al- and cr-rich chromitites from the mayar[iacute]-baracoa...

Economic Geology Vol. 94, 1999, pp. 547•566

A1- and Cr-Rich Chromitites from the Mayarf-Baracoa Ophiolitic Belt (Eastern Cuba): Consequence of Interaction between Volatile-Rich Melts and

Peridotires in Suprasubduction Mantle

JOAQUIN PROENZA, Departamento de Geologia, Instituto Superior Minero Metahirgico de Moa (Cuba), and Departament de Cristallografia,

Mineralogia i DepOsits Minerals, Facultat de Geologia, Universitat de Barcelona, Martl i Franques s/n, 08028 Barcelona, Spain

FERNANDO GERVILLA, {

Instituto Andaluz de Ciencias de la Tierra, Universidad de Granada-CSIC, Facultad de Ciencias, Avda. Fuentenueva s/n, 18002 Granada, Spain

JOAN CARLES MELGAREJO, Departament de Cristallografia, Mineralog•a i Dep6sits Minerals, Facultat de Geolog•a, Universitat de Barcelona,

Marti i Franques s/n, 08028 Barcelona, Spain

AND JEAN LOUIS BODINIER URM C,5569, Gdofluides-Bassins-Eau, ISTEEM, CNRS et Universitd de Montpellier II, Cc 57, Place Eugene Bataillon,

34095 Montpellier Cedex 05, France

Abstract

The Mayar/-Baracoa belt occupies the easternmost part of the east-west-trending Cuban ophiolitic belt. It comprises two large, chromite-rich massifs: Mayarf-Cristal and Moa-Baracoa. Neither of these massifs show a complete ophiolite sequence, but they consist of a part of an ideal section made up of(1) harzburgites grading upward into interlayered harzburgites and dunires, (2) interlayered harzburgites (with minor dunites) and gabbros, (3) gabbros, microgabbros, dolerites, and diabase dikes, and (4) pillowed basalt, cherts, and radiolarites.

Chromite deposits can be grouped into three mining districts according to the chemistry of chromite ore: the Mayarf district and the Sagua de T•namo district, both in the Mayarf-Cristal massif, and the Moa-Baracoa district in the Moa-Baracoa massif. The latter is the most important as it contains more than 5.5 million tons of ore. All chromitites mainly exhibit massive texture, show a pseudotabular, lenticular shape, and are concordant with the foliation of the enclosing harzburgites. In Moa-Baracoa they tend to occur in the mantle-crust transition zone, commonly contain dunitc and gabbro bodies oriented parallel to the elongation of the lenses, and are cut by late pegmatitic gabbro dikes. By contrast, in Mayarf, and to some extent in Sagua de Tanarno, chromitites occur deeper in the mantle tectonites and are cut by websteritc dikes. Intergranular minerals are olivine, serpentine, and chlorite. Chromite has abundant, randomly distributed solid inclusions of olivine and pargasite, and minor pyroxene, laurite, and millerite. Toward the contact with the included gabbros, chromitite from Moa-Baracoa shows increasing amounts of gabbro-related alteration products. Abundant clinopyroxene, partly altered plagioclase, and rutile occur as inclusions in the chromite. The composition of the chromite ore varies from typical refractory grade (Al rich) at Moa-Baracoa to metallurgical grade (Cr rich) at Mayaif, where the Cr no. ranges between 0.41 and 0.75, the Mg no. between 0.57 and 0.81, and the TiO2 content between 0.09 and 0.52 wt percent. At Moa-Baracoa, the Cr no. of chromite decreases and TiO2 content increases from harzburgite to dunitc and massive chromitite, positively correlated with the forsteritc content of coexisting olivine. At Mayari, both the Cr no. and TiO•. content of chromite, and the forsteritc content of olivine increase from harzburgite to dunitc and chromitite. Bulk platinum-group element abundances in chromitite vary from 20 to 538 ppb and show a broad positive correlation with Cr•O3 percent of chromite. The latter correlation is strongest in the Sagua de TSnamo district.

Structural, textural, mineralogical, and chemical characteristics of the studied chromite deposits, as well as the lithophile trace element geochemistry of their host rocks, support a genetic model based on the crystallization of chromite from different types of melts (from back-are basin basalts to boninitic andesites) at around 1,200øC, at variablefo•. Chromite formed when talc-alkaline melts, formed by melt-rock reactions at increasing melt volume, percolated through subhorizontal, porous dunitic channels and mixed with oxidized melts formed by low degrees of hydrous melting and low-temperature melt-rock reactions in suprasubduction zone mantle. Mixing of these two melts generated a hybrid melt whose bulk composition fell within the chromite liquidus field in the P-T-fo•. space (Hill and Roeder, 1974). Percolation of the hybrid melt through the dunitic channels promoted dissolution of preexisting silicate minerals and chromite crystallization. The Al-rich chromitites formed at the mantle crust transition zone at highfo• (-- logfo• = -7), whereas Cr-rich chromitites formed deeper in mantle tectonites under more reducing conditions, at logfo2 -- -10, depending on Cr contents of the parental magma.

*Corresponding author, email: [email protected]. es

0361-0128/99/2070/547-20 $6.00 547

548 PROENZA ET AL.

Introduction

POD•ORM chromitites from ophiolitic complexes are the only source of refractory-grade chromite (chromite ore with Al203 > 20 wt %, low Fe and (Cr203 + A12Os) > 60 wt %) and commonly occur within the same complex with chromitites of metallurgical-grade chromite (chromite ore with Cr•Os > 40 wt % and Cr/Fe = 2.2-4.0). Although both types of chromi- tires can be locally interspersed, they tend to occur at different depths in the mantle tectonites, with the Al-rich chromitites being the most shallow (Leblanc and Violette, 1983). The origin of this bimodal compositional distribution has been related to the crystallization of chromite from parental melts generated by (1) different degrees of partial melting (Arai, 1992), (2) progressive fractionation (Graham et al., 1996), and (3) partial melting or melt-rock reactions related to the percolation of melts through variably depleted peridotites (Zhou et al., 1994; Melcher et al., 1997; Zhou and Robinson, 1997). Most of the relevant research is focused on understanding both the nature and origin of the parental melts of chromitites. However, Arai and Yurimoto (1994), Zhou and Robinson (1997), and Zhou et al. (1994, 1996) have presented a model based on the mixing of a differentiated basaltic melt and a primitive melt in open fractures at shallow levels of the lithospheric mantle, and Melcher et al. (1997) suggest that the crystallization of chromite takes place at relatively low temperatures (<1,100øC) from highly oxidized melts (at fo• values i to 2 units above quartz-fayalite-mag- netitc). The former model can be only applied to Cr-rich chromitites, and the conditions inferred for the crystallization of chromite by Melcher et al. (1997) are far from the liquidus region of chromite in the P-T-fo2 space, even considering different melt compositions (Hill and Roeder, 1974).

In the Mayar/-Baracoa ophiolitic belt (eastern Cuba) there are abundant podiform chromitite deposits with calculated reserves up to 6.5 Mr. Their chemical composition is relatively homogeneous within the individual mining districts but exhibits a rather continuous variation from west to east (e.g., Cr•Os varies from 56-34 wt %), providing an ideal setting for investigating, in a single ophiolite complex, the origin of both refractory- and metallurgical-grade chromitites. The observed trend in chromite composition is correlated with increasing abundance of concordant gabbro layers within the chromite orebodies (especially in those with highly refractory ore) and with changes in the mineralogy and chemistry of late dikes that crosscut the chromitite lenses. The latter dikes vary from websteritc in the west to pegmatitic gabbro in the east.

In this paper we present and discuss the structural, petro- logical, and geochemical features of the representative chromitites and associated ultramafic rocks from the Mayar/- Baracoa ophiolitic belt in order to develop a genetic model in- tegrated with the geodynamic evolution of Cuban island arcs. This model aims to explain both the origin of the different parental melts and the mechanism of crystallization of Cr- and Al-rich chromitites.

The Mayar/-Baracoa Ophiolitic Belt The term "Mayar/-Baracoa belt" refers to the easternmost

part of the east-west-trending Cuban ophiolitic belt, which crops out for more than 1,000 km in the northern part of

Cuba and is thought to be Upper Jurassic-Lower Cretaceous in age (Iturralde-Vinent, 1996).

The Mayar/-Baracoa belt is an allochthonous pseudotabular body that is 170 km long, 10 to 20 km wide, and on average ca. 1,000 m thick, that was strongly faulted after its eraplace- ment. According to Iturralde-Vinent (1996) and Proenza (1998), it overlies island arc-related volcanic and sedimentary rocks of Cretaceous age and, in the easternmost part, schist, quartzite, marble, and amphibolite of the Asuncion terrane and the Gtiira de Janco Complex. Locally, it is covered by Paleogene island arc-related volcanic and sedimentary rocks and younger sandstone, shale, and limestone of the Bahamas platform. The Mayar/-Baracoa belt comprises two large, chromite-rich massifs: the Mayar/-Cristal massif and the Moa-Baracoa massif, and the small, barren massif of the Sierra del Convento. None of these massifs shows a complete ophiolite sequence but they consist of portions of an idealized ophiolitic section made up of (Iturralde-Vinent, 1996) (1) partly serpentinized harzburgites grading upward into interlayered harzburgites and dunites; (2) intedayered harzburgites (with minor dunires) and gabbros, olivine gabbros, and troctolites; (3) massive gabbros, microgabbros, dolerites, and diabase dikes; and (4) pillowed basalt flows interbedded with limestones and radiolarites.

The MayalS-Cristal massif extends over an area of 1,200 km • in the western part of the Mayari-Baracoa belt (Fig. 1). The western part of the massif is made up of partly serpentinized ultramafic rocks consisting of harzburgites (with minor lherzolites, dunites, and wehrlites) that vary upward into interlayered harzburgites and dunires. The latter contain lenticular bodies of chromitites which are locally cut by dikes of websterites. Ultramafic rocks are thrusted over by a crustal unit composed of scarce gabbros and microgabbros and abundant diabase dikes. The central and eastern parts display a very complex structure characterized by the presence of highly fractured monomictic and polimictic melanges where different ophiolitic rock types are mixed with volcanic and sedimentary rocks from the arc (Sagua de T•mamo region) or with pegmatites, granitoids, and variably metamorphosed mafic rocks (La Corea melange).

