Journal of Petrology

Journal of Petrology Advance Access originally published online on July 8, 2004
Journal of Petrology 2004 45(8):1689-1723; doi:10.1093/petrology/egh030

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Journal of Petrology 45(8) © Oxford University Press 2004; all rights reserved

A Metamorphosed Early Cambrian Crust–Mantle Transition in the Eastern Alps, Austria

FRANK MELCHER^1,* and THOMAS MEISEL²

¹ GERMAN FEDERAL INSTITUTE OF GEOSCIENCES AND NATURAL RESOURCES, STILLEWEG 2, 30655 HANNOVER, GERMANY
² GENERAL AND ANALYTICAL CHEMISTRY, UNIVERSITY OF LEOBEN, 8700 LEOBEN, AUSTRIA

RECEIVED JULY 10, 2002; ACCEPTED MARCH 9, 2004

	ABSTRACT

TOP ABSTRACT INTRODUCTION GEOLOGICAL BACKGROUND ANALYTICAL METHODS PETROGRAPHY MINERAL CHEMISTRY WHOLE-ROCK CHEMICAL COMPOSITION DISCUSSION SUMMARY AND CONCLUSIONS REFERENCES

 
In the Speik Complex (Eastern Alps, Austria), highly melt-depleted, metamorphosed harzburgites with abundant pods and layers of chromitite are interlayered with a suite of metamorphosed orthopyroxenites, clinopyroxenites and gabbros. Coarse-grained orthopyroxenites occur as centimetre- to metre-wide veinlets and pods, but also as intrusive plugs several tens of metres wide. Intimately associated metaclinopyroxenite and metagabbro are present as bodies up to several metres thick at a distinct stratigraphic level within the complex. In the ultramafic rocks, relict magmatic olivine, orthopyroxene, clinopyroxene and spinel have been overprinted by a metamorphic assemblage of forsterite, diopside, tremolite, anthophyllite, chlorite, serpentine, talc and Cr–Fe-rich spinel. Hornblende, epidote, zoisite and chlorite dominate the metamorphic paragenesis in metagabbros, in addition to rare relicts of clinopyroxene and two phases of Ca-rich garnet. The polymetamorphic evolution of the Speik Complex includes rarely preserved pre-Variscan (400 Ma) eclogite-facies conditions, Variscan (330 Ma) amphibolite-facies conditions (600–700°C, >5 kbar) and Eoalpine (100 Ma) greenschist- to amphibolite-facies conditions reaching 550°C and 7–10 kbar. Orthopyroxenites are characterized by high concentrations of SiO₂, MgO and Cr, and by U-shaped chondrite-normalized rare earth element (REE) patterns similar to those of their harzburgite hosts. The REE patterns of the clinopyroxenites are flat to slightly enriched in light REE. Metagabbro compositions are variable, but generally characterized by low SiO₂ and high mg-numbers (61–78). Their REE patterns all have Gd_N/Yb_N > 1; some samples have large positive Eu anomalies implying the original presence of cumulus plagioclase. In the orthopyroxenites, clinopyroxenites and some peridotites, Pt, Pd and Re are distinctly enriched compared with Os, Ir and Ru, whereas most harzburgites have unfractionated to slightly fractionated platinum-group element (PGE) patterns with respect to average upper mantle. The Re–Os isotope compositions of the pyroxenites define an errorchron at 550 ± 17 Ma and a supra-chondritic ¹⁸⁷Os/¹⁸⁸Os of 0·179 ± 0·003. An isochron age of 554 ± 37 Ma with _Nd(i) +0·7 is indicated by the Sm–Nd isotope compositions of whole-rock pyroxenite and gabbro samples, whereas the harzburgites plot on an errorchron of 745 ± 45 Ma and _Nd(i) +6. The pyroxenites and gabbros probably represent a cogenetic suite of magmatic dykes intruded into uppermost, highly depleted, suboceanic mantle below the crust–mantle transition zone in an oceanic basin close to the northwestern margin of Gondwana.

KEY WORDS: pyroxenite; metagabbro; geochemistry; Re–Os isotopes; Sm–Nd isotopes

	INTRODUCTION

TOP ABSTRACT INTRODUCTION GEOLOGICAL BACKGROUND ANALYTICAL METHODS PETROGRAPHY MINERAL CHEMISTRY WHOLE-ROCK CHEMICAL COMPOSITION DISCUSSION SUMMARY AND CONCLUSIONS REFERENCES

 
The mantle–crust transition zones of ophiolites are dominantly composed of harzburgite and dunite, with various types of segregations including gabbros, pyroxenites and chromitites (e.g. Boudier & Nicolas, 1995; Jousselin & Nicolas, 2000). Orthopyroxenite (‘bronzitite’) is a minor, but fairly common and conspicuous, rock type in the transition zone, and has been described, among others, from ophiolites in Albania (Neziraj, 1992), Cyprus (Büchl et al., 2002), Thailand (Orberger et al., 1995), New Caledonia (Podvin et al., 1985), Newfoundland (Edwards, 1990; Tanguay et al., 1990; Varfalvy et al., 1996, 1997), Tasmania (Peck & Keays, 1990) and China (Mei et al., 1988). Clinopyroxenite, orthopyroxenite and websterite occur both as cumulates and as non-cumulate dykes (e.g. Ahmed, 1991; Varfalvy et al., 1996, 1997; Bédard & Hébert, 1998; Nicolas et al., 2000). Wehrlitic intrusions into crustal gabbros have been reported from the mantle–crust transition zone in the Oman (e.g. Benn et al., 1988; Boudier & Nicolas, 1995) and Troodos ophiolites (e.g. Benn & Laurent, 1987), where they are regarded as the result of impregnation by basaltic melt. Pyroxenites are also documented from the modern seafloor: ultramafic mantle xenoliths from supra-subduction zone seamounts within the Bismarck microplate (Papua New Guinea) reveal a strong metasomatic overprint, including the formation of crosscutting orthopyroxene-, clinopyroxene-, phlogopite- and hornblende-bearing veins (McInnes et al., 1999; Franz et al., 2002).

The genesis of pyroxenites is enigmatic, because they have mineralogical and chemical features in common with both their ultramafic host rocks and the ultramafic–mafic cumulate sequences that overlie the mantle residue. Orthopyroxenites may represent cumulates that crystallized from fractionating melts in magma chambers above the petrological Moho. However, they may also form by melt–rock reactions or metasomatic reactions within residual mantle, or as products of metamorphic reactions. Kelemen et al. (1992) have pointed out that orthopyroxenites containing more than 32% pyroxene cannot represent mantle residues from previous melt extraction, and may have formed by reaction of olivine-rich peridotite with siliceous melt according to the following reaction:

Accumulations of chromite as ‘podiform’ and ‘stratiform’ chromitites are common in the upper mantle, mantle transition zone, and ultramafic cumulates of many ophiolite complexes (e.g. Leblanc & Nicolas, 1992; Melcher et al., 1997, 1999a; Burgath et al., 2002). The genesis of mantle-hosted chromitites is still much debated, although some consensus has been reached that they may form from melt–rock reaction (Zhou et al., 1994). Recent experimental data show that typical textures of podiform chromitites may result from magma mingling (Ballhaus, 1998) or from fluid-rich melts (Matveev & Ballhaus, 2002).

In the Kraubath Massif (Fig. 1), Eastern Alps, chromitites and coarse-grained orthopyroxenites have long been recognized as pods and veins in serpentinized harzburgite (Clar, 1929; Meixner, 1937; Angel, 1964). Metamorphosed, amphibole-rich rocks (‘hornblendites’), interpreted here as original clinopyroxenite, websterite and gabbro, have also been found, but their relations to the orthopyroxenites and harzburgites have never been investigated. In this paper, we present petrological and chemical data [major and trace elements including rare earth elements (REE) and platinum-group elements (PGE), and Sm–Nd, Rb–Sr and Re–Os isotope data] from metamorphosed harzburgite, chromitite, orthopyroxenite, clinopyroxenite and gabbro, to discuss the origin of this distinctive rock assemblage.

View larger version (69K): [in this window] [in a new window]   Fig. 1. Geological sketch maps illustrating the position of the Speik Complex in the Middle Austroalpine basement of Austria. (a) Major tectonic units of Austria after Frank (1987). (b) Generalized map of the Austroalpine units in Styria after Neubauer (1988). (c) Simplified map of the Kraubath Massif showing locations of orthopyroxenite (‘bronzitite’), clinopyroxenite and chromitite.

 
We propose that the ultramafic–mafic association represents refractory upper mantle intruded by a spectrum of dykes formed from a variety of melts, in a supra-subduction zone geotectonic setting during an Early Cambrian collision event. Because of the polymetamorphic history of the ultramafic and mafic protoliths in the Speik Complex, radiometric data have not previously been reported. Here, we present the first Sm–Nd and Re–Os isotopic data for this part of the Austroalpine basement complex.

	GEOLOGICAL BACKGROUND

TOP ABSTRACT INTRODUCTION GEOLOGICAL BACKGROUND ANALYTICAL METHODS PETROGRAPHY MINERAL CHEMISTRY WHOLE-ROCK CHEMICAL COMPOSITION DISCUSSION SUMMARY AND CONCLUSIONS REFERENCES

 
The Kraubath, Pernegg and Hochgrössen Massifs are part of the Speik Complex in the Middle Austroalpine basement nappe, Styria, Austria (Fig. 1a and b) (e.g. Tollmann, 1977; Frank, 1987). The Speik Complex rests tectonically on the Core Complex, a late Proterozoic to Early Palaeozoic magmatic arc assemblage, and is in turn overthrust by the Micaschist–Marble Complex of presumed Early Palaeozoic age (Neubauer & Frisch, 1993). Lithologically, the Speik Complex is composed of variably serpentinized ultramafic rocks associated with orthopyroxenites and clinopyroxenites, garnet–zoisite and plagioclase amphibolites (metagabbroic rocks), garnet-bearing and banded amphibolites (metabasaltic rocks), locally exposed eclogites (Faryad et al., 2002a), thin augengneiss lenses, and some metasedimentary rocks. The total estimated thickness is several hundred metres.

