1

2

3

4

5

6

7

8

9

10

41.91 0.02 2.39 0.00 2.12 39.97 0.00 0.06 2.02 0.23 0.00 0.01 13.03 101.75

41.99 0.02 2.36 0.02 2.14 40.14 0.02 0.07 1.99 0.16 0.00 0.00 13.06 101.96

42.12 0.00 2.21 0.02 2.00 40.34 0.00 0.00 2.41 0.15 0.00 0.02 13.09 102.36

42.34 0.01 2.17 0.02 1.99 40.33 0.01 0.02 1.94 0.24 0.00 0.00 13.10 102.18

41.92 0.01 2.32 0.00 2.17 40.24 0.00 0.08 2.10 0.17 0.02 0.00 13.06 102.09

44.53 0.02 0.07 0.00 0.02 40.20 0.06 0.00 2.41 0.35 0.00 0.01 12.98 100.66

41.87 0.01 2.02 0.02 0.22 40.43 0.03 0.01 2.68 0.18 0.01 0.03 12.88 100.37

42.63 0.03 1.36 0.00 0.15 40.22 0.00 0.06 3.16 0.10 0.01 0.02 12.91 100.66

43.21 0.02 0.64 0.00 0.17 41.12 0.07 0.06 1.60 0.91 0.03 0.00 12.96 100.79

43.11 0.04 0.67 0.00 0.11 41.20 0.08 0.02 1.90 0.76 0.04 0.01 12.96 100.89

3.86 0.00 0.26 0.00 0.15 5.48 0.00 0.00 0.16 0.02 0.00 0.00

3.86 0.00 0.26 0.00 0.16 5.50 0.00 0.01 0.16 0.01 0.00 0.00

3.86 0.00 0.24 0.00 0.15 5.51 0.00 0.00 0.18 0.01 0.00 0.00

3.88 0.00 0.23 0.00 0.14 5.51 0.00 0.00 0.15 0.02 0.00 0.00

3.85 0.00 0.25 0.00 0.16 5.51 0.00 0.00 0.16 0.01 0.00 0.00

4.12 0.00 0.01 0.00 0.00 5.54 0.00 0.00 0.19 0.03 0.00 0.00

3.90 0.00 0.22 0.00 0.02 5.62 0.00 0.00 0.21 0.01 0.00 0.00

3.96 0.00 0.15 0.00 0.01 5.57 0.00 0.00 0.25 0.01 0.00 0.00

4.00 0.00 0.07 0.00 0.01 5.67 0.01 0.01 0.12 0.07 0.00 0.00

3.99 0.00 0.07 0.00 0.08 5.68 0.01 0.00 0.15 0.06 0.00 0.00

1–5: serpentine in chromitite (cations calculated on the basis of 14 oxygens). 6–10: serpentine in serpentinite (cations calculated on the basis of 14 oxygens).

656

J.A. Proenza et al. / Journal of South American Earth Sciences 16 (2004) 649–666

significant difference between them is that the chlorites in serpentinites are more impoverished in Cr2O3 than those in the chromitites (Table 3). 5.3. Chromite alteration

Fig. 5. Chemical composition of matrix chlorites in Tehuitzingo chromitites: (a) Fe/(Fe þ Mg) versus Si, and (b) Cr versus Si. Chlorite classification after Hey (1954).

is the result of subsolidus Mg –Fe exchange between Mgsilicate and chromite after cooling. In addition, the accessory chromites exhibit slightly higher contents in MnO (up to 0.52 wt%) and ZnO (up to 0.46 wt%) than chromites from chromitites. 5.2.2. Serpentine and chlorite The composition of serpentines from serpentinites is variable, probably due to the presence of different serpentine generations (Gonza´lez-Mancera, 2001). The serpentines show 41.34– 45.88 wt% SiO2, 1.25– 6.08 wt% FeO, 0.05– 2.02 wt% Al2O3, and 0.0 –1.58 wt% Cr2O3. The serpentines from serpentinites show lower Al2O3 and Cr2O3 contents than from the matrix of chromitites (Table 2). Chlorites in the serpentinites show very similar compositions to those in the associated chromitites, and are also classified as clinochlore after Hey (1954). The only

Detailed compositional profiles, obtained by electron probe microanalyses (EPMA), in chromite grains from chromitites display two main compositional zones from core to rim: (1) core chromite zones that normally retain the primary chromite composition, and (2) ‘ferritchromite’ rims. The alteration to ‘ferritchromite’ (or ferrian chromite) in chromitites is characterized by a progressive enrichment in the total iron content (Fe3þ þ Fe2þ) and a strong depletion in Al and Mg (Fig. 6; Table 4). The outer most part of ‘ferritchromite’ rims are also depleted in Cr (Fig. 6). In addition, the ‘ferritchromite’ zones are relatively enriched in Mn and Zn compared to the relict primary chromite in grain cores (Fig. 6; Table 4). In contrast, the EPMA profiles in accessory chromite grains from serpentinites show three different compositional zones from core to rim: (1) core chromite zones that in some cases retain the primary chromite composition; (2) intermediate ‘ferritchromite’ zones; and (3) magnetite rims. The accessory chromites also show a zoning pattern characterized, from core to rim, by a significant increase in Fe3þ and Fe2þ, and a decrease in Mg and Al (Fig. 7a and b; Table 4). However, two different compositional ranges of accessory chromite were found in the grain cores (Fig. 7a and b). Type-I cores have lower MnO (, 0.5 wt%), ZnO (, 0.4 wt%) and FeO contents (, 18 wt%), and higher MgO contents (. 12 wt%) than type-II cores (up to 1.6 wt% MnO, . 0.73 wt% ZnO, . 21 wt% FeO, and , 8.5 wt% MgO). This may be due to modification of the primary chemical composition of type-II cores by metamorphic reactions and the formation of ‘ferritchromite’. Finally, the magnetite rims show strong systematic enrichments in Fe3þ, Fe2þ and Ni, and depletion in Al, Cr, Mg, Mn and Zn with respect to cores (Fig. 7a and b).

