American Mineralogist, Volume 86, pages 1130–1142, 2001
Compositional zoning and element partitioning in nickeloan tourmaline from a metamorphosed karstbauxite from Samos, Greece DARRELL J. HENRY* AND BARBARA L. DUTROW Department of Geology and Geophysics, Louisiana State University, Baton Rouge, Louisiana 70803, U.S.A.
ABSTRACT Blue-green nickeloan tourmaline from a micaceous enclave of a marble from Samos, Greece, contains unusually high concentrations of Ni (up to 3.5 wt% NiO), Co (up to 1.3 wt% CoO), and Zn (up to 0.8 wt% ZnO). The polymetamorphic karstbauxite sample has an uncommon assemblage of nickeloan tourmaline, calcite, zincian staurolite, gahnite, zincohögbomite, diaspore, muscovite, paragonite, and rutile. The complex geologic history is reflected in multi-staged tourmaline growth, with cores that represent detrital fragments surrounded by two-staged metamorphic overgrowths. Zone-1 metamorphic overgrowths, which nucleated next to detrital cores, are highly asymmetric and exhibit compositional polarity such that narrow overgrowths of brown schorl developed at the (–) c-pole are enriched in Mg, Ti, and F, and depleted in Al, Fe, and X-site vacancies (X■) relative to wider, gray-blue schorl-to-foitite overgrowths developed at the (+) c-pole. Volumetrically dominant Zone-2 overgrowths are strongly zoned nickeloan dravites with a continuous increase in Mg, Co, Ca, and F at the expense of Fe, Zn, Cr, and V from the Zone-1 interface to the outermost rim. Within Zone 2, Ni reaches a maximum of 0.5 apfu before decreasing in the outer 20–40 µm. Zone-2 overgrowths also exhibit compositional polarity such that, at the (–) c-pole, overgrowths are enriched in Mg, F, Na, Ca, and Cr relative to overgrowths at the (+) c-pole that are, in turn, enriched in Al, Fe, Ni, Co, and X■. Element partitioning involving tourmaline rims and coexisting minerals indicates that relative partitioning of Ni is tourmaline >> staurolite > gahnite; Co is tourmaline > staurolite > gahnite; and Zn is gahnite > staurolite >> tourmaline.
INTRODUCTION Natural tourmalines contain a wide range of elements, but those with substantial amounts of Ni and Co are relatively unusual. In most cases, tourmaline with high Ni contents also has high Cr levels because the tourmaline-bearing rocks are associated with meta-ultramafic rocks (e.g., Challis et al. 1995; Michailidis et al. 1995). For instance, a chromian dravite (8.5 wt% Cr2O3), from a calcareous rock in contact with serpentinite, has the highest Ni level previously recorded in tourmaline, 0.75 wt% NiO [0.1 atoms per formula unit (apfu)] (Jan et al. 1972; Henry and Dutrow 1996). The recorded concentrations of Co in tourmaline are even lower, with the maximal reported Co being 45 ppm in a tourmaline from an aplite (Power 1968; Henry and Dutrow 1996). Tourmaline with substantial Zn contents is a much more common feature. Those tourmalines with ZnO levels >1 wt% are most prevalent in the internal zones of fractionated pegmatites (e.g., Jolliff et al. 1986; Federico et al. 1998). The maximum reported Zn content is 2.85 wt% ZnO (0.35 Zn apfu) for elbaite coexisting with amblygonite in a pegmatite (Jedwab 1962). However, high levels of Zn have also been noted in tourmaline that coexists with sphalerite from sulfide deposits
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(Griffen et al. 1996; Slack et al. 1999). The existence of natural Ni- or Co-tourmaline end-members is possible, based on experimental synthesis of Na-Ni and Na-Co tourmalines. Taylor and Terrell (1967) hydrothermally synthesized both pale-green Ni tourmaline and mauve-pink Co tourmaline. Recently, Gourdant and Robert (1997) synthesized a complete solid-solution series between dravite and “Nidravite,” and between foitite and “Ni-foitite.” This result reflects the similar octahedrally coordinated ionic radii of Ni2+ (0.77 Å) and Mg2+(0.80 Å). Structurally, Ni tends to cluster as 3 Ni2+ cations at the octahedral Y site in tourmaline (Gourdant and Robert 1997). Consequently, given the appropriate bulk compositions, mineral assemblages, and conditions of formation, Ni- or Co-tourmalines can potentially form naturally. This paper describes tourmaline that crystallized in a geochemical environment that was unusually enriched in Ni, Co, and Zn: the bauxite-limestone interface of a metamorphosed karstbauxite from eastern Samos, Greece. This area underwent multiple metamorphic events, and the metamorphosed karstbauxites developed a fascinating Zn-rich mineral assemblage that coexists with tourmaline. Significantly, the tourmaline retains textural and chemical information from all stages of its growth, such that it can be related to the geological history of the region. In addition, the relative partitioning of Ni, Co, and Zn at minor-to-major element-concentration levels can be established between tourmaline and coexisting minerals.