The Kraubath Massif (Fig. 1c), situated along the Mur valley on the northern flank of the Gleinalm domal structure (Fig. 1b; Neubauer et al., 1995), represents the largest outcrop of ultramafic rocks in the Eastern Alps (e.g. Clar, 1929; Hiessleitner, 1953; Angel, 1964; El Ageed et al., 1980; Haditsch, 1981; Haditsch et al., 1981). The ultramafic rocks are less serpentinized than in most of the other ultramafic massifs of the Speik Complex, and reveal abundant relics of primary minerals. Various deformation features are observed, including an early S₁ foliation attributed to high-temperature mantle deformation. This S₁ foliation dips approximately 45–70° towards the north in most parts of the Massif (Fig. 2a), but it is folded around ENE-trending axes in the easternmost part of the Massif (Haditsch et al., 1981). Porphyroclastic orthopyroxene- and amphibole-bearing meta-harzburgite with low concentrations of Al₂O₃ (0·66 wt % on average; Petersen-Krauß, 1979) is the dominant rock type; minor dunite is also present. Concentrations of chromian spinel either as small, often concordant veins, irregular lenses or as disseminations locally form small stratiform (Fig. 2b) and podiform schlieren-type deposits (Hiessleitner, 1953; El Ageed, 1979). Orthopyroxenites form discrete lenses, dykes and irregular bodies from several centimetres to tens of metres thickness. They are usually conformable with the foliation, and have been deformed together with their harzburgitic host rocks (Fig. 2a). Transposition of late intrusions into the foliation is possible only under high-temperature, plastic deformation. Some orthopyroxenite plugs are compositionally zoned with a zone of olivine-rich rock containing small accumulations of chromian spinel mantling olivine-free orthopyroxenite cores. Spatially, the orthopyroxenites appear to be associated with chromitite in the so-called ‘Main Dunite Zone’ (Hiessleitner, 1953), which is a thick and elongated zone mainly composed of metamorphosed harzburgite along the northern margin of the Massif (Fig. 1c). Towards the hanging wall, breccia-like textures with ‘dunite’ clasts enclosed by pyroxenite were noted by Hiessleitner (1953). Clar (1929) assumed that pyroxenites once formed a continuous, now tectonically disrupted, layer. Quantification of the percentage of orthopyroxenite at Kraubath is difficult; from 227 randomly distributed ultramafic rock samples chemically analysed by Petersen-Krauß (1979), only six were classified as olivine pyroxenite (2·5%), and eight as olivine–hornblende pyroxenite (3·5%). Hornblende–(diopside–garnet–zoisite) and diopside–olivine–tremolite rocks occur in a narrow zone towards the southern margin of the Kraubath Massif only.

View larger version (168K): [in this window] [in a new window]   Fig. 2. Field relations of rocks in the Speik Complex. (a) Boudinaged orthopyroxenite vein (Opx) in harzburgite (Hb), Windberg, Kraubath Massif. (b) Banded chromitite (Chr) in dunite in a mine adit, Sommergraben, Kraubath Massif. Size of compass is 7 cm. (c) Podiform chromitite (Chr) veinlets in harzburgite (Hb), Hochgrössen Massif.

 
In the Pernegg Massif, situated about 25 km ENE of Kraubath (Fig. 1b), a mafic–ultramafic rock assemblage rests tectonically on top of banded amphibolites of the Core Complex. Neubauer (1988) assumed that the ultramafic rocks and hornblende–zoisite and diopside–tremolite rocks, forming layers 1–10 m thick, represent a cumulate sequence. In the field, such rocks can be followed as thin layers in a laterally continuous zone over a distance of about 1 km along strike, 200 m above the tectonic base of the Massif. They are overlain by massive, serpentinized harzburgite similar in composition to harzburgite at Kraubath. Blocks of chromite-bearing dunite and orthopyroxenite have been found only in the northern and eastern portions of the Massif.

The Hochgrössen Massif forms the westernmost extension of the Speik Complex (Fig. 1b), which is sandwiched here between the Palaeozoic Greywacke Zone towards the north and the Micaschist–Marble Complex towards the south. A rootless body of strongly foliated serpentinite structurally overlies paragneiss and amphibolite of the Core Complex (El Ageed, 1979). Retrogressed eclogites form lenses subparallel to the strike of the foliation in eclogite and surrounding serpentinite (Faryad et al., 2002a). Stringers and pods (Fig. 2c) of massive chromite are locally common in the serpentinites (El Ageed, 1979; Thalhammer et al., 1990), which range in chemical composition from harzburgite to lherzolite carrying rarely preserved clinopyroxene (Melcher et al., 2002). Orthopyroxenites are notably lacking.

The magmatic and metamorphic evolution of the Speik Complex is complex. Whereas Clar (1929), Hiessleitner (1953) and Haditsch et al. (1981) regarded the Kraubath Massif as a differentiated ultramafic intrusion, El Ageed (1979), El Ageed et al. (1980) and Stumpfl & El Ageed (1981) recognized that the Hochgrössen and Kraubath Massifs resembled parts of ophiolite complexes. According to their interpretation, Hochgrössen represents an ultramafic cumulate sequence, whereas Kraubath represents residual mantle. This interpretation is based on the more Mg-rich chemistry of spinel at Hochgrössen, and on the REE patterns of the ultramafic and mafic rocks. The geochemical characteristics of metabasic rocks in the Pernegg and Gleinalm areas were later interpreted as a pre-Silurian, back-arc basin ophiolite with a ‘subduction-related component’ (Neubauer et al., 1989). Based on an extensive set of whole-rock geochemical data, Melcher et al. (2002) concluded that the Speik Complex probably represents the oldest oceanic crust preserved in the Eastern Alps, and that the harzburgites, lherzolites and pyroxenites record a two-stage magmatic history, including an ancient (Proterozoic) melt depletion event followed by second-stage melting in a supra-subduction zone setting during the Eocambrian or Cambrian.

	ANALYTICAL METHODS

TOP ABSTRACT INTRODUCTION GEOLOGICAL BACKGROUND ANALYTICAL METHODS PETROGRAPHY MINERAL CHEMISTRY WHOLE-ROCK CHEMICAL COMPOSITION DISCUSSION SUMMARY AND CONCLUSIONS REFERENCES

 
Samples of metaperidotite, orthopyroxenite, chromitite, metaclinopyroxenite and metagabbro from the Kraubath, Pernegg and Hochgrössen Massifs were selected for geochemical and mineral-chemical studies; the locations and mineralogical compositions of the samples are listed in Table 1. All samples were studied under the microscope in both transmitted and reflected light. Mineral compositions were determined using an ARL-SEMQ electron microprobe at the University of Leoben and a CAMECA SX100 electron microprobe at BGR, Hannover. Compositions of representative minerals are listed in Tables 2 and 3.

View this table: [in this window] [in a new window]   Table 1: Description and location of samples

 

View this table: [in this window] [in a new window]   Table 2. Representative electron microprobe analyses (in wt %) of major rock-forming minerals in orthopyroxenite

 

View this table: [in this window] [in a new window]   Table 3: Representative electron microprobe analyses (in wt %) of major silicates in metaperidotite, metaclinopyroxenite and metagabbro

 
Major and trace element analyses were obtained using X-ray fluorescence (XRF) and inductively coupled plasma mass spectrometry (ICP-MS) techniques. Analytical details have been described by Meisel et al. (2002) and Melcher et al. (2002). More whole-rock chemical compositions have been published by Melcher et al. (2002), and only new results are presented in Table 4. Trace element and sulphur concentrations of samples K84 and K102 were determined by the Service d'Analyse des Roches et des Minéraux du CNRS (CRPG–CNRS, Nancy).

View this table: [in this window] [in a new window]   Table 4: Major and trace element concentrations in ultramafic and mafic rocks

 
Rhenium and PGE concentrations were obtained on the same sample aliquots by isotope dilution ICP-MS at the University of Leoben following the technique described by Meisel et al. (2001a, 2001b, 2003a). Whole-rock samples and chromite concentrates obtained after electric-pulse disintegration were used. The powders were spiked and sealed with a mixture of concentrated HNO₃–HCl (5 and 2–3 ml, respectively) with Teflon tape in quartz glass vessels. The vessels were heated to 300°C at c. 125 bar for 3–5 h in a high-pressure asher (HPA-S, Anton Paar–Perkin Elmer Instruments). Osmium isotope ratios were either sparged directly as OsO₄ into the ICP quadrupole mass spectrometry (ICP-QMS) system at the University of Leoben or osmium was extracted from the solution using CCl₄ extraction (Cohen & Waters, 1996), followed by microdistillation in PFA vials (Birck et al., 1997). The remaining PGE were purified and enriched by anion exchange chromatography and measured by either ICP-QMS or ICP sector field mass spectrometry (ICP-SFMS) at the Institute of General and Analytical Chemistry at the University of Leoben. The concentrations and isotopic compositions of the extracted Os and the Re cuts of replicate digestions were measured by means of multi-collector ICP-SFMS (NU Plasma, NU instruments) at the Isotope Geology Group of the University of Bern (Schoenberg et al., 2000). Total procedural blanks of the HPA-S digestions and anion exchange separations were lower than 10 pg and thus negligible. As a result of Re contamination in the ICP-MS system, Re corrections on ¹⁸⁷Os had to be applied in some cases (PE21 and K84), causing an increase in measurement uncertainty of ¹⁸⁷Os/¹⁸⁸Os. Uncertainties for the isotope ratios have been given by Schoenberg et al. (2000). Two replicate analyses (samples K84 and MG45h) were carried out at CRPG–CNRS (Nancy) following low-pressure acid digestion procedures and Os extraction with Br₂ described by Birck et al. (1997) and Meisel et al. (2003b).

To determine Sm–Nd isotope compositions, whole-rock samples were processed at the Institute of Precambrian Geology and Geochronology in St. Petersburg, Russia. Approximately 1 g of the sample was weighed, spiked with mixed spikes (¹⁴⁶Nd–¹⁴⁹Sm) and decomposed for 3–5 days in a mixture of HF–HNO₃ in an oven at 120°C. Samarium and Nd separation was carried out according to the standard method of two-stage ion-exchange and extraction chromatography (Richard et al., 1976). All measurements were performed on a Finnigan MAT-261 mass spectrometer equipped with eight collectors in a static mode. The ¹⁴³Nd/¹⁴⁴Nd ratio was normalized within run to ¹⁴⁸Nd/¹⁴⁴Nd = 0·241570. The 2 errors for ¹⁴³Nd/¹⁴⁴Nd reflect in-run precision and demonstrate only the quality of the analyses. The blank levels for Sm and Nd were 50 ppt and 100 ppt, respectively. The data obtained for 11 runs of BCR-1 during the course of this analytical work are [Sm] = 6·49 ppm, [Nd] = 28·5 ppm, ¹⁴³Nd/¹⁴⁴Nd = 0·512644 ± 0·000008 and ¹⁴⁷Sm/¹⁴⁴Nd = 0·1382, and for the La Jolla standard were ¹⁴³Nd/¹⁴⁴Nd = 0·511830 ± 0·000012 (n = 22). Reproducibility of the ¹⁴⁷Sm/¹⁴⁴Nd ratio was 0·5%, and the precision for concentration measurements of Sm and Nd was 0·5%. One replicate analysis of orthopyroxenite (97So20) was carried out at CRPG–CNRS (Nancy). The Sm and Nd concentration data obtained on the same sample in two different laboratories via isotope dilution thermal ionization mass spectrometry (TIMS) and external calibration ICP-MS, respectively, compare well. Relative differences of the Sm/Nd ratio are below 4%, except for 12% for 97So20, which has the lowest Sm concentration (13 ppb).

	PETROGRAPHY

TOP ABSTRACT INTRODUCTION GEOLOGICAL BACKGROUND ANALYTICAL METHODS PETROGRAPHY MINERAL CHEMISTRY WHOLE-ROCK CHEMICAL COMPOSITION DISCUSSION SUMMARY AND CONCLUSIONS REFERENCES

 
Most ultramafic rocks of the Speik Complex are variably serpentinized (from 100% to <10% serpentine) metamorphosed harzburgites. In the Hochgrössen Massif, they are completely converted into serpentinite; in the Kraubath and Pernegg Massifs, porphyroclastic harzburgites and rare granoblastic dunites carry partly preserved pre-metamorphic mineral assemblages. Primary mineral relics include olivine (Ol), orthopyroxene (Opx) and chromian spinel (Spl). The metamorphic assemblages consist of forsteritic olivine (abundant in the Hochgrössen Massif, less abundant at Pernegg and Kraubath), serpentine (Srp; lizardite, chrysotile, rare antigorite), abundant tremolite (Tr), talc (Tc), chlorite (Chl), rare anthophyllite (At), and carbonates (calcite, magnesite). Metamorphic olivine was observed only in highly deformed, sheared serpentinites. A metaperidotite sample from Pernegg (PE31) carries relics of olivine, orthopyroxene and pargasitic amphibole, replaced by tremolite, anthophyllite, chlorite and serpentine (Fig. 3a).