6. Distribution of platinum-group elements The results of bulk-rock PGE and Au analyses in chromitites and serpentinites from the Tehuitzingo area are listed in Table 5 and respectively plotted in Fig. 8a and b. The results show low PGE contents in the chromitites, ranging from 102 to 303 ppb. The distribution of PGE is strongly dependent on texture, bulk PGE contents increasing with increasing modal proportion of chromite. Thus, the sample with the highest total PGE content (CR-07, 303 ppb PGE) is a massive, whereas chromitite samples with disseminated textures have the lowest total PGE contents. The chondrite-normalized PGE patterns of these samples are characterized by an enrichment in Ir-subgroup elements

J.A. Proenza et al. / Journal of South American Earth Sciences 16 (2004) 649–666

657

Table 3 Representative electron microprobe analyses of chlorite from the chromitites and serpentinites of the Tehuitzingo area 1

2

3

4

5

6

7

8

9

10

SiO2 TiO2 Al2O3 V2O3 Cr2O3 MgO CaO MnO FeO NiO Na2O K2O H2O Total

32.44 0.03 12.66 0.02 4.00 35.57 0.00 0.03 1.92 0.26 0.00 0.00 12.59 99.52

28.77 0.07 20.51 0.00 2.91 33.24 0.00 0.02 1.86 0.33 0.02 0.00 12.74 100.47

29.78 0.02 18.56 0.00 2.87 34.24 0.03 0.05 1.52 0.38 0.00 0.00 12.72 100.17

27.97 0.02 19.14 0.03 5.08 33.41 0.00 0.04 1.75 0.20 0.03 0.00 12.62 100.31

30.78 0.01 16.63 0.04 3.36 34.71 0.00 0.00 1.85 0.15 0.00 0.02 12.72 100.28

30.38 0.00 19.54 0.04 1.00 33.45 0.02 0.00 3.20 0.21 0.00 0.00 12.79 100.63

29.37 0.07 19.56 0.03 1.03 33.85 0.03 0.01 3.19 0.20 0.05 0.01 12.68 100.08

30.93 0.05 17.67 0.05 1.68 35.55 0.00 0.00 0.93 0.24 0.00 0.04 12.78 99.92

32.32 0.01 16.95 0.06 1.04 36.19 0.00 0.00 1.09 0.23 0.02 0.01 12.94 100.86

28.87 0.00 21.63 0.02 2.47 33.41 0.00 0.02 1.33 0.36 0.00 0.01 12.86 100.98

Si Ti Al V Cr Mg Ca Mn Fe Ni Na K

6.18 0.01 2.84 0.00 0.60 10.10 0.00 0.00 0.31 0.04 0.00 0.00

5.42 0.01 4.55 0.00 0.43 9.33 0.00 0.00 0.29 0.05 0.01 0.00

5.62 0.00 4.13 0.00 0.43 9.62 0.01 0.01 0.24 0.06 0.00 0.00

5.32 0.00 4.29 0.01 0.76 9.46 0.00 0.01 0.28 0.03 0.01 0.00

5.81 0.00 3.70 0.01 0.50 9.76 0.00 0.00 0.29 0.02 0.00 0.00

5.70 0.00 4.32 0.01 0.15 9.35 0.00 0.00 0.50 0.03 0.00 0.00

5.56 0.01 4.36 0.01 0.15 9.54 0.01 0.00 0.50 0.03 0.02 0.00

5.80 0.01 3.91 0.01 0.25 9.94 0.00 0.00 0.15 0.04 0.00 0.01

5.99 0.00 3.70 0.01 0.15 10.00 0.00 0.00 0.17 0.04 0.01 0.00

5.38 0.00 4.75 0.00 0.36 9.29 0.00 0.00 0.21 0.05 0.00 0.00

1–5: chlorite in chromitite (cations calculated on the basis of 28 oxygens). 6–10: chlorite in serpentinite (cations calculated on the basis of 28 oxygens).

(IPGE ¼ Os, Ir, Ru) relative to those of the Pd-subgroup (PPGE ¼ Rh, Pt, Pd). In addition, all chromitite samples show negative slopes from Ru to Pd [(Os þ Ir þ Ru)/(Pt þ Pd) ¼ 4.78 – 14.13]. These patterns, and the low PGE abundances, are typical of ophiolitic chromitites elsewhere (Page and Talkington, 1984; Leblanc, 1991; Zhou et al., 1998; Proenza et al., 1999; Ahmed and Arai, 2002). The Pd/Ir ratio, an indicator of PGE fractionation (Naldrett et al., 1979), is relatively constant, varying from 0.045 to 0.250. A significant feature of the PGE distribution in the chromitites is the similarity of their chondrite-normalized PGE patterns regardless of their degree of alteration (Fig. 8a). For example, sample M-106 (a massive chromitite) shows extensive alteration of primary chromite to ‘ferrichromite’, yet has a chondrite-normalized PGE pattern similar to those obtained in the other analyzed massive chromite samples that show low degrees of alteration to ‘ferrichromite’. The serpentinite samples from Tehuitzingo also show low PGE contents (SPGE ¼ 49– 55 ppb, Table 5). However, they exhibit nearly flat PGE-normalized patterns (Pd/ Ir ¼ 1.11 –1.67) and very slight enrichments in Pd (Fig. 8b). These patterns are comparable with those obtained from serpentinites and mantle peridotites elsewhere (e.g. Barnes et al., 1985; Leblanc, 1991; Lorand et al., 1993).

7. Discussion 7.1. Primary chromite compositions in chromitites and serpentines, and their tectonic implications The results of experimental crystallization estudies (Maurel and Maurel, 1982; Roeder and Reynolds, 1991) suggest that chromite compositions are controlled mainly by the composition and oxygen fugacity of the melt, and are only weakly dependent on temperature and pressure. The Cr/Al ratio of chromite in equilibrium with a given melt is controlled by the total concentration of Cr2O3 and Al2O3 in the melt, whereas the Cr content in chromite shows a negative correlation with the Al content in the melt. The chemical composition of the primary chromites from the Tehuitzingo chromitites is typical of those of ophiolitic chromitites (Irvine, 1967; Leblanc and Nicolas, 1992). With the exception of sample CR-15, the Tehuitzingo chromitites have high Al2O3 contents, typical of refractory-grade chromitites (chromite ores with Al2O3 . 20 wt%, low Fe and [Cr2O3 þ Al2O3] . 60 wt%). In ophiolitic complexes containing important chromitite bodies, Al-rich chromitites tend to occur at the shallowest levels of the upper mantle, within the so-called Moho Transition Zone, very close to the lower layered-gabbro unit. Peridotites hosting such Al-rich

658

J.A. Proenza et al. / Journal of South American Earth Sciences 16 (2004) 649–666

chromitites are also less depleted than those hosting Cr-rich chromitites (e.g. Leblanc and Violette, 1983; Leblanc and Nicolas, 1992; Zhou and Robinson, 1994; Proenza et al., 1999). Al-rich chromitites are considered to form from tholeiitic melts, whereas Cr-rich chromitites are considered to form from boninitic melts (Zhou and Robinson, 1997). The chromites from the Tehuitzingo chromitites have Cr# similar to those of MORB- and BABB-derived chromites, as defined by Dick and Bullen (1984); Arai (1992). Accordingly, we calculated the Al2O3 content of the parental melts in equilibrium with the Tehuitzingo chromitite bodies using the equation proposed by Maurel and Maurel (1982): ðAl2 O3 ÞSp ¼ 0:035ðAl2 O3 Þ2:42 Liquid

Fig. 6. Backscattered-electron (BSE) image, and profile of electronmicroprobe data across a compositionally zoned chromite grain in chromitite.