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GEOLOGIC SETTING AND METAMORPHIC HISTORY The metamorphosed karstbauxite sample was collected from the eastern portion of the island of Samos, Greece, approximately 4 km east of the town of Samos. The island of Samos is located between differentially extending areas of the Aegean Attic-Cycladic Metamorphic Complex and the Menderes Complex in SW Turkey (Ring et al. 1999). Each of these complexes is composed primarily of late Paleozoic and Mesozoic sequences of sediments that were metamorphosed during Alpine times (e.g., Schliestedt et al. 1987; Okrusch and Bröcker 1990). Thick Mesozoic neritic carbonate units are common in each of the complexes, and locally contain metamorphosed karstbauxite deposits that mark periods of uplift, exposure and chemical weathering of the carbonate units (Feenstra 1997). In this region, there are extensive metabauxite deposits along the southern and western margins of the Menderes Massif, on Naxos and on Samos (Feenstra 1985, 1996). In all of these areas, metabauxite units typically develop as 2–5 meter-thick lenses confined to specific stratigraphic horizons within thick-bedded calcite marbles (Feenstra 1997). Metabauxites in these areas are corundum- or diaspore-bearing, depending on the local grade of metamorphism. The karstbauxite sample is from the Ampelos nappe, one of the Cycladic blueschist nappe complexes of eastern Samos (Ring et al. 1999). According to Ring et al. (1999), Cycladic blueschist units of central and eastern Samos have a complex metamorphic history. Initially, blueschist-facies metamorphism (M1) developed at 420–500 °C and 11–15 kbar during and after initial deformation (D1) associated with nappe stacking at roughly 50 Ma (see also Chen 1995; Will et al. 1998). Porphyroblasts statically overgrew the S1 foliation. During compressional deformation (D2), the M1 high-pressure assemblages of the Cycladic blueschist units were locally replaced by transitional blueschist-greenschist assemblages (M2) at 35– 40 Ma. However, D2 and M2 overprints are generally not welldeveloped in the eastern portion of Samos. A Barrovian-type metamorphic overprint (M3) occurred during extension-related decompression of the nappe complex (450–490 °C and 6–7 kbar) at 18–25 Ma. However, Feenstra (1997) obtained lower temperatures of 270–320 °C for M3, based on coexisting muscovite and paragonite compositions in the metabauxites and local occurrence of kaolinite. D3 structures are associated with greenschist-facies mineral assemblages (M3) and reflect the dominant ductile deformation observed in the rocks, commonly modifying and transposing earlier structures. Superimposed on the ductile fabrics of D1–D3 are late ductile/brittle deformational features (D4 and D5) that likely developed in the last 9 Ma (Ring et al. 1999). One of the most interesting features of the metamorphosed karstbauxites is the basal unit at the interface between previously dissolving limestone host and overlying bauxite. Karst depressions associated with limestones are generally filled with argillaceous or sandy sediments that can undergo bauxitization, with clay-rich bauxites most concentrated at the base (Bardossy 1982). The interface of bauxite with the underlying limestone is typically nonuniform, being undulose and locally powdery due to alteration and dissolution of the footwall limestone (Bardossy 1982). During the bauxitization process, Al is greatly
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enriched and Si, Mg, Ca, Na, and K are generally depleted in the material undergoing bauxitization. In addition, based on trace-element profiles of a number of Cycladic karstbauxites, Feenstra (1985) suggested that Ni, Co, Cu, Y, La, Be, Zn, and Pb are mobile and enriched near the carbonate footwall by factors of 10–100 relative to the bulk composition of the bauxite. In similar metamorphosed karstbauxites from Naxos, Feenstra (1985) found that, at the interface with the carbonate footwall, whole-rock Zn levels commonly exceed 1000 ppm, Ni levels were over 500 ppm, and Co was up to 200 ppm. The enrichment of mobile elements at the base of the karstbauxite was attributed to the carbonate footwall acting as a pH barrier for meteoric water percolating downward during