View larger version (149K): [in this window] [in a new window]   Fig. 3. Back-scattered electron images. (a) Hornblende replaced by tremolite and chlorite in lherzolite, Pernegg Massif, sample PE31. (b) Orthopyroxene with relict exsolution lamellae; metamorphic olivine; tremolite with spinel inclusion rimmed by anthophyllite; orthopyroxenite, Kraubath, sample K84. (c) Sulphide (pyrrhotite and pentlandite) and ilmenite interstitial to magmatic orthopyroxene, and metamorphic tremolite, chlorite, olivine; orthopyroxenite, Kraubath, sample K84. (d) Euhedral chromite in orthopyroxene, replaced by talc and chlorite; orthopyroxenite, Kraubath, sample MG45h. At, anthophyllite; Chl, chlorite; Hbl, hornblende; Ilm, ilmenite; Ol, olivine; Opx, orthopyroxene; Pn, pentlandite; Po, pyrrhotite; Spl, spinel; Srp, serpentine; Tc, talc; Tr, tremolite.

 
Podiform, schlieren-type chromitite lenses and veinlets, a few centimetres thick, composed of massive chromite with included olivine, clinopyroxene, amphibole and chlorite are abundant in metamorphosed harzburgites and dunites in the northern part of the Kraubath Massif (Hiessleitner, 1953) and in the Hochgrössen Massif (El Ageed, 1979) (Fig. 2c). More continuous bands up to 30 cm thick in the Sommergraben area at Kraubath (Fig. 2b) are composed of densely disseminated chromite; this chromite type carries abundant Fe–Ni–Cu sulphide inclusions and a PGE mineral assemblage considerably different from the schlieren-type chromitite (Thalhammer et al., 1990, 2001; Melcher & Mali, 1998; Melcher et al., 1998; Malitch et al., 2001, 2003a, 2003b). Texturally primary, usually euhedral to subhedral spinel grains (chemically unzoned or zoned) are distinguished from secondary, metamorphic spinel occurring as two types, (1) overgrowing or (2) replacing primary spinel together with chlorite and serpentine, resulting in typical intergrowth textures.

The orthopyroxenites exhibit polygonal granular textures with frequent 120° triple junctions between pyroxene grains. Mineral assemblages are (1) pre-metamorphic, magmatic Opx ± Ol ± Spl ± Ilm (ilmenite) ± Fe–Cu–Ni sulphides, and (2) a metamorphic assemblage of Tr + Tc + Srp + Chl ± Ol ± At ± Fe–Cu–Ni sulphides. Orthopyroxene forms large (up to 5 cm) clasts, which are frequently deformed (kinked). Tremolitic amphibole is the only calcium-bearing phase, forming euhedral crystals of less than 1 mm size either interstitial to, or replacing orthopyroxene along cleavage planes (Figs 3a, c, and 4a, b). Anthophyllite (At) occurs in close association with tremolite, talc, olivine and carbonate, as coronas around tremolite (Figs 3b and 4b), as fracture fillings, or in massive amphibole–(talc–olivine–antigorite) schist considered as altered orthopyroxenite (Schantl, 1979, 1982). Trails of inclusions (spinel, chromian chlorite) parallel to crystallographic planes, as well as single subhedral to euhedral chromian spinel and/or anhedral sulphide grains are common in orthopyroxene (Fig. 3c). Along cleavage planes and grain boundaries, orthopyroxene is replaced by talc (Figs 3d and 4a). Olivine is occasionally present as small, elongated single grains or chains of grains along pyroxene grain boundaries (Fig. 3b and c). Chromian spinel forms small (<0·2 mm), subhedral or euhedral grains interstitial to or included in pyroxene (Fig. 3b and d). Accumulations of chromian spinel have been observed in the olivine-rich contact zone of orthopyroxenite with harzburgite (K64). Ilmenite is present in the most Fe-rich orthopyroxenite (K84; Fig. 3c). Small sulphide grains (bornite, pentlandite) are ubiquitous in most orthopyroxenites, usually occurring as texturally late mobilizates in serpentine-rich domains. The most massive orthopyroxenites carry abundant sulphides as intergranular and intragranular blebs consisting of intergrown pyrrhotite, pentlandite (Fig. 3c) and chalcopyrite.

View larger version (154K): [in this window] [in a new window]   Fig. 4. Photomicrographs (crossed polars) of metapyroxenites and metagabbro. (a) Growth of metamorphic tremolite and talc along grain boundaries of primary orthopyroxene; orthopyroxenite, Kraubath, sample 97So20. (b) Tremolite rimmed by anthophyllite in orthopyroxenite, Pernegg, sample PE20. (c) Clinopyroxene (Cpx) relics replaced by diopside neoblasts (Di) and tremolitic amphibole (Tr) in metaclinopyroxenite, Pernegg, sample PE27. (d) Garnet (Grt), relict clinopyroxene (Cpx) and two types of amphibole (Hbl 1, Hbl 2) in metagabbro, Kraubath, sample KR25.

 
Thin layers of strongly foliated, locally mylonitic diopside–olivine–tremolite schist are interlayered with serpentinized harzburgite in the southern portions of the Kraubath and Pernegg Massifs. In this rock type, clinopyroxene (1–2 mm in diameter) and rare relict olivine (K73) form a magmatic cumulate texture (Fig. 4c), with interstitial metamorphic tremolite, olivine (PE27) and diopside. The rocks are rich in chlorite and serpentine, and contain chromian magnetite with rare cores of chromite. In the following discussion, these rocks are referred to as metaclinopyroxenites.

Medium- to coarse-grained amphibole-rich rocks are associated with metaclinopyroxenites. In one sample (KR25), relict clinopyroxene and garnet are preserved (Fig. 4d). Most of the samples are completely recrystallized to coarse-grained (partly >5 mm in size) amphibole rocks. Amphibole may be aligned along a dominant foliation, or forms large porphyroblasts that show evidence of dynamic recrystallization. Zoisite/clinozoisite is common, but plagioclase is notably absent in the samples studied. Additional phases are chlorite, epidote, and ilmenite rimmed by sphene. In the following discussion, the amphibole-rich rocks lacking plagioclase are referred to as metagabbros.

	MINERAL CHEMISTRY

TOP ABSTRACT INTRODUCTION GEOLOGICAL BACKGROUND ANALYTICAL METHODS PETROGRAPHY MINERAL CHEMISTRY WHOLE-ROCK CHEMICAL COMPOSITION DISCUSSION SUMMARY AND CONCLUSIONS REFERENCES

 
Olivine
The forsterite content (Fo) in olivine from the metamorphosed ultramafic rocks of the Speik Complex ranges from 81 to 99 (Fig. 5a, Tables 2 and 3). Three compositional types are distinguished. (1) In most orthopyroxenites, olivine interstitial to orthopyroxene varies from Fo₈₇ to Fo₈₁ and NiO varies between 0·25 and 0·08 wt % (Table 2, Fig. 5a). In a chromite- and olivine-rich rock (K64) mantling a large orthopyroxenite plug, olivine has an mg-number of 84 and <0·2 wt % NiO. (2) Relict olivine grains in metaperidotite, peridotite-hosted chromitite and metaclinopyroxenite from Kraubath and Pernegg range from Fo₈₉ to Fo₉₅ and have between 0·05 and 0·9 wt % NiO (Fig. 5a). Lowest mg-numbers (Fo₈₉) are found in a fertile peridotite from the Pernegg Massif (PE 31). Texturally, olivine types (1) and (2) may represent relict primary phases. However, their highly variable NiO concentrations suggest that metamorphic equilibration probably has taken place. (3) In serpentinites from the Hochgrössen and in a few samples from the Pernegg Massif, olivine intergranular to coarse antigorite has high mg-numbers (>96; Fig. 5a). Their texture suggests a metamorphic origin. Metamorphic olivine in textural equilibrium with tremolite in metaclinopyroxenite has mg-numbers of 92–93 and <0·2 wt % NiO.

View larger version (35K): [in this window] [in a new window]   Fig. 5. Chemistry of olivine and pyroxenes in metaperidotite, chromitite, orthopyroxenite, metaclinopyroxenite and metagabbro of the Speik Complex. (a) mg-number vs wt % NiO in olivine. (b) mg-number vs wt % Al2O3 in orthopyroxene; the most Al-rich orthopyroxenes (up to 3·8 wt % Al2O3 at mg-numbers of 90) from sample PE31 have been omitted from this plot. (c) mg-number vs wt % CaO in orthopyroxene. (d) mg-number vs wt % Al2O3 in clinopyroxene. HG, Hochgrössen Massif; KR, Kraubath Massif.

 
Orthopyroxene
Orthopyroxene is considered to be a primary magmatic phase in both orthopyroxenites and metaperidotites based on preserved high-temperature deformation textures. mg-numbers vary between 84 and 93 (Tables 2 and 3; Fig. 5b and c), and wollastonite components (wo) are <0·7 mol %. High mg-numbers (>90) are recorded in harzburgite and in narrow orthopyroxenite dykes in harzburgite, whereas massive pegmatoidal orthopyroxenites have mg-numbers as low as 84. Al₂O₃ concentrations are below 0·5 wt %, and CaO is below 0·35 wt % in orthopyroxenites (Fig. 5b and c); between 1·0 and 3·8 wt % Al₂O₃ is recorded in opx from more Al-rich metaperidotites of the Pernegg Massif (Table 3).

Clinopyroxene
Clinopyroxene is rarely observed in metaperidotite, whereas it is more abundant as inclusions in podiform chromitite and as relict grains in metaclinopyroxenite and metagabbro. Porphyroclasts of clinopyroxene in an Al- and Ca-rich serpentinite of the Hochgrössen Massif (HG4) are diopsides with cores having higher Al₂O₃ and lower mg-numbers than rims (Table 3, Fig. 5d). Chromite-hosted diopside in chromitite has low Al₂O₃ concentrations (up to 0·4 wt %) and high mg-numbers (95–96).

In metaclinopyroxenite, two compositional varieties of clinopyroxene are present: (1) cores have mg-numbers of 86–93, up to 1 wt % Al₂O₃ and <0·3 wt % Na₂O; (2) metamorphic rims and neoblasts are pure diopside with mg-numbers of 97–98 and very low concentrations of Na, Al and Ti. Clinopyroxene forming relict cores in metamorphic amphibole in a metagabbro sample from Kraubath (KR25) is characterized by high mg-numbers (85–87), 1·0–3·1 wt % Al₂O₃, 0·1–0·3 wt % TiO₂ and 0·2–0·5 wt % Na₂O (Table 3).

Amphibole
The dominant amphibole in metamorphosed ultramafic rocks of the Kraubath and Pernegg Massifs is Al-poor to Al-bearing (0·3–3·6 wt % Al₂O₃) tremolite or tremolitic hornblende with high mg-numbers (94–98) and 7·0–8·0 Si a.p.f.u. (Tables 2 and 3; Fig. 6). In amphiboles enclosed by chromite in chromitite, concentrations of Al, Ti, Cr and Na are higher than in orthopyroxenite and metaharzburgite. In metaclinopyroxenite, tremolite has a restricted compositional range with mg-numbers of 94–98, and low concentrations of Al₂O₃ (0·12–0·38 wt %), TiO₂ and Na₂O (Fig. 6).