The Moa-Baracoa massif occupies the eastern part of the MayalS-Baracoa belt and underlies an area of ca. 1,500 km 2 (Fig. 1). It consists of a lower unit of mantle tectonites composed of harzburgites with minor dunites, plagioclase-bearing dunites, wehdites, pyroxenites, and scarce lherzolites. Toward the top of this unit, harzburgites contain increasing amounts of dunitc, gabbro, and chromitite, all of which occur as elon- gate, pseudotabular bodies oriented parallel to the foliation of the host harzburgites, as well as discordant dikes of pegmatitic gabbro (Thayer, 1942; Guild, 1947; Rios and Cobleila, 1984; Fonseca et al., 1985, 1992; Torres, 1987; Garcia and Fonseca, 1994, Kov•tcs et al., 1997). Although most of these authors interpret this upper zone as the lower part of the cumulate section of the ophiolite, it can also be considered to be part of the transition zone, in the sense of Nicolas and Prinzhofer (1983). The gabbroic complex always occurs in normal fault contact with the mantle tectonites and has a

lower unit of layered gabbro grading upward into massive gabbro. The latter locally contains tectonic blocks rich in diabase dikes (Torres, 1987). Basalts (locally with pillow

CHROMITITES, MAYAR[-BARACOA OPHIOLITIC BELT (EASTERN CUBA) 549

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FIG. 1. a. Distribution of ophiolite-related rocks in Cuba and location of the Mayaif- Cristal and the Moa-Baracoa massifs, which constitute the Mayaff-Baracoa ophiolitic belt. b. Sketch map of the Mayaff-Cristal massif, showing the distribution of chromite deposits in the Mayaft and Sagua de TSnamo districts (modified from Kravchenko and VSzquez, 1985; and Nekrasov et al., 1989). c. Sketch map of the Moa-Baracoa massif, showing the distribution of chromite deposits (modified from Nagy et al., 1976).

structure), chert, and limestone occur at the top of the crustal sequence (Iturralde-Vinent, 1996). Commonly, this sequence is tectonically inverted, where the upper volcanosedimentary unit is thrusted over by the gabbro complex and/or the mantle tectonites (Rios and Cobiella, 1984).

Chromite Deposits The chromite deposits of the Mayarf-Baracoa ophiolitic

belt can be grouped into three mining districts: (1) the Mayarf district, (2) the Sagua de Tanarno district, and (3) the Moa- Baracoa district. The Mayaft district is located in the western part of the Mayari-Cristal massif and includes two large deposits with more than 200,000 t of metallurgical-grade ore (Caledonia and Casimba), five deposits with slightly more than 10,000 t of ore, and 32 minor deposits (Lavaut et al., 1994). The Sagua de Tanarno district lies in the easternmost part of the Ma•varf-Cristal massif, in an area with a complex structure characterized by the imbrication of different tectonic sheets of ophiolite-related, mainly serpentinized ultramafic rocks (Fig. 1). This district contains 35 small deposits of both refractory (25 deposits) and metallurgic (10 deposits) chromite ore (Murashko and Lavandero, 1989). The Moa- Baracoa district is the most important from an economic point of view. It comprises the whole Moa-Baracoa massif and includes more than 100 chromite deposits of refractory-grade ore (Ostrooumov, 1986). Most of these deposits are of small

size; however, four (Amores, Loro, Yarey, and Piloto) contain more than 100,000 t of ore and the calculated reserves of the Mercedira mine exceed 5 Mr. In addition, the Cayo Guam and Potosf deposits were extensively exploited in the past (see Thayer, 1942; Guild, 1947) and they produced more than 800,000 t of ore.

Chromitite bodies have irregular, tabular to lenticular shapes (they are called lenses by the Cuban miners; Fig. 2) and are extremely variable in size. The central body of Mer- cedira is the largest, with a lateral extent of about 600 m, a width of 200 to 250 m, and a thickness of up to 20 m. These bodies are usually cut and rotated by normal faults giving rise to a variable set of orientations from northeast-southwest to east-west and southeast-northwest. More variable orientations

are observed in the Sagua de Tanarno district. The individual chromitite bodies studied are all parallel to the foliation and lineation of the enclosing peridotires, but Kovacs et al. (1997) mentioned that some others clearly cut these structures.

Chromitite lenses in Mayaft are always located within the interlayered harzburgite and dunitc unit, however, at Moa- Baracoa they occur within the mantle-crust transition zone, characterized by interlayered ultramarie and gabbroie rocks. Chromitites display variably thick (from some centimeters to several meters) dunitc envelopes and are hosted in harzburgites. Nevertheless, the chromitite can be locally in direct contact with harzburgite (Fig. 2). The contact between

550 eaOENZa ET

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FIG. 2. Vertical cross sections of the Mercedita chromite lens 1, drawn from drill-core data, showing the structural relationships between chromitite, dunitc, gabbros, and harzburgite. Drill holes are also drawn starting from the roofs of the mine galleries upward.

chromitite and dunitc is rather sharp or takes place along a zone of disseminated ore. Figure 2 also illustrates that some chromitites contain concordant tabular bodies of gabbro. The size of these bodies varies from a few centimeters up to 100 m long and 3 m thick, and-their contacts with the host chromitites are transitional, as described by Guild (1947) in the Cayo Guam deposit. This feature is characteristic of most chromite deposits of the Moa-Baracoa district, as well as of other deposits of refractory-grade ore like those from Cam- agiiey in central Cuba (Flint et al., 1948) and Coto in the Zambales ophiolite, Philippines (Leblanc and Violette, 1983). Dunitc lenses are also common in chromitite. In the Moa-

Baracoa district, gabbro (or troctolite) bodies occur within dunitc enclosing chromitite and in harzburgite and are always

parallel to the foliation of the enclosing rocks (Fig. 2). These concordant gabbros grade laterally into plagioclase- and clinopyroxene-bearing dunitc and, rarely, to plagioclase-lherzolite that display a transitional contact between the gabbro and the host peridotitc.

Several generations of gabbro (pegmatitic gabbro) occur as crosscutting dikes in the Moa-Baracoa chromitites and in some refractory-grade chromitites from Sagua de Tanarno. They are variable in thickness (up to 1.5 m), coarse grained (Guild (1947) describes a 90-cm single crystal of augitc), and, although they cut different rock types, are commonly restricted to chromitite bodies, as described by Flint et al. (1948) in the Camagiiey district in central Cuba. In contrast to the gabbro dikes, chromitites are cut by late dikes of


olivine-bearing, Cr-rich websterites in the Mayarf district. These pyroxenite dikes are generally thin (<20-30 cm) and show characteristic cumulate texture. They have equilibrium temperatures of 1,066 ø to 1,129øC according to Wood and Banno's (1973) geothermometer (Proenza, 1998).

Mineralogy and Texture of Chromite Ore and Distribution of Solid Inclusions in Chromite

In the inner part of the orebodies, chromite represents more than 90 vol percent and is subhedral to anhedral with grain sizes ranging from 0.5 to i cm. Intergranular minerals are olivine, serpentine (mainly antigorite), and chlorite with minor amounts of disseminated Fe-Ni sulfides, Fe-Ni alloys, native Cu and Au, and laurite. Chromite grains contain abundant solid, frequently biphase inclusions with variable sizes from 30 to 200/xm. The inclusions mainly consist of olivine and pargasite and, less frequently, clinopyroxene (in Mayarl chromitites) and platinum group minerals (mainly laurite). Platinum group mineral inclusions display rounded to euhedral (cubic or pentagonododecahedric) shape and are smaller than the silicate inclusions (Proenza, 1998).

At the contact between chromitite and dunitc, chromite grains show abundant inclusions of olivine which preserve the same optical orientation as the intergranular olivine. Toward the contact with gabbro inclusions, in the Moa-Baracoa chromitites, chromite contains increasing amounts of an intergranular material composed of a mixture of very fine grained phyllosilicates and some epidote group minerals, that represent the alteration products of the gabbro. Slightly away from the chromitite-gabbro contact, intergranular fresh clinopyroxene and partly altered plagioclase are present. Chromite has abundant minute inclusions (<100 /xm) of clinopyroxene (with the same optical orientation as that from the matrix) and partly altered plagioclase, both commonly associated with pargasite, as well as crystallographically oriented lamellae of rutile (following the {111} planes of the host chromite).

Primary chromitite textures are typically massive but locally grade to disseminated and commonly display pull-apart fractures. However, late shearing and faulting, as well as the intrusion of pegmatitic gabbros, have partly obliterated this texture, giving rise to mylonitic, cataclastic, and brecciated textures. Banded and nodular ores are also present, and many of them show no evidence of plastic deformation.

Enclosing peridotites are usually highly serpentinized, but some samples remain relatively fresh, with less than 20 vol percent of serpentine. Less altered harzburgite exhibits por- phyroclastic texture. In contrast, dunitc from the chromitite wall is strongly recrystallized and displays a coarse-grained granoblastic texture. Orthopyroxene porphyroclasts from harzburgite can be equigranular with grain sizes from 5 to 7 ram, although very commonly they are smaller and strongly deformed by intracrystalline gliding, showing kink bands and corroded grain boundaries within a recrystallized, undeformed olivine matrix. Similar textures have been interpreted by Dick (1977a and b), Quick (1981), Berger and Vannier (1984), and Kelemen (1990) as the result of melt-rock interaction involving dissolution of pyroxenes and crystallization of olivine. Chromite occupies less than 2 vol percent of the harzburgite and mainly occurs as subhedral, disseminated

grains in the olivine matrix and included in the orthopyroxene crystals, or as vermieular-shaped grains, sympleetitieally in- tergrown with orthopyroxene. The dunitie wall of the ehromitite contains up to 4 vol percent chromite showing rounded to euhedral habits and, less frequently, vermieular or anhedral shapes. In this dunite, olivine crystals can reach up to 1 em across.