bauxitization. However, the limestone below the bauxite interface was not subject to similar enrichment and has significantly lower levels of mobile elements (Feenstra 1985). The magnitude of enrichment depends on the chemical and mineralogical composition of the bauxitized material, the thickness of the deposit, the duration of bauxitization process, and the climatological and hydrological conditions. Furthermore, Feenstra (1985) argued that, once metamorphism of the karstbauxites commences, there is little further mobilization of any of the previously mobile elements. Consequently, gradational or irregular contacts between bauxite and limestone are likely to have retained a sharp concentration gradient of the mobile elements, which will influence the chemistry of the subsequent metamorphic minerals. At the eastern Samos locality, Feenstra (1997) described some of the minerals associated with Zn-enriched metamorphosed karstbauxite, including diaspore, calcite, zincian staurolite, chloritoid, two types of gahnite, zincohögbomite, paragonite, margarite, muscovite, rutile, and tourmaline. The “first-type” gahnite is an extremely Zn-rich (nearly end-member) colorless spinel, but “second-type” gahnite is blue and may be associated with rims of “first-type” gahnite or developed as separate grains near zincian staurolite, possibly due to staurolite breakdown. Feenstra (1997) suggested that zincian staurolite, chloritoid, “first-type” gahnite, diaspore, muscovite, rutile, margarite, paragonite, tourmaline, and calcite were early phases associated with the early high-P metamorphism (M1), but that “second-type” gahnite, zincohögbomite, and kaolinite developed during and after the medium-P overprint (M3).
SAMPLE DESCRIPTION AND TEXTURAL INTERPRETATIONS
The sample is a metamorphosed and complexly deformed marble with micaceous enclaves that locally contain significant amounts of tourmaline, zincian staurolite, and gahnite. The marble component is brown and granoblastic with local crosscutting veins of coarse, colorless calcite. Within the brown marble, there are multiply deformed micaceous bands. Unlike most metabauxites, the micaceous enclaves are significantly enriched in tourmaline (cf., Feenstra 1985). Consequently, the sample is interpreted to be a mixed lithology representing the deformed interface of limestone and bauxite protoliths of a karstbauxite environment. The calcite-marble component consists of calcite (>95%) + diaspore + muscovite + kaolinite + rutile + hematite + pyrrho-
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tite + apatite + monazite + gahnite. Because of the polymetamorphic and polydeformational nature of the terrain, equilibrium is localized. Optically, the dominant mineral, calcite, appears to have developed in multiple stages with at least one early generation of granoblastically recrystallized calcite (100–300 µm), marked by local Fe staining, cut by a later generation of coarser grained (300–500 µm) recrystallized calcite veins. However, cathodoluminescence microscopy (CL) reveals that there are three distinct generations of calcite. The first generation, typified by abundant fine-grained brown hematite inclusions and irregularly-shaped fluid inclusions, has bright orange, relatively uniform luminescence. Minerals associated with generation-1 calcite include muscovite, diaspore, colorless gahnite, paragonite, rutile, pyrrhotite, and apatite. Generation-1 calcite, probably developed during M1, is locally replaced by two subsequent calcite generations. Generation-2 calcite has fewer fluid inclusions and coarser hematite inclusions, is well recrystallized and exhibits complex CL zoning involving dull orangered luminescent to non-luminescent bands. Generation-2 calcite contains transposed and fractured pre-existing silicate and oxide minerals, and is probably related to the M3-D3 overprint. Locally, kaolinite and hematite are associated with generation2 calcite. Generation-3 calcite is weakly to non-luminescent, coarser-grained calcite crystallized in veins that cut all preexisting structures. Vein calcite seems to have grown into open spaces and is likely related to late D4-D5 brittle deformation. The three different generations of calcite establish the heterogeneous nature of deformation and complex