View larger version (30K): [in this window] [in a new window]   Fig. 6. Compositional variation (atoms Si per formula unit vs mg-number and Si vs (Na + K)A of calcic amphiboles in ultramafic rocks, pyroxenites and metagabbros of the Speik Complex.

 
Anthophyllite, which rims tremolite in some orthopyroxenites (Figs 3b and 4b), has mg-numbers of 81–83 and low CaO concentrations (Table 2). It was also observed as part of the metamorphic assemblage in a lherzolitic sample (PE31). In addition, this sample carries magnesio-hastingsite to pargasite as relicts in tremolite. These amphiboles have 3 wt % Na₂O, 6·4–6·6 Si a.p.f.u., mg-numbers of 86–89 and considerable concentrations of Cl (0·2–0·5 wt %, Table 3).

In metagabbro, amphiboles are characterized by mg-numbers ranging from 45 to 80, Si from 5·8 to 7·0 a.p.f.u., 1·2–3·0 wt % Na₂O and 0·1–0·6 wt % K₂O (Fig. 6). They are classified as pargasitic to edenitic hornblende. Distinct prograde and weak retrograde zoning patterns are recognized. A garnet- and clinopyroxene-bearing sample (KR25) contains two different amphibole generations (Table 3): amphibole I with 6·4–6·8 Si a.p.f.u. and mg-numbers of 56–77 forms cores to more Al-rich amphibole II with 5·8–6·1 Si a.p.f.u. and mg-numbers of 45–52; amphibole II is also found rimming clinopyroxene (Fig. 4d). In a coarse-grained hornblende–zoisite rock (PE22), cores of magnesiohornblende (6·7–6·9 Si a.p.f.u., mg-number 72) are mantled by narrow rims of tschermakitic hornblende (6·2–6·5 Si a.p.f.u., mg-number 65).

Chlorite
Mg-rich chlorite [mg-numbers 93–97, 5·5–6·8 Si a.p.f.u.; clinochlore and penninite according to the Hey (1954) classification] is present in all metamorphosed ultramafic rocks studied. In orthopyroxenite, metaharzburgite and chromitite of the Kraubath and Pernegg Massifs, chlorite has high Cr₂O₃ concentrations (3–4 wt %; Table 2), whereas in clinopyroxenites (Pernegg Massif) and serpentinites of the Hochgrössen Massif, only up to 1 wt % Cr₂O₃ is recorded. Chlorite in metagabbro spans a large range in mg-numbers (67–95; pycnochlorite to clinochlore), depending on the phases it replaces.

Garnet
Two compositional types of garnet were identified in a metagabbro sample from Kraubath (KR25): garnet I is high in Mg (pyrope component), whereas garnet II is rich in Ca (grossularite) and higher in Mn (spessartine) than garnet I (Table 3). The composition of garnet I is Alm_36–47Grs_13–30Prp_16–39Adr_0–11Sps_1–6, and of garnet II is Alm_33–46Grs_38–51Prp_0–11Adr_0–8Sps_3–11. Garnet II crosscuts garnet I along fractures of preferred orientation and texturally coexists with amphibole II and chlorite. Garnet I is generally mantled either by garnet II, or by amphibole II and chlorite, and is not in textural equilibrium with amphibole I or clinopyroxene.

Spinel
Compositions of chromian spinel in metaperidotites and chromitites of the Speik Complex are fairly similar. Their cr-numbers [100 x Cr/(Cr + Al); 55–95, with most analyses between 75 and 90] and mg-numbers [100 x Mg/(Mg + Fe²⁺); 20–70] span a wide compositional range (Fig. 7a). However, a general grouping according to their mg-numbers is possible: (1) spinel from the Kraubath and Pernegg Massifs is more Fe-rich, having mg-numbers between 20 and 50; (2) spinel from the Hochgrössen Massif is more Mg-rich, having mg-numbers of 45–70. The most distinctive feature of spinel in the Speik Complex is zoning: cores are more Al- and Mg-rich than rims, which consist of ferrite–chromite and chromian magnetite (Thalhammer et al., 1990). Concentrations of TiO₂ are usually <0·2 wt %, but exceed 1 wt % in places (Fig. 7b). Highly deformed and metamorphically recrystallized spinel has higher concentrations of MnO (up to 3 wt %) and the ratio Fe³⁺/R³⁺ [100 x Fe³⁺/(Fe³⁺ + Al + Cr)] is >15.

View larger version (31K): [in this window] [in a new window]   Fig. 7. Spinel composition in ultramafic rocks (fields) and pyroxenites (circles) of the Speik Complex. Grey fields delineate different types of spinel (T-I, II, IV) from the Kraubath Massif, hatched fields those from the Hochgrössen Massif. (a) Fe2+/(Mg + Fe2+) vs Cr/(Cr + Al). (b) Fe3+/(Fe3+ + Cr + Al) vs wt % TiO2. (c) Fe2+/(Mg + Fe2+) vs Fe3+/(Fe3+ + Cr + Al).

 
Melcher et al. (1999b) and Melcher (2000) distinguished four textural and chemical types of spinel in chromitites of the Speik Complex (Table 5). Types 1 and 2 form small orebodies, which mainly differ in their shape (deformed schlieren vs planar banded type), cr-number, and inclusion assemblage. Type 3 chromite, replacing or overgrowing magmatic chromite, is characterized by a wide range of cr-numbers and mg-numbers (‘metamorphic type’). All chromites analysed from the Hochgrössen and Pernegg Massifs follow this trend. Type 4 chromite is associated with orthopyroxenite.

View this table: [in this window] [in a new window]   Table 5: Textural–chemical classification of chromite types in the Speik Complex

 
Most orthopyroxenites have chromian spinel as a primary magmatic phase, which forms either interstitial to pyroxene, as inclusions, or along trails enclosed by pyroxene. In general, cr-numbers are high (77–97; Fig. 7a) and mg-numbers are low (33–15); Fe³⁺/R³⁺ [=100 x Fe³⁺/(Fe³⁺ + Al + Cr)] ranges between 3 and 18 (Fig. 7b and c) (Table 2). The concentrations of TiO₂ reach 3·8 wt % in the most Fe-rich orthopyroxenites, but range between 0·2 and 1 wt % in the remaining samples (Fig. 7b). Although within-sample variations are considerable, spinel compositions tend to develop from more Al-, Mg-rich to more Cr-, Fe-, Ti-rich compositions with decreasing mg-numbers in their host rocks. In a diagram of cr-number vs 100 x Fe²⁺/(Mg + Fe²⁺), spinel compositions follow trends of metamorphically equilibrated spinel (Fig. 7a), but plot outside the field typically given for ophiolitic chromites.

In metaclinopyroxenites, Cr-rich spinel (up to 45 wt % Cr₂O₃, <3·5 wt % MgO, <7·2 wt % Al₂O₃) is rarely present as relict cores in chromian magnetite that contains up to 13 wt % Cr₂O₃, <1·3 wt % MgO and <0·1 wt % Al₂O₃.

Ilmenite
Ilmenite is present in addition to Ti-rich spinel in the most Fe-rich orthopyroxenites and in a lherzolite sample (PE31); it contains minor MnO (up to 1·5 wt %) and MgO (up to 4 wt %). Small inclusions of ilmenite within chromitite associated with orthopyroxenite have similar composition. Ilmenite rimmed by sphene occurs in some metagabbros.

	WHOLE-ROCK CHEMICAL COMPOSITION

TOP ABSTRACT INTRODUCTION GEOLOGICAL BACKGROUND ANALYTICAL METHODS PETROGRAPHY MINERAL CHEMISTRY WHOLE-ROCK CHEMICAL COMPOSITION DISCUSSION SUMMARY AND CONCLUSIONS REFERENCES

 
Major and trace elements
The chemical composition of metagabbros, and of some orthopyroxenite and metaclinopyroxenite samples is presented in Table 4. A large set of additional data on ultramafic rocks of the Speik Complex was published by Melcher et al. (2002).

Metaperidotites in all three massifs studied have mg-numbers [100 x Mg/(Mg + Fe^tot)] of 89–91 and average 0·66 wt % Al₂O₃ (0·1–3·5 wt %) and 0·42 wt % CaO (0·01–2·6 wt %), respectively. Orthopyroxenites from Kraubath differ from their host harzburgites by having higher SiO₂ (52–57 wt % vs 35–44 wt % in peridotites), and lower MgO (31–35 wt % vs 34–46 wt %). Concentrations of Al₂O₃ (0·5–0·7 wt %), CaO (<1·5 wt %), TiO₂ (<0·05 wt %), MnO (0·13–0·19 wt %), Na₂O, K₂O and P₂O₅ are low and do not differ significantly from values in associated harzburgite and dunite (Fig. 8). Relatively low values for loss on ignition (0·7–5·4% compared with 5–10% for harzburgites) are explained by the relative freshness of the rocks resulting from the fact that there is only minor alteration of orthopyroxene compared with olivine. Chromium concentrations are high in peridotites and orthopyroxenites, ranging from 2200 to 3300 ppm (Fig. 8), whereas Ni is significantly lower in orthopyroxenite than in harzburgite (<1000 ppm vs an average of 2100 ppm for Kraubath peridotites). mg-numbers range from 85·5 to 91·8, and thus are within the range of harzburgite and dunite (Melcher et al., 2002; Fig. 8). Most orthopyroxenites and harzburgites have sulphur concentrations below 100 ppm; however, the two most massive orthopyroxenites carry 2300 and 2700 ppm sulphur, along with elevated Cu (221–276 ppm) and Re (4·4–6·4 ppb). Some trace elements such as Sc, Th, V, Y, and rare earth elements (REE) are slightly higher in orthopyroxenite than in harzburgite.

View larger version (15K): [in this window] [in a new window]   Fig. 8. Whole-rock geochemical data: wt % SiO2 vs ppm Cr, wt % CaO, wt % Al2O3 and mg-number for orthopyroxenite, metaclinopyroxenite and metagabbro. Metaperidotites (harzburgites, dunites, lherzolites) are summarized in a field [data from Melcher et al. (2002)].

 
The metaclinopyroxenites have higher CaO (5·5–17·3 wt %), Sc, V, and Sr, slightly higher TiO₂ (0·07–0·11 wt %) and lower SiO₂ (41–50 wt %) than orthopyroxenites. Concentrations of Al₂O₃ (0·8–1·2 wt %), alkalis, and mg-numbers (82–85) are similar, or correspond to the most Fe-rich orthopyroxenites (Fig. 8). High Cr (2000–3000 ppm) contents are noteworthy.

Metagabbros vary considerably in SiO₂ (41–48 wt %), Al₂O₃ (9·2–19·4 wt %), MgO (9·5–14·9 wt %), and Fe₂O₃ (8·4–13·2 wt %); mg-numbers range from 61 to 78 (Fig. 8). The variations in CaO (12·1–16·2 wt %), Na₂O (1·5–2·7 wt %), and TiO₂ (0·3–0·7 wt %) are smaller. Metagabbros have variable concentrations of Cr (200–1100 ppm), V (320–650 ppm), Ni (52–280 ppm) and Sr (80–560 ppm), and low Zr (3–27 ppm).