Concordant gabbro bodies in dunite and harzburgite are typically undeformed, display eumulate textures, and frequently are finely layered comprising millimeter-thick, olivine-rieh layers alternating with olivine-free gabbro layers. They evolve laterally to elinopyroxene- and altered plagioelase-rieh dunite (rarely harzburgite). Clinopyroxene displays undeformed, coarse-grained crystals (1-3 ram), spinel has rounded to euhedrie habits, and both olivine and the seame orthopyroxene show typical curved grain boundaries (Proenza, 1998).

Mineral Chemistry of Chromitites Mineral compositions were determined by electron micro-

probe, using a CAMECA SX50 instrument in wavelength dis- persive mode, at the Serveis Cientifieo-T6enies of the Uni- versity of Barcelona (Spain). Excitation voltage was 20 kV, sample current 20 nA, and counting time 10 s per element. Ko• lines were monitored for all the analyzed elements. The standards used for ebromite analyses were ebromite (Cr, A1, Fe), perielase (Mg), rhodonite (Mn), ruffle (Ti), NiO, and metallic V. For silicate analyses, the standards used were or- thoelase (Si, AI, K), perielase (Mg), wollastonite (Ca), rhodonite (Mn), albite (Na), ruffle (Ti), Fe203, NiO, and Cr203. Minimum detection limits were 0.1 wt percent for all the analyzed elements. Fe•Oa in chromite was calculated by stoiehiometry, following the procedure of Carmichael (1967).

The chemical composition of ebromite from the different mining districts has been plotted in Figure 3A for the Cr no. (Cr/(Cr + A1)) vs. the Mg no. (Mg/(Mg + Fe)), and representative analyses are listed in Table 1. These results show that ebromite composition is highly variable (Fig. 3A). As a whole, the Cr no. varies from 0.41 to 0.75, the Mg no. from 0.41 to 0.81, and TiO• content from 0.1 to 0.52 wt percent (Fig. 3B).

The chemical composition of ehromitites from Mayar/ overlap the Cr-rieh part of this field: the Cr no. varies from 0.70 to 0.75, the Mg no. from 0.60 to 0.77, Fe•Oa content from 2.4 to 5.1 wt percent, and TiO• content from 0.1 to 0.2 wt percent (Fig. 3). No compositional variations have been observed between massive and nodular ehromitites, but disseminated ehromitites tend to have lower Mg nos.

At Sagua de T•namo, the composition of ebromite from the three deposits studied plots in a very large compositional field overlapping those of the Mayar/and Moa-Baraeoa ehromitites (Fig. 3). In ehromitites from the Santa Isabel deposit the Cr no. ranges from 0.63 to 0.72, the Mg no. from 0.57 to 0.67, Fe2Oa content from 0.44 to 3.90 wt percent, and TiO• content from 0.19 to 0.33 wt percent. These compositions resemble those of ehromites from the Mayar/district. In contrast, the chemical composition of chromite ores from Rupertina is similar to that of Moa-Baraeoa ehromitites: the Cr no. varies

from 0.46 to 0.48, the Mg no. from 0.71 to 0.74, Fe•Oa content from 2.0 to 3.01 wt percent, and TiO• content from 0.1 to 0.15 wt percent. Chromite ores from the Negro Viejo

552 PROENZA ET AL.

TABLE 1. Representative Electron Microprobe Analyses of Chromite from the Different Deposits Studied

Sample I G12-14 G15-27 L2-8 L6-1 AM-7 CG-8 PO-1 RU-1 NV-1 SI-1 CS-6 CA-BB CL-5 CP-24

TiO2 0.12 0.35 0.40 0.27 0.26 0.20 0.47 0.15 0.19 0.30 0.13 0.17 0.16 0.17 AlcOa 32.39 29.94 28.79 25.93 32.55 28.97 24.61 30.72 23.43 16.23 13.81 13.36 12.94 14.76 V2Oa 0.12 0.10 0.18 0.16 0.08 0.13 0.17 0.16 0.15 0.07 0.07 0.08 0.11 0.08 Cr•Oa 36.75 37.07 36.63 41.61 35.96 38.66 42.25 38.33 45.15 53.02 53.72 56.49 56.05 54.41 Fe•Oa 1.28 4.34 4.08 4.39 3.26 3.72 3.91 2.00 2.65 1.47 4.34 2.67 3.19 4.66 MgO 16.00 16.62 15.24 15.97 17.86 16.70 I5.23 16.56 14.61 12.43 13.23 14.69 14.09 16.54 MnO 0.16 0.27 0.17 0.27 0.21 0. I6 0.21 0.17 0.21 0.24 0.23 0.18 0.23 0.16 FeO 12.30 11.30 12.83 11.65 9.85 10.79 12.55 11.25 13.13 15.29 13.49 11.24 12.01 8.88 NiO 0.11 0.16 0.18 0.16 0.19 0.08 0.16 0.18 0.20 0.27 0.07 0.07 0.10 0.20 Total 99.24 100.14 98.51 100.41 100.21 99.40 99.57 99.51 99.71 99.32 99.09 98.96 98.88 99.85

Atomic proportions

Ti 0.02 0.06 0.07 0.05 0.05 0.04 0.08 0.03 0.03 0.06 0.02 0.03 0.03 0.03 A1 8.92 8.24 8.13 7.26 8.80 8.04 7.01 8.47 6.72 4.88 4.18 4.02 3.91 4.33 V 0.02 0.02 0.03 0.03 0.01 0.02 0.03 0.03 0.03 0.01 0.02 0.02 0.02 0.02 Cr 6.79 6.85 6.94 7.82 6.52 7.20 8.07 7.09 8.69 10.69 10.91 11.39 11.38 10.71 Fe a* 0.23 0.76 0.74 0.78 0.56 0.66 0.71 0.35 0.48 0.28 0.84 0.51 0.62 0.87 Mg 5.57 5.79 5.44 5.66 6.10 5.87 5.48 5.77 5.30 4.72 5.07 5.59 5.39 6.13 Mn 0.03 0.05 0.03 0.05 0.04 0.03 0.04 0.03 0.04 0.05 0.05 0.04 0.05 0.03 Fe •+ 2.40 2.21 2.57 2.32 1.89 2.13 2.54 2.20 2.67 3.26 2.88 2.40 2.58 1.85 Ni 0.02 0.03 0.03 0.03 0.04 0.02 0.03 0.03 0.04 0.06 0.02 0.01 0.02 0.04

Cr no. 0.43 0.45 0.46 0.52 0.43 0.47 0.54 0.46 0.56 0.69 0.72 0.74 0.74 0.71 Mg no. 0.70 0.72 0.68 0.71 0.76 0.73 0.68 0.72 0.66 0.59 0.64 0.70 0.68 0.77

Note: The complete set of analyses (around 500 spot analyses) may be requested from the corresponding author Samples location: Arnores = AM-7; Caledonia = CL-5 and CP-24; Casimba = CS-6 and CA-BB; Cayo Guam = CG-8; Mercedita = G12-14, G15-27,

L2-8 and L6-1; Negro Viejo = NV-1; Potosi = PO-1; Rupertina = RU-1; Santa Isabel = SI-1

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FIG. 3. A. Composition of chromite in terms of Cr no. (Cr/(Cr + A1)) vs. Mg no. (Mg/(Mg + Fe)). B. Cr no. vs. TiO2 content. Filled fields = chromitite from the three mining districts studied, thick squares = dunite from Moa-Baracoa,; squares = harzburgite from Moa-Baracoa, thick triangles = dunite from Mayari, triangles = harzburgite from Mayar/. Arrows connect the composition of accessory chromite in dunite with the associated chromitite.

0.60


deposit have intermediate compositions with the Cr no. between 0.51 and 0.58, the Mg no. between 0.61 and 0.68, Fe203 content between 1.82 and 3.27 wt percent, and TiO2 content between 0.14 and 0.22 wt percent.

Chromite composition of Moa-Baracoa chromitites plots in the Al-rich part of the compositional spectrum defined by chromitites from the entire Mayar/-Baracoa ophiolitic belt (Fig. 3). In the Moa-Baracoa chromitites, the Cr no. varies from 0.41 to 0.54, the Mg no. from 0.41 to 0.81, TiO• content from 0.05 to 0.52 wt percent, and Fe•O3 content from 0 to 4.8 wt percent. The TiO2 values are relatively high compared to the most common ophiolitic chromitites (< 0.25 wt % according to Leblanc and Nicolas, 1992) but are similar to other refractory-grade chromitites like those from Coto, Phillipines (0.26-0.41 wt % TiO•; M. Leblanc, pers. commun.) and Tari- Misaka, Japan (0.2-0.4 wt % TiO2; Matsumoto et al., 1997). At the contact between the Moa-Baracoa chromitites and

both the enclosing and included dunitc, the Mg no. of chromite decreases, probably due to subsolidus equilibrium between chromite and olivine during cooling. At the contact between chromite and gabbro lenses, Cr and TiO• contents of

chromite decrease, and A1 contents increase, progressively away from the contact to the chromitite (Fig.4A). These trends indicate that the crystallization of chromite was controlled by the presence of plagioclase as described by Leblanc and Violette (1983) in similar situations at Coto. Chromite composition varies in the opposite way near the contact between chromitite and the late, crosscutting pegmatitic gabbros (Fig. 4B). In this case, the intrusion of a volatile-rich basaltic melt is interpreted to have caused A1 enrichment and Cr depletion in the nearby chromite but did not affect the chromite Mg no.

Disseminated chromite from the enclosing dunitc at Ma- yar/has a Cr no. (0.70-0.74) similar to the associated chromite from the chromitite but lower Mg no. (0.42-0.53) and slightly lower TiO• contents (0.09-0.13). Accessory chromite from the associated harzburgite shows a lower Cr no. (0.56-0.69) and TiO2 contents (0.01-0.08 wt %), and a slightly higher Mg no. (0.48-0.59). These variations in Cr no. and TiO• content in chromite from harzburgite through dunitc to chromitite are similar to the trends generally described in ophiolitic tectonites (see Leblanc et al., 1980; Zhou et al., 1996; Arai,

Gabbro sill

8.5 I

• 8.0 -- .2

• 7.5 -

7.0

6.5

1 2 3 4 5 6 7 8 9 101112131415

Analyzed points

0.70 '

0.60

•o.5o ß

•: 0.40

0.20 / I I [ I I I I I I I I I 1 2 3 4 5 6 7 8 9 101112131415

Analyzed points

9.0 --

7.5 - _

7.0 - _

6.5 -• --

6.0

Cr

1 2 3 4 5 6 7 8 9 101112131415

Analyzed points

0.50 --

0.45

0.40

0.35

0.30

1 2 3 4 5 6 7 8 9 101112131415

Analyzed points

FIG. 4. Compositional variations of chromite in chromite lenses of the Mercedita mine from the contact with the enclosed, concordant gabbros (above) and the discordant gabbro dikes (below). The sketch drawing on the left side represents single thin sections. Mg and Fe proportions are not represented because they do not show significant variations in the different sections analyzed.