fluid-influx history. The heterogeneous nature of deformation is further illustrated by framboids of pyrite, replaced by hematite, that may be either undeformed or deformed and disaggregated in different parts of the sample. Micaceous enclaves containing muscovite (30–50%) + calcite (10–30%) + tourmaline (5–30%) + zincian staurolite (1– 5%) + diaspore (1–5%) + rutile (1–5%) + hematite (1–3%) + gahnite (tr) + paragonite (tr) + zincohögbomite (tr) + pyrrhotite (tr) + zircon (tr) + monazite (tr) + apatite (tr) are folded into the marble. Muscovite and paragonite laths, 100–300 µm long, define a transposed foliation (D3). An earlier deformation (D1) is preserved by aligned fine-grained rutile inclusions in muscovite, tourmaline, gahnite, and zincian staurolite. Calcite associated with the micaceous zones is typically generation-2 calcite exhibiting complex CL zoning. Zincian staurolite is weakly pleochroic from pale violet to pale blue to pale greenblue, with color zoning marked by a darker blue rim. In addition to abundant rutile inclusions, there are uncommon inclusions of tourmaline, zircon, and pyrrhotite in the staurolite. In a few of the larger staurolite grains, there are inclusionrich cores with inclusion-poor rims, suggesting two-stage growth. Locally, some of the staurolite grains are fractured and filled with generation-2 calcite. Similar to the karstbauxite described by Feenstra (1997), two generations of gahnite are present. Generation-1 gahnite, the coarser-grained of the two generations (200–300 µm), is colorless, locally fractured, and exhibits blue-green luminescence. Generation-2 gahnite crystals are nonluminescent, subhedral-to-anhedral medium-blue grains (50–100 µm) that typically develop at, or near, zincian staurolite, but are not in-
cluded in zincian staurolite. Generation-2 gahnite locally includes rutile, calcite, and tourmaline. In some cases, brown zincohögbomite overgrows portions of the blue-green generation-2 gahnite (see Feenstra 1997 for an extended discussion of this type of occurrence). Muscovite + diaspore + rutile + staurolite (interior) + generation-1 gahnite + tourmaline (interior) + generation-1 calcite + pyrrhotite + zircon + monazite + apatite likely represent the M1 assemblage. The generation-2 blue gahnite and the rims of staurolite and tourmaline may represent either later stages of M1 or an M3 overprint. The M1 assemblage of minerals is generally transposed relative to the current foliation, which is primarily defined by micas, is fractured, and is joined by hematite, generation-2 calcite, paragonite, and zincohögbomite that formed during the M3D3 overprint. Optically zoned blue-green tourmaline, 20–100 µm in crosssection and 50–250 µm in length, is common in the micaceous enclaves. Some of the tourmaline grains contain distinctive colorless, pale-brown or pale-green cores with subangular-tosubrounded morphologies. These morphologies are similar to detrital tourmalines found as nucleation sites for metamorphic tourmaline from other metamorphic terrains (e.g., Henry and Guidotti 1985; Henry and Dutrow 1992; Sperlich et al. 1996). In sections oriented roughly parallel to the c-axis of tourmaline, the colorless, brown, or green cores are notably surrounded by asymmetric color-zoned overgrowths, with the narrow side being brown and the wider side being gray-green (Fig. 1). By analogy with other metamorphic tourmaline, the narrow brown, high-Ti overgrowth develops at the (–) c-pole and the wider gray-green overgrowth develops at the (+) c-pole (e.g., Henry and Dutrow 1996). This initial generation of asymmetric overgrowth is termed Zone-1 metamorphic tourmaline. Within Zone-1 overgrowths, there are fine-grained, randomly oriented rutile grains. Blue-green outer metamorphic overgrowths, termed Zone-2, are volumetrically dominant and exhibit very weak pleochroism. Zone-2 tourmaline overgrowths contain calcite and coarser rutile needles that define an earlier foliation, but are commonly rotated relative to the orientation of the matrix micas. The outer portions of Zone-2 overgrowths contain slightly coarser rutile than the inner portion of Zone-2 overgrowths. Also, tourmalines with Zone-2 overgrowths are locally included in the outer rim of zincian staurolite and generation-2 gahnite.
