Journal of Petrology Advance Access originally published online on January 7, 2005
Journal of Petrology 2005 46(4):717-747; doi:10.1093/petrology/egh095
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Metamorphic Evolution of Iron-rich Mafic Cumulates from the Ötztal–Stubai Crystalline Complex, Eastern Alps, Austria
1 INSTITUT FÜR MINERALOGIE UND PETROGRAPHIE, UNIVERSITÄT INNSBRUCK, INNRAIN 52, A-6020 INNSBRUCK, AUSTRIA
2 RESEARCH SCHOOL OF EARTH SCIENCES, THE AUSTRALIAN NATIONAL UNIVERSITY, CANBERRA 0200, A.C.T., AUSTRALIA
3 INSTITUT FÜR GEOLOGISCHE WISSENSCHAFTEN, UNIVERSITÄT WIEN, GEOZENTRUM, ALTHANGASSE 14, A-1090 VIENNA, AUSTRIA
RECEIVED NOVEMBER 15, 2003; ACCEPTED NOVEMBER 5, 2004
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ABSTRACT |
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TOP
ABSTRACT
INTRODUCTIONOlivine-rich rocks containing olivine + orthopyroxene + spinel + Ca-amphibole ± clinopyroxene ± garnet are present in the central Ötztal–Stubai crystalline basement associated with eclogites of tholeiitic affinity. These rocks contain centimetre-sized garnet layers and lenses with garnet + clinopyroxene ± corundum. Protoliths of the olivine-rich rocks are thought to be olivine + orthopyroxene + spinel dominated cumulates generated from an already differentiated Fe-rich (
) tholeiitic magma that was emplaced into shallow continental crust. Protoliths of the garnet-rich rocks are interpreted as layers enriched in plagioclase and spinel intercalated in a cumulate rock sequence that is devoid of, or poor in, plagioclase. U–Pb sensitive high-resolution ion microprobe dating of zircons from a garnet layer indicates that emplacement of the cumulates took place no later than 517 ± 7 Myr ago. After their emplacement, the cumulates were subjected to progressive metamorphism, reaching eclogite-facies conditions around 800°C and >2 GPa during a Variscan metamorphic event between 350 and 360 Ma. Progressive high-P metamorphism induced breakdown of spinel to form garnet in the olivine-rich rocks and of plagioclase + spinel to form garnet + clinopyroxene ± corundum in the garnet layers. Retrogressive metamorphism at T 
650–680°C led to the formation of Ca-amphibole, chlorite and talc in the olivine-rich rocks. In the garnet layers, högbomite formed from corundum + spinel along with Al-rich spinel, Ca-amphibole, chlorite, aspidolite–preiswerkite, magnetite, ilmenite and apatite at the interface between olivine-rich rocks and garnet layers at P < 0·8 GPa. Progressive desiccation of retrogade fluids through crystallization of hydrous phases led to a local formation of saline brines in the garnet layers. The presence of these brines resulted in a late-stage formation of Fe- and K-rich Ca-amphibole and Sr-rich apatite, both characterized by extremely high Cl contents of up to 3·5 and 6·5 wt % Cl, respectively. KEY WORDS: cumulates; Variscan metamorphism; SHRIMP dating; högbomite; saline brines
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INTRODUCTION |
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Variscan high-P metamorphic rocks are present throughout the Austroalpine and Penninic basement of the Eastern Alps (e.g. Droop, 1983
; Zimmermann & Franz, 1988) but are particularly widespread within the Austroalpine basement west of the Tauern Window represented by the Ulten zone (Obata & Morten, 1987; Godard et al., 1996, and references therein; Hauzenberger et al., 1996; Tumiati et al., 2003) and the Ötztal–Stubai and Silvretta Crystalline Complexes (ÖSC and SC) (Miller, 1974; Maggetti & Galetti, 1988; Maggetti & Flisch, 1993; Miller & Thöni, 1995; Schweinehage & Massonne, 1999; Ladenhauf et al., 2001). Within the Austroalpine basement, evidence for high pressures is almost exclusively preserved in metabasic rocks of predominantly mid-ocean ridge basalt (MORB) affinity (Maggetti & Galetti, 1984; Maggetti et al., 1987; Miller & Thöni, 1995; Schweinehage & Massonne, 1999) that are present as lenses and elongate bodies of mostly garnet amphibolites, which may reach kilometre size and are enclosed in metapelitic country rocks. Eclogitic garnet + omphacite + quartz assemblages are locally preserved within these bodies as a result of incomplete re-equilibration during exhumation subsequent to the peak of Variscan metamorphism. In rare cases, relict gabbros and gabbroic textures still indicate the nature of the protoliths (Miller, 1974; Miller & Thöni, 1995). P–T conditions for the Variscan high-P metamorphism range from 1·5 ± 0·5 GPa and 600–850°C for the Ulten zone (Godard et al., 1996) to 2·5–2·9 GPa at 600–730°C in the ÖSC and SC (Maggetti & Galetti, 1988; Miller & Thöni, 1995; Schweinehage & Massonne, 1999; Ladenhauf et al., 2001). The timing of the high-P event was determined for metabasic eclogites of the ÖSC and SC by Miller & Thöni (1995) and Ladenhauf et al. (2001) using Sm–Nd mineral–whole-rock isochrons and sensitive high-resolution ion microprobe (SHRIMP) U–Pb dating of zircons. Those workers obtained consistent ages in the range 350–360 Ma, which are somewhat older than U–Pb zircon ages of 332 and 336 Ma and Sm–Nd garnet ages of 326–336 Ma obtained from garnet pyroxenites and lherzolites of the Ulten zone (Gebauer & Grünenfelder, 1978; Godard et al., 1996).
Rocks other than metabasites only rarely preserve evidence for a Variscan high-P event. Examples are garnet–spinel peridotites and their pelitic country rocks from the Ulten zone (Obata & Morten, 1987; Godard et al., 1996; Tumiati et al., 2003). In the ÖSC, non-eclogitic rocks that have preserved evidence for high-P metamorphism are olivine-dominated Fe-rich mafic cumulates associated with eclogites in the northern ÖSC. In this paper, we present results of a petrological, geochemical and geochronological study on the evolution of these cumulates. The aims of this study are: (1) to trace the effects of eclogite-facies P–T conditions and the subsequent exhumation on an Fe-rich ultramafic bulk composition; (2) to constrain the time of emplacement of the cumulates and their metamorphic overprint; (3) to use geochemical characteristics of the cumulates in combination with data from the metabasic rocks to constrain the geotectonic environment during cumulate emplacement.
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GEOLOGICAL BACKGROUND |
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The ÖSC is part of the Austroalpine nappe system west of the Tauern Window and was thrust on top of Penninic units during the late Cretaceous–early Tertiary (Fügenschuh et al., 2000
; Sölva et al., 2001) (Fig. 1). The dominant ÖSC rock types are metapelitic to metapsammitic gneisses and schists with intercalations of metacarbonates. Meta-igneous rocks are present throughout the entire ÖSC as (garnet-) amphibolites and granitic orthogneisses cropping out in concordant layers or lenses. Structures in the ÖSC are dominated by a strong east–west-trending schistosity and by a kilometre-scale folding around steep axes in its northern and southern parts (Sölva et al., 2001, and references therein).
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Petrological and geochronological data indicate at least three major periods of metamorphism or thermal activity in the ÖSC (Miller & Thöni, 1995
; Hoinkes et al., 1997, 1999; Klötzli-Chowanetz et al., 1997; Thöni, 1999): (1) a Cadomian to Caledonian high-grade event at 530–440 Ma that led to the formation of migmatites and the intrusion of granites and gabbros; (2) a dominant Variscan eclogite- to amphibolite-facies event at 390–330 Ma; (3) an Eo-Alpine event at 100–65 Ma. The degree of Eo-Alpine metamorphic overprint increases in the ÖSC from the NW to the SE and culminates in a zone of epidote-amphibolite- to eclogite-facies conditions at 550–600°C and 0·8 to >1·2 GPa. The Eo-Alpine overprint can also be traced by a continuous rejuvenation of mica ages from the NW to the SE, leading to a complete Eo-Alpine resetting in an area approximately coinciding with the zone of staurolite formation in pelitic lithologies. |
SAMPLE DESCRIPTION AND PETROGRAPHY |
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In the north–central ÖSC, high-P rocks are present in two parallel kilometre-sized elongated bodies consisting of garnet amphibolites, eclogites and olivine-rich rocks enclosed in pelitic gneisses, forming the northern and southern eclogite zones (NEZ, SEZ) (Miller, 1974
; Koziol & Oberhänsli, 1995). The investigated samples (Table 1) were collected from the SEZ at the locality Untere Pollesalm (Fig. 1) where large lenses of eclogites and olivine-rich rocks—in part retrogressed to diablastic garnet amphibolites and chlorite–talc–tremolite rocks—are present.