ðAl2 O3 in wt%Þ

The results (excluding sample CR-15) suggest that the melt had an average Al2O3 content of 15.3 wt%. This value is similar to the Al2O3 content of mid-ocean ridge and backarc basin basalts (Wilson, 1989; Fryer et al., 1990), and comparable to the Al2O3 contents of the parental melts for the Al-rich chromitites of the Moa-Baracoa massif in eastern Cuba (Proenza et al., 1999). The tectonic setting for chromitite crystallization is still a subject of debate (Lago et al., 1982; Robert, 1988; Leblanc and Nicolas, 1992). However, the most recent interpretations of the genesis of ophiolitic chromitites suggest a suprasubduction-zone setting. By contrast, no ophiolitic chromitites are thought to form in mature spreading centers, such as mid-ocean ridges (Robert, 1988; Arai and Abe, 1994; Arai and Yurimoto, 1995; Robinson et al., 1997; Schiano et al., 1997; Zhou and Robinson, 1997; Zhou et al., 1998; Proenza et al., 1999; Edwards et al., 2000). From experiments in water-oversaturated basalts, Matveev and Ballhaus (2002) concluded that ophiolitic chromite deposits would form only where there are primitive melts saturated in olivine-chromite and rich in water. Such conditions are most likely to occur in suprasubduction-zone environments because this is the only tectonic setting that might generate primitive basalts with water contents high enough to exsolve a fluid phase necessary for chromite formation. Proenza et al. (1999) interpreted the Al-rich chromitite bodies from eastern Cuba to be the result of chromite-forming, olivinedissolving melt/rock reactions produced when back-arc basin basalt melts migrated by porous flow through dunitic channels, and mixed with oxidized, volatile-rich melts in the suprasubduction mantle. Hence, given the chemical compositions of the Al-rich primary igneous chromite from the Tehuitzingo chromitites, it is likely that they too crystallized from melts with BABB affinity in the suprasubduction mantle. The Cr# of accessory chromite in peridotites has been used extensively as an indicator of the degree of melting in the upper mantle, high-Cr chromites correlating with the highest degree of melting and, hence, the greater degree of

J.A. Proenza et al. / Journal of South American Earth Sciences 16 (2004) 649–666

659

Table 4 Representative electron microprobe analyses of ‘ferritchromite’ from the chromitites and serpentinites of the Tehuitzingo area 1

2

3

4

5

6

7

8

9

10

SiO2 TiO2 Al2O3 V2O3 Cr2O3 Fe2O3 MgO MnO FeO NiO ZnO Total

0.22 0.22 0.50 0.20 45.51 22.05 2.63 0.67 26.80 0.24 0.31 99.35

0.20 0.29 0.18 0.19 42.54 25.57 2.35 0.91 27.16 0.28 0.27 99.94

0.14 0.14 8.53 0.09 44.71 15.24 4.08 0.77 26.21 0.08 0.45 100.44

0.11 0.33 0.31 0.18 43.23 24.98 2.32 0.77 27.58 0.25 0.15 100.21

0.17 0.11 8.72 0.05 45.18 14.71 4.02 0.71 26.43 0.15 0.46 100.72

0.13 0.27 0.79 0.14 32.23 36.60 4.42 1.14 23.71 0.35 0.06 99.85

0.15 0.26 0.32 0.06 31.60 37.78 3.91 1.29 24.34 0.38 0.07 100.17

0.20 0.52 0.36 0.10 45.06 21.66 0.98 1.75 28.72 0.10 0.44 99.88

0.17 0.58 0.25 0.15 43.52 22.86 0.78 1.59 29.08 0.07 0.52 99.58

0.23 0.59 0.25 0.12 43.98 22.66 0.84 1.78 28.92 0.05 0.53 99.93

Si Ti Al V Cr Fe3þ Mg Mn Fe2þ Ni Zn

0.06 0.05 0.17 0.05 10.66 4.91 1.16 0.17 6.64 0.06 0.07

0.06 0.06 0.06 0.05 9.96 5.70 1.04 0.23 6.72 0.07 0.06

0.04 0.03 2.80 0.02 9.84 3.19 1.69 0.18 6.10 0.02 0.09

0.03 0.07 0.11 0.04 10.09 5.55 1.02 0.19 6.81 0.06 0.03

0.05 0.02 2.85 0.01 9.91 3.07 1.66 0.17 6.13 0.03 0.09

0.04 0.06 0.27 0.03 7.45 8.05 1.93 0.28 5.80 0.08 0.01

0.05 0.06 0.11 0.02 7.33 8.34 1.71 0.32 5.97 0.09 0.02

0.06 0.12 0.13 0.03 10.64 4.87 0.43 0.44 7.17 0.02 0.10

0.05 0.13 0.09 0.04 10.34 5.17 0.35 0.40 7.31 0.02 0.12

0.07 0.13 0.09 0.03 10.40 5.10 0.37 0.45 7.23 0.01 0.12

Cr# Mg# Fe3þ#

0.98 0.15 0.31

0.99 0.13 0.36

0.78 0.22 0.20

0.99 0.13 0.35

0.78 0.21 0.19

0.96 0.25 0.51

0.99 0.22 0.53

0.99 0.06 0.31

0.99 0.05 0.33

0.99 0.05 0.33

1–5: ‘ferritchromite’ in chromitite. 6–10: ‘ferritchromite’ in serpentinite. Cations calculated on the basis of 32 oxygens.

depletion of peridotites (Dick and Bullen, 1984; Arai, 1992). Accessory chromites in the Tehuitzingo serpentinites plot in the Cr-rich part of the compositional range of backarc basin peridotites, as well as in the lower part of the compositional field of the fore-arc basin peridotites from the Mariana Trench (Fig. 9). According to Ohara et al. (1996, 2002), back-arc basin peridotites (e.g. Vela Basin and Mariana Trough) contain accessory chromite with Cr# # 0.55. In contrast, fore-arc peridotites (e.g. Mariana Trench) contain accessory chromites with high Cr# (up to 0.79; Ohara and Ishii, 1998). Primary accessory chromites from the serpentinites at Tehutizingo have Cr# ranging 0.53 –0.57, close to those described in back-arc basin peridotites. Accordingly, we suggest that the Tehuitzingo serpentinites and the associated chromitites are fragments of oceanic lithosphere formed or modified in a back-arc environment, and represent the mantle sequence of a suprasubduction zone ophiolite in the sense of Pearce et al. (1984). This is consistent with the chemical composition of the primary chromites in the serpentinites and the abundance of chromitites in the area (Robert, 1988; Arai and Abe, 1994; Arai and Yurimoto, 1995; Zhou and Robinson, 1997; Proenza et al., 1999; Matveev and Ballhaus, 2002).