ANALYTICAL AND NORMALIZATION PROCEDURES Tourmaline, staurolite, gahnite, muscovite, and paragonite were analyzed quantitatively by wavelength-dispersive spectrometry (WDS) using the automated JEOL 733 electron microprobe at LSU. WDS analyses were done at an accelerating potential of 15 kV and 5–20 nA using a 1–5 µm diameter electron beam, depending on the mineral. Lower sample current and wider beam diameters were used for mica analyses. Standards were well-characterized synthetic and natural minerals, including andalusite (Al), diopside (Ca, Mg, Si), fayalite (Fe), chromite (Cr), synthetic glasses (V, Ni, Co), kaersutite (Ti), rhodonite (Mn), willemite (Zn), albite (Na), sanidine (K), and apatite (F). Several well-characterized tourmalines and staurolites served as secondary standards to ensure good-quality
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FIGURE 1. Backscattered-electron (BSE) image of a micaceous enclave rich in tourmaline. Tourmaline grains are in various orientations. Two examples of detrital cores are shown; detrital-core 1 is dark gray and detrital-core 2 (outlined by a white line) is medium gray. The tourmaline grain at the bottom of the image is oriented roughly parallel to the c-axis and is cut through the center of the grain, exposing the detrital core, asymmetric Zone-1 overgrowth (outlined in white) and Zone-2 overgrowth. The thicker black line represents the trace of a compositional traverse (see Fig. 2). Muscovite, calcite, and rutile (white inclusion and matrix grains) are also shown.
analyses (Dutrow et al. 1986, 1999; Holdaway et al. 1986; Henry and Dutrow 1990). On the basis of replicate analyses of several secondary standards, analytical precision associated with counting statistics is estimated to be ± 0.21% for SiO2, ± 0.13% for Al2O3, ± 0.06% for FeO, ± 0.11% for MgO, ± 0.02% for CaO, and ± 0.03% for Na2O. Minimum detection limits in wt% are considered to be 0.08 for NiO, 0.04 for CoO, 0.03 for ZnO, 0.03 for Cr2O3, 0.03 for V2O5, and 0.03 for K2O. The most appropriate normalization procedures for tourmaline were carefully considered. The general formula for tourmaline can be expressed as XY 3 Z 6 (T 6 O 18 )(BO 3 ) 3 V 3 W (Hawthorne and Henry 1999). The X site contains Na, Ca, K, and vacancies (X■); the two distinct octahedral sites, Y and Z, are occupied by a variety of divalent, trivalent, and tetravalent cations; and the tetrahedral T site is dominated by Si with lesser Al. The triangular B site is occupied exclusively by B. There
are two anion sites, V and W. Both sites can contain OH– and O2–, but the W site exclusively contains F–, and O2– tends to partition into this site. Tourmaline formulae were normalized on the basis of 15 cations exclusive of Na, Ca, and K, i.e., assuming no vacancies at the tetrahedral or octahedral sites and insignificant Li content (Henry and Dutrow 1996; Henry et al. 1999). Insignificant amounts of Li in tourmaline are probable because Li tends to strongly partition into any coexisting staurolite or micas relative to tourmaline (Dutrow et al. 1986; Henry and Dutrow 1996). The amount of B2O3 necessary to produce three B cations in the structural formula was calculated from stoichiometric constraints. Because of the weak pleochroism of the tourmaline and the antithetic nature of Fe with the other octahedral divalent cations, it is assumed that the Fe is essentially all Fe2+ (see discussion below). Consequently, chargebalance constraints could be used to estimate the amounts of H
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and O associated with the V and W anion sites in the structural formula, although it is subject to significant uncertainties (Henry and Dutrow 1996; Dutrow and Henry 2000). It is further assumed that F and calculated O are preferentially partitioned at the W site (Hawthorne and Henry 1999). These assumptions permit grouping of tourmaline in accordance with the classification scheme proposed by Hawthorne and Henry (1999). A number of potential tourmaline species have been identified that are currently not recognized as tourmaline species. These hypothetical species are designated by quotation marks, e.g., “oxy-dravite.” The minerals coexisting with tourmaline were normalized using several approaches. Staurolite was normalized on the basis of 25.53 (Si + Al), in accordance with the suggestion of Holdaway et al. (1986). Low total amounts of divalent cations (2.4–2.7 R2+) and low analytical totals (96–97%) suggest the likelihood of significant amounts of unanalyzed Li and/or H (Dutrow et al. 1986; Hawthorne et al. 1993a). Consequently, the (Si + Al) = 25.53 cation normalization procedure is considered to be more appropriate than a fixed oxygen basis normalization. Micas and gahnite were normalized on the basis of 11 and 4 oxygen atoms, respectively.