Most ultramafic rocks of the Speik Complex have very low REE concentrations [2·6–36 ppb Yb, 0·01–0·16 times chondritic composition (x CN)] and display U-shaped REE patterns with usually positively sloping, concave-upward middle REE–heavy REE (MREE–HREE) parts to the curves (Fig. 9; Melcher et al., 2002). Orthopyroxenites have on average slightly higher HREE concentrations (maximum of 64 ppb Yb; Table 4) than the harzburgites, but the shape of chondrite-normalized REE patterns is similar (Fig. 9a). A few metaperidotites, such as samples PE31 and HG4, have REE patterns typical of fertile lherzolites (Fig. 9b).

View larger version (36K): [in this window] [in a new window]   Fig. 9. Chondrite-normalized REE diagrams for (a) pyroxenites and (b) metagabbros. REE concentrations of pyroxenite samples (PE27, PE28, K102, 97So20, MG45h), and of metaperidotites from the Speik Complex are taken from Melcher et al. (2002). Metaperidotites are grouped into chondritic (<2) and suprachondritic samples (>5) using their Os(T) values. The harzburgite field in (a) covers most metaperidotites from Kraubath, Hochgrössen and Pernegg. The normalizing values for CI-chondrites of Boynton (1984) have been used.

 
Metaclinopyroxenites have distinctly different, almost flat, chondrite-normalized REE patterns with 86–305 ppb Yb (Fig. 9a). La_N/Yb_N, La_N/Sm_N, and Gd_N/Yb_N all are greater than unity (1·04–1·77), and all samples have a slight negative Eu anomaly.

Metagabbros display variable types of REE patterns. All have Gd_N/Yb_N >1 (1·2–2·0) (Fig. 9b). Two samples (PE22, PE24) are light REE (LREE)-enriched (La_N/Sm_N, 4·4–5·2) and show positive Eu anomalies, but differ considerably in their MREE–HREE concentrations. Sample PE22 has lower HREE, in the range of metaclinopyroxenite (<2 x CN). The MREE–HREE curves of PE24 (amphibole only) and of KR25 (amphibole + garnet + clinopyroxene) are virtually identical (about 5–6 x CN), but these patterns differ in their LREE, with KR25 being LREE-depleted (La_N/Sm_N, 0·8). The sample with the highest REE concentrations (PE29) has a ‘hump-shaped’ pattern with La_N/Sm_N <1, Gd_N/Yb_N >1 and no Eu anomaly.

Concentrations of PGE
Concentrations of PGE and Re have been determined in a number of peridotites, pyroxenites (Table 6) and chromitites (Table 7). Repeated measurements have been performed for most samples using different chemical procedures. Averages of concentrations are normalized to primitive mantle values (McDonough & Sun, 1995) and plotted in Figs 10 and 11.

View larger version (39K): [in this window] [in a new window]   Fig. 10. Platinum-group element and Re concentrations of metaperidotites and pyroxenites normalized to silicate Earth-‘pyrolite’ (McDonough & Sun, 1995). (a) Metaperidotites group 1; (b) metaperidotites group 2; (c) orthopyroxenites; (d) clinopyroxenites.

 

View larger version (21K): [in this window] [in a new window]   Fig. 11. Platinum-group element and Re concentrations of chromitites normalized to silicate Earth-‘pyrolite’ (McDonough & Sun, 1995). Chromitites are grouped according to their curve shapes, as discussed in the text. (a) Kraubath, type 1 chromitites; (b) Kraubath, type 2, 3 and 4 chromitites; (c) Hochgrössen chromitites.

 

View this table: [in this window] [in a new window]   Table 6: Re–Os isotopic composition and PGE concentrations (in ppb) of ultramafic rocks

 

View this table: [in this window] [in a new window]   Table 7: Re–Os isotopic composition and PGE concentrations (in ppb) in chromite fractions and whole rock from chromitite and harzburgite

 
PGE totals in metaperidotites range from <10 to >100 ppb, with most samples having between 14 and 34 ppb. Two types of chondrite-normalized PGE patterns are observed. (1) The first type has flat patterns at almost mantle concentrations without significant fractionations from Os to Pd (HG4, HG56, PE31, MG41, K59); Re_N/Os_N is <1. A dunite (MG40) close to a chromitite has IPGE (Os, Ir, Ru) concentrations similar to average harzburgite, but higher PPGE (Rh, Pt, Pd); Pd_N/Os_N = 4·7 and Re_N/Os_N = 0·3 (Fig. 10a). Some harzburgites are depleted in Pd and Re compared with the IPGE. (2) The second type has positively sloping patterns with low concentrations of the IPGE (0·1 times mantle), but significant fractionation between IPGE and PPGE and Pd_N/Os_N ratios of 10–50 are observed in samples LO3, HG23, HG52 and PE21; Re_N/Os_N = 2–17 (peridotite group 2; Fig. 10b). The PPGE are also significantly fractionated, usually with Pd_N/Pt_N >1. Re is depleted relative to Pd in most, but not in all peridotites analysed.

Orthopyroxenites have between 15 and 200 ppb PGE and are characterized by positively sloping PGE patterns (Fig. 10c) with high Pd_N/Os_N (10 to >100) and Pt_N/Ir_N (8–35), but variable Pd_N/Pt_N (0·5–24). The low concentrations of the IPGE, especially of Os (0·1 ppb), are puzzling, because most orthopyroxenites carry accessory chromite, which is considered a common carrier of the IPGE (e.g. Shirey & Walker, 1998). Elevated Pd and Re concentrations correlate with S and Cu, which are mainly hosted by interstitial sulphides. Rhenium concentrations range from 0·03 to 4·6 ppb, respectively from values <0·1 to 20 times the average mantle value. The mantle-normalized PGE patterns of metaclinopyroxenites are even more fractionated than those of orthopyroxenites (Pd_N/Os_N = 12–925, Pd_N/Pt_N = 4–12); PGE totals range from 17 to 240 ppb (Fig. 10d).

Chromitites from the Kraubath and Hochgrössen Massifs have a variety of normalized PGE patterns (Melcher et al., 1999b; Malitch et al., 2001, 2003a), and PGE concentrations range from <50 ppb to 4 ppm. The PGE are enriched compared with primitive mantle, whereas Re is mostly depleted by a factor of 10. Many chromitites have negative slopes typical for ophiolitic–podiform chromitite (many type 1 and type 3 chromitites; Fig. 11a and c). Also present are flat patterns parallel to some of their host harzburgites (many type 2 chromitites; Fig. 11b), zig-zag patterns with positive peaks of one or more of the PGE (Fig. 11b), and positive slopes similar to metaperidotite group (2) and to the pyroxenites (type 4 chromitite; Fig. 11b). Re_N/Os_N ratios are below 0·2 in all samples, except for chromitite associated with orthopyroxenite (K64) and spinel fractions separated from completely serpentinized harzburgite in the Pernegg Massif.

Re–Os and Sm–Nd isotope systematics
Re–Os isotopes
In most metaperidotites of the Speik Complex, ¹⁸⁷Os/¹⁸⁸Os ratios vary from 0·1185 to 0·1288 (Fig. 12a). Samples with positively sloping (Pd_N/Ir_N > 1) normalized PGE patterns (Fig. 10b) have significantly lower Os (<1 ppb) and/or higher Re concentrations (up to 0·8 ppb); these samples (metaperidotite group 2) yield radiogenic ¹⁸⁷Os/¹⁸⁸Os ratios ranging from 0·17 to 0·77 (Table 6). They show similar LREE enrichment to the unradiogenic peridotites (Fig. 9b). The Os isotopic data of the metaperidotites contain no or little age information, because of (1) large scatter in ¹⁸⁷Os/¹⁸⁸Os, and (2) low Re/Os ratios. The ¹⁸⁷Os/¹⁸⁸Os ratios of most metaperidotites are within the range obtained for most peridotite-hosted chromitites at Kraubath and Hochgrössen (0·1236–0·1270; Fig. 12b, Table 7). Few chromitites have suprachondritic ¹⁸⁷Os/¹⁸⁸Os ratios (0·13–0·16). ¹⁸⁷Re/¹⁸⁸Os ratios are lower than 0·05, except in type 4 chromitite hosted by orthopyroxenite.

View larger version (24K): [in this window] [in a new window]   Fig. 12. Re–Os isotope diagrams for peridotites, pyroxenites and chromitites of the Speik Complex. (a) 187Re/188Os vs 187Os/188Os diagram. The bulk of the data plot within the indicated range. (b) Log 187Re/188Os vs 187Os/188Os diagram for samples with low Re concentrations. Most chromitites and peridotites have very low Re–Os ratios; the x-scale is therefore logarithmic. The range of in situ LA–ICP-MS measurements for laurite included in chromite is also given (Malitch et al., 2003b). (c) 187Re/188Os vs 187Os/188Os diagram showing an errorchron for orthopyroxenites only. (d) Internal isochron calculated for orthopyroxenite sample 97So20.

 
Radiogenic ¹⁸⁷Os/¹⁸⁸Os ratios are encountered in orthopyroxenite (Fig. 12c and d), metaclinopyroxenite and chromitite associated with orthopyroxenite (Tables 7 and 8). Replicate runs of Os isotopic compositions and Os concentrations of these samples are not very reproducible. This variance is not caused by uncertainties in the analytical procedure but by the nugget effect of PGE-rich trace mineral phases (Meisel et al., 2001b), as demonstrated by the linear relationship between ¹⁸⁷Os/¹⁸⁸Os and 1/Os (Fig. 13). Osmium concentrations range from c. 0·1 to 0·8 ppb. On the other hand, Re concentrations are very reproducible, indicating that Re is mainly located in other mineral phases, probably in interstitial sulphides. All samples have suprachondritic, sometimes highly radiogenic, ¹⁸⁷Os/¹⁸⁸Os (0·14 to >10) and high ¹⁸⁷Re/¹⁸⁸Os (0·7 to >400).

View larger version (13K): [in this window] [in a new window]   Fig. 13. 187Os/188Os vs 1/Os for different splits of samples 97So20 and MG45h.

 

View this table: [in this window] [in a new window]   Table 8: Results of Sm–Nd isotope analyses on whole-rock samples from the Speik Complex

 
For the orthopyroxenites, the replicate analyses plot on a model-1 errorchron (MSWD = 32) yielding 550 ± 17 Ma and ¹⁸⁷Os/¹⁸⁸Os_(i) = 0·179 ± 0·003 (Fig. 12c). This age is dominated by Re-rich mineral phases, whereas the suprachondritic ¹⁸⁷Os/¹⁸⁸Os is dominated by Os-rich nuggets (Meisel et al., 2001b). A model-1 age (Ludwig, 1999) was calculated to give the data on the lower end of the errorchron more weight. This is necessary as the initial ratio is actually fixed through the replicate analysis of 97So20. The replicates of 97So20 define an internal isochron with a model-1 solution defining a slightly younger age of 519 ± 13 Ma (¹⁸⁷Os/¹⁸⁸Os_(i) = 0·1798 ± 0·0006; Fig. 12d). The sample heterogeneity of orthopyroxenite MG45h, a thin pod enclosed by harzburgite, is expressed by the wide range in its Os isotopic composition (¹⁸⁷Os/¹⁸⁸Os 0·13–0·145) and by the positive trend in an ¹⁸⁷Os/¹⁸⁸Os vs 1/Os plot (Fig. 13). The isotopic composition is significantly lower than that of the other orthopyroxenite samples and is closer to values typical for the peridotites from the Speik Complex or the primitive upper-mantle value (PUM, 0·1296; Meisel et al., 2001c).