554 PROENZA ET AL.

1997). In contrast, chromite from the dunite walls in the different deposits of the Moa-Baracoa district shows a higher Cr no. (0.44-0.58), a lower Mg no. (0.49-0.66) and lower TiO2 content (0.14-0.38) than that from the enclosed chromitites. In addition, accessory chromite from the associated harzburgite shows an even higher Cr no. (0.45-0.65), and lower Mg no. (0.39-0.64) and TiO2 (0-0.15) than those from the dunites (Fig.3). Variations in the Mg no. depend on the chromite/ olivine ratio since they are the result of subsolidus Mg-Fe exchange between olivine and chromite on cooling. In fact, equilibrium temperatures based on Mg-Fe exchange between olivine and chromite, using the calibrations by Fabrics (1979), Roeder et al. (1979), and Ballhauss et al., (1991a, b), show that Mg and Fe diffusion was effective down to temperatures as low as 600øC (Proenza, 1998). On the contrary, the variations of Cr no. and TiO2 contents of chromite from harzburgite through dunitc to chromitite in Moa-Baracoa (Fig. 3) have, to our knowledge, never been reported in ophiolitic complexes and cannot be explained in terms of partial melting or fractional crystallization.

Forsteritc proportions in olivine range between 90 and 97 mole percent, increasing progressively from harzburgite to chromitite, and correlate positively with NiO, which varies from 0.30 to 0.81 wt percent, following a trend similar to that described by Leb]anc et al. (1984) in New Caledonia. The composition of orthopyroxene from harzburgite contains more molar enstatite (90.4-92.9 mole %), and slightly less CaO (0.6-1.5 wt %) and A1•O3(1.3-1.5 wt %) at Mayalff than at Moa-Baracoa (87.2-90.3 mole % enstatite, 0.8-2.5 wt % CaO, and 1.3-2.6 wt % A1203).

Distribution of Platinum-Group Elements Whole-rock, platinum-group elements (PGE), and Au

analyses were performed by Activation Laboratories Ltd. (ACTLABS, Ontario, Canada). The procedure involves INAA after nickel sulfide fire assay collection. Detection limits were 0.1 ppb Ir, 0.5 ppb Au, i ppb Rh, 2 ppb Os and Pd, and 5 ppb Ru and Pt.

Because of the economic importance of Mercedira relative to the other deposits of the Mayar/-Baracoa belt, we have analyzed 16 samples, each of which is representative of several kilograms of massive chromite from the central body and

lens 7. The results show that these chromitites have low total

PGE tenors that range from 55 to 166 ppb and average 90 ppb. Other small samples from Mercedira (32.4 and 114 ppb), Arnores (20.3 and 44.4 ppb), Cayo Guam (100 ppb), and Potos/(89.5 ppb), have total PGE contents either below or within the described range of Mercedira (Table 2). PGE abundances of Moa-Baracoa chromitites are in the range 0.004 to 0.07 x chondritic values, and the chondrite-normalized patterns have a negative anomaly in Ir with respect to Os and Ru (except in sample L2-8 that shows a relatively fiat pattern from Os to Ru) and variable negative slopes from Ru to Pd (Fig. 5).

In the four samples analyzed from Mayar/, the distribution of PGE is strongly dependent on the texture and on the chromite/olivine ratio. Sample CA-B is a nodular textured ore that has the lowest total PGE tenors (48.5 ppb, Table 2). This sample shows a nearly fiat chondrite-normalized pattern from Os to Rh (with a small, positive anomaly in Ru) and a negative slope from Rh to Pd (Fig.5). In the other three samples, both the bulk PGE contents and the ratio (Os + Ir + RU)N/•PGE N increases with increasing modal proportions of chromite. The sample with the highest total PGE tenor (CS- 6 = 227.2 ppb, Table 2) has a massive texture and shows a negative-sloped, chondrite-normalized pattern from Os (0.2 x chondritic values) to Pd (0.0037 x chondritic values).

Total platinum-group element contents of the three samples analyzed from Sagua de TSnamo district are very different: 18.3 ppb from Rupertina, 105.5 ppb from Negro Viejo, and 537.5 ppb from Santa Isabel (Table 2). The increase in the bulk PGE content correlates with a variation in the shape of the chondrite-normalized pattern (Fig. 5B), mainly due to an increase in the ratio (Os + Ir + RU)N/•PGE N (0.048, through 0.33-1.83), which is reflected by the negative slope of the pattern from Ru to Pd. A comparison of Figures 3 and 5 shows that the Cr no. of chromite ore has a positive correlation with total PGE tenors. This correlation is a characteris-

tic feature of chromite ore from Sagua de TSnamo and to a lesser extent of the chromitites of the entire Mayarf-Baracoa ophio]itic belt.

Gold is systematically increased with respect to Pd in the analyzed samples. The relative enrichment of gold increases from the massive-textured samples through the nodular to

TABLE 2. Whole-Rock Platinum-Group Element Contents (ppb) of Representative Samples from the Different Deposits Studied

Sample• Mercedita G15-27 L2-8 AM-7 AM-9 CG-8 PO-1 NV-1 RU-1 SI-2 CS-6 CA-B CL-5 CL-3

Os 19.7 25.0 7.0 3.0 11.0 13.0 18.0 17.0 3.0 110.0 99.0 8.0 11.0 19.0 Ir 8.0 14.0 6.4 2.2 4.8 9.9 15.0 19.0 2.8 210.0 55.0 9.5 18.0 24.0 Ru 45.6 46.0 10.0 7.0 20.0 40.0 45.0 55.0 5.0 120.0 61.0 20.0 48.0 25.0 Rh 3.1 7.3 2.0 1.1 1.6 5.1 3.3 4.5 0.5 42.5 2.2 4.0 5.3 6.8 Pt 10.3 5.0 5.0 5.0 5.0 22.0 5.0 5.0 5.0 43.0 8.0 5.0 5.0 10.0 Pd 3.8 17.0 2.0 2.0 2.0 10.0 2.0 5.0 2.0 12.0 2.0 2.0 2.0 11.0 Au 6.1 5.1 1.7 0.5 3.5 2.3 1.2 1.3 0.5 1.0 0.6 0.9 2.5 8.7

Total PGE 90.4 114.3 32.4 20.3 44.4 100.0 89.5 106.8 18.8 538.5 227.8 48.5 89.3 95.8

(Os + Ir + Ru)N/ PGEN 2 0.79 0.66 0.68 0.59 0.78 0.61 0.84 0.80 0.62 0.74 0.94 0.69 0.78 0.65

1 Mercedita: average of 16 samples from lenses 1 and 7 of Mercedita mine; Other sample locations: Amores = AM-7 and AM-9; Caledonia = CL-5 and CL-3; Casimba = CS-6 and CA-B; Cayo Guam = CG-8; Mercedita = G15-27 and L2-8; Negro Viejo = NV-1; Potos/= PO-1; Rupertina = RU-1; Santa Isabel = SI-2

'2(Os + Ir + Ru)N/PGE N indicates the ratio between the normalized values of (Os +Ir + Ru) and the normalized bulk PGE content


i

LU 1E-2-

v

o

1E-3

Moa-Baracoa district

G15-27 •A Average Mer•ta ' /,

, •AM-7 ß

Sagua de TSnamo district

ß

SI-2

RU-1

Mayari district

"'CA-Ba

.

I I

Au Os Ir

A

&

I I I I I I I I I I I I I I I

Os Ir Ru Rh Pt Pd Au Os Ir Ru Rh Pt Pd Ru Rh Pt Pd Au

FIG. 5. Chondrite-normalized diagrams for the platinum-group element and gold contents of representative chromitite samples from the different mining districts of the Mayar/-Baracoa ophiolitic belt. Sample locations are listed in Table 2. Nor- malization values are from Naldrett and Duke (1980).

the disseminated and correlates positively with modal pro- portion of serpentinized olMne in the rock. This correlation agrees with gold abundances in serpentinites (Leblane, 1991) and with the presence of minute grains of gold in the serpentine matrix of the studied ehromitites (Proenza, 1998). The Pd increase in the Au-rieh samples is related to the existence of significant amounts of secondary sulfides (pentlanditc, heazlewoodite, and millerite) associated with the serpentine minerals.

Lithophile Trace Elements Twelve samples representative of different chromitite-

associated peridotites from the Mayarf district and the Mer- cedita mine were analyzed for REE and other trace elements (Sc, Rb, Sr, Zr, Nb, Ba, Hf, Th, and U) by inductively coupled plasma-mass spectrometry (ICP-MS). These analyses were performed at Montpellier, on a VG PlasmaQuad instrument, with the analytical procedure described by Ionov et al. (1992). The results are shown in Figures 6 and 7 on chondrite- and primitive mantle-normalized diagrams. The whole data set is available from the corresponding author.