7.2. Alteration of chromite in chromitites and serpentinites As described above, chromite grains in the chromitites and serpentinites show systematic compositional variation, either by their rims or through cracks, from chromite to ‘ferritchromite’ or, in the case of accessory chromites, from chromite to magnetite. ‘Ferritchromite’ is the most common alteration phase observed in Crspinels elsewhere. The alteration trend, from core to rim, is usually characterized by a decrease in Al, Mg, and Cr contents, accompanied by an increase in Fe3þ and Fe2þ. The origin of ‘ferritchromite’ is related to metamorphic and/or hydrothermal processes. According to Bliss and MacLean (1975), ‘ferritchromite’ represents a product of prograde metamorphic reactions between Cr-spinel cores and magnetite rims. Abzalov (1998) similarly demonstrated, using Cr-spinels from the Pechenga intrusions (Kola Peninsula, Russia), that the alteration induced by prograde metamorphism has a major control on the composition of Cr-spinels in the amphibolite metamorphic facies. Evans and Frost (1975) likewise concluded that the Cr content of spinels increases with increasing grade in the amphibolite facies. On the other hand, the results obtained by Burkhard (1993) suggest the opposite trend, and Roeder (1994) showed

660

J.A. Proenza et al. / Journal of South American Earth Sciences 16 (2004) 649–666

Fig. 7. Backscattered-electron (BSE) image, and profile of electron-microprobe data across a compositionally zoned accessory chromite grain in serpentinite. (a) type-I accessory chromite in sample CR-3, (b) type-II accessory chromite in sample P-1.

‘ferritchromite’ to be the product of a reaction between Cr-spinels and chlorite from the host rocks. Hence, the origin of ‘ferritchromite’ remains an unsolved problem, the resolution of which is beyond the scope of this paper.

The chemical compositions of all analyzed spinels from Tehuitzingo (primary chromite, ‘ferritchromite’ and magnetite) are plotted on a triangular Fe3þ –Cr –Al diagram in Fig. 10, together with the spinel compositional fields from

J.A. Proenza et al. / Journal of South American Earth Sciences 16 (2004) 649–666

661

Table 5 Whole-rock analyses on platinum-group elements (PGE) and Au (ppb) of representative samples of chromitites and serpentinites from the Tehuitzingo area Sample Os Ir Ru Rh Pt Pd Au

CR-05 52 37 55 6 9 4 6

Total PGE Pd/Ir (Os þ Ir þ Ru)/(Pt þ Pd) (Os þ Ir þ Ru)N/PGEN

163 0.108 11.08 0.84

CR-07

CR-09

CR-12

CR-13

CR-15

M-106

M-107

P-3

P-5

63 107 96 10 16 11 ,5

18 25 44 4 8 3 ,5

65 78 69 7 10 5 ,5

22 20 44 3 13 5 ,5

37 43 75 6 15 8 ,5

30 89 38 13 9 4 ,5

6 8 12 3 11 9 ,5

6 6 14 2 13 10 ,5

6 9 13 3 14 10 ,5

303 0.103 9.85 0.84

102 0.120 7.91 0.81

234 0.064 14.13 0.87

107 0.250 4.78 0.8

184 0.186 6.74 0.81

183 0.045 12.08 0.77

49 1.125 1.30 0.51

51 1.667 1.13 0.51

55 1.111 1.17 0.5

Chromitites: CR-05, CR-07, CR-09, CR-12, CR-13, CR-15, M-106. Serpentinites: M-107, P-3, P-5.

different metamorphic facies where compositional changes in spinel have been recorded with increasing metamorphic grade (Evans and Frost, 1975; Frost, 1991; Suita and Streider, 1996; Barnes and Roeder, 2001). No compositional field for spinels in the eclogite facies is shown because their alteration during high-pressure metamorphism is not well established. The spinels in the chromitites and serpentinites from Tehuitzingo show an alteration trend characterized by low Al contents, with the resulting spinels (‘ferritchromite’ and magnetites) plotting along the Cr – Fe3þ join. This compositional trend is typical of chromites altered at low temperatures (Roeder, 1994). As shown in Fig. 10, the compositions of Cr-spinels from Tehuitzingo lie outside the compositional fields of Cr-spinels from the upper-amphibolite to granulite facies. Rather, their compositions plot mainly within the greenschist field like those of Cr-spinels from serpentinized peridotites and chromitites that have experienced ocean-floor metamorphism alone. Based on the available data, we suggest two explanations for this alteration trend: (1) the Cr-spinels from Tehutizingo experienced no significant compositional variations during the high-pressure metamorphic event, or (2) what compositional variations the Cr-spinels experienced

Fig. 8. Chondrite-normalized PGE patterns (C1 chondrites; Naldrett and Duke, 1980) of representative samples from the Tehuitzingo area: (a) chromitites, (b) serpentinites.

during the high-pressure event were obliterated by retrograde metamorphism in the lower-amphibolite and greenschist facies. We favor the first interpretation since the chemical composition of relict primary chromite cores (Fig. 4) and their PGE distribution patterns (Fig. 8) are consistently similar to ophiolites that have not experienced eclogite facies metamorphism. In Cr-spinels from ultramafic complexes in Brazil that has been metamorphosed (700 –750 8C and 6 – 7 kbars) under variable PH2O conditions, Candia and Gaspar (1997) found that when PH2O ¼ Ptot, complete metamorphic reequilibration occurred, but when PH2O , Ptot, relict igneous cumulate textures were preserved. Hence, the

Fig. 9. Comparative #Cr versus #Mg plot for accessory chromite in the Tehuitzingo serpentinites and those in peridotites from: (1) Mariana Trench (fore-arc basin; Ohara and Ishii, 1998), (2) Vela Basin and Mariana Trough (back-arc basins; Ohara et al., 1996, 2002).