RESULTS Mineral chemistry of associated minerals The silicate and oxide minerals that coexist with nickeloan tourmaline have relatively unusual compositions that reflect the geochemical environment and metamorphic processes by which they developed. Mineral chemistry is summarized below. Staurolite. Considering only divalent cations (R2+), staurolite is exceptionally Zn-rich, with XZn = 0.71–0.81, XFe = 0.12– 0.18, XCo = 0.02–0.05, XMg = 0.02–0.03, and XNi = 0.02–0.03, where Xcation = divalent cation/(sum of the divalent cations) (Table 1). Staurolite is zoned, with higher amounts of Zn (~0.2 apfu) in the core than in the rim. The darker blue, inclusionpoor rim is correspondingly enriched in Fe, Co, and Mg relative to the core, implying that zoning primarily reflects the simple homovalent exchange vector (Fe, Co, Mg, Ni)(Zn)–1. One of the unusual features of these staurolites is that the sum of the divalent cations is very low, 2.44–2.69 total R2+, relative to the range found in most staurolite, 3.17–4.77 total R2+ (Holdaway et al. 1986). Low total R2+ typically indicates that H contents are probably high, i.e., >4 H atoms per formula unit (apfu), and/or that significant amounts of Li are present (e.g., Dutrow et al. 1986; Holdaway et al. 1986, 1991; Hawthorne et al. 1993a). Unanalyzed Li and H could account for relatively low oxide totals (96–97 wt%) found in these staurolites. Hawthorne et al. (1993b) reported a similar zincian staurolite from Barrhorn, Switzerland with 11.82 wt% ZnO that has high Li (0.45 wt% Li2O). High Li and H in staurolite are most likely governed by the exchange vector (LiOH)(FeO)–1 (Hawthorne et al. 1993b). Gahnite. The two generations of gahnite have significant differences in chemistry (Table 1). The colorless generation-1 spinel is nearly end-member gahnite with XZn = 0.987, XFe = 0.007, XNi = 0.004, XMg = 0.002, and XCo = 0.001. The blue generation-2 gahnite that typically develops at, or near, the mar-
gins of staurolite exhibits some compositional zoning, with the highest Zn levels in the cores. Generation-2 gahnite is relatively enriched in Fe and Co, with XZn = 0.869–0.931, XFe = 0.041–0.050, XNi = 0.005–0.012, XMg = 0.009–0.012, and XCo = 0.025–0.048. The darker blue color, relative to generation-1 gahnite, is attributed to higher concentrations of Co. The change in composition from generation-1 to generation-2 gahnite, as well as the zoning in generation-2 gahnite, essentially reflects the simple homovalent exchange vector (Fe, Co, Mg, Ni)(Zn)–1, similar to staurolite. The nonluminescent character of generation-2 gahnite is a consequence of significant levels of Fe2+ that quench the CL response. The generations-1 and -2 gahnite compositions are very similar to the “first-type” and “secondtype” gahnites described in the metabauxites from this area by Feenstra (1997). Muscovite and paragonite. The white micas are dominated modally by muscovite. Muscovite is moderately paragonitic, with Na/(Na + K + Ca) ratios of 0.123–0.135 (Table 1). Interestingly, Si values of 2.988–2.997 and Σ (Fe + Mn + Mg) of 