Similar to other refractory peridotites from ophiolitic complexes (e.g., Prinzhofer and All•gre, 1985; Sharma and Wasserburg, 1996), the analyzed samples have low REE abundances, with HREEN (chondrite-normalized heavy REE) < 0.3, and LREEN and MREEN (chondrite-normalized light and middle REE, respectively) _< 0.05 (Figs. 6 and 7). The samples from Mayarf (Fig. 6) are characterized by marked U-shaped chondrite-normalized REE patterns, and are significatively more depleted in MREE than in HREE and LREE. This signature is typical of refractory peridotites

from various suites of mantle rocks, ophiolites, orogenic lherzolite massifs, and several suites of mantle xenoliths (McDo- nough and Frey, 1989, and references therein), as well as refractory peridotires from fore-arc environments (e.g., Parkinson and Pearce, 1998). The samples from Mercedira (Fig. 7) show REE patterns with a flat segment for LREE and MREE (from La to Gd) and a positive slope for HREE (from Gd to Lu). Similar to some refractory peridotires from orogenic lherzolite massifs (Van der Wal and Bodinier, 1996), several mantle xenoliths (e.g., Alard et al., 1996; Bedini et al., 1997; Lenoir et al., 1997) and fore-arc peridotires (Parkinson and Pearce, 1998). The studied samples (Figs. 6 and 7) are enriched in highly incompatible elements (Cs, Rb, Ba, Th, U, and Nb: LILE) relative to REE (e.g., Thm/La m = 0.6-8.0). Furthermore, they are selectively enriched in U relative to Th (UN/Th m = 1.3-13.3), and some of them are also enriched in Sr relative to LREE (SrN/CeN = 0.7--9.8). The high U/Th signature has been recently shown to be typical of many mantle xenoliths (Alard et al., 1998) and some peridotires, pyroxen- ires, and mineral separates (including olivine and spinel) from the Ronda orogenic lherzolite (Garrido, 1995; Garrido and Bodinier, 1999; Gervilla et al., current study).

The differences between harzburgites and dunires from Mayar/are that dunires are relatively enriched in LREE and MREE relative to HREE and in highly incompatible elements relative to REE. LaN/YbN is in the range 0.6 to 0.8 and Ths/Lam is in the range 2.7 to 4.5 for the dunires, whereas these values are 0.3 to 0.4 and 1.3 to 1.6, respectively, for the harzburgites. In contrast, in the samples from Mercedira these differences are less important for REE but significant for some highly incompatible elements. In particular, the

556 PROENZA ET AL.

MAYARI HARZBURGITES

1

MAYARI DUNITES

0.1 0.1

0.01 0.01

La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

10 10

MAYARI HARZBURGITES

c- 1 1

._> • 0.1 0.1 E

.B

0 0.01 0.01 o

MAYARI DUNITES

0.001 0.001

CsRbBaTh U NbLaCePr Sr NdZrHfSmEuGdTbDy Y HoErTmYbLu CsRbBaTh U NbLaCePr Sr NdZrHfSmEuGdTbDy Y HoErTmYbLu

FIG. 6. Chondrite-normalized REE diagrams (above) and primitive mantle-normalized trace element diagrams (below) of some dunites and harzburgites from the Casimba mine (Mayarf district). Normalization values are from Sun and McDonough (1989).

dunites are less enriched in Th relative to REE. ThN/LaN is in the range 0.6 to 2.7 for the dunires, compared to 3.8 to 8.0 for the harzburgites. Note that this variation tends to be the opposite of the one observed in MayarL

Discussion

Structural and textural constraints

The structural and textural data reported above clearly show that the studied chromitite bodies from the Mayarf- Baracoa ophiolitic belt are all concordant with the foliation and lineation of the enclosing peridotites. This allows for their classification within the concordant structural type defined by Cassard et al. (1981). These authors suggest that such structures derive from crosscutting chromitite pods originating at a midoceanic ridge (see Lago et al., 1982), which were later stretched and transposed by plastic flow, thereby becoming concordant to the fabric of the enclosing peridotites. This interpretation is in agreement with the presence of pull-apart fractures in most massive chromitites. However, it does not explain the abundance of weakly deformed (or undeformed) olivine at the contact between chromitite and its dunitic wall, considering that plastic deformation is concentrated mainly in the silicate minerals. Olivine exhibits a coarse-granular,

recrystallized texture in the dunitic walls and in most harzburgites. The latter have strongly deformed and partly dissolved orthopyroxene crystals included in a coarse-granular recrystallized olivine matrix, suggesting that harzburgites formed by melt-rock reactions that involved dissolution of pyroxenes and recrystallization of olivine. These melt-rock reactions post- dated deformation. In addition to these objections, the chromitites from the Moa- Baracoa district enclose concor-

dant gabbro bodies and dunitc lenses with gradational contacts and, locally, the chromitites cut the dunite-gabbro contact (Fig. 2), suggesting that the crystallization of the chromitite overprinted preexisting lithologies. Moreover, the presenee of large gabbro bodies within the ehromitites makes it diffieult to envisage a meehanism for the formation of ehromitite and gabbro in an open dike as well as the ereation and maintenanee of an elutriation eell within the dike as suggested by Lago et al. (1982).

Furthermore, the petrology and mierostruetures of eoneor- dant gabbros in dunite and harzburgite suggest that late melts whieh impregnated the preexisting refraetory peridotite gave rise to melt-roek reaetions at deereasing melt volume (asso- eiated with eonduetive heat loss) and resulted in the dissolution of olivine and the erystallization of elinopyroxene and plagioelase according to Bodinier (1988) and Kelemen et al.


1 1

MERCEDITA HARZBURGITES MERCEDITA DUNITES

(.• 0.1 0.1 (.,3 o

O.Ol O.Ol

La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

lO lO

' MERCEDITA HARZBURGITES MERCEDITA DUNITES

•'• 0.1 0.1

• 0.01 0.01 0.001 - 0.001

Cs Rb BaTh U Nb LaCe Pr Sr Nd Zr Hf Srn Eu Gd Tb Dy Y Ho ErTrn Yb Lu Cs Rb BaTh U Nb LaCe Pr Sr NdZr Hf Srn Eu GdTb Dy Y Ho ErTrnYb Lu

FIG. 7. Chondrite-normalized BEE diagrams (above) and primitive mantle-normalized trace element diagrams (below) of some dunites and harzburgites from the Mercedita mine (Moa-Baracoa district). Normalization values are from Sun and McDonough (1989).

(1995b). A similar process has been described in the Ronda ultramafic massif (Spain) by RemaYdi (1993), Van der Wal and Bodinier (1996), and Bodinier et al. (unpub. data), and in the East Pacific Rise mantle by Dick and Natland (1996) and Arai and Matsukage (1996).

All these features seems to indicate that chromite forma-

tion was coeval with or occurred slightly after the recrystallization of the host harzburgites and dunires, which overprinted most of the microstructures related to plastic deformation during mantle flow. Recrystallization could be associated with the intergranular percolation of deep-seated melts that reacted with the host peridotire first dissolving pyroxenes and crystallizing olivine (forming orthopyroxene-poor harzburgite and dunite) and later, during the cooling of the system, dissolving olivine and crystallizing clinopyroxene and plagioclase.

Chemical variations of chromite in chromitite and peridotite The compositional variations of chromite from the Moa-

Baracoa district have not previously been reported, whereas those from the Mayar/district are similar to most trends described in ophiolitic complexes (Fig.3; see Irvine, 1967; Leblanc and Nicolas, 1992; Arai, 1997). Chemical variations of chromite in the Mayar/-Baracoa belt can be seen better by

comparing the average Cr no. and TiO2 content of chromite with the average forsterite (Fo) content in olivine (Fig. 8). This figure shows: (1) that different lithologies have different TiO2-Fo relationships, chromitites having the highest Fo and, excluding some samples from Mercedita, TiOs contents; (2) no significant correlation between the Cr no. of chromite and Fo content in olivine at the Mayaft district; and (3) a negative correlation between the Cr no. of chromite and Fo content in

olivine at the Mercedira mine, which is consistent with the chemical variations of chromite from harzburgite to dunire and to chromitite in most deposits of the Moa-Baracoa district (Fig. 3A). The variations of Fo content in primary olivine had to be smaller at high temperatures, before subsolidus reequilibration, suggesting that the chemical variations in the crystallizing chromite took place at almost constant Fo content of coexisting olivine.

Considering the geochemical behavior of A1, Cr, Ti, and Mg during magmatic processes, and assuming a genesis of dunite by partial melting of harzburgite, a progressive increase in Fo content in olivine should be expected as the associated chromite was impoverished in A1 and Ti. In contrast, fractional crystallization should produce a decrease of Fo content in olivine with a progressive increase of A1 and Ti in chromite. This implies that to understand the chemical disequilibrium

558 PROENZA ET AL.

0.5

0.4

0.3

09 0.1

0.0

0.8

0.7

0.6 o

rj') 0.5

0.4

[111.

I I I I

[]

.lilt I I I I

0.90 0.92 0.94 0.96 0.98

Olivine Mg/(Mg+Fe) FIC. 8. Average Cr/(Cr + AI) and TiO.2 contents of chromite grains vs. the

average Mg/(Mg + Fe) of the nearby olivine grains in harzburgites (open symbols), dunites (thick symbols), and ehromitites (filled symbols). Squares = samples from Mereedita, triangles = samples from Mayarf.

between chromitite and its host peridotire, it is necessary to consider magmatie processes others than partial melting and fractional crystallization. According to McKenzie (1984), the composition of a melt that migrates intergranularly and reacts with the percolated peridotire will have its compatible element contents (e.g., MgO) buffered by the peridotire matrix, whereas the incompatible element abundances (e.g., TiO0 will be controlled by the composition of the infiltrated melt. This mechanism is similar to the ehromatographie effect proposed for trace elements (Navon and Stolper, 1987; Bodinier et al., 1990). In addition, Kelemen (1990) has shown that melt- peridotite reactions at increasing melt volume produce the dissolution of pyroxenes and the formation of Fo-rieh olivine in equilibrium with a new, SiOn.-, Al•.Oa-, and TiO2-enriehed

melt of eale-alkaline affinity. As the reactions progress toward the upper part of the mantle section, the new eale-alkaline melt becomes selectively enriched in the most incompatible elements relative to the less incompatible elements by a ehromatographie effect (Navon and Stolper, 1986; Bodinier et al., 1990).

Assuming that chromite was in equilibrium with its parental melt, the chemical changes in chromite reflect those of the melt. Thus, the composition of dunite from Moa-Bara- eoa can be interpreted by the reaction at increasing melt volume of the harzburgite with a percolating, A1- and Ti-rieh melt near the upper part of the mantle section (just below the gabbroie eumulates). The formation of ehromitite implies a further supply of A1 and Ti to the melt in equilibrium with dunite. At Mayar/, dunite can be formed either by partial melting or, most probably, by melt-rock reactions at deeper levels of the mantle section. The parental melt of ehromitite had essentially the same composition (slightly enriched in TiO0 as that equilibrated with dunite. Parental melts

The experimental results of Hill and Roeder (1974), Mau- rel and Maurel (1982), and Roeder and Reynolds (1991) have shown that neither pressure changes nor temperature changes significatively affect the chemistry of the crystallizing chromite, but fro strongly controls the solubility of chromium in a given basic melt (Roeder and Reynolds, 1991).