Sm–Nd isotopes
¹⁴³Nd/¹⁴⁴Nd values in harzburgite and lherzolite whole-rock samples from the Speik Complex vary from 0·51251 to 0·51364, and ¹⁴⁷Sm/¹⁴⁴Nd from 0·107 to 0·388 in (Table 8). The peridotites define a model-3 solution age (Ludwig, 1999) of 745 ± 44 Ma with an initial _Nd(i) of +6 (Fig. 14a). The peridotites, although sampled from three different ultramafic massifs within the Speik Complex, appear to be cogenetic. _Nd(T) varies from 4·6 to 7·4 in the harzburgites, and depleted mantle model ages are between 631 and 962 Ma.

View larger version (15K): [in this window] [in a new window]   Fig. 14. Sm–Nd isochron diagram for (a) pyroxenites and metagabbros and (b) peridotites from the Speik Complex.

 
Orthopyroxenites from Kraubath have low ¹⁴³Nd/¹⁴⁴Nd and ¹⁴⁷Sm/¹⁴⁴Nd ratios (Table 8, Fig. 14b), whereas metaclinopyroxenite (PE27) from Pernegg has higher values. Two analyses of garnet-bearing LREE-poor metagabbro from Kraubath (KR25) gave high ¹⁴⁷Sm/¹⁴⁴Nd and ¹⁴³Nd/¹⁴⁴Nd = 0·5129; the REE-enriched metagabbro from Pernegg (PE29) has lower ¹⁴⁷Sm/¹⁴⁴Nd, and relatively high ¹⁴³Nd/¹⁴⁴Nd (0·5128). A well-defined isochron can be calculated if pyroxenites and gabbros are grouped on the basis of their sample locations. The three samples from the Kraubath Massif and their replicates plot on a model-1 solution isochron (Ludwig, 1999) giving 554 ± 37 Ma, with an initial ¹⁴³Nd/¹⁴⁴Nd_(i) of 0·51196 (=_Nd(i) +0·7). The reason why the two samples from Pernegg have somewhat higher ¹⁴³Nd/¹⁴⁴Nd is unclear; however, the REE pattern of sample PE29 is unusual and its Sm–Nd ratio may have been considerably disturbed after crystallization. The model-1 isochron age for Kraubath pyroxenites is significantly different from the ‘apparent’ age and initial isotopic composition of the host peridotites.

	DISCUSSION

TOP ABSTRACT INTRODUCTION GEOLOGICAL BACKGROUND ANALYTICAL METHODS PETROGRAPHY MINERAL CHEMISTRY WHOLE-ROCK CHEMICAL COMPOSITION DISCUSSION SUMMARY AND CONCLUSIONS REFERENCES

 
Origin and emplacement of the harzburgite–pyroxenite–gabbro suite
Evaluation of whole-rock geochemical data of metaperidotites from the Speik Complex shows that most samples may be considered as residual material from an ancient melting event (Melcher et al., 2002). The MREE to HREE parts of chondrite-normalized REE curves (Fig. 9) can be modelled by large degrees (of the order of 40%) of near-fractional melting of PUM; alternatively, between 20 and 30% ‘second-stage’ melting of already slightly melt-depleted peridotite [using the composition of lherzolite PE31 as starting material; Table 1 and Melcher et al. (2002)] may be assumed. The LREE enrichment in harzburgites, which is a feature common to many depleted mantle peridotites in ophiolite settings (McDonough & Frey, 1989), is most probably related to metasomatic enrichment (‘cryptic metasomatism’) or to infiltration of an LREE-rich component (Melcher et al., 2002). Thus, REE patterns of depleted peridotites from the Speik Complex record both partial melting and melt impregnation. Trace element patterns normalized to PUM resemble those of residual mantle formed above subduction zones by multistage melting processes (Parkinson & Pearce, 1998).

Pyroxenites within ophiolites may have a variety of origins. They may represent cumulate material crystallized from fractionating melts in magma chambers above the petrological Moho; pyroxenite sheets of this type of hundreds of metres to a kilometre thick are found at the base of exposed island-arc sections (Müntener et al., 2001). Pyroxenites may form as high-pressure crystal segregations at the margins of magma conduits within residual mantle (Bodinier et al., 1987; Spray, 1989). They may also be regarded as products of melt–rock reactions or metasomatic reactions within residual mantle (Kelemen et al., 1992; Varfalvy et al., 1996, 1997), or as products of metamorphic reactions (Ibarguchi et al., 1999). Geological and geochemical features of pyroxenites and peridotites in the Speik Complex closely resemble situations reported from other, ancient and modern ophiolite settings. In the Bay of Islands Complex, Newfoundland, Varfalvy et al. (1996, 1997) described coarse-grained pyroxenite dykes a few centimetres to 1 m thick, ranging in composition from orthopyroxenite to websterite and clinopyroxenite. They are regarded as conduits along which melts, generated by multistage crystal–melt segregation from different sources, migrated within the upper mantle. The extremely depleted nature of the melts is explained by multistage melting of ultra-depleted mantle sources at shallow depths. Pyroxenites acquired the signatures of variably depleted pyroxenites by reacting with large volumes of peridotites during the migration of these magmas through the mantle. The variety of pyroxenite compositions is explained by a combination of fractional crystallization of olivine + spinel, assimilation of dissolved clinopyroxene from wallrocks, and magma mixing, between a primary low-Ti magnesian tholeiitic component and low fractions of melt produced by partial remelting of surrounding peridotites.

Consequently, pyroxenites and gabbros of the Speik Complex may either represent crystallization products of melts trapped en route during their ascent from depleted mantle to the surface, or products of melt–rock reaction. Two approaches are used to estimate possible liquid compositions in equilibrium with pyroxenites and gabbros: (1) mineral chemical data; (2) whole-rock geochemical data.

From chemical compositions of orthopyroxene, equilibrium magma compositions may be calculated using experimentally determined mineral–melt equations of Barnes (1986a, 1986b). Provided that the highest Al and Cr concentrations in orthopyroxene measured in orthopyroxenites of the Speik Complex still represent primary magmatic compositions, melts in equilibrium with these orthopyroxenes would have low Al₂O₃ concentrations (<11 wt %) and between 0·2 and 0·3 wt % Cr₂O₃. The mineral chemical data display a trend of increasing Al and Cr in orthopyroxene with decreasing mg-number (Fig. 5a and b). This either suggests that melts evolved to more Al-rich compositions with time, or reflects the degree of reaction with the wall rocks.

Chromites from harzburgite, dunite and associated chromitite in the Kraubath Massif are Cr-rich (cr-numbers >74) and Mg-poor (mg-numbers <45) (Fig. 7). During metamorphism, chromite grains developed ferrite–chromite and magnetite rims, which have high cr-numbers and very low mg-numbers (Thalhammer et al., 1990). The cores occasionally preserve a primary(?) low-Al, low-Mg signature. The most Al-rich spinel compositions (about 10 wt % Al₂O₃) would correspond to about 10·3 wt % Al₂O₃ in the melt, using the equation relating spinel and liquid composition of Maurel & Maurel (1982); this is in accordance with the Al₂O₃ concentration derived from the composition of orthopyroxene.

As a result of the strong metamorphic overprint, whole-rock chemical data instead of mineral data may be more useful to calculate the composition of liquids in equilibrium with pyroxenites and gabbros, and the degree of interaction between possible liquids and residual mantle. Melts oversaturated in silica, and having mg-numbers, Ni and Cr concentrations higher than primitive mid-ocean ridge basalt (MORB), can be inferred taking into account published mineral/melt partition coefficients (D values) for these elements (Roeder & Emslie, 1970; Budahn, 1986). Both boninites and high-Mg andesites would meet these criteria. Trace element diagrams of harzburgites, pyroxenites (see Melcher et al., 2002) and gabbros normalized to Bulk Silicate Earth (BSE) reveal negative Zr, Hf, Nb and Th anomalies. Such features most probably are related to the source and not to alteration; they resemble melts derived from highly depleted mantle in a supra-subduction zone setting. In residual mantle peridotites high La/Nd ratios are believed to reflect the amount of the slab-derived supra-subduction component added to the depleted mantle wedge (Pearce & Parkinson, 1993).

Mass-balance calculations allow determination of liquid compositions that were in equilibrium with pyroxenite and gabbro cumulates (e.g. Bédard, 1994). This method requires an accurate whole-rock trace element analysis, knowledge of modal abundances and a consistent set of D values [e.g. compilation of Bédard (1999)]. A major source of error in such calculations is the poor quality of the D values for orthopyroxene. Absence or presence of a trapped melt fraction (TMF) in the rocks is of particular importance. As the TMF is usually expelled during compaction and crystallization, it is difficult to quantify. In the following calculations, TMF is treated as an additional phase with D = 1. Trace element patterns shown in Fig. 15 are derived from mass-balance equations of the general form (Bédard, 1994)

where the concentration C of element i in a mineral or phase (A, B, C, ...) is calculated based on the relation of the concentration of i in bulk rock, the modal abundances of (A, B, C,...), and the D values of i for (A, B, C,...). Trapped melt tm is treated as an additional phase.

View larger version (37K): [in this window] [in a new window]   Fig. 15. Chondrite-normalized incompatible element abundance patterns showing calculated magmatic liquids in equilibrium with pyroxenite and metagabbro of the Speik Complex. (a) Liquids in equilibrium with orthopyroxenite (shaded grey) and metaclinopyroxenite (striped field) assuming 1–3% trapped melt fraction. Curves representing low- and intermediate-Ti boninites are averages from the Betts Cove ophiolite (Bédard, 1999). (b) Liquids in equilibrium with metagabbros assuming 30% trapped melt fraction (TMF). Modal compositions used for the calculations (sample number, % cpx, % ol, % plg, % spl, % TMF): KR25, 25·9, 15·4, 25·9, 2·8, 30; PE22, 12·6, 16·8, 38·5, 2·1, 30; PE24, 35·0, 11·9, 21·0, 2·1, 30.

 
To calculate the composition of liquids in equilibrium with orthopyroxenite, modal abundances of orthopyroxene (94–97%), olivine (0–3%), spinel (0–2%) and TMF (1–3%) have been estimated. Calculated liquids have very low concentrations of most incompatible elements (between 0·6 and 4 times BSE) with the exception of the most incompatible elements, which are significantly enriched in the melt (Fig. 15a).

Modal abundances of the original phases are impossible to estimate for the metaclinopyroxenites of the Speik Complex, because of the near-complete metamorphism to amphibolite facies. Therefore, a CIPW-normative composition was calculated based on the major element data, being aware of the fact that (1) element mobility may severely change the concentrations of mobile elements and (2) calculation of a CIPW-normative composition may differ from the magmatic modal composition. As with the orthopyroxenites, calculated liquids in equilibrium with clinopyroxenite/websterite (Fig. 15a), assuming an arbitrary value of 1% TMF, overlap the incompatible element patterns of average boninites for many elements, except for the negative anomalies of Zr and Hf. Again, the most incompatible elements are strongly enriched in the calculated melt.