0.055–0.084 apfu are low relative to more celadonitic muscovite found in other Cycladic blueschist metamorphic rocks (e.g., Feenstra 1985). Muscovite also has low Mg / (Mg + Fe) ratios of 0.32–0.39. There are small, but significant, amounts of V (0.006–0.007 apfu), and very small amounts of Ni, Co, and Zn in muscovite, that are at, or below, detection limits. Paragonite is commonly in contact with staurolite, and partially replaces the margins of some staurolite grains. Paragonite has significant muscovite and margarite components [K/(Na + K + Ca) = 0.014–0.028, and Ca/(Na + K + Ca) = 0.043–0.101]. Relative to coexisting muscovite, paragonite has higher concentrations of Zn (0.009–0.022 apfu). Tourmaline chemistry and zoning characteristics Metamorphosed karstbauxite tourmaline grains have three segments with distinct morphology and chemistry. Tourmaline grains generally have cores, interpreted as detrital grains that are surrounded by two-stage metamorphic overgrowths: small, asymmetric Zone-1 overgrowths and larger blue-green Zone-2 overgrowths. A detailed compositional traverse, roughly parallel to the c-axis, clearly reveals the chemical relations among these different segments (Fig. 2). Additional analyses were obtained in tourmalines with different orientations to establish complete ranges and most representative compositions for different tourmaline segments (Tables 2 and 3). Detrital cores. Detrital cores of four different grains have chemically distinct compositions relative to each other, and to the metamorphic overgrowths (Table 2). All four detrital-core grains are alkali-group tourmalines with relatively high X-site vacancies, and generally fall into the hydroxy-subgroup tourmalines, with the exception of grain 2, an oxy-subgroup tourmaline (Fig. 3). The detrital tourmaline grains range from aluminous “oxy-schorl” to dravite [Mg/(Mg + Fe) = 0.38–0.64] (Fig. 4). Zone-1 overgrowths. Zone-1 overgrowths typically nucleate proximal to detrital tourmaline cores. They are highly asymmetric and exhibit distinct compositional polarity, i.e., distinct compositional differences at the (+) and (–) c-poles of the grains (Figs. 1 and 2). In the direction of the (–) c-pole, narrow brown
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TABLE 1. Representative compositions of minerals coexisting with tourmaline Mineral Comment SiO2 (wt%) Al2O3 Cr2O3 V 2O 5 TiO2 FeO MnO NiO CoO MgO ZnO CaO Na2O K 2O F TOTAL
Gahnite Gen. 1 n.d. 55.14 0.48 0.06 0.03 0.27 0.00 0.15 0.05 0.04 43.75 n.d. n.d. n.d. n.d. 99.97
Gahnite Gen. 2 n.d. 56.46 0.23 0.00 0.03 1.82 0.00 0.49 1.97 0.26 39.11 n.d. n.d. n.d. n.d. 100.37
Norm basis Si IV Al
O=4 – –
O=4 – –
Al Cr V5+ Ti Fe2+ Mn2+ Ni Co Mg Zn Ca Na K F
1.985 0.012 0.001 0.001 0.007 0.000 0.004 0.001 0.002 0.987 – – – –
2.003 0.005 0.000 0.001 0.046 0.000 0.012 0.048 0.012 0.869 – – – –
17.530 0.090 0.004 0.012 0.333 0.005 0.059 0.042 0.053 2.082 – – – –
0.880 0.047 0.012 0.012 0.049
0.809 0.129 0.021 0.023 0.016
Zn/R2+ Fe/R2+ Mg/R2+ Ni/R2+ Co/R2+ Note: n.d. = not
0.986 0.007 0.002 0.004 0.001 detected.