The chromite from the studied deposits in the Mayari- Baracoa belt exhibits a limited range of Mg no. values and very large variations in Cr no. values (Fig. 3). This compositional range of chromite mostly overlaps that of ophiolitic chromitites (Irvine, 1967) and extends from the field of chromite crystallized from midocean ridge basalt (MORB)- type melts to that of chromite crystallized from boninites in the Cr no. vs. TiO•. diagram of Arai (1992). A further insight into the chemistry of the parental melts for the individual chromites studied can be made using, respectively, the equations by Maurel and Maurel (1982) and Maurel (1984, in Aug& 1987):

(A1.2Oa wt %)sp = 0.035 (A1203 wt %)liquid 2'42, and

ln(FeO/MgO)s r = 0.47-1.07XAl•r + 0.64XFe•j + ln(FeO/MgO)•q,ia, where

XAlsp = [A1/(A1 + Cr + Fea+)] and XF%af = [FeS+/(A1 + Cr + Fe3+)]. These equations show that the melts in equilibrium with

the chromitites from the MayaN district had an average AI•.O 3 content of 11.9 wt percent and an FeO/MgO ratio equal to 0.74 (Table 3). This Al•.O3 value is lower than that of the Santa Isabel deposit from the Sagua de T•namo district (19..9 wt %) but both are comparable with the Al•.O3 content of boninitic melts (Bloomer and Hawkins, 1987; Wilson, 1987) and that of the melts in equilibrium with the chromitites from Nan Ut- taradit, Thailand, which have been interpreted as crystal seg- regates from boninitic melts (Oberger et al., 1995). Although the values of the FeO/MgO ratio are affected by the subsolidus equilibrium between chromite and olivine, the results listed in Table $ are in agreement with the interpretation


TABLE 3. Average A1203 Contents and FeO/MgO Ratio of Melts in Equilibrium with the Chromitites of the Mayarf-Baracoa Belt and Other Ophiolitic Complexes, as well as of Various Arc-Related Magmas

Location A12031 FeO/MgO • References

Moa-Baracoa district 16.4 0.9

Sagua de Tfinamo district Rupertina 16.0 1.0 Negro Viejo 14.8 0.8 Santa Isabel 12.9 0.7

Mayaft district 11.9 0.7

Oman 11.4-14.4 0.6 Nan Uttaradit 11.6-12.0 Loubusa 9.8-11.0

Kempirsai Main ore field 9.0-12.2 0.3-0.8

Batamshinsk type 13.5-16.7 0.8-1.0

Boninita 10.6-14.4 0.7-1.4 MORB 16.0 1.2-1.6 Back-arc basin basalts > 16.0

Auge (1987) Oberger et al. (1995) Zhou et al. (1996) Melcher et al. (1997)

Wilson (1987), Bloomer and Hawkins (1987) Wilson (1987) Fryer et al. (1990)

1 Maurel and Maurel (1982) 2 Maurel (1984), in Augd (1987)

given above. The average A1203 contents of the melts in equilibrium with the ehromitites from the Moa-Baraeoa district

(16.4 wt %) and Rupertina (16 wt %) are comparable with those of the parental melts of Al-rieh ehromitite of the Batamshinsk type in the Kempirsai massif (Meleher et al., 1997) but mainly with those of midocean ridge and back-are basin basalts. However, their FeO/MgO ratio is lower than that of MORB, possibly suggesting a back-are basin basalt affinity (Table 3).

The presence of abundant inclusions of pargasite in chromite irrespective of the chemical composition of the host chromite shows that the formation of the different deposits of the Mayarf-Baraeoa belt took place from volatile-rich melts, in agreement with the interpretations of Johan et al. (1983), MeElduff and Stumpfl (1991), and Meleher et al. (1997) for other ehromitite deposits. The presence of volatile-rich, probably hydrous melts varying in composition from back-are basin basalts (or MORB) to boninites is a characteristic feature of suprasubduetion zones, which is in agreement with the current interpretation of the northern ophiolites of Cuba (Iturralde-Vinent, 1986). Furthermore, suprasubduetion zones are notably favorable for the development of melt pereolation-reaetion systems in upper mantle peridotites giving rise to the formation of highly refractory rocks, associated eale-alkaline magma series (Kelemen, 1990), and ehromitite bodies (Roberts, 1988; Arai and Yurimoto, 1994). In fact, the only reported example of basalt-borne xenoliths of mantle ehromitites come from back-are alkali basalts in the Japan island ares (Arai and Abe, 1994).

Trace element geochemistry Similar to many refractory peridotites from various geody-

namic settings, the studied peridotires have a LREE-enriched signature hardly consistent with standard melting models such as batch, fractional, continuous, and incremental melting (Cast, 1968; Shaw, 1970; Langmuir et al., 1977; John- son et al., 1990). On the other hand, the hypothesis that LREE were increased during secondary alteration is very

unlikely, as demonstrated by (1) harzburgitie mantle xenoliths with no trace of alteration, which are nevertheless LREE enriched and show U-shaped REE patterns very similar to those observed, for instance, in the Mayari peridotites (e.g., A_lard et al., 1996; Lenoir et al., 1997); (2) heterogeneous suites of mantle rocks showing very contrasted REE patterns in refraetory peridotites and coexisting lherzolites in spite of their similar degree of serpentinization (LREE enriched vs. LREE depleted, respectively; McDonough and Frey, 1989); (3) acid- leached pyroxenes separated from serpentinized refractory peridotites showing similar LREE enrichment as whole rocks (RemaXdi, 1993).

Most often, the U-shaped REE patterns have been explained by two-stage processes involving partial melting followed by interaction with LREE-enriehed melt-fluid (i.e., cryptic metasomatism). Navon and Stolper (1987) have demonstrated with a ehromatographie numerical approach that intergranular melt pereolation may result in strongly fraetionated REE compositions. However, some authors have suggested that the selective enrichment of LREE in refractory peridotites is a character which is too widely distributed to be just incidental, i.e., tributary of late enrichment processes. They also pointed out that the two-stage model does not account for the negative correlation between LREE/HREE ratios and CaO observed in several suites of

mantle rocks (McDonough and Frey, 1989). This correlation suggests that the selective LREE enrichment and the individualization of the refractory peridotites are coeval. In the single-stage models, the U-shaped REE patterns are considered to reflect MREE and HREE depletion coupled with depletion of basaltic constituents, whereas LREE tend to remain constant (Godard et al., 1995). With the exception of the disequilibrium melting mechanism proposed by Prinzhofer and Allbgre (1985), the processes envisioned to explain the relatively narrow range of LREE variations in U-shaped, residual peridotites call upon the effect of melt transport. Diffuse migration of melt is considered to buffer the most incompatible elements, whereas the other elements are

560 PROENZA ET AL.

controlled par melting-reaction processes (Bodinier et al., 1991). Single-stage processes include melt extraction by pervasive porous flow (Verni•res et al., 1997), open-system melting (Johnson and Dick, 1992; Ozawa and Shimizu, 1995; Suhr et al., 1998), and melt-rock reactions (Godard et al., 1995; Verni•res et al., 1997).

Whereas the two-stage model is relevant to the case of strongly LREE-enriched mantle xenoliths with variable LREE content and depleted isotopic signature, the single- stage model is more justified in the case of refractory peridotites with U-shaped REE patterns and a restricted range of LREE variations, such as those analyzed for this study. In suprasubduction mantle, single-stage processes may reflect melting or melt-rock reaction triggered and/or enhanced by the inf'fitration of slab-derived, volatile-rich melts (Kelemen, 1990, and references herein). It is worth noting that the single-stage model may also account for the increase of highly incompatible elements (Cs, Rb, Ba, Th, U, and Nb) relative to REE, either because slab-derived melts are originally enriched in these elements, or because they become enriched during diffuse percolation of small melt fractions (McKenzie, 1989; Bedini et al., 1997). Although the distribution of these elements may have been somewhat modified during secondary alteration, their overall enrichment is likely of primary, mantle origin. This is shown, for instance, by the significant increases observed for Th and Nb, two elements that are commonly considered as rather insensitive to alteration. On the other hand, the positive anomalies observed for U and Sr--two mobile elements--might be related to secondary processes. Yet, recent ICP-MS data indicate that moderate anomalies of these elements (e.g., UN/ThN < 50) are commonly observed in mantle xenoliths, even when they are de- void of alteration products (Ionov et al., 1995; Alard et al., 1998). In contrast, altered, abyssal peridotites have much more pronounced U anomalies, with UN/ThN ratios up to 5,000 (Godard and Bodinier, 1998). On the basis of experimental works (Keppler and Wyllie, 1990; Brenan et al., 1995), the U (ñ Sr) anomaly could be ascribed to interaction with C1- and/or COs-bearing, water-rich, small melt fractions, probably derived from a subducted slab (Keppler, 1996).