662

J.A. Proenza et al. / Journal of South American Earth Sciences 16 (2004) 649–666

Fig. 10. Compositional changes in spinels from Tehuitzingo (primary chromite, ‘ferritchromite’, and magnetite) expressed in a triangular Fe3þ-Cr-Al plot. (a) Spinels in chromitite, (b) spinels in serpentinites. Spinel compositional fields from different metamorphic facies are from Purvis et al., (1972); Evans and Frost (1975), and Suita and Streider (1996).

metamorphic modifications of Cr-spinel were controlled by the fluid phases. In a high-pressure metamorphic environment, the degree of fluid/rock interaction is still poorly understood. However,

recent work has shown that the stable isotope compositions of ophiolitic rocks metamorphosed at high pressures are similar to those of non-metamorphosed ophiolites, indicating that fluid flow is not pervasive during high-

J.A. Proenza et al. / Journal of South American Earth Sciences 16 (2004) 649–666

pressure metamorphism (Cartwright and Barnicoat, 1999). Fru¨h-Green et al. (2001) similarly found the stable isotope (H –O) variations of ultramafic rocks under high pressure to be comparable to the variations observed in modern oceanic rocks and non-metamorphosed ophiolites. This suggests that high-pressure metamorphism induces local fluid flow at low water/rock ratios, and so displays closed-system behavior (Fru¨h-Green et al., 2001). Accordingly, we tentatively suggest that at high-pressure conditions there is no resetting of the primary igneous composition of chromites. The main alteration features observed in the chromites at Tehuitzingo are consequently considered products of a low-P and low-T retrograde alteration history. 7.3. PGE and Au geochemistry The PGE abundances and chondrite-normalized patterns obtained from the chromitites at Tehuitzingo are similar to those found in most chromitite bodies of ophiolitic complexes (Fig. 11). The latter are characterized by relatively low PGE contents, with Os, Ir and Ru abundances of 0.05 – 1.0 times chondritic values, and Pt and Pd abundances of about 0.01 times chondritic values. Steep negative slopes from Ru to Pt are also characteristic of ophiolitic chromitites (Page and Talkington, 1984; Leblanc 1991, 1995; Zhou et al., 1998; Proenza et al., 1999; Ahmed and Arai, 2002), and have been attributed to the early crystallization of the most refractory Os-, Ir- and Ru-bearing minerals prior to or coeval with chromite. This induces the subsequent partitioning of Rh, Pt and Pd to the residual melt (Barnes et al., 1985, 1988). According to Proenza et al. (2001), on the other hand, the chondrite-normalized pattern of ophiolitic chromitites could be explained by the higher solubility in basaltic melts of Pd, Pt and Rh, relative to Ir,

Fig. 11. Comparison between chondrite-normalized PGE patterns of chromitites from Tehuitzingo (shaded area) and those from chromitites in non-metamorphosed ophiolites: (1) OM, Oman (in Leblanc, 1991); (2) V, Vourinos, Greece (Economou, 1986), and (3) MB, Moa-Baracoa, Cuba (Proenza et al., 1999).

663

with increasing oxygen fugacity at low sulfur fugacities (Amosse´ et al., 1990; Vatin-Perignon et al., 1998). As the PGE contents in chromite depend mainly on the PGE concentration in the parental melt (Barnes et al., 1985, 1988), the chromitites from Tehuitzingo presumably crystallized from a PGE-poor parental magma. The low PGE contents of the Tehuitzingo chromitites, as well as other Al-rich chromitites in ophiolites, can be explained by crystallization from tholeiitic melts, (Zhou et al., 1998, 2001), provided the tholeiitic magmas are saturated in S and depleted in PGE (Hamlyn et al., 1985). The chondrite-normalized PGE pattern of podiform chromites is a good indicator of secondary processes (Malitch et al., 2002). These authors studied two massive ophiolitic chromitites, one unaltered and the other highly altered, in the Kraubath ultramafic massif, Austria. They found that massive highly altered chromitites are PPGE rich and display chondrite-normalized PGE patterns unlike to those found in typical unaltered ophiolitic chromitites. The unusual PPGE enrichment of altered chromitites was attributed to hydrothermal processes during polyphase amphibolite facies regional metamorphism during the Variscan and Alpine orogenies. Thus, they concluded that altered massive chromitites of the Kraubath type may be good targets for PGE exploration. In Tehuitzingo, however, all chromitite samples exhibit similar chondrite-normalized PGE patterns to those of unmetamorphosed ophiolitic chromitites (Fig. 11). This indicates that the PGE distribution has not been significantly affected by high-pressure (eclogite facies) metamorphism. We therefore conclude that the Tehuitzingo chromitites experienced no significant redistribution (or concentration) of PGE either during metamorphism or during other alteration processes. If some form of PGE mobilization did occur, it was restricted to individual chromitite bodies and did not change the bulk-rock PGE compositions. The absence of aqueous fluids in the eclogitic event could also account for the lack of PGE redistribution. Similar results were obtained in the stratiform chromitites of the Santa Cruz massif, Brazil (Ipanema mafic-ultramafic Complex; Angeli et al., 2001), where PGE accumulated during chromite crystallization were not remobilized from primary chromite during granulite- and amphibolite-facies metamorphic events. Likewise, the chondrite-normalized PGE patterns obtained from the serpentinites at Tehuitzingo indicate that the PGE distribution in the parental peridotites was not significantly affected by the various serpentinization events associated with their metamorphic history. Instead, the patterns obtained in the Tehuitzingo serpentinites are similar to those of partially serpentinized peridotites from mantle sequences of non-metamorphosed ophiolites (Ahmed and Arai, 2002). These results match previous studies that suggest PGE are immobile during serpentinization (Groves and Keays, 1979; Prichard and Tarkian, 1988). However, according to Leblanc (1991), serpentinites are

664

J.A. Proenza et al. / Journal of South American Earth Sciences 16 (2004) 649–666

usually enriched in Au realtive to primary peridotites, and this Au enrichment is related to serpentinization processes. For example, in sulfide- and/or arsenide-bearing serpentinites, the Au content is up to 100 times the average content in serpentinites (Leblanc, 1991). Yet, the Au contents in all the analyzed chromitite and serpentinite samples from Tehuitzingo are below detection limits (5 ppb). Our data suggest therefore that, if any concentration of Au occurred during serpentinization, it was leached by hydrothermal fluids during the alteration history.

8. Conclusions The data presented here suggest that the serpentinites and associated chromitites of Tehuitzingo represent a fragment of oceanic lithosphere that formed and/or was modified in an arc/back-arc environment, equivalent to an ophiolitic mantle sequence from a supra-subduction zone. The primary igneous chromite composition of this fragment of oceanic lithosphere remained essentially unchanged during the main alteration and metamorphic events that affected the Acatla´n Complex. The PGE accumulated during chromite crystallization were not remobilized from the chromitite during subsequent high-pressure metamorphism in the eclogitic facies, or during other alteration processes.