Staurolite core 28.68 55.16 0.42 0.02 0.06 1.46 0.02 0.27 0.19 0.13 10.35 n.d. n.d. n.d. n.d. 96.76
Staurolite rim 28.03 54.82 0.37 0.01 0.18 2.03 0.06 0.30 0.63 0.18 9.19 n.d. n.d. n.d. n.d. 95.80
Muscovite near st 45.33 37.75 0.23 0.15 0.03 0.67 0.00 0.03 0.01 0.18 0.04 0.16 1.04 9.95 0.10 95.67
Paragonite near st 42.69 41.92 0.12 0.00 0.03 0.20 0.00 0.01 0.01 0.02 0.46 0.62 7.36 0.34 0.12 93.90
O = 11 2.997 1.003
O = 11 2.791 1.209
17.530 0.081 0.002 0.037 0.468 0.014 0.067 0.139 0.074 1.870 – – – –
1.939 0.012 0.007 0.001 0.037 0.000 0.002 0.001 0.018 0.002 0.011 0.133 0.839 0.021
2.021 0.006 0.000 0.001 0.011 0.000 0.001 0.001 0.002 0.022 0.043 0.933 0.028 0.025
0.711 0.178 0.028 0.026 0.053
0.033 0.617 0.300 0.033 0.017
0.595 0.297 0.054 0.027 0.027
Formula Proportions Si+Al=25.53 Si+Al=25.53 7.815 7.725 0.185 0.275
FIGURE 2. Detailed quantitative traverse of a tourmaline oriented parallel to the c-axis, exposing the detrital core and two-stage metamorphic overgrowths (Zones 1 and 2). (a) Compositional zonation of the major elements on the tourmaline traverse (see BSE image in Fig. 1). Note the difference in size and composition of the Zone-1 overgrowth at the opposite poles of the tourmaline grain. (b) Compositional zonation of the minor elements ( gahnite; Co = tourmaline > staurolite > gahnite; Zn = gahnite > staurolite >> tourmaline; Mg = tourmaline >> staurolite > gahnite; and Fe2+ = tourmaline > staurolite > gahnite (Table 4). Both staurolite and gahnite have tetrahedral sites that preferentially accommodate Zn, and this factor has a significant effect on element partitioning, with gahnite and staurolite taking up Zn much more effectively than tourmaline (e.g., Griffen 1981; Hawthorne et al. 1993a; Koch-Muller et al. 1999). Furthermore, gahnite partitions Zn preferentially relative to staurolite, with a partition coefficient of 2.98. This finding is generally consistent with the partition coefficient of 1.8 obtained experimentally at 700 °C and 5 kbars (Bubenik and KochMüller 1996). Cobalt is known to occur in significant concentrations in staurolite and spinel, forming lusakite (up to 8.5 wt% CoO) and cochromite (up to 17.5 wt% CoO) (Skerl and Bannister 1934; DeWaal 1978). Cobalt is primarily in tetrahedral coordination in staurolite and spinel-group minerals, but the higher partitioning of Co into tourmaline relative to these minerals suggests that Co is accommodated readily in octahedral coordination in tourmaline as well (Table 4). Magnesium and Ni are strongly partitioned into tourmaline relative to coexisting minerals in a similar manner, probably due their similar ionic size and crystallochemical responses in tourmaline. The high relative partitioning of Ni and Co into tourmaline suggests that tourmaline with much higher Ni and Co levels may exist in the appropriate petrologic environments, e.g., metamorphosed paleosols developed from ultramafic rocks (e.g., DeWaal 1978) or Co-rich aluminous rocks similar to those ˇ that host lusakite ( Cech et al. 1981). Furthermore, because of the relative partitioning, Ni and Co concentrations should be determined in tourmaline in any rock that has even a minor enrichment in these elements.
TABLE 4. Partition coefficients for coexisting minerals or poles of minerals Partition Cations: Tur(+) – Tur(–)* Zone 2 outer rim Tur(+) – Tur(–) Zone 2 inner rim
KD Zn–R 0.799
KD Fe–R 1.085
KD Mg–R 0.895
KD Ni–R 1.087
KD Co–R 1.088
0.670
0.968
0.843
1.466
1.860
Tur(+) – St Tur(–) – St
0.003 0.002
1.556 1.434
44.72 49.95
4.821 4.444
1.222 1.120
Tur(+) – Gah Tur(–) – Gah
0.001 0.001
6.830 6.302
106.1 118.5
10.59 9.76
1.327 1.217
Gah – St 2.980 0.227 0.422 0.455 0.921 * Mineral pair used for calculation of the partition coefficients. Notes: (+) and (–) represent average compositions of tourmaline at the (+) and (–) c-pole of the tourmaline. The partition coefficients are calculated in accordance with the general expression: KDα–βX–R2+ = (X/R2+)α/(X/R2+)β where X is a divalent cation. Abbreviations: Tur = tourmaline, St = staurolite and Gah = gahnite.
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HENRY AND DUTROW: NICKELOAN TOURMALINE
ACKNOWLEDGMENTS We acknowledge Rick Young for thin section preparation and Xiaogang Xie for assistance with microanalysis. This work was supported by NSF grant EAR-9405747 to DJH and NSF grant EAR-9814418 to BLD. We thank Frank Hawthorne, Dave London, and Bob Tracy for very useful reviews of the manuscript.
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