Significant differences in LREE and highly incompatible element enrichments according to localities (Mercedita vs. Mayari) and rock types (harzburgites vs. dunites) are indica- tive of variations either in the nature and degree of peridotitc melt interaction, or in the composition of infiltrated melt. Ex- cept for one sample (86-2-23, from Mercedita), all Mayari peridotites and Mercedita harzburgites define a negative correlation between ThN/LaN and LaN/SmN, from ThN/LaN = 7.5 and LaN/Sm N = 0.4 to ThN/LaN = 1.3 and LaN/SmN = 2.25 (Fig. 9). This correlation can be explained in several ways, all consistent with melt processes in suprasubduction mantle. The simplest interpretation is to consider that the different peridotites were percolated by melt fractions that were originally homogeneous, but melt-rock ratios were higher at Ma- yari compared to Mercedita. In other words, either Mayari peridotites have undergone higher melt flow, or percolation lasted longer in this locality. This model, which involves no reaction, may account for significant trace element fractionation provided that melt fractions are small and the distance of percolation long enough (Navon and Stolper, 1987; Bodinier

5 MAYARI

z harzb'urgite$ • "• J I •".• 4 / •dL,•ite$ e' MERCED•ITA • •-- 3 '•... • '• harzburgites

2 du3iteSo•o •""• 1

0 i i i i

0.5 1 1.5 2 2.5

LaN/SmN

FIC. 9. Variations of ThN/LaN vs. LaN/Sm• in the analyzed harzburgites and dunRes from Mayar/and Mercedita.

et al., 1990). A further alternative, based on the chromatographic approach, is to envisage infiltrated melts of variable compositions, due to previous percolation through deeper mantle rocks. In this scheme, melts inœdtrated in Mercedita harzburgites had percolated through larger volume of mantle rocks, compared with Mayari melts. More complex--and probably more realistic--models involve open-system melting or melt-rock reactions coupled with porous flow (Kele- men, 1990; Kelemen et al., 1993). Based on numerical mod- eling of reactive porous flow (Godard et al., 1995; Verni•res et al., 1997), the negative correlation between Th/La and La/Sm (Fig. 9) may be explained by spatial (and/or temporal) evolution between two end-member processes:

1. Migration of volatile-rich, small volume melts across relatively cold mantle (Mercedita harzburgites). This process involves near-solidus reactions, whereby percolating melts preserve their primary (slab-derived) enriched composition. Alternatively, they may be secondarily enriched due to reactions at decreasing melt mass (Bedini et al., 1997).

2. Reactive flow of larger melt fractions in relatively hot mantle (Mayari harzburgites). During this process, hyper- solidus reactions at increasing melt mass (= open-system melting) are responsible for a depletion of the most incompatible elements (Verni•res et al., 1997).

Mercedita dunites have lower ThN/LaN and LaN/SmN ratios than the other samples and plot off the negative correlation between these ratios (Fig. 9). This depleted signature may simply result from the olivine-forming reaction at increasing melt mass responsible for dunitc formation (Kelemen, 1990). On the other hand, Mayari dunites plot on the negative correlation and were probably percolated by (and reequilibrated with) relatively enriched melts after their individualization.

Mechanism of chromite crystallization Although the crystallization of accessory chromite during

the fractionation of basic and ultrabasic melts is a frequent and well-known process (e.g., Arai, 1994; Roeder, 1994), the


formation of large volumes of chromitite in ophiolitic complexes is still debated. The crystallization of nearly monomineralic chromite rocks in ophiolites has been explained by (1) corectic crystallization of chromite + olivine followed by me- chanical separation (e.g., Thayer, 1969; Lago et al., 1982), (2) magma mixing (Arai and Yurimoto, 1994; Zhou et al., 1996; Zhou and Robinson, 1997), and (3) changes in fo2 stressing the role of volatiles (Johan et al., 1983; McElduff and Stumpfl, 1991; Melcher et al., 1997).

Cotectic crystallization of chromite + olivine followed by their physical separation should produce dunitic rocks having disseminated chromite with the same mineral chemistry as the chromitite. This is not the case for the chromitite deposits of the Moa-Baracoa district where dunitc and chromitite

were in equilibrium with different melts. In the magma-mixing models, interaction between a perco-

lating basaltic melt and the host harzburgite gives rise to a dunitic residuum (the dunitc walls) and a silica-rich, calc-alkaline melt. The mixing of this calc-alkaline melt with new batches of a more primitive basaltic melt in open conduits produces a hybrid melt within the chromite + liquid region of the olivine-silica-chromite system rich in Cr (Irvine, 1977), promoting the crystallization of monomineralic chromite. This model could explain the described disequilibrium between accessory chromite in peridotitc and associated chromitite. However, it can be applied only for Cr-rich chromifttes. In contrast, this model does not explain the formation of giant, monomineralic bodies of M-rich chromitite like that from Mercedita, because in the condensed system olivine-silica-chromite is rich in AI (Irvine, 1977, see also Arai and Abe, 1995); the corectic phase boundary between olivine and spinel is almost straight--only slightly convex toward the olivine-quartz joint. Thus, if a basaltic melt on this olivine- spinel corectic boundary mixes with a silica-rich melt, the resulting hybrid melt comes back rapidly to the corectic proportions by the crystallization of a small amount of chromite. New batches of primitive melt progressively shift the composition of the hybrid melt to the olivine-chromite cotecftc by the segregation of a very small amount of chromite.

The role of fo2 on chromium solubility in basaltic melts and on chromite crystallization has been experimentally studied by Roeder and Reynolds (1991) and Hill and Roeder (1974), respectively, who show that (1) chromium solubility strongly decreases with increasing fo•, (2) monomineralic chromite rocks can be formed above 1,200øC and 10 -1ø atmospheres of partial pressure of oxygen from Cr-rich basaltic melts, and at slightly lower temperature and higherfo• (log10 fo• from -7 to -5) from Cr-poor melts, and (3) a pressure increase slightly expands the chromite + liquid field toward lower values offo•. However, these conditions are difficult to measure in natural chromitite because the published oxygen geobarometers (e.g., Ballhans et al., 1991a, b) are all based on the ratio Fe3VFe 2+ of the chromite which must be calculated assuming chromite stoichiometry. These calculations are affected by the instrumental error of the microprobe, the presence of structural defects and vacancies, and by the fact that most chromian spinels are not stoichiometric (Kamperman et al., 1996). In addition, oxygen geobarometers are based on the coexistence of chromite and olivine with orthopyroxene, which is a very unusual mineral in ophiolitic chromitite and if

present is usually included in and partly reequilibrated with chromite.

The mineralogical and geochemical data described for the chromite deposits of the Mayarf-Baracoa ophioliftc belt indicate that they crystallized from different types of volatile-rich melts, probably under the conditions established by Hill and Roeder (1974). Normal MORB- or back-arc basin basalt-type magmas at 1,200øC need an additional supply of a fluid com- ponent to enter the chromite + liquid field at around log10 foa : -7. If these conditions are reached, the resulting melt is not in equilibrium with olivine, pyroxene, or plagioclase, and chromite is the only liquidus mineral. Thus, during percolation of such a melt through duniftc channels it will tend to dissolve the preexisting olivine and crystallize chromite. The ef- fectiveness of this process to form giant chromite deposits like the central body of Mercedira depends on the continuity in the supply of both magma and fluid to maintain thermal and fo2 conditions. The location of this process within dunitic channels is justified by the fact that intergranular porous flow in mantle rocks is favored by the abundance of olivine in the matrix (Toramaru and Fujji, 1986). However, this could also be used to argue that the crystallization of small amounts of chromite should makes the rock impermeable to porous flow, as pyroxene does. Although no experimental data on the surface energy of chromite or on the value of dihedral angles between chromite and melt are available, the characteristic rounded habits of ophiolitic chromite (Leblanc, 1980) re- duces its surface anisotropy and suggests that the dihedral angles are relatively low, allowing for the creation and mainte- nance of an interconnected, intergranular melt film during the formation of chromitite. In this model, intergranular olivine and olivine inclusions in chromite constitute relicts of

the preexisting dunitc, and amphibole-bearing inclusions are the products of reactions between trapped hydrous melts and olivine, as suggested by Lorand and Cottin (1987) and Melcher et al. (1997).

If gabbro bodies were formed from MORB- or back-arc basin basalt-type melts within or at the borders of the dunitic channels before the arrival of fluids, they would react with the chromite-forming melt only to a very limited extent because gabbros are impermeable to porous flow (Kelemen et al., 1997). Consequently the dissolution of clinopyroxene and plagioclase, and the crystallization of chromite would be restricted to a few centimeters into the gabbro. However, melt flow could locally cut the gabbro body, especially if it was not totally consolidated, thus explaining the existence of chromi- tire veins cutting the dunite-gabbro contact at Mercedira.

The chromitites from Mayar/can be explained by the same process. However, it must be taken into account that the composition of the parental melts at Mayarf had a boninitic affinity and, consequently, was probably relatively rich in chromium. This means that these melts needed, for the same temperature (1,200øC), a lesser supply of fluids to enter in the liquidus field of chromite (log10 foa -> -10; Hill and Roeder, 1974). Chromitite from Sagua de Tanarno represents an intermediate situation.

Concentration and fractionation of noble metals The relative increase in PGE in the metallurgic chromites

from Mayarf and Santa Isabel with respect to the refractory

562 PROENZA ET AL.

ores from Moa-Baracoa and Rupertina is in agreement with the nature of their parental melts, as boninitic-type melts are richer in PGE than MORB (Hamlyn, et al., 1985; Barnes et al., 1988). The same trend has been described by Leblanc (1995) at Poum (New Caledonia), also showing a decrease in the Pd/Ir value and an increase in Os-Ir-Ru contents from Al- to Cr-rich chromitites.

The negative-sloped, chondrite-normalized PGE pattern of ophiolitic chromitites has been ascribed to the early crystallization of the most refractory Os, Ir, and Ru minerals, to- gether with chromite, and the segregation of Rh, Pt, and Pd with the residual melt (Barnes et al., 1988). Nevertheless, experimental work by Amoss• et al. (1990) shows that the solubility of Ir, and to a lesser extent, Pt in basaltic melts, at low fs.2 conditions, strongly depends on fo2. Other factors such as SiO2 activity can also play an important rode in the precipitation of platinum group minerals (Peck et al., 1992). These re- suits are in good agreement with the mechanism of chromi- rite crystallization proposed in this paper, since the increase in fo• necessary to form chromite could be produced by mixing of a fluid- and SiOn-rich melt and the parent basaltic melt. This should reduce the solubility of Ir (and probably of Ru), giving rise to the precipitation of platinum group minerals and the partitioning of Rh, Pt, and Pd to the residual melt. Preliminary data indicate that an example of such a residual melt remained trapped at the borders of a chromitite pod at Potos/(Moa-Baracoa district). These rocks are made up of re- cystallized chromite crystals with intergranular pyrrhotite and pentlandite and show fiat chondrite- normalized PGE patterns

from Os to Pd (Proenza, 1998). The latter probably results from the combination of the negative pattern of chromite and the positive pattern of the intergranular sulfides, as Gervilla and Leblanc (1990) describe in chromite-arsenide ores from the Ronda and Beni Bousera ultramafic massifs.