Acknowledgements The XRD, SEM-EDS and EPMA analyses carried out in the Serveis Cientı´fico-Te`cnics of the Universitat de Barcelona. Special thanks are given to Xavier Llovet for his assistance with the electron-microprobe analyses. We also acknowledgements the help of J. Sole´ and R. Corona with the field work. Moreover we want express our gratitude to Damian Nance, Ohio University, and Jarda Dostal, Saint Mary’s University, for critical reviewing and their constructive comments. This paper was supported by CONACyT (Mexico) project J 32506-T, and the Spanish project BTE2001-3308. In addition, this work received funds from UNAM-DGAPA, project PAPIIT IN 107999 to Fernando Ortega Gutie´rrez.

References Abzalov, M.Z., 1998. Chrome-spinel in gabbro-wehrlite intrusions of the Pechenga area, Kola Peninsula, Russia: emphasis on alteration features. Lithos 43, 109–134. Ahmed, A.H., Arai, S., 2002. Unexpectedly high-PGE chromitite from the deeper mantle section of the northern Oman ophiolite and its tectonic implications. Contributions to Mineralogy and Petrology 143, 263– 278. Amosse´, J., Allibert, M., Fischer, W., Piboule, M., 1990. Experimental study of the solubility of platinum and iridium in basic silicate

melts-implications for the differentiation of platinum-group elements during magmatic processes. Chemical Geology 81, 45–53. Angeli, N., Fleet, M.E., Thibault, Y., Candia, M.A.F., 2001. Metamorphism and PGE-Au content of chromitite from the Ipanema mafic/ultramafic Complex, Minas Gerais, Brazil. Mineralogy and Petrology 71, 173 –194. Arai, S., 1992. Chemistry of chromian spinel in volcanic rocks as a potential guide to magma chemistry. Mineralogical Magazine 56, 173 –784. Arai, S., Abe, N., 1994. Possible presence of podiform chromitite in the arc mantle: chromitite xenoliths from the Takashima alkali basalt, southwest Japan arc. Mineralium Deposita 29, 434–438. Arai, S., Yurimoto, H., 1994. Podiform chromitites of the Tari-Misaka ultramafic complex, Southwestern Japan, as mantle-melt interaction products. Economic Geology 89, 1279–1288. Arai, S., Yurimoto, H., 1995. Possible sub-arc origin of podiform chromitites. The Island Arc 4, 104–111. Auzende, A.L., Devouard, B., Guillot, S., Daniel, I., Baronnet, A., Lardeaux, J.M., 2002. Serpentinites from Central Cuba: petrology and HRTEM study. European Journal of Mineralogy 14, 905–914. Barnes, S.J., Roeder, P.L., 2001. The range of spinel compositions in terrestrial mafic and ultramafic rocks. Journal of Petrology 42, 2279–2302. Barnes, S.J., Naldrett, A.J., Gorton, M.P., 1985. The origin of the fractionation of platinum group elements in terrestrial magmas. Chemical Geology 53, 303–323. Barnes, S.J., Boyd, R., Korneliussen, A., Nilsson, L.P., Often, M., Pedersen, R.B., Robins, B., 1988. The use of mantle normalization and metal ratios in discriminating between the effects of partial melting, crystal fractionation and sulfide segregation on platinum-group elements, gold, nickel and copper: examples from Norway. In: Prichard, H.M., Potts, P.J., Bowles, J.F.W., Cribb, S.J. (Eds.), Geoplatinum 87, Elsevier, London, pp. 113–144. Bebout, G.E., Barton, M.D., 1989. Fluid flow and metasomatism in a subduction zone hydrothermal system: Cataline Schist terrane, California. Geology 17, 976 –980. Berman, R.G., Engi, M., Greenwood, H.J., Brown, T.H., 1986. Derivation of internally-consistent thermodynamic data by the technique of mathematical programming: A review with application to the system MgO– SiO2 –H2O. Journal of Petrology 27, 1331–1364. Bliss, N.W., MacLean, W.H., 1975. The paragenesis of zoned chromite from central Manitoba. Geochimica et Cosmochimica Acta 39, 973 –990. Burkhard, D.J.M., 1993. Accessory chromium spinels: Their coexistence and alteration in serpentinites. Geochimica et Cosmochimica Acta 57, 1297–1306. Candia, M.A.F., Gaspar, J.C., 1997. Chromian spinels in metamorphosed ultramafic rocks from Mangabal I and II complexes, Goia´s, Brazil. Mineralogy and Petrology 60, 27–40. Carballido-Sa´nchez, E.A., Delgado-Argote, L.A., 1989. Geologı´a del cuerpo serpentinı´tico de Tehuitzingo, Estado de Puebla. Interpretacio´n preliminar de su emplazamiento. Revista Mexicana de Ciencias Geolo´gicas 8(2), 134–148. Carmichael, I.S.E., 1967. The iron-titanium oxides of salic volcanic rocks and their associated ferromagnesian silicates. Contributions to Mineralogy and Petrology 14, 36–64. Cartwright, I., Barnicoat, A.C., 1999. Stable isotope geochemistry of Alpine ophiolites: a window to ocean-floor hydrothermal alteration and constraints on fluid-rock interaction during high-pressure metamorphism. International Journal of Earth Sciences 88, 219– 235. Dick, H.J.B., Bullen, T., 1984. Chromian spinel as a petrogenetic indicator in abyssal and alpine-type peridotites and spatially associated lavas. Contributions to Mineralogy and Petrology 86, 54 –76. Economou, M.I., 1986. Platinum-group elements (PGE) in chromite and sulphide ores within the ultramafic zone of some Greek ophiolite complexes. In: Gallagher, M.J., Ixer, R.A., Neary, C.R., Prichard, H.M