In the metagabbros, calculations based on major element data using the CIPW-norms indicate a variation in the proportions of original plagioclase (which is not preserved in the rocks; however, the positive Eu anomalies indicate significant plagioclase fractionation), clinopyroxene, and olivine. In Fig. 15b, liquids have been calculated based on an arbitrary assumption of 30% TMF in the rock. The incompatible element pattern of the liquid in equilibrium with sample PE22 is similar to low-Ti boninite (Fig. 15b), with the exception of the mobile, highly incompatible elements. The remaining samples are in equilibrium with melts having higher concentrations of incompatible elements than boninites, but similar positive Ti, Sr, and negative Nb anomalies (Fig. 15b). In general, element concentrations in calculated liquid profiles tend to have lower absolute values with increasing fraction of trapped liquid. Any absolute value for concentrations of the more compatible elements can be modelled by simply varying TMF; the general curve shapes remain constant.

Calculated liquid compositions from orthopyroxenites and clinopyroxenites correspond to high-Mg melts. Considerably higher concentrations of the more mobile, highly incompatible elements may be attributed either to mobility during metamorphism, or to addition via a ‘subduction zone component’ (Bédard, 1999).

PGE and Re–Os isotope characteristics of the mantle and percolating melts
Siderophile elements, including the PGE, and the Re–Os system are powerful tools to detect mantle petrogenetic processes and crustal contributions in melts, and to model melt–rock interaction processes (e.g. Rehkämper et al., 1999; Hart et al., 2002; Walker et al., 2002; Woodland et al., 2002). In general, flat PGE patterns are expected in fertile upper mantle and PPGE-depleted patterns are found in melt-depleted mantle (Rehkämper et al., 1999). Some PGE, including Os, may be mobile as a result of fluid transfer within the mantle wedge (Brandon et al., 1996).

The flat, unfractionated PGE patterns and Os isotope compositions of most peridotites (Fig. 10a) in the Speik Complex fall within the range of abyssal peridotites (Snow & Reisberg, 1995), depleted MORB mantle (DMM) and PUM (Meisel et al., 2001c). Partial melting had not completely dissolved residual mantle sulphides, and thus a near-chondritic PGE distribution had been retained. PPGE depletion in some samples may point to high degrees of partial melting with all sulphides consumed by the melt; in such cases, mantle residues become depleted, and melts become enriched in PPGE (Büchl et al., 2002). Most Speik Complex chromitites are characterized by IPGE > PPGE, Re/Os <1 (Fig. 11), and chondritic ¹⁸⁷Os/¹⁸⁸Os ratios (Fig. 12) with Os(T) close to or slightly above zero (–0·8 to +3); they fall within the range established for ophiolitic chromitites (Walker et al., 2002). The chromite data are in general agreement with recent in situ laser-ablation work on platinum-group minerals (PGM) included in such chromitites (laurite, ruarsite), which give a large range of slightly suprachondritic to subchondritic ¹⁸⁷Os/¹⁸⁸Os values (0·1158–0·1244; Malitch et al., 2003b; Fig. 12b). The most radiogenic ¹⁸⁷Os/¹⁸⁸Os_(i) in peridotite-hosted chromitite from Kraubath is 0·127, which is slightly above the chondritic value for the estimated time of formation (for T = 550 Ma). The large range of Os(T) in chromitites and peridotites (–4 to +2) is reminiscent of situations reported from supra-subduction zone mantle elsewhere (Büchl et al., 2002). The least radiogenic ¹⁸⁷Os/¹⁸⁸Os_(i) in harzburgite gives a model age of 1257 Ma; this is also the sample with the lowest concentrations of S and Al, and thus the most melt-depleted sample analysed. In the peridotites, ¹⁸⁷Os/¹⁸⁸Os is weakly correlated with S and Al.

Despite the overall similarity in trace element signatures between orthopyroxenites and their harzburgite hosts (Fig. 9), mantle-normalized PGE patterns are markedly different with higher Pd_N/Ir_N and Pt_N/Ir_N than in most metaperidotites (Fig. 10c), and also higher Re/Os. Some peridotites (Fig. 10b), however, have strongly fractionated PGE patterns with PPGE > IPGE, Re/Os > 1, and radiogenic, suprachondritic ¹⁸⁷Os/¹⁸⁸Os ratios with high Os(T) (Table 6). Osmium concentrations are rather variable in orthopyroxenites (Meisel et al. 2001b): in small bodies that are Os poor, such as MG45h, the Os isotopic composition is dominated by the host peridotites, which have higher Os abundances. Suprachondritic Re–Os isotopes in orthopyroxenites, clinopyroxenites, some peridotites and few chromitites, coupled with PGE fractionation, indicate that the PGE and Re budget in these rocks is no longer controlled by residual mantle, but by a melt having a radiogenic Os isotope composition. Such melts probably record a subduction zone component; suprachondritic Re/Os, Pt/Os and Pd/Os is typical of mantle melts (Büchl et al., 2002). Melts interact with their host rocks because they are not in equilibrium with harzburgite; depending on their size, pyroxenite pods and veinlets record smaller or larger contributions of a radiogenic component.

Fractionated PGE patterns similar to those from the pyroxenites have been reported from pyroxenite cumulates of the Thetford Mines ophiolite, Quebec (Tanguay et al., 1990). Similar PGE patterns have been reported from chromitite hosted by orthopyroxenite in the Xumai ultramafic massif, China (Mei et al., 1988). Platiniferous chromitite from the cumulate sequence at l'Estrie-Beauce, Appalachians, southern Quebec, contains up to 690 ppb Pd and 2000 ppb Pt (Pt_N/Ir_N <0·1 to 20, Pd_N/Ir_N <0·1 to 94) (Gauthier et al., 1990). Suprachondritic Os isotope compositions are known from a number of similar settings. In the Troodos ophiolite, orthopyroxenite veins are much more radiogenic (¹⁸⁷Os/¹⁸⁸Os_(i) = 0·1740) than the harzburgite host (0·129–0·131; Büchl et al., 2002). The veins, however, have much higher Os concentrations (7·8 ppm) and very low Re (0·03 ppm) compared with orthopyroxenites from Kraubath. A websterite vein at Troodos has a slightly suprachondritic initial ¹⁸⁷Os/¹⁸⁸Os (0·1332), similar to dunitic melt channels explained as reaction zones between harzburgite and through-flowing melts; such values are close to the composition of island-arc lavas (Hart et al., 2002; Woodland et al., 2002). A suprachondritic initial ¹⁸⁷Os/¹⁸⁸Os (0·17851, Os = +40) is also reported from an orthopyroxenite xenolith sampled from the mantle beneath Lihir island, Papua New Guinea (McInnes et al., 1999); harzburgites yield near-chondritic values (Os = –4·3 to +5·0). The radiogenic value for the orthopyroxenite requires a slab-derived Os component; McInnes et al. (1999) assumed a common origin for the Os in orthopyroxenite, clinopyroxenite, alkalic arc lavas and volcanic-hosted gold ores. O–Os mixing models indicate that mixing 9% of a subduction component (95% oceanic crust, 5% sediments) with depleted mantle will create radiogenic Os isotope signatures as observed on Lihir island.

There is some evidence for magmatic sulphur saturation in orthopyroxenites from the Kraubath Massif, e.g. high background levels of sulphur (up to 2700 ppm) and the presence of interstitial sulphides. In a reduced magmatic system, excess sulphide will be present within an immiscible sulphide melt, and Re and the PGE will preferentially enter the sulphide melt. The mechanism for local sulphur saturation in orthopyroxenites is open to speculation.

Protolith age and metamorphic evolution of the ultramafic–mafic rock assemblage
In the Speik Complex, ¹⁴³Nd/¹⁴⁴Nd covaries with ¹⁴⁷Sm/¹⁴⁴Nd in orthopyroxenites, metaclinopyroxenites and metagabbros. The samples from the Kraubath Massif define an isochron of 554 ± 37 Ma with _Nd = 0·7 (Fig. 14a) that is interpreted as a crystallization age of melts intruded into depleted mantle. In contrast, a whole-rock errorchron of the host peridotites gives 745 ± 44 Ma with an initial value of _Nd(i) +6 (Fig. 14b). An apparent Re–Os age of 550 ± 17 Ma and ¹⁸⁷Os/¹⁸⁸Os_(i) = 0·178 ± 0·003 (Fig. 12) from orthopyroxenite whole-rock samples overlaps with the Sm–Nd age. Together with _Nd lower than in harzburgites, the suprachondritic initial Os composition points to influence of a crustal component in the formation of the melts crystallizing the pyroxenites. T_DM model ages calculated for chromitites range from young, Mesozoic, ages to Proterozoic (Table 6); however, many type 2 chromitites from the Kraubath Massif have model ages between 500 and 600 Ma.

The 550 Ma age derived from two independent isotopic methods is within the uncertainty of an Rb–Sr errorchron of 518 ± 50 Ma (initial ⁸⁷Sr/⁸⁶Sr = 0·7044) from a plagioclase-rich orthogneiss considered to be a member of a bimodal volcanic suite in the Core Complex outcropping in the Gleinalm Dome (Fig. 1b) (Frank et al., 1976). Cambrian magmatic protolith ages have been reported from localities throughout the Alps (Thöni, 1999), and have been explained by the existence of several Cambrian ocean and back-arc basins that developed between different microcontinents along their continuous subduction underneath the northern Gondwana margin (Miller & Thöni, 1997; von Quadt et al., 1997; von Raumer, 1998; Schaltegger & Gebauer, 1999; Ladenhauf et al., 2001; Schaltegger et al., 2002; von Raumer et al., 2002; Eichhorn et al., 2003; Schulz et al., 2003). Early Palaeozoic ages (520–490 Ma) have recently become known from zircons within orthogneiss boulders of transgressive meta-conglomerate from the Kaintaleck metamorphic complex, which is part of the Upper Austroalpine basement nappes that tectonically overlie the Middle Austroalpine Speik and Core Complexes (Neubauer & Frisch, 1993; Neubauer et al., 2002).

Ultramafic rocks within the Speik Complex are variably serpentinized, ranging from 100% (Hochgrössen) to <20% (Kraubath). A number of metamorphic index minerals (from lower to higher temperatures: forsterite, diopside, tremolite, anthophyllite) may be used to constrain metamorphic conditions. The polymetamorphic history of the Austroalpine basement in this part of the Eastern Alps poses a major problem to correct interpretation of the mineral assemblages. A petrogenetic grid (Fig. 16) summarizes the P–T data derived for the Pre-Alpine and Alpine evolution of the Speik Complex.

View larger version (40K): [in this window] [in a new window]   Fig. 16. P–T diagram for metamorphic assemblages in ultramafic and mafic rocks of the Speik Complex. Reaction lines in the system SiO2–MgO–CaO–H2O are from Spear (1993). Dashed lines are calculated from chemical compositions of At + Ol + Tc + Srp + Tr using the program THERMOCALC 3.21 for different water activities (XW = 0·5, 0·75 and 1). Fields characterize metamorphic conditions of M1 (>400 Ma; Faryad et al., 2002a), M2 (330 Ma) and M3 (90–110 Ma) in different lithological units (Am, amphibolite; Ec, eclogite; G, gabbro; Gn, gneiss; P, pyroxenite; Ph, phyllite) calculated from silicate equilibria. Eoalpine (M3) conditions are different for the studied massifs (KR, Kraubath and Pernegg; HG, Hochgrössen); MC, Micaschist–Marble Complex (Faryad & Hoinkes, 2002). Bold black lines give ranges obtained from garnet I–hornblende and garnet I–clinopyroxene Fe–Mg exchange geothermometry in metagabbro from Kraubath (KR25).

 
In the Hochgrössen Massif, early or pre-Variscan eclogite-facies conditions (M₁) are documented in metabasites (700°C, 18–22 kbar; Faryad et al., 2002a). ⁴⁰Ar/³⁹Ar dating of edenitic hornblende in textural equilibrium with omphacite in eclogite from the Hochgrössen Massif gave a plateau age of 397·3 ± 7·8 Ma (Faryad et al., 2002a), which is a minimum age for M₁ metamorphism in the area. Handler et al. (1999) reported similar, mid-Palaeozoic (430–405 Ma) ⁴⁰Ar/³⁹Ar hornblende cooling ages from the Upper Austroalpine Kaintaleck metamorphic complex.

Neubauer (1988) has postulated thrusting of the Speik Complex over the Core Complex during Variscan medium-grade metamorphism (M₂). Pre-Alpine (Late Carboniferous, about 330 Ma, M_2a) metamorphic conditions have been estimated to be 540 ± 15°C and 6·6 ± 0·8 kbar in the Rappold unit of the Micaschist–Marble Complex (Faryad & Hoinkes, 2002), which is a tectonically higher nappe in the Austroalpine basement complex, exposed a few kilometres to the NW and west of the Gleinalm dome (Fig. 1a). A Permian (246 Ma; Schuster & Thöni, 1996) low-pressure (<5 kbar), medium-temperature metamorphic–magmatic event (M_2b) documented in the Wölz Unit of the Micaschist–Marble Complex can not yet be verified in the Speik Complex.

Most textural features of the pre-Alpine events have probably been obliterated during Eo-Alpine (M₃) metamorphism (100 Ma). In Austroalpine units, M₃ is characterized by low geothermal gradients of 10–16°C/km, and rocks are now exposed within south-dipping thrust units. Deeper portions in the Koralpe Complex record P–T conditions of >18 kbar and 700°C (Hoinkes et al., 1999; Faryad & Hoinkes, 2002), whereas 600–650°C and 10–11 kbar are reached in the Micaschist–Marble Complex, and 510 ± 10°C and 8 kbar have been calculated for Permian quartzites near the top of the Hochgrössen Massif. ⁴⁰Ar/³⁹Ar dating of amphibole from amphibolite (Core Complex) and metagabbro (Speik Complex) from the Pernegg, Kraubath and Hochgrössen Massifs yielded total gas ages ranging from 277 Ma to 80 Ma (Puhl, 2000). Exclusively Eo-Alpine ages (95–107 Ma) have been reported only for hornblende from metabasites in the southern part of the Core Complex (Neubauer et al., 1995). The pressure-dominated metamorphism probably resulted from the eastward motion of the Apulian plate, causing closure of the Meliata–Hallstatt ocean and the opening of the Penninic ocean in the Cretaceous.

High-pressure conditions (M₁) are neither preserved in the ultramafic rocks of the Speik Complex nor in metagabbros of the Kraubath and Pernegg Massifs. Instead, a two-stage P–T evolution can be inferred from mineral chemical data. Antigorite serpentinites at Hochgrössen carry metamorphic forsterite (Fo > 96) and rare diopside, but lack tremolite and talc, indicating conditions of the upper greenschist facies only (425–525°C at 5 kbar; Spear, 1993). In the Kraubath and Pernegg Massifs, metaperidotites carry tremolite, chlorite and talc, but lack metamorphic orthopyroxene; this indicates a temperature window of 550–700°C.

In orthopyroxenites and in some lherzolites, anthophyllite is observed rimming tremolite and replacing orthopyroxene (Figs 3d and 4b). Formation of anthophyllite from olivine and talc takes place at temperatures above 650°C and pressures <7 kbar (Evans & Trommsdorff, 1970). Reaction temperatures have been calculated for the assemblage At + Ol + Tc + Srp + Tr using the program THERMOCALC 3.21 (Powell & Holland, 1988) and are plotted for a pressure range of 2–10 kbar and various activities of H₂O; they average to 600°C (Fig. 16) (Puhl, 2000). For orthopyroxenite carrying metamorphic olivine and spinel (Fig. 3c and d), temperatures ranging from 640–660°C for Kraubath to 670°C for the Pernegg Massif have been calculated using the olivine–orthopyroxene–spinel geothermometer–oxygen fugacity barometer (Ballhaus et al., 1991) for an estimated maximum pressure of 10 kbar. Oxygen fugacities calculated using the Ballhaus et al. (1991) calibration are between 1 and 2 log units above the fayalite–magnetite–quartz (FMQ) buffer curve.

Mineral assemblages in most of the studied metagabbros are unsuitable to derive P–T conditions, because of the lack of sensitive minerals. In a sample from the Kraubath Massif (KR25), garnet I (pyrope-rich) originally may have equilibrated with relict clinopyroxene, although no contacts are preserved. The garnet–clinopyroxene Fe–Mg exchange geothermometer [calibrations of Ellis & Green (1979) and Powell (1985)] yields temperatures between 520 and 670°C for this assemblage. Garnet I and amphibole I give between 600 and 767°C [calibrations of Graham & Powell (1984) and Perchuk et al. (1985)]. Calculations using cpx, grt I and amph I and the Powell & Holland (1988) dataset result in an average temperature of 800 ± 250°C at 3–10 kbar. The large uncertainty probably indicates that equilibrium between garnet cores, clinopyroxene relics and amphibole cores is no longer achieved. Texturally, (grossularite-rich) garnet II coexists with Al-rich amphibole II. An intersection of five reactions between garnet II, amphibole II and chlorite was found at 531 ± 22°C and 7·6 ± 3·0 kbar using the THERMOCALC 3.21 software.

These data are in general accordance with estimates for Variscan conditions (M₂) of 713 ± 7°C at >7 kbar, and Eo-Alpine metamorphic conditions (M₃) of 550°C and 10·2 kbar recorded in metabasites of the Kraubath Massif (Faryad et al., 2002b). Hornblende–plagioclase assemblages in amphibolite of the Core Complex at Kraubath and Pernegg gave temperatures of 550–650°C and pressures of 5–8 kbar (Puhl, 2000).

	SUMMARY AND CONCLUSIONS

TOP ABSTRACT INTRODUCTION GEOLOGICAL BACKGROUND ANALYTICAL METHODS PETROGRAPHY MINERAL CHEMISTRY WHOLE-ROCK CHEMICAL COMPOSITION DISCUSSION SUMMARY AND CONCLUSIONS REFERENCES

 
Field relations, as well as whole-rock and mineral chemical data from mantle peridotites, pyroxenites and associated metagabbros of the Austroalpine Speik Complex show many similarities to ophiolite complexes. However, polyphase metamorphism and tectonism complicate reconstruction of the original relations between different ultramafic–mafic massifs in the Speik Complex, and between different lithological units within single massifs. Although it is possible that the scattered outcrops of amphibole-rich basic rocks in the Speik Complex represent mafic cumulates precipitated in magma chambers, a number of arguments support the hypothesis of crosscutting ultramafic–mafic dykes in the mantle transition zone of a supra-subduction zone complex.

Geochemical modelling indicates that both ortho- and clinopyroxenites are genetically related. They may have formed from melts of similar composition, which first fractionated opx + spinel, and later cpx and plagioclase. Metaclinopyroxenites are spatially closely related to metagabbros, which may represent cumulates after fractional crystallization of olivine, clinopyroxene and plagioclase, with variable proportions of trapped melt. Trace element modelling has shown that the pyroxenites can be explained by fractional crystallization from high-Mg, siliceous melts that have been contaminated with a crustal component (radiogenic Os, unradiogenic Nd) and extensive reactions with peridotite wall rocks. The melts are considered as re-melts of refractory harzburgitic mantle, because their trace element signature—except for the PGE—is similar to the composition of the host harzburgites. These features are consistent with supra-subduction zone settings, where hydrous metasomatism is widespread. PGE concentrations and Re–Os isotopes reflect the influence of variable degrees of melt–rock reaction in harzburgites.

The existence of subduction zones in oceanic basins at the northwestern margin of Gondwana has been postulated by several workers (Neubauer et al., 1989; von Raumer, 1998). The ultramafic–mafic association of the Speik Complex most probably developed in a fore-arc to island-arc transitional setting above subducted oceanic crust. During subduction, melts generated by hydrous melting of the mantle wedge, probably significantly modified by chemical contributions of the subducted plate, migrated upward into the transition zone between highly depleted mantle and crust, forming the observed suite of pyroxenites and gabbros. The migrating melts were highly magnesian and silica-rich, and chemical variability within the pyroxenites may be explained by a combination of fractional crystallization processes and reactions with harzburgitic wall rocks.

Thus, the Speik Complex in the Austroalpine basement east of the Tauern Window represents remnants of the oldest oceanic crust–mantle transition preserved in the Eastern Alps. The presence of orthopyroxenite bodies of variable size and shape intruding harzburgite distinguishes the Speik Complex from other Eastern Alpine ultramafic complexes (Melcher et al., 2002). From Sm–Nd and Re–Os isotope analyses of whole-rock samples, an age of formation of about 550 Ma can be inferred.

The following scenario best meets the available geochemical, petrological and isotopic constraints: (1) during a first stage, partial melting of relatively undepleted mantle at some time during the Proterozoic formed moderately depleted harzburgite, probably in a mid-ocean ridge system. (2) Second-stage melting in a supra-subduction zone setting during the Early Cambrian resulted in formation of highly depleted residual mantle (harzburgite–dunite) and high-(Si, Mg) mantle melts; from such melts, the intrusive orthopyroxenite/clinopyroxenite/gabbro suite formed under the influence of variable melt–rock reaction. (3) Subduction- or collision-related high-pressure metamorphism (M₁, older than 400 Ma) was probably related to the final closing of an oceanic basin that had formed in the interval from 520 to 485 Ma (Neubauer et al., 2002). (4) Variscan medium- to high-grade metamorphism (M₂) is related to thrusting and intrusion of granitoids. (5) Compressional Eo-Alpine metamorphism (M₃) mostly affected the southern portion of the Austroalpine nappe system, whereas greenschist-facies conditions at moderate to high pressures are preserved in the northern part.

	ACKNOWLEDGEMENTS

 
This study was financed by Austrian Science Fund (FWF) grants P12323-CHE (to F.M.) and P12322-CHE (to T.C.M.). The authors would like to acknowledge support by the following persons: Dr Fritz Koller, Dr Jürgen Puhl, Helmut Mühlhans and Johann Seiser. J. Raith critically commented on a first draft of the manuscript. T.C.M. is indebted to Jan Kramers (University of Bern) and Ronny Schönberg for the help with the NU instrument. Some of the measurements were made possible by the attribution of a CNRS visiting scientist post (chercheur associé) to T.C.M. An earlier version of the manuscript benefited from thoughtful reviews by J. Pearce and J. Snow. The constructive comments of the editor, M. Wilson, are gratefully acknowledged.

	FOOTNOTES

 

^* Corresponding author. Telephone: +49 511 643 2562. Fax: +49 511 643 3664. E-mail: F.Melcher{at}bgr.de

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