The precipitation of Os minerals under high fo• is very problematic because of the mobility of this element in oxidized fluids (Brandon et al., 1996). However, considering that chromite crystallizes from a hybrid melt resulting from the mixing of fluid and melt in our model, it is possible to account for a partial reduction of the Os dissolved in the fluid compo- nent saturating the hybrid melt in Os. This could lead Os to behave similarly to Ir and Ru during chromite crystallization.

Genetic Model

The geology and tectonic evolution of the northern Caribbean margin indicate that during Upper Cretaceous, eastern Cuba was located in the hanging wall of an intrao- ceanic subduction zone that consumed young Cretaceous crust along a southward-dipping Benloft zone (Lewis and Draper, 1990; Iturralde-Vinent, 1996). This suggests the existence of a high geothermal gradient in the suprasubduction zone. A possible scenario is shown in Figure 10.

During the subduction of the oceanic crust, large volumes of fluids from the breakdown of antigorite (Ulmer and Trommsdorff, 1995) can easily cross the 1,000øC isotherm (above the hydrous solidus of peridotite; O'Hara, 1967), promoting very small rates of partial melting of the overlying mantle, thus enhancing the capacity of the generated, highly

3o km

6O km

.9O km

F•c. 10. Geodynamic model for the genesis of the chromite deposits in the Mayar/-Baracoa ophiolitic belt. The picture represents the stages just before the opening of a back-arc basin. Continuous lines -- isotherms, discontinuous lines -- corner flow. Abbreviations: C •- SiO2- and MgO-rich calc-alkaline melts, H -- highly hydrated melts, MB = Moa-Baracoa deposits, MY -• Mayar/deposits, ST -- Sagua de T•namo deposits.


hydrated melts to percolate upward intergranularly (Fig. 10). Below 1,200øC the low temperature of the peridotitie mantle matrix does not allow extensive highly hydrated melt-rock reactions. This produces no important variations in the H20 vol percent of highly hydrated melts during its upward migration and, consequently, no or only small changes in the oxidation state of the melt. Thus, highly hydrated melts mostly preserve the geochemical signature of the slab-derived fluids. As the temperature of the matrix increases, which takes place in the back-are mantle away from the are vertical (Fig. 10), highly hydrated melt-rock reactions take place at progressively increasing melt volume, giving rise to progressively depleted residua and eale-alkaline melts progressively enriched in SiO2 and MgO (Kelemen, 1990).

Upward migration of the SiOn- and MgO- rich eale-alkaline melts generated above 1,200øC should tend to crystallize below 1,150 ø to 1,100øC (tentative liquidus temperatures depending on the composition of the melt at around 10 kbars). Thus, these isotherms can behave as impermeable barriers to porous flow. According to Spiegelman (1993), if the freezing rate is high enough, pressure gradients generated by the combination of compaction pressure (normal to the impermeable barrier) and buoyancy of incoming melt flux make the melt move laterally, along a channel parallel to the impermeable region. Me'It movement is also (or mainly) favored by the plastic flow (corner flow) of the matrix which tends to transport the melt parallel to flow lines assuming that melt distribution is controlled by strain (Daines, 1997). This gives rise to a subhorizontal transport of melt along a porous, probably dunitic channel.

SiO• and MgO-rich calc-alkaline melts moving along the dunitic channel can mix with upward migrating highly hydrated melts, producing a hybrid melt with higherfo• and, to a lesser extent, a higher SiO• content than the former calc-alkaline melts. Assuming a melt temperature around 1,200øC and no or very small changes in temperature, mixing produces a hybrid melt composition, which lies within the liquidus field of chromite, as shown by the experimental results of Hill and Roeder (1974) and Irvine (1975). Under these physicochemical conditions the interstitial hybrid melt will tend to dissolve olivine from the matrix and crystallize chromite. M-rich chromite from Moa-Baracoa should be

formed in the distal parts of percolation channels (Fig. 10, MB) by mixing of the calc-alkaline melts (relatively rich in A12Oa) with oxidized hydrated fluids rich in AlcOa and TiO•, and with a very high H•O/silicate melt ratio. Toward the proxo imal parts of the percolation channels (Fig. 10, MY), calc-alkaline melts (rich in Cr•Oa and depleted in TiO2, approaching the composition of boninitic melts) mix with AlcOa- and TiO•- impoverished hydrated melts with a lower H•O/silicate melt ratio (due to the buffering effect of the percolated peridotitc matrix), giving rise to a moderate increase in fo•, sufficient to promote chromite crystallization because the chromite liquidus field expands toward lowfo2 conditions as the Cr•Oa content of the melt increases (Hill and Roeder, 1974). This explains the formation of Cr-rich chromitites at Mayafl.

The presence of higher A1 contents in chromite from dunitc relative to chromite in associated harzburgite in the Moa Baracoa district can be explained by equilibrium with the percolating basaltic melt at higher melt/rock ratios associated with magma channeling (Kelemen et al., 1995a). This can explain,

also, the overall Ti enrichment in dunites relative to harzburgites (el. spinel composition). However, the Mayar/dunites compared to dunites from Mereedita, experienced relatively late pereolation of Ti-poor (boninitielike) melts during the waning stages of the pereolation system, which is in agreement with the trace element distribution discussed above. A

similar evolution was recently suggested by Garrido and Bod- inier (1999) for the Ronda peridotitc.

In Moa Baraeoa (Fig. 10, MB) mixing of hydrated and eale- alkaline melts results in a further increase of AlcOa and TiO• in the hybrid melt, thus explaining the higher A120 a and TiO2 contents of chromite from chromitite relative to chromite

from dunitc. Toward the proximal parts of the percolation channel (from Moa-Baracoa to Mayar/) the ascending highly hydrated melts react with an increasingly hotter peridotitc matrix giving rise to melt-rock reactions at increasingly higher melt volume and to melts progressively rich in SiO2 and MgO. Consequently, hydrated and calc-alkaline melts tend to have similar compositions toward the proximal parts of the percolation channel. Thus, mixing of highly hydrated melts with calc-alkaline melts in areas such as Mayari, does not produce important changes in the composition of calc-alkaline melts and therefore the Cr no. of chromite in chromitite is similar

to that of accessory chromite from the host dunitc. Neverthe- less, evidence of such a mixing is recorded in the chromite: the TiO• content of chromite is higher in massive ore than in dunitc, indicating that the parental melt of chromite was slightly richer in TiO2.

Gabbro sills and the surrounding clinopyroxene- and plagioclase- bearing peridotites could have been generated in cool upper parts of the percolation channels during periods of nonsupply of highly hydrated melts, assuming that hydrated and calc-alkaline melts percolated episodically because of the migration of compaction-porosity waves (see Spiegelman, 1993, and references herein). As in chromitite, olivine and, to a lesser extent, orthopyroxene from gabbro sills are residual phases from the host peridotitc, whereas clinopyroxene and plagioclase are magmatic phases formed by melt-rock reaction at decreasing melt mass, as described by Burg et al. (1998) beneath the Kohistan arc. This mechanism explains the frequent layered structure of gabbro sills where thin olivine-rich layers alternate with olivine-free gabbro layers. Because magmatic flow (both of hydrated and calc-alkaline melts) is sporadic, the relative position of the isotherms and the permeability barrier can locally shift, modifying the location of dunitic channels. This means that early formed gabbro layers can be included in a new dunitic channel. Because gabbros are impermeable to porous flow (Kelemen et al., 1997) and are in equihbrium with the percolating calc-alkaline melts, those included in the dunitic channel do not suffer any kind of reaction unless a new supply of highly hydrated melts gives rise to a new hybrid melt that crystalhzes chromite. In this case, the hybrid melt partly reacts with the gabbro (at a centimetric scale because of the impermeability of the gabbro) dissolving clinopyroxene and plagioclase and crystallizing chromite. This explains the presence of concordant, tabular gabbro bodies within chromite lenses.

Discordant pegmatitic gabbro dikes from Moa-Baracoa and websteritc dikes from Mayaft can be the result of the fractional crystallization of either hydrated or residual melts generated

564 PROENZA ET AL.

after the crystalhzation of chromite. The latter is supported by the fact that in some deposits (Potosi) of the Moa-Baracoa district (as well as in the Camaguey district in central Cuba; Flint et al., 1948), gabbro dikes follow the direction of pull- apart fractures and are restricted to the chromitite body. On the contrary, the chemical variation of gabbro and pyroxenite dikes from Moa-Baracoa to Mayar/is consistent with the proposed chemical variation of highly hydrated melts from low- to high-temperature regions in the mantle wedge. Thus, whereas at Moa-Baracoa gabbro dikes mainly consist of coarse-grained plagioclase and clinopyroxene, indicating crystallization from volatile-rich melts, at Mayar/they consist of medium-grained Cr-rich clinopyroxene and orthopyroxene probably crystallized from high Mg andesitc or boninite-type melts at around 1,066 ø to 1,129øC. This temperature is shghtly lower than that estimated for the crystallization of chromite. Furthermore, the Cr depletion and Al enrichment in chromite at the contact with discordant gabbro dikes is also consistent with the fact that parental melts of pegmatitic gabbros (highly hydrated melts) were depleted in Cr and enriched in A1 with respect to the parental melt of chromitite (a hybrid melt from the mixing of hydrated and calc-alkahne melts).

Acknowledgments The authors thank the personnel of the Mercedira mine,

and especially D. Rev6 and G. Rodr/guez for their assistance during field work. We also thank W. Lavaut for his constructive comments on the chromite deposits of the Mayar/district, and J. Garcia Veigas and X. Llovet for their help during electron microprobe analyses. We are indebted to S. Arai and M. Leblanc for their constructive comments and criticisms of

an early version of this manuscript, as well as to P. Robinson and D. Peck who carefully reviewed a later version. All their comments have greatly improved the paper. This research was financially supported by the Instituto de Cooperaci6n Iberoamericana and the Fundaci6 Sohdaritat of Barcelona

University through different grants to J.P., as well as the Span- ish DGES through project PB97-1211. This paper is a contribution to International Geophysical Correlation Programme projects 427.

August 4, 1998; February 16, 1999

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al- and cr-rich chromitites from the mayar[iacute]-baracoa...

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