J.A. Proenza et al. / Journal of South American Earth Sciences 16 (2004) 649–666 (Eds.), Metallogeny of Basic and Ultrabasic Rocks, Institute of Mining and Metallurgy, London, UK, pp. 441– 454. Edwards, S.J., Pearce, J.A., Freeman, J., 2000. New insights concerning the influence of water during the formation of podiform chromite. In: Dilek, Y., Moores, E.M., Elthon, D., Nicolas, A. (Eds.), Ophiolites and oceanic crust: new insights from field studies and the ocean drilling program: Boulder, vol. 349. Geological Society of America Special Paper, Colorado, pp. 139 –147. Elı´as-Herrera, M., Ortega-Gutie´rrez, F., 2002. Caltepec fault zone: An Early Permian dextral transpressional boundary between the Proterozoic Oxaca and Paleozoic Acatla´n complexes, Southern Mexico, and regional tectonic implications. Tectonics 21(3), 1–4. Evans, B.W., Frost, B.R., 1975. Chome-spinel in progressive metamorphism—a preliminary analysis. Geochimica et Cosmochimica Acta 39, 959–972. Frost, B.R., 1991. Stability of oxide minerals in metamorphic rocks. In: Lindsley, D.H., (Ed.), Oxide Minerals: Petrologic and Magnetic Significance. Review Mineral, vol. 25., pp. 469–487. Fru¨h-Green, G.L., Scambelluri, M., Vallis, F., 2001. O –H isotope ratios of high pressure ultramafic rocks: implications for fluid sources and mobility in the subducted hydrous mantle. Contributions to Mineralogy and Petrology 141, 145–159. Fryer, P., Taylor, B., Langmuir, C.H., Hochstaedter, A.G., 1990. Petrology and geochemistry of lava from the Sumisu and Torishima backarc rifts. Earth and Planetary Science Letters 100, 161 –178. Gonza´lez-Mancera, G., 2001. Mineralogı´a y petrologı´a de las serpentinitas del cuerpo ultrama´fico de Tehuitzingo, Estado de Puebla. Unpublished MSc thesis, Instituto de Geologı´a, Universidad Nacional Auto´noma de Me´xico, 103 pp. Groves, D.I., Keays, R.R., 1979. Mobilisation of ore-forming elements during alteration of dunites, Mt Keith-Betheno, Western Australia. The Canadian Mineralogist 17, 373–389. Hamlyn, P.R., Keays, R.R., Cameron, W.E., Crawford, A.J., Waldron, H.M., 1985. Precious metals in magnesian low-Ti lavas: implication for metallogenesis and sulfur saturation in primary magmas. Geochimica et Cosmochimica Acta 49, 1797–1811. Hey, M.H., 1954. A new revision of the chlorite. Mineralogical Magazine 30, 277 –292. Irvine, T.N., 1967. Chromian spinel as a petrogenetic indicator; Part II, Petrologic applications. Canadian Journal of Earth Sciences 4, 71–103. Keppie, J.D., Solari, L.A., Ortega-Gutie´rrez, F., Elias-Herrera, M., Nance, R.D., 2003. Paleozoic and Precambrian rocks of southern Me´xicoAcatla´n and Oaxaca complexes. In: Geologic Transects Across Cordilleran Mexico, Guidebook for the Field Trips of 99th Geological Society of American Cordilleran Section Annual Meeting, Puerto Vallarta, Jalisco, Mexico, April 4–11, 2003. Mexico, D.F. Universidad Nacional Auto´noma de Me´xico, Instituto de Geologı´a, Publicacio´n Especial 1, Field Trip 11, p. 281 –314. Lago, N.L., Rabinowicz, M., Nicolas, A., 1982. Podiform chromite orebodies: a genetic model. Journal of Petrology 23, 103 –125. Leblanc, M., 1991. Platinum-group elements and gold in ophiolitic complexes: distribution and fractionation from mantle to oceanic floor. In: Peters, T.J., Nicolas, A., Coleman, R.G. (Eds.), Ophiolite Genesis and Evolution of Oceanic Lithosphere, Kluwer Academic Publisher, Dordrecht, pp. 231 –260. Leblanc, M., 1995. Chromite and ultramafic rock compositional zoning through a paleotransform fault, Poum, New Caledonia. Economic Geology 90, 2028–2039. Leblanc, M., Nicolas, A., 1992. Les chromitites ophiolitiques. Chronique de la Recherche Minie`re 507, 3–25. Leblanc, M., Violette, J.F., 1983. Distribution of aluminium-rich and chromium-rich chromite pods in ophiolite peridotites. Economic Geology 78, 293–301. Lorand, J.P., Keays, R.R., Bodinier, J.L., 1993. Copper and noble metal enrichments across the lithosphere-asthenosphere boundary of mantle diapirs: Evidence from the Lanzo lherzolite massif. Journal of Petrology 34, 1111–1140.

665

Malitch, K.N., Melcher, F., Mu¨hlhans, H., 2002. Palladium and gold mineralization in podiform chromitite at Kraubath, Austria. Mineralogy and Petrology 73, 427–477. Malone, J.R., Nance, R.D., Keppie, J.D., Dostal, J., 2002. Deformation history of part of the Acatla´n Complex: Late Ordovician– Early Silurian and Early Permian orogenesis in southern Mexico. Journal of South American Earth Sciences 15, 511 –524. Matveev, S., Ballhaus, C., 2002. Role of water in the origin of podiform chromitite deposits. Earth and Planetary Science Letters 203, 235–243. Maurel, C., Maurel, P., 1982. E´tude expe´rimentale de la distribution de l’aluminium entre bain silicate´ basique et spinelle chromife`re. Implications pe´troge´ne´tiques: teneur en chrome des spinelles. Bulletin de Mine´ralogie 105, 197–202. Mountain, B.W., Wood, S.A., 1988. Solubility and transport of platinumgroup elements in hydrothermal solution: thermodynamic and physical chemical constraints. In: Prichard, H.M., Potts, P.J., Bowles, J.F.W., Cribb, S.J. (Eds.), Geoplatinum 87, Elsevier Applied Science, London, UK, pp. 57 –82. Naldrett, A.J., Duke, J.M., 1980. Platinum metals in magmatic sulfide ore. Science 280, 117–1428. Naldrett, A.J., Hoffman, E.L., Green, A.H., Chou, C.L., Naldrett, S.R., Alcock, R.A., 1979. The composition of Ni-sulfide ores, with particular reference to their content of PGE and Au. The Canadian Mineralogist 17, 403–415. Ohara, Y., Ishii, T., 1998. Peridotites from the southern Mariana forearc: Heterogeneous fluids supply in mantle wedge. The Island Arc 7, 541– 558. Ohara, Y., Kasuga, S., Ishii, T., 1996. Peridotites from the Parece Vela Rift in the Philippine Sea upper mantle material exposed in an extinct backarc basin. Proceedings Japan Acad. 72, 118–123. Ohara, Y., Stern, R.J., Ishii, T., Yurimoto, H., Yamazaki, T., 2002. Peridotites from the Mariana Trough: first look at the mantle beneath an active back-arc basin. Contributions to Mineralogy and Petrology 143, 1–18. Ortega-Gutie´rrez, F., 1978. Estratigrafı´a del Complejo Acatla´n en la Mixteca Baja, estados de Puebla y Oaxaca. Revista Mexicana de Ciencias Geolo´gicas 2(2), 112 –131. Ortega-Gutie´rrez, F., 1981. Metamorphic belts of southern Me´xico and their tectonic significance. Geofı´sica Internacional 20(3), 177– 202. Ortega-Gutie´rrez, F., 1993. Tectonostratigraphic analysis and significance of the Paleozoic Acatla´n Complex of southern Mexico. In: OrtegaGutie´rrez, F., Centeno-Garcı´a, E., Mora´n-Zenteno, D., Go´mezCaballero, A. (Eds.), Terrane Geology of Southern Mexico, Universidad Nacional Auto´noma de Me´xico, Instituto de Geologı´a, First Circum-Atlantic Terrane Conference, Guanajuato, Mexico, pp. 54– 60, Guidebook of Field Trip B. Ortega-Gutie´rrez, F., Ruiz, J., Centeno-Garcı´a, E., 1995. Oaxaquia, a Proterozoic microcontinent accreted to North America during the Late Paleozoic. Geology 23, 1127–1130. Ortega-Gutie´rrez, F., Elı´as-Herrera, M., Reyes-Salas, M., Macı´as-Romo, C., Lo´pez, R., 1999. Late Ordovician–Early Silurian continental collisional orogeny in southern Mexico and its bearing on GondwanaLaurentia connections. Geology 27, 719 –722. Page, N.J., Talkington, R.W., 1984. Palladium, platinum, rhodium, ruthenium and iridium in peridotites and chromitites from ophiolite complexes in Newfoundland. The Canadian Mineralogist 22, 137–149. Pearce, J.A., Lippard, S.J., Roberts, S., 1984. Characteristics and tectonic significance of suprasubduction zone ophiolites. In: Kokelaar, B.P., Howells, M.F. (Eds.), Marginal Basin Geology, 16. Geological Society of London, Special Publication, pp. 77 –94. Prichard, H.M., Tarkian, M., 1988. Platinum and palladium minerals from two PGE-rich localities in the Shetland Ophiolite Complex. The Canadian Mineralogist 26, 979–990. Proenza, J., Gervilla, F., Melgarejo, J.C., Bodinier, J.L., 1999. Al- and Crrich chromitites from the Mayarı´-Baracoa Ophiolitic Belt (eastern Cuba): consequence of interaction between volatile-rich melts and peridotite in suprasubduction mantle. Economic Geology 94, 547–566.

666

J.A. Proenza et al. / Journal of South American Earth Sciences 16 (2004) 649–666

Proenza, J., Gervilla, F., Melgarejo, J.C., Vera, O., Alfonso, P., Fallick, A., 2001. Genesis of sulfide-rich chromite ores by the interaction between chromitite and olivine-norite dikes in the Potosı´ Mine (MoaBaracoa ophiolitic massif, eastern Cuba). Mineralium Deposita 36, 658– 669. Purvis, A.C., Nesbitt, R.W., Hallberg, J.A., 1972. The geology of part of the Carr Boyd Complex and its associated nickel mineralization, Western Australia. Economic Geology 67, 1093–1113. Robert, S., 1988. Ophiolitic chromitite formation: a marginal basin phenomenon. Economic Geology 83, 1034– 1036. Robinson, P.T., Zhou, M.F., Malpas, J., Bai, W.J., 1997. Podiform chromitites: Their composition, origin and environment of formation. Episodes 20(4), 247–252. Roeder, P.L., 1994. Chromite: from the fiery rain of chondrules to Kilauea lava lake. The Canadian Mineralogist 32, 729–746. Roeder, P.L., Reynolds, I., 1991. Crystallization of chromite and chromium solubility in basaltic melts. Journal of Petrology 32, 909–934. Rollinson, H., 1995. Composition and tectonic settings of chromite deposits through time-a discussion. Economic Geology 90, 2091– 2098. Schiano, P., Clocchiatti, R., Lorand, J.P., Massare, D., Deloule, E., Chaussidon, M., 1997. Primitive basaltic melts included in podiform chromites from the Oman ophiolite. Earth and Planetary Science Letters 146, 489–497. Stowe, C.W., 1994. Compositions and tectonic setting of chromite deposits through time. Economic Geology 89, 528 –546. Suita, M.T., Streider, A.J., 1996. Cr-spinels from Brazilian mafic-ultramafic complexes: metamorphic modifications. International Geology Review 38, 245 –267. Thalhammer, O.A.R., Prochaska, W., Mu¨hlhans, H.W., 1990. Solid inclusions in chrome-spinels and platinum group element concentrations form the Hochgro¨ssen and Kraubath ultramafic massifs (Austria): their relationship to metamorphism and serpentinization. Contributions to Mineralogy and Petrology 105, 66– 80.

Ulmer, P., Trommsdorff, V., 1995. Serpentine stability to mantle depths and subduction-related magmatism. Science 268, 858–861. Vatin-Perignon, N., Amosse´, J., Radelli, L., 1998. Platinum group element behavior in a supra-subduction zone setting: the UB-B complex and the basic dyke swarm of the Vizcaino Peninsula, Baja California Sur, Mexico. 8th International Platinum Symposium. The Geological Society of South Africa and the South African Institute of Mining and Metallurgy, Symposium Series S18, 419 –421. Weber, B., Meschede, M., Ratschbacher, L., Frisch, W., 1997. Structure and kinematic history of the Acatla´n Complex in the Nuevos Horizontes-San Bernanrdo region. Puebla. Geofı´sica Internacional 36, 63 –76. Wilson, M., 1989. Igneous petrogenesis, Unwin Hyman, London, pp. 466. Ya´n˜ez, P., Ruiz, J., Patchett, P.J., Ortega-Gutie´rrez, F., Gehrels, G.E., 1991. Isotopic studies of the Acatla´n Complex, southern Mexico: implications for Paleozoic North American tectonics. Geological Society of America Bulletin 103, 817– 828. Zhou, M.F., Robinson, P.T., 1994. High-Cr and high-Al podiform chromitites, Western China: relationship to partial melting and melt/ rock reaction in the upper mantle. International Geology Reviews 36, 678 –686. Zhou, M.-F., Robinson, P.T., 1997. Origin and tectonic environment of podiform chromite deposits. Economic Geology 92, 259–262. Zhou, M.-F., Sun, M., Keays, R.R., Kerrich, R.W., 1998. Controls on platinum-group elemental distributions of podiform chromitites: A case study of high-Cr and high-Al chromitites from Chinese orogenic belts. Geochimica et Cosmochimica Acta 62, 677–688. Zhou, M.-F., Robinson, P.T., Malpas, J., Aitchison, J., Sun, M., Bai, W.-J., Hu, X.-F., Yang, J.-S., 2001. Melt/mantle interaction and melts evolution in the Sartohay high-Al chromite deposits of the Salabute ophiolite (NW China). Journal of Asian Earth Sciences 19, 517 –534.