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Macroscopically, the olivine-rich rocks are fine grained and homogeneous with a grain size
5 mm. Elongate lenses or layers of garnet up to 2 cm in diameter may be present as intercalations within the olivine-rich rocks. All garnet layers show alteration to hydrous phases along their rims and along cracks within the garnet layers.
Microtextures of olivine-rich rocks
Olivine-bearing rocks show the assemblage olivine + orthopyroxene + spinel + Ca-amphibole ± garnet ± clinopyroxene ± chlorite ± talc ± apatite + accessories (Table 1) with olivine as the modally dominant phase (>80 vol. %) and orthopyroxene and Ca-amphibole accounting for most of the remaining modal mineralogy.
Olivine appears in four microstructural settings, three of which show distinct compositions: (1) olivine-I: large rounded grains (up to 5 mm) with deformation lamellae and inclusions of orthopyroxene, spinel, amphibole and chlorite; (2) olivine-II: small subhedral grains (<100 µm) surrounding olivine-I or filling cracks therein, producing porphyroclastic but well-equilibrated textures with triple points; (3) olivine-III: narrow olivine rings separating garnet coronae from spinel grains, the latter being present as inclusions in large olivine porphyroclasts; (4) olivine-IV: small irregular grains formed by retrogressive breakdown of orthopyroxene, coexisting with talc ± chlorite (Fig. 2a–c).
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Orthopyroxene forms porphyroclasts and small recrystallized grains, the latter commonly associated with recrystallized olivine (Fig. 2a). Large orthopyroxenes are usually zoned, with cores showing needle-like oxide exsolution and narrow discontinuous rims (<20–40 µm) free of exsolution. Spinel forms rounded to xenomorphic aggregates either as inclusions in olivine or interstitial between olivine and orthopyroxene. With very few exceptions, all spinel aggregates consist of green and brown grains. Clinopyroxene was detected by electron microprobe as tiny inclusions in Ca-amphibole only.
Ca-amphiboles appear as discontinuously zoned subhedral grains irregularly scattered in the olivine + orthopyroxene matrix (Fig. 2a), showing dark green cores with sagenite-like oxide exsolution and up to three optically and chemically distinct growth zones.
Garnet is present as coronae around spinel (Fig. 2b) and as small irregular inclusions in Ca-amphibole. In samples 386 and 402, garnet is modally significant and present as spinel coronae and also as irregular grains interstitial between olivine and orthopyroxene. In both cases, garnets show extremely fine-grained symplectite rims with Ca-amphibole–spinel (sample 402) or Ca-amphibole–gedrite (sample 386).
All samples are retrogressively altered as indicated by varying amounts of chlorite, talc, Ca-amphibole and trace amounts of carbonates growing at the expense of olivine and orthopyroxene across the fabric or forming overgrowths on pre-existing grains (Fig. 2d). As an additional secondary phase, Cl- and Sr-rich apatite may be present.
Microtextures of garnet layers
Garnet layers contain abundant inclusions comprising clinopyroxene, spinel, corundum, högbomite, ilmenite, rutile and OH-apatite (Table 1). Phases exclusively found as inclusions in garnet are corundum, högbomite and OH-apatite. Corundum is present as rounded to irregular grains in part containing garnet and/or högbomite inclusions and with strong brownish-yellow pleochroism when rich in Cr. Högbomite, when present, is invariably associated with corundum (Fig. 2e). Clinopyroxene forms irregular grains in garnet but has almost always partly reacted to Ca-amphibole. In some cases garnet and clinopyroxene show myrmekite-like intergrowths. OH-apatite is present as brownish grains with abundant fluid inclusions within garnet as opposed to secondary Cl- and Sr-rich apatite that is present as colourless euhedral grains associated with Ca-amphibole, chlorite or magnetite. In sample JK3/00, numerous rounded grains of zircon up to 200 µm x 50 µm in size are present as inclusions in garnet.
All garnet layers show retrogressive alteration to a suite of secondary phases forming rims around or irregular patches and veinlets within the garnet layers (Fig. 2f). The most abundant secondary phases are Ca-amphibole and chlorite with accessory phlogopite, aspidolite, margarite, gedrite, green spinel, Cl- and Sr-rich apatite, ilmenite, magnetite, clinozoisite, allanite, plagioclase and zircon. Zircon was found in two samples (CM36/00 and CM37/00) as isolated grains of 5–10 µm associated with amphibole within the secondary rims around garnet layers.
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ANALYTICAL PROCEDURES |
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The composition of minerals was determined with an ARL-SEMQ electron microprobe using both wavelength- and energy-dispersive analytical modes with 15 kV and 20 nA sample current on a benitoite standard. Whole-rock major and trace element analyses were performed by the Service d'Analyses des Roches et des Minéraux du CNRS (SARM) using inductively coupled plasma atomic emission spectrometry (ICP-AES) and inductively coupled plasma mass spectrometry (ICP-MS). In situ trace element analysis of individual mineral phases was carried out using laser-ablation (LA)-ICP-MS at the Consiglio Nationale delle Ricerche Istituto di Geoscienze e Georisorse (CNR-IGG), Sezione di Pavia, Italy (Tiepolo et al., 2003
).
For U–Pb isotopic analysis, the thin section JK3/00 was cut and combined into epoxy mounts together with Research School of Earth Sciences (RSES) reference zircons FC1 and SL13. Spots suitable for analysis were selected on the basis of cathodoluminescence (CL) analysis. All U–Pb isotopic analyses were carried out using SHRIMP II, each analysis consisting of six scans through the relevant mass range. Data reduction was performed using the SQUID Excel Macro by Ludwig (2000). U/Pb ratios have been normalized relative to a value of 0·1859 for the 206*Pb/238U ratio of FC1 reference zircons, equivalent to an age of 1099 Ma (Paces & Miller, 1993). Uncertainties given for individual analyses (ratios and ages) are at the 1
level; however, uncertainties in the calculated weighted mean ages are reported with 95% confidence limits. Concordia plots were made and weighted mean age calculations carried out using Isoplot/Ex (Ludwig, 1999).
Bulk-rock Nd, Sm, Rb and Sr concentrations were determined by isotope dilution using mixed 147Sm–150Nd and 87Rb–84Sr spikes, respectively. Nd and Sr isotopic compositions were measured on a Finnigan MAT 262 multicollector mass spectrometer using a Re double filament. Within-run isotope fractionation was corrected for 146Nd/144Nd = 0·7219 (Nd) and 86Sr/88Sr = 0·1194 (Sr). 143Nd/144Nd and the 86Sr/88Sr ratios for the La Jolla (Nd) and the NBS 987 (Sr) international standards were 0·511860 ± 13 and 0·71023 ± 2, respectively. The following model parameters were used for the calculation of depleted mantle (DM) ages: 147Sm/144Nd = 0·222, 143Nd/144Nd = 0·513114 (Michard et al. 1985). A linear evolution of the Nd isotope composition of the DM is assumed throughout geological time; 
Nd values are calculated relative to a chondritic uniform reservoir (CHUR).
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BULK-ROCK COMPOSITION |
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The olivine-rich rocks are characterized by high MgO contents (29–38 wt %) combined with low SiO2 (38–43 wt %), Al2O3 (2·5–6·1 wt %) and CaO (1·2–4·3 wt %), and variable FeO contents (13–18 wt %), yielding bulk Mg-numbers [100 Mg/(Mg + Fetotal)] of 78–82 (Table 2). In contrast, sample CM23/00, which represents a garnet layer, is distinctly lower in MgO and higher in Al2O3 and CaO. On the Al2O3–TiO2 diagram of Pearce (1983)
, all samples plot in the cumulate field whereby olivine-rich rocks and garnet layer CM23/00 plot close to olivine-bearing and plagioclase–clinopyroxene gabbroic rocks of the NEZ, respectively (Fig. 3). (FeO/MgO)ol/(FeOtot/MgO)whole-rock values for olivine-rich rocks are in the range 0·6–1·0, which is considerably lower than the value of 3·3 determined experimentally for the Fe–Mg partitioning between olivine and basaltic liquid (Roeder & Emslie, 1970; Ulmer, 1989). These low values clearly indicate olivine accumulation. Al2O3, CaO, TiO2 and Sc are inversely correlated with MgO. Ni (800–1441 ppm), Co (105–162 ppm) and Cr (1900–4920 ppm) are positively correlated with MgO, albeit with some scatter in the data (Fig. 4), indicating olivine and clinopyroxene fractionation.
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, gabbroic rocks from the NEZ (Miller & Thöni, 1995
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All analysed samples have chondrite-normalized rare earth element (REE) patterns enriched in light rare earth elements (LREE), with LaN/YbN = 2·4–16·5 and relatively unfractionated heavy REE (HREE). Most of the samples lack Eu anomalies, except for 386 and JK5/00, which have negative Eu anomalies with Eu/Eu* = 0·81 and 0·72, respectively, and garnet layer CM23/00, which shows a small positive Eu anomaly (Eu/Eu* = 1·15). By comparison, olivine-bearing metagabbros from the NEZ are depleted in LREE and characterized by marked positive Eu anomalies (Fig. 5a).
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The primitive-mantle normalized diagram for incompatible elements (Fig. 5b) shows: (1) variable large ion lithophile element (LILE) concentrations, which are not correlated with the REE; (2) distinct negative Sr anomalies not only in sample JK5/00 with a negative Eu anomaly, but also in sample CM23/00 which is characterized by a positive Eu anomaly; (3) minor negative Nb–Ta–Ti anomalies.
Two olivine-rich rocks and one garnet layer were analysed for Sr–Nd isotopes. Present-day Nd values are negative for all samples, corresponding to depleted mantle model ages in the range 620–1659 Ma (Table 3). When recalculated for 517 Ma, samples JK5/00 and CM23/00 plot close to the mantle array, i.e. within the quadrant for depleted reservoirs, and are similar to NEZ metagabbros and eclogites (Fig. 6). In contrast, sample 386 has an unusual initial isotopic composition with Nd(t) = –5·2, but with 87Sr/86Sr = 0·70289, thus plotting within the quadrant characterized by the enrichment parameters fSm < 0 and fRb < 0 (Fig. 6) (see McCulloch & Wasserburg, 1978).
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COMPOSITION OF MINERALS |
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Olivine
Olivine-I shows a relatively wide range in XMg values (0·79–0·87) most commonly between 0·80 and 0·84. Within individual samples, olivines are generally homogeneous in composition. In rare cases, they may show a slight zoning with decreasing XMg from core to rim. Regardless of XMg, Ni contents are low and range between 0·10 and 0·20 wt % NiO (Table 4). Although clearly different in terms of textures, olivine-II is indistinguishable from olivine-I by composition. Olivine-III shows consistently higher XMg compared with olivine-I, and olivine-IV is characterized by distinctly higher MnO and XMg compared with olivine-I in a given sample (Table 4).
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Olivine and coexisting orthopyroxene, chlorite and garnet show systematic variations in Fe/Mg as a function of bulk-rock Fe/Mg. For coexisting olivine–orthopyroxene (core analyses of matrix olivine and orthopyroxene or olivine and orthopyroxene inclusions),
and averaged ln KD [KD = (Fe/Mg)ol/(Fe/Mg)opx] values ranging from 0·05 to 0·20 are positively correlated with 
, which is consistent with results from experimental studies (von Seckendorff & O'Neill, 1993, and references therein) (Fig. 7a). For olivine–chlorite, a regression through data from six samples yields a KD [KD = (Fe/Mg)ol/(Fe/Mg)chl] of 2·4 (Fig. 7b), which is in good agreement with data from more Mg-rich rocks (e.g. Trommsdorff & Evans, 1969; Müntener et al., 2000).
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Orthopyroxene
Orthopyroxenes are essentially enstatite–ferrosilite solid solutions and, like olivine, show a considerable range in XMg (0·79–0·86). Cr2O3, TiO2 and Na2O contents are invariably low and rarely exceed 0·25 wt % (Table 5). Large matrix orthopyroxene porphyroblasts commonly show a discontinuous compositional zoning with Al- and Ca-rich cores and narrow rims that approach the composition of pure enstatite. Small recrystallized orthopyroxene grains associated with olivine-II share the compositional characteristics of these rims.
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Porphyroblast cores have Al2O3 contents in the range 1·7–2·7 wt % (0·07–0·11 Al a.p.f.u.) but may reach values as high as 3·4 wt %, whereas rims of discontinuously zoned porphyroblasts do not contain more than 0·5 wt % Al2O3 (0·02 Al a.p.f.u.). CaO contents of porphyroblast cores range between 0·2 and 0·4 wt %, and drop to values <0·05 wt % in the Al-poor rims.
Trace element concentrations of orthopyroxenes are uniformly low and mostly close to or below the detection limits with the exception of Cr, V and Co (see Table 13, below). REE patterns show a moderate relative enrichment in HREE consistent with data from the literature (e.g. Tribuzio et al., 1999). A low-Al/low-Ca orthopyroxene rim measured in sample 376 shows significantly lower concentrations in Cr, V, Ti, Sc, Zr and Hf compared with the core containing exsolution lamellae (Table 13). In spite of the fact that spots for LA-ICP-MS analyses were carefully selected to avoid any visible inclusions, an irregularity in count rates of the LILE masses in both samples 376 and 401b-3 might indicate the presence of undetected inclusions.
Clinopyroxene
All analysed clinopyroxenes are Ca–Mg rich (73–93 mol % diopside) with a maximum jadeite content of 13 mol %. All clinopyroxenes except those from 401b-1 are characterized by Na > Al(VI) and Na/Altot 
1, which indicates the presence of an acmite component and implies a very low Ca-Tschermak component. Most clinopyroxenes have minor Cr2O3 (0·1–0·9 wt %) and TiO2 contents (0·2 wt %) (Table 6). Fe/Mg ratios of coexisting clinopyroxene and garnet show systematic covariation (Fig. 7d), with high KD indicative of retrogressive equilibration (see below).
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Trace element patterns of clinopyroxene from a garnet layer (sample JK3/00) are characterized by a positive Sr anomaly and pronounced negative anomalies for Zr, Hf and Nb. REE patterns are moderately LREE enriched with LaN/YbN = 8–18. Two out of three analysed clinopyroxenes show positive Eu anomalies with Eu/Eu* = 1·25 and 1·31, respectively (Fig. 8a and b; Table 13).
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Garnet
Garnet appears in two textural varieties that also show significant differences in composition, indicating different formation mechanisms. Garnet forming coronae around spinel (e.g. JK2/00) or irregularly shaped interstitial grains between olivine and orthopyroxene (e.g. 386, 402) are dominantly pyrope–almandine solid solutions with a compositional range py43–57alm30–36gross3–14 and a negligible spessartine component. Major variations in garnet compositions are restricted to the Ca/Mg ratio without systematic compositional zoning within individual grains. With the exception of sample 402, all garnet analyses from olivine-rich samples have minor Cr2O3 contents (0·2–1·1 wt %) (Table 7). Layer-forming garnet is more calcic with a compositional range py23–47alm26–46 gross10–42 and with a higher within-sample compositional variation (Fig. 9; Table 7). The latter is due to an irregular patchy zoning involving Ca/Mg ratios ascribed to retrogressive overprint of the garnet layers. Spessartine contents are generally low (1–3 mol %) but may locally reach 20 mol % as a result of Mn enrichment caused by advanced amphibolitization. With the exception of CM23/00, Cr2O3 values are below 0·05 wt %. The REE spectrum of garnet from garnet layer JK3/00 is characterized by pronounced enrichment in middle REE (MREE), a positive Eu anomaly (Eu/Eu* = 1·41) and a uniformly negative MREE–HREE slope (Fig. 8a, Table 13).
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Spinel
Both olive–green and brown types of spinel may show considerable compositional within-sample variation but no overlap in composition (Fig. 10, Table 8). Green spinels are Al rich with a compositional range chromite0·05–0·33magnetite0·03–0·17spinel0·27–0·62hercynite0·20–0·32 with minor ZnO (0·4–1·4 wt %) and TiO2 (
0·4 wt %). Brown spinel is characterized by higher Fe and lower Al contents with a compositional range chromite0·05–0·36magnetite0·37–0·88spinel0·03–0·12hercynite0–0·09. Compared with green spinel, TiO2 contents are consistently higher (0·5–2·6 wt %). In addition, only brown spinels may show minor V contents (0·7 wt % V2O3).
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Spinels present as inclusions within garnet layers, or within the amphibole–chlorite mantles developed along the interfaces between garnet layers and olivine-rich rocks, appear in two distinct varieties: (1) Al-rich and Cr-poor (chromite0·0–0·11magnetite0·0–0·06spinel0·32–0·78hercynite0·20–0·56) grains with highly variable ZnO contents ranging from <0·3 to 15 wt %; (2) almost pure magnetites in part intergrown with ilmenite. Magnetite is restricted to strongly retrogressed samples and may contain up to 3·2 wt % V2O3 as the only significant minor component (Table 8).
Spinels present within symplectitic spinel–amphibole coronae (sample 402) are Al-rich spinel–hercynite solid solutions and can be distinguished from the other types of Al-rich spinels by their very low Cr contents (Table 8).
Amphibole
In the olivine-rich rocks, amphiboles are calcic and predominantly pargasitic in composition, with a narrow range in XMg (Fig. 11). Individual grains show discontinuous compositional zoning with relatively Ti-rich cores characterized by sagenite-like exsolution of ilmenite and K-rich rims. In strongly retrogressed samples (e.g. 399), amphiboles may show up to three discontinuous growth zones with decreasing Al, Ti and K from core to rim, whereby the outermost rims are FeMg-amphibole (analyses 6–9, Table 9).
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Compared with olivine-rich rocks, Ca-amphiboles in garnet layers along cracks or forming part of the amphibole + chlorite ± aspidolite mantles span a much wider compositional range from sadanagaite to (ferro)-pargasite to magnesio-hornblende and tremolite (Fig. 11; Table 9). Zoning patterns of Ca-amphiboles from garnet layers are irregular, but in some cases with Al-poor cores and Al-rich rims. In sample CM23/00, taramite was found in addition to Ca-amphibole associated with aspidolite as a product of retrograde metamorphism. Some Ca-amphiboles of the garnet layers are characterized by unusually high Cl contents with up to 3·45 wt % Cl or 48 mol % Cl-pargasite end-member. They appear as small euhedral crystals interstitial between, or more commonly as narrow discontinuous outermost rims on pargasitic amphiboles, and represent the latest amphibole generation. Cl is positively correlated with Fe and K and negatively correlated with Mg and the number of A-site vacancies (Fig. 12), whereby the most Cl-rich amphiboles reach 0·43 a.p.f.u. K and XMg values as low as 0·27. In sample CM36/00, Cl-rich Ca-amphiboles may also contain significant amounts of V2O3 (0·5–1·1 wt %).
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The trace element composition of low-Cl pargasitic Ca-amphiboles was analysed in two olivine-rich rocks and one garnet layer (Table 13). All analyses are characterized by positive anomalies for Ba and Sr, and LILE and high field strength elements (HFSE) in the range 0·2–20 times primitive mantle (Fig. 8c). With the exception of one analysis from JK3/00, positive Ba and Sr anomalies are combined with negative anomalies for the HFSE Nb, Ta, Zr and Hf. REE patterns are moderately LREE enriched with LaN/LuN in the range 1·95–3·71 (Fig. 8d). In the case of JK3/00, the maximum REE enrichment is shifted towards Pr, Nd and Sm. The HREE enrichment in 376 amphiboles (Fig. 8d) is thought to be inherited from precursor garnet. Six out of eight amphibole analyses show positive Eu anomalies withEu/Eu* = 1·05–1·73, the significance of which will be discussed below.
FeMg-amphiboles may appear as the outermost rims of multiply zoned Ca-amphiboles (e.g. sample 399) or as Na-rich gedrite (0·72–0·84 Na a.p.f.u.) in narrow alteration rims around garnet blasts (e.g. sample 386).
Mica and chlorite
In olivine-rich rocks, phlogopite is present as small euhedral laths interstitial between olivine and enstatite and commonly associated with chlorite. Its composition is variable with Na/(K + Na) in the range 0·03–0·25 and Al(VI) between 0·10 and 0·35 a.p.f.u. but with constant XMg (0·90–0·91). In chlorite–amphibole mantles around garnet layers, trioctahedral micas of the aspidolite–preiswerkite series are present, intergrown with phlogopite or as individual grains. They are characterized by Na/(K + Na) = 0·82–0·96 and Al(VI) in the range 0·39–0·59, thus showing substantial preiswerkite solid solution (Table 10). In some cases, patches of phlogopite are present within larger aspidolite grains indicating a metasomatic reaction phlogopite–aspidolite. Compared with phlogopite in the olivine-rich rocks, XMg for phlogopite–aspidolite is lower and more variable in the range 0·72–0·85.
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In olivine-rich rocks, chlorite shows little intra-sample chemical variation, with most analysed grains close to clinochlore composition (Table 10). However, in garnet layers clinochlore may show strong discontinuous zoning, with Fe-rich cores (XMg = 0·75–0·80) and Mg-rich rims (XMg= 0·89–0·91). The only minor element detected in chlorites from all samples is Cr; its concentration ranges from <0·05 to 1·2 wt % Cr2O3.
Corundum and högbomite
All analysed corundum grains contain Fe2O3 and/or Cr2O3 as minor components. Both Cr2O3 and Fe2O3 contents are variable between samples and also within individual samples and range from <0·05 to 3·7 wt % Cr2O3, and from 0·4 to 1·1 wt % Fe2O3, respectively (Table 11). In sample 401b-1, individual corundum grains may display a core-to-rim increase in Cr with a maximum difference of 1·1 wt % Cr2O3.
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Högbomites show distinct ranges in Ti, Al, Cr and XMg in the three samples investigated. The lowest Ti is found in CM25/00 with 0·17–0·21 atoms per eight oxygens, intermediate values of 0·21–0·36 a.p.f.u. in JK3/00 and the highest values of 0·28–0·42 are present in 401b, the last being amongst the highest reported in the literature (Zakrzewski, 1977
; Teale, 1980; Coolen, 1981; Mancktelow, 1981; Ackermand et al., 1983; Beukes et al., 1986; Gieré, 1986; Grew et al., 1987, 1990; Rammlmair et al., 1988; Petersen et al., 1989; Yalçin et al., 1993; Armbruster et al., 1998). Cr contents are negatively correlated with Al. The observed compositional range is from <0·05 to 6·1 wt % Cr2O3 (0·25 a.p.f.u.) in 401b, which is the highest Cr content reported in the literature. Cr contents of högbomite and coexisting corundum and spinel are well correlated (Fig. 13a and b). Whereas no literature data are available for Cr partitioning between högbomite and corundum, that between högbomite and spinel is in good agreement with data reported by Coolen (1981). XMg values of högbomites enclosed in garnet are in the range 0·53–0·70 and show a systematic variation with XMg of coexisting garnet (Fig. 13c) and spinel with 
. For högbomite–garnet and högbomite–spinel pairs, averaged 
(Fetot = FeO) values of 1·37 ± 0·19 and 0·90 ± 0·15 are obtained, the latter being in good agreement with data from the literature.
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Apatite
Two apatite generations can be distinguished based on textures and composition (Table 12), both containing significant amounts of Cl and Sr but no F. A texturally early generation is present as brown irregular grains included in garnet (samples 401b-1 and JK3/00), with 0·5–0·7 w% Cl [Cl/(Cl + OH) = 0·07–0·10] and 0·8–2·0 wt % SrO. A second generation of mostly euhedral colourless apatite associated with amphibole, chlorite and magnetite in both garnet layers and olivine-rich rocks is characterized by high Cl concentrations (2·7 and 6·5 wt %) corresponding to XCl of 0·38–0·98 along with SrO contents between 0·5 and 3·9 wt %. Individual grains are mostly homogeneous in composition, and only in sample JK3/00 could a patchy irregular zoning be observed ranging from 2·7 to 3·4 wt % Cl in a single grain.
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ELEMENT PARTITIONING AND GEOTHERMOBAROMETRY |
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Element partitioning was used to assess the attainment of equilibrium and where possible to constrain metamorphic P–T conditions. For example, Fe–Mg partitioning between olivine and orthopyroxene is not suitable for determination of equilibration temperatures because of the large variation in experimentally determined KD values (Fonarev, 1981
; Koch-Müller et al., 1992; von Seckendorff & O'Neill, 1993) and the small variation of KD with T at a given XFe. Nevertheless, the systematic variation in Fe/Mg observed in the present study serves to indicate equilibration between the originally most abundant phases at least for cations with diffusivities as fast as Fe 
Mg. Likewise, the systematic variation in Fe/Mg in coexisting olivine and chlorite indicates cooling that was slow enough to maintain equilibrium with respect to Fe–Mg exchange.
Constraints on metamorphic temperatures
Al in orthopyroxene
In this study, cores of the largest orthopyroxene grains available in the matrix were used along with orthopyroxene inclusions in olivine. The latter were analysed when present irrespective of grain size. To obtain an integrated composition for exsolved cores, a rastered beam was used for the analyses (see Obata, 1980; Müntener et al., 2000) and samples with coarse exsolution lamellae that would have yielded non-representative compositions were excluded. Using the formulation of Witt-Eickschen & Seck (1991), temperatures between 770 and 800°C are obtained with little intra-sample variation (Table 14). No correlation between temperature and bulk composition or XMg in olivine could be observed. It has to be noted that Al in orthopyroxene was used to constrain metamorphic temperatures under the assumption that a homogeneous Mg–Al-rich spinel was present under peak metamorphic conditions, which exsolved and recrystallized upon slow cooling. Textural evidence for coexisting olivine–orthopyroxene–spinel is ubiquitous in the form of coexisting (composite) spinel–orthopyroxene inclusions in olivine.
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Ca in orthopyroxene
The coexistence of clinopyroxene with orthopyroxene under peak metamorphic conditions cannot be proven unambiguously for the rocks of this study because of a lack of clinopyroxene inclusions in orthopyroxene cores and the absence of matrix clinopyroxene from olivine-rich rocks. In sample 376, however, clinopyroxene inclusions are present in Al-poor rims of orthopyroxenes. If a P range of 0·5–1·5 GPa is assumed for post-peak formation of Al-poor rims, then temperatures of 645–682°C are obtained based on averaged Ca a.p.f.u. for orthopyroxene rims and the formulation of Brey & Köhler (1990)
.
Fe–Mg exchange thermometers
For samples 402 and 386, averaged compositions of cores of garnet and orthopyroxene were used to calculate temperatures based on the calibration of Harley (1984a). For a P range of 2·0–2·7 GPa (for justification see below), temperatures in the range 710–743 and 750–787°C are obtained for samples 402 and 386, respectively. Temperatures obtained with the THERMOCALC program and the dataset of Holland & Powell (1998) are higher: 730–810°C and 790–870°C.
For coexisting garnet–olivine-I and garnet–olivine-III that follow distinct partitioning trends (Fig. 7c) high ln values of 1·15–1·25 indicate substantial retrogressive equilibration to <600°C (see O'Neill & Wood, 1979). Similarly high ln KD values are obtained if averaged core compositions of olivine and large garnet grains are used from samples 386 and 402. Likewise, high values in the range 6·9–9·2 indicate retrogressive equilibration between coexisting garnet and clinopyroxene. The trend followed by coexisting Fe/Mg ratios (Fig. 7d) is interpreted to reflect the bulk composition rather than an equilibration at different temperatures.
Constraints on metamorphic pressures
Eclogite-facies conditions are implied by the presence of MORB-derived eclogites in both the NEZ and SEZ. Miller & Thöni (1995) derived P–T conditions of 2·7 GPa and 730°C from Ky-eclogites of the NEZ. This estimate is consistent with the absence of coesite from the eclogites. Averaged P calculations for Fe-rich quartz eclogites of the SEZ yield 2·2–2·4 GPa (Konzett et al. in preparation). Based on these results, peak metamorphic pressures are assumed to be >2·0–2·7 GPa and were used to calculate metamorphic temperatures.
A P estimate for the Fe-rich cumulates is hampered by a lack of suitable assemblages. Although garnet and orthopyroxene are present in a number of investigated samples, they do not share grain boundaries. Moreover, corona textures indicate that equilibrium between garnet and orthopyroxene was not attained on a thin-section scale. For samples 386 and 402, which contain large interstitial garnet, metamorphic pressures were estimated based on Al in orthopyroxene according to Harley (1984b) and Brey & Köhler (1990). If temperatures at the time of garnet–orthopyroxene equilibration are assumed to be 750–850°C, then pressures in the range of 1·4–1·9 GPa and 1·1–1·6 GPa are obtained for samples 386 and 402, respectively. These are substantially lower than pressures derived from the NEZ and SEZ eclogites. A possible explanation for this discrepancy would be re-equilibration of orthopyroxene with spinel, thus leading to anomalously high Al in orthopyroxene.
The presence of plagioclase in Ca-amphibole + chlorite mantles around garnet layers implies amphibolite-facies conditions during retrogression. Assuming <600–650°C, the stability of plagioclase with compositions in the range An57–73 and in the absence of kyanite + quartz indicates pressures <0·8 GPa (Goldsmith, 1982).
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INTERPRETATION OF PHASE RELATIONS |
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Olivine-rich rocks
Primary assemblages
The protolith of the olivine-rich rocks is interpreted to be a cumulate of olivine + orthopyroxene + spinel derived from a tholeiitic magma. After its emplacement and prior to high-P metamorphism, this cumulate lacked plagioclase as a major constituent. Arguments for this assumption are: (1) a lack of positive or the presence of negative Eu anomalies in all but one of the investigated samples indicating plagioclase fractionation rather than accumulation; (2) the absence of typical high-P breakdown products of plagioclase such as Na-rich clinopyroxene, kyanite or zoisite (Goldsmith, 1982
; Mottana et al., 1990, and references therein). The composite spinel grains are interpreted as products of exsolution and re-equilibration upon slow cooling of an originally homogeneous spinel. Evidence for this interpretation is provided by a regular alignment of spinel compositions close to the spinel gap in the Al–Cr–Fe3+ diagram (Fig. 10) (see Loferski & Lipin, 1983; Barnes & Roeder, 2001).
The role of clinopyroxene and amphibole in the olivine-rich rocks cannot be unambiguously assessed based on textural information. Clinopyroxenes are poor in Ti and Cr compared with gabbroic clinopyroxenes of the NEZ from rocks with similar bulk TiO2 (Miller, 1974), indicating a metamorphic origin. This assumption is further supported by a very low Ca-Tschermak component (see Kushiro & Thompson, 1972; Bender et al., 1978; Fujii & Bougault, 1983). Clinopyroxene inclusions in amphibole and in Al-poor rims of orthopyroxene indicate that this phase was present during at least part of the retrogressive P–T path. Indirect evidence for the presence of a Ca phase during the prograde P–T path is provided by Ca-bearing garnet coronae around (composite) spinel grains that are thought to have formed by continuous spinel breakdown through the reaction
 | (1) |
, 1970; O'Neill, 1981; Obata & Morten, 1987), which is consistent with the presence of eclogites associated with the olivine-rich rocks. In olivine-bearing assemblages, garnet may form through a number of reactions other than (1) but these involve plagioclase (e.g. Miller, 1974; Johnson & Essene, 1982). However, there is no evidence for plagioclase as a reactant for garnet formation. Nevertheless, the presence of clinopyroxene under peak metamorphic conditions cannot unambiguously be demonstrated because of a lack of matrix clinopyroxene or clinopyroxene inclusions in orthopyroxene cores.
Secondary assemblages
The retrogressive evolution of olivine-rich rocks is characterized by the formation of Ca-amphibole and chlorite. Large discontinuously zoned amphiboles may contain garnet, clinopyroxene and orthopyroxene relicts. In addition, some amphiboles contain myrmekitic intergrowths of Al-rich spinel + Ca-amphibole. These textures are consistent with retrogressive hydration through the reaction (Obata & Thompson, 1981)
 | (2) |
 | (3) |
) or by a subsolidus formation through retrogressive Sr–Ba-enriched fluids with a positive Eu anomaly (Michard, 1989; Douville et al., 2002). In the former case, Ca-amphibole formation would have removed all plagioclase as there is no evidence for involvement of this phase in the metamorphic evolution of the olivine-rich rocks. Textures and the P–T stability of pargasitic amphibole in gabbroic bulk compositions do not preclude the presence of at least some Ca-amphibole under peak metamorphic conditions originally inherited from the magmatic precursor (see Niida & Green, 1999; Coogan et al., 2001). However, a subsolidus origin of amphibole would be more consistent with its comparatively low REE contents as compared with igneous amphiboles (Fig. 8d) (e.g. Gillis, 1996). The depletion in HFSE of amphiboles (Fig. 8c) can be explained by preferential partitioning of these elements into coexisting ilmenite and/or zircon.
Al-poor discontinuous rims on orthopyroxene blasts, too, probably formed during retrogressive metamorphism. A possible mechanism of formation would be reaction of an infiltrating SiO2-rich fluid with olivine as proposed by Morishita et al. (2001b) for low-Al orthopyroxene from the Horoman Peridotite, Japan. In any case, Al concentrations in orthopyroxene rims do not mirror equilibrium distribution between spinel and orthopyoxene rims but rather a change from spinel to amphibole as coexisting Al phase.
Amphibolite-facies hydration of the olivine-rich rocks is documented in samples 399 and CM36/00-2 by the formation of talc from orthopyroxene according to the reaction
 | (4) |
, reaction (4) can be located between 630 and 650°C at P 1·0 GPa. This would be consistent with experimental results on the stability of olivine + talc obtained by Ulmer & Trommsdorff (1995).
The presence in the same sample of anthophyllite can be explained only by the topology proposed by Trommsdorff (1983). In this case, anthophyllite must have formed at <630–650°C by either of the following reactions after overstepping of reaction (4):
 | (5) |
 | (6) |
Garnet layers
Primary assemblages
Textures indicate a peak metamorphic assemblage garnet + clinopyroxene ± corundum ± spinel + Ti phase overprinted by a retrogressive infiltration of fluids. Corundum in mafic or ultramafic rocks is rather uncommon and mainly known from peraluminous mantle eclogites (e.g. Dawson & Carswell, 1990). Corundum-bearing mafic rocks were reported by Morishita & Arai (2001) from the Horoman peridotite complex, Japan. Corundum-bearing garnet clinopyroxenites have been described from the Beni Bousera, Malenco and Ronda peridotite massifs (Kornprobst et al., 1990; Müntener & Hermann, 1996; Müntener, 1997; Morishita et al., 2001a) and the Cabo Ortegal complex in Spain (Girardeau & Gil Ibarguchi, 1991). Garnet + clinopyroxene can form as a primary near-solidus assemblage in pyroxenitic rocks (e.g. Adam et al., 1992) but the additional presence of corundum was ascribed by the workers mentioned above to subsolidus reactions involving plagioclase.
For the garnet layers described here it is suggested that garnet + clinopyroxene ± corundum has formed by prograde subsolidus reactions in spinel + plagioclase-rich layers intercalated in a plagioclase-poor body of olivine + spinel + orthopyroxene cumulates according to reactions such as
 | (7) |
 | (8) |
from corundum-bearing garnet pyroxenites in the Malenco peridotite and ascribed to enrichment during subsolidus spinel + plagioclase breakdown.
Although no plagioclase was found in the garnet layers, its presence in the protoliths would be consistent with the following observations. (1) The bulk composition of CM23/00 is similar to that of plagioclase–clinopyroxene cumulate NEZ gabbroic rocks in terms of SiO2–Al2O3–CaO–TiO2. (2) Two out of three clinopyroxenes and the garnet analysis from JK3/00 have positive Eu anomalies with Eu/Eu* = 1·25–1·31 and 1·41, respectively (Fig. 5a). In addition, clinopyroxene analyses are characterized by positive Sr anomalies (Fig. 8b) that could be explained by a formation involving plagioclase as a reactant. (3) A comparison between garnet compositions from layers and from coronae in olivine-rich rocks reveals significantly higher grossular contents in layer garnets (Fig. 9), which would also be consistent with a formation from a Ca-rich phase such as plagioclase. (4) The bulk Sm/Nd ratio of 0·0890 for sample CM23/00 is lower than any value reported so far from ÖSC metabasites and cannot easily be inherited from a plagioclase-free precursor assemblage because common gabbroic phases except plagioclase show Sm/Nd ratios significantly higher than 0·10 (see Miller & Thöni, 1995; Miller et al., 2003). Sm/Nd ratios of plagioclase from ÖSC gabbros, on the other hand, are in the range 0·0815–0·0999 (Miller & Thöni, 1995). Thus, a significant amount of plagioclase in the CM23/00 precursor would be consistent with the low bulk Sm/Nd ratio.
A magmatic formation of garnet is unlikely because tholeiitic magmas do not have coexisting plagioclase and garnet as solidus phases (Elthon, 1993, and references therein). A metamorphic origin of garnet would also be consistent with the unusually HREE-depleted REE pattern of garnet from sample CM23/00, which contains zircon. This depletion is similar to that reported from eclogite-facies garnets by Rubatto & Hermann (2003) (Fig. 8a) and, following these workers, may be ascribed to disequilibrium REE partitioning between garnet formed during eclogite-facies metamorphism and magmatic zircon relics. Further arguments against a magmatic garnet + clinopyroxene + corundum assemblage are the positive Eu anomalies and the low Ca-Tschermak contents (<4 mol %) found in clinopyroxenes from the garnet layers.
Conditions of formation of the garnet–clinopyroxene–corundum assemblages cannot be assessed with confidence because of retrogressive re-equilibration. This is demonstrated by the heterogeneity of garnet compositions (Fig. 9) and high garnet–clinopyroxene values (6·9–9·2, Fig. 7d), which yield low- to mid-amphibolite-facies temperatures.
Secondary assemblages
Fluid infiltration into the garnet layers during uplift led to an assemblage högbomite, Al-rich spinel, amphibole, chlorite, phlogopite–aspidolite, Sr- and Cl-rich apatite, ilmenite and magnetite along the interface between garnet layers and olivine-rich rocks.
Because of a lack of suitable assemblages, only order-of-magnitude estimates can be made for the P–T conditions of retrogression: plagioclase + zoisite coexisting with amphibole + chlorite in sample CM37/00 indicates P < 0·8 GPa at T < 600–650°C (Goldsmith, 1982). The very low Ti contents of magnetite coexisting with ilmenite would be consistent with retrogression extending to greenschist-facies conditions of <400–450°C. Apart from P and T, the composition of the infiltrating fluid and local variations in the bulk chemistry are major factors controlling the succession and composition of retrogressive phases, which is demonstrated by the formation of högbomite and amphibole.
Högbomite is typically present as an accessory phase in aluminous rocks of the upper amphibolite to lower granulite facies (e.g. Grew et al., 1987, 1990; Rammlmair et al., 1988; Petersen et al., 1989). In the rocks studied here, högbomite is restricted to garnet layers and it is invariably intergrown with or present in the immediate vicinity of corundum. Because högbomite has not been reported so far as a stable matrix phase from eclogite-facies environments, its presence in garnet layers is ascribed to retrogression during uplift after the peak of Variscan metamorphism under upper amphibolite-facies conditions and locally controlled by the stability of corundum. A retrograde formation would also be consistent with högbomite and sapphirine in symplectites around kyanite reported from retrogressed eclogites by Liati & Seidel (1994). Textures and phase compositions indicate the formation of högbomite from spinel, rutile and corundum:
 | (9) |
Amphiboles show a wide compositional range and multiple discontinuous zoning, which in many cases is not consistent with a formation under decreasing (P)–T (e.g. tremolite cores with discontinuous pargasite rims) and, thus, rather monitors the composition of the infiltrating fluid. Textures indicate an evolution of amphibole composition from Mg-rich, K- and Cl-poor to Fe-, K- and Cl-rich with a strong decrease in the modal abundance of the Cl-rich amphiboles (Fig. 2f). Textures and the range of Cl/OH ratios of amphiboles from the garnet layers may be explained by infiltration of Cl-bearing fluids along the interface between garnet layers and olivine-rich rocks during uplift and exhumation. Amphibole (+ chlorite) formation through reaction of the fluid with garnet and clinopyroxene then led to a progressive enrichment of the fluid in Cl as a result of preferential partitioning of H into amphibole (see Markl et al., 1998). Conditions of retrogressive hydration of the garnet layers are estimated to be 600–650°C and 
0·8 GPa.
A further indicator for Cl-rich fluids during retrograde metamorphism is Cl-rich apatite. Although it is mostly associated with Cl-rich amphibole, its rare presence as tiny (<5 µm) grains in olivine-rich rocks implies the presence of at least small amounts of saline fluids throughout the entire cumulate body. Although there is clear evidence for penetration of Cl-rich fluids into the garnet layers, the texturally early generation of apatite preserved their low Cl contents and, thus, did not equilibrate with the saline fluids.
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GEOCHRONOLOGY |
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SHRIMP U–Pb dating of zircons
Zircons are present both in olivine-rich rocks and garnet layers (Table 1). In sample JK1/00, two zircons appear as isolated inclusions in a large olivine-I not associated with any cracks or other inclusions, suggesting a primary magmatic origin. Unfortunately, these zircons are too small to be analysed with the SHRIMP.
Zircons suitable for SHRIMP dating are present in sample JK3/00 as inclusions in a garnet layer. They form anhedral, rounded to ovoid grains, usually 30–40 µm in diameter, and in rare cases elongate grains up to 100 µm x 50 µm size. In transmitted light, the zircons are clear and not obviously structured. CL imaging, however, reveals some internal zoning, which may appear as diffuse and cloudy areas (Fig. 14a) or subhedrally shaped zones (Fig. 14b) that are frequently cut and overgrown by zircon of more uniform CL intensity. In rare cases, zircons may show oscillatory zoning (Fig. 14c). These internal structures and the oscillatory zoning are interpreted as variably preserved magmatic textures, whereas the rims with uniform CL intensity are attributed to metamorphic recrystallization or growth. The development of markedly rounded terminations of the elongate zircons could also be attributed to metamorphic modification.
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Eleven analyses were carried out on different zircons (Table 15) targeting cloudy cores and uniformly structured rims on a concordia plot (Fig. 14d). No single age can be calculated from the data as they spread down the concordia, indicating significant but variable Pb loss. No relationship between U contents and degree of discordance could be found. The least discordant and apparently oldest analyses comprise a group of data for which a weighted mean 206*Pb/238U age of 517 ± 7 Ma [mean squared weighted deviates (MSWD) = 0·54, probability = 0·65] can be calculated; we interpret this as the original age of crystallization of the zircons and the enclosing rock. This age would be consistent with a late Cambrian magmatic episode in the ÖSC that also led to the emplacement of MORB-type gabbros and basalts around 530–521 Ma as inferred by Miller & Thöni (1995)
using Sm–Nd isotope data. The zircon ages in the range 492–376 Ma may be due to partial resetting as a result of a later thermal overprint. The most likely event to cause this resetting is the Variscan high-P metamorphism, which not only overprinted the Fe-rich cumulates described here but also led to the formation of eclogite-facies assemblages in MORB-type gabbros and basalts around 370–340 Ma (Miller & Thöni, 1995). Recent U–Pb SHRIMP data for zircons from eclogites of the Silvretta nappe (Ladenhauf et al., 2001) confirm a widespread high-P event in the Austroalpine crystalline basement west of the Tauern Window around 350 Ma. It should be noted that the data from the Silvretta nappe also hint at a thermal event around 440 Ma that might also have contributed to the resetting of ages from JK3/00. The significance of this event, however, has to await an assessment on the basis of additional data.
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Constraints from Sm–Nd isotopes on the post-emplacement history of the mafic cumulates
Fe-rich eclogites of the NEZ show an unusually large range in whole-rock 147Sm/144Nd ratios (0·194–0·718) and, for this reason, were used by Miller & Thöni (1995)
to calculate an errorchron based on whole-rock Sm–Nd data. Although the resulting age of 362 ± 29 Ma displays a considerable uncertainty with MSWD = 15·7, the similarity of the initial Nd isotopic ratio derived from whole-rock data and that derived from mineral–whole-rock isochrons of individual samples led Miller & Thöni (1995) to conclude that the errorchron age of 362 Ma, in spite of the large uncertainty, represents a meaningful metamorphic date. This implies that the Sm–Nd isotopic system was largely reset during the Variscan high-P metamorphic event.
A comparison between whole-rock Sm–Nd data from cumulate rocks of this study and Fe-rich eclogites of the NEZ reveals that two out of three samples plot close to the 362 ± 29 Ma errorchron, extending the range of 147Sm/144Nd ratios to values as low as 0·089 (Fig. 15). A regression through data from these two samples (JK5/00 and CM23/00) and data for Fe-eclogites presented by Miller & Thöni (1995, table 8) yields a regression age of 372 ± 22 Ma (MSWD = 84). This coincidence—although insufficient to establish beyond doubt a thermal event in the age range given by the errorchron—would at least be consistent with the assumption that the Sm–Nd systems in both eclogites of the NEZ and mafic cumulates of the SEZ were (partially) reset as a result of Variscan high-P metamorphism provided that both protoliths originated from the same isotopic reservoir. A partial resetting of isotopic systems would also be consistent with the observed spread in U–Pb zircon ages.
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NATURE OF THE PROTOLITH |
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The CIPW-normative compositions of the Pollesalm ultramafic rocks plot in the fields of mela-troctolite, mela-olivine norite, mela-olivine gabbro and troctolite. The data can be divided into two groups: (1) Mg-rich ultramafic rocks with CIPW-normative assemblages of 60–89 olivine (containing 75–79 mol % Fo), 7–20 plagioclase, 0–7 clinopyroxene and 0–18 orthopyroxene; (2) Al–Ca-rich sample CM23/00. The high whole-rock XMg of 78·6–81·7 points to a cumulate nature of the olivine-rich samples. The Mg-numbers of olivine-I are too low to have been in equilibrium with a mantle-derived primitive magma, suggesting some olivine + Cr-spinel fractionation prior to cumulate crystallization. The LREE-enriched signature of the olivine-rich samples (Fig. 5a) would be consistent with the presence of some intercumulus melt. By comparison, the garnet-rich sample CM23/00 is more evolved, with a whole-rock XMg of 55, and contains approximately 58% CIPW normative plagioclase, 35% olivine and 2% clinopyroxene. Its REE pattern is subparallel to those of the olivine-rich samples but its REE concentration is distinctly higher, suggesting that it crystallized from a cogenetic melt enriched in incompatible elements. The small positive Eu anomaly of CM23/00 would be consistent with the former presence of plagioclase in this layer. An alternative explanation for the LREE enrichment that cannot be excluded based on the available data would be late fluid infiltration that was also responsible for amphibole and/or apatite growth in these samples.
The flat HREE patterns and the positive correlations between Nb and Y (and Yb) (Figs 4 and 5) indicate that Y and the HREE were incompatible during the processes responsible for within-suite variations in the Pollesalm ultramafic rocks, suggesting that garnet was initially not present as a residual or fractionating phase. The lack of evidence for magmatic clinopyroxene and the observed Eu anomalies are thought to be a consequence of low-P (<0·5 GPa) plagioclase formation involving fractionation–accumulation of a moderately differentiated tholeiitic magma with 
in the range 0·57–0·65.
The Nd and Sr isotopic signatures of samples JK5/00 and CM23/00 (Fig. 6, Table 2) suggest that the protolith material could have been derived from a depleted asthenospheric mantle source similar to that of the NEZ protolith in spite of contrasting trace element patterns, which show LREE enrichment in the SEZ samples but LREE depletion in the NEZ melagabbros (Fig. 5a). The deviation of sample 386 from the 517 Ma mantle array (Fig. 6), however, is striking. Its time-integrated 87Sr/86Sr and 143Nd/144Nd ratios are lower than Bulk Earth values. This isotopic signature could have been generated by substantial contamination of mantle-derived magmas with material characterized by low Rb/Sr and low Sm/Nd, such as granulite-facies material stored in the lower continental crust (e.g. Carter et al., 1978). This indicates magma generation through decompression during continental rifting rather than a MOR-type scenario involving thin oceanic crust.
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ACKNOWLEDGEMENTS |
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We are greatly indebted to Peter Ulmer and Othmar Müntener for incisive reviews of an early version of the manuscript. Alberto Zanetti is thanked for performing LA-ICP-MS analyses. Reviews by Reto Gieré, Tomoaki Morishita and John Schumacher are gratefully acknowledged. Part of this work was financially supported by a University of Innsbruck research grant.
* Corresponding author. Telephone: +43-(0)512-507-5506. Fax: +43-(0)512-507-2926. E-mail: juergen.konzett{at}uibk.ac.at
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REFERENCES |
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Ackermand, D., Widley, B. F. & Herd, R. K. (1983). Magnesian högbomite in a sapphirine-bearing rock from the Fiskenæsset region, W. Greenland. Mineralogical Magazine 47, 555–561.[Web of Science]
Adam, J., Green, T. H. & Day, R. A. (1992). An experimental study of two garnet pyroxenite xenoliths from the Bullenmerri and Gnotuk maars of western Victoria, Australia. Contributions to Mineralogy and Petrology 111, 505–514.[CrossRef][Web of Science]
Armbruster, T., Bermanec, V., Zebec, V. & Oberhänsli, R. (1998). Titanium and iron poor zincohögbomite-16H, Zn14(Al,Fe3+,Ti,Mg)8Al24O62(OH)2, from Nezilowo, Macedonia: occurrence and crystal structure of a new polysome. Schweizerische Mineralogische und Petrographische Mitteilungen 78, 469–477.
Barnes, S. J. & Roeder, P. L. (2001). The range of spinel compositions in terrestrial mafic and ultramafic rocks. Journal of Petrology 42, 2279–2302.
Bender, J. F., Hodges, F. N. & Bence, A. E. (1978). Petrogenesis of basalts from the project FAMOUS area: experimental study from 0–15 kbars. Earth and Planetary Science Letters 41, 277–302.[CrossRef][Web of Science]
Beukes, G. J., van Zyl, V. C., Schoch, A. E., De Bruiyn, H., van Aswegen, G. & Strydom, D. (1986). A högbomite–spinel–gedrite–paragenesis from northern Bushmanland Namaqua mobile belt, South Africa. Neues Jahrbuch für Mineralogie, Abhandlungen 155, 53–66.
Boynton, W. V. (1984). Cosmochemistry of the rare earth elements: meteorite studies. In: Henderson, P. (ed.) Rare Earth Element Geochemistry. Amsterdam: Elsevier, pp. 63–114.
Brey, G. P. & Köhler, T. (1990). Geothermobarometry in four-phase lherzolites II: new thermobarometers, and empirical assessment of existing thermobarometers. Journal of Petrology 31, 1353–1378.
Carter, S. R., Evensen, N. M., Hamilton, P. J. & O'Nions, R. K. (1978). Neodymium and strontium isotope evidence for crustal contamination of continental volcanics. Science 202, 743–746.
Coogan, L. A., Wilson, R. N., Gillis, K. M. & MacLeod, C. J. (2001). Near-solidus evolution of oceanic gabbros: insights from amphibole geochemistry. Geochimica et Cosmochimica Acta 65, 4339–4357.[CrossRef][Web of Science]
Coolen, J. J. (1981). Högbomite and aluminium spinel from some metamorphic rocks and Fe–Ti ores. Neues Jahrbuch für Mineralogie, Monatshefte 8, 374–384.
Cortesogno, L., Gaggero, L. & Zanetti, A. (2000). Rare earth and trace elements in igneous and high-temperature metamorphic minerals of oceanic gabbros (MARK area, Mid-Atlantic Ridge) Contributions to Mineralogy and Petrology 139, 373–393.[CrossRef][Web of Science]
Dawson, J. B. & Carswell, D. A. (1990). High-temperature and ultra-high pressure eclogites. In: Carswell, D. A. (ed.) Eclogite Facies Rocks. Glasgow: Blackie, pp. 315–349.
Douville, E., Charlou, J. L., Oelkers, E. H., Bienvenu, P., Jove Colon, C. F., Donval, J. P., Fouquet, Y., Prieur, D. & Appriou, P. (2002). The rainbow vent fluids (36°14'N, MAR): the influence of ultramafic rocks and phase separation on trace metal content in Mid-Atlantic Ridge hydrothermal fluids. Chemical Geology 184, 37–48.[CrossRef][Web of Science]
Droop, G. T. R. (1983). Pre-Alpine eclogites in the Pennine basement complex of the Eastern Alps. Journal of Metamorphic Geology 1, 3–12.[Web of Science]
Elthon, D. (1993). The crystallization of mid-ocean ridge basalts at moderate and high pressures. European Journal of Mineralogy 5, 1025–1037.
Fonarev, V. I. (1981). Thermodynamic functions of olivine and orthopyroxene solid solutions. Geochemistry International 8, 142–157.
Fujii, T. & Bougault, H. (1983). Melting relations of a magnesian abyssal tholeiite and the origin of MORBs. Earth and Planetary Science Letters 62, 283–295.[CrossRef][Web of Science]
Fügenschuh, B., Mancktelow, N. S. & Seward, D. (2000). Cretaceous to Neogene cooling and exhumation history of the Ötztal–Stubai basement complex, Eastern Alps: a structural and fission track study. Tectonics 19, 905–918.[CrossRef][Web of Science]
Gebauer, D. & Grünenfelder, M. (1978). U–Pb zircon dating of Alpine-type garnet peridotites, example Val d'Ultimo (eastern Alps, northern Italy). US Geological Survey, Open-File Report 78-101, 135–137.
Gieré, R. (1986). Zirconolite, allanite and hoegbomite in a marble skarn from the Bergell contact aureole: implications for mobility of Ti, Zr and REE. Contributions to Mineralogy and Petrology 93, 459–470.[CrossRef][Web of Science]
Gillis, K. M. (1996). Rare earth element constraints on the origin of amphibole in gabbroic rocks from site 894, Hess Deep. In: Mével, C., Gillis, K. M., Allan, J. F. & Meyer, P. S. (eds) Proceedings of the Ocean Drilling Program, Scientific Results 147. College Station, TX: Ocean Drilling Program, pp. 59–75.
Girardeau, J. & Gil Ibarguchi, J. I. (1991). Pyroxenite-rich peridotites of the Cabo Ortegal Complex (northwestern Spain): evidence for large-scale upper-mantle heterogeneity. Journal of Petrology Special Lherzolites Issue, 135–154.
Godard, G., Martin, S., Prosser, G., Kienast, J. R. & Morten, L. (1996). Variscan migmatites, eclogites and garnet peridotites of the Ulten zone, Eastern Austroalpine system. Tectonophysics 259, 313–341.[CrossRef][Web of Science]
Goldsmith, J. R. (1982). Plagioclase stability at elevated temperatures and water pressures. American Mineralogist 67, 653–675.[Abstract]
Grew, E. S., Abraham, K. & Medenbach, O. (1987). Ti-poor hoegbomite in kornerupine–cordierite–sillimanite rocks from Ellammanvilpatti, Tamil Nadu, India. Contributions to Mineralogy and Petrology 45, 21–31.
Grew, E. S., Hiroi, Y. & Shiraishi, K. (1990). Högbomite from the Prince Olav Coast, East Antarctica: an example of oxidation–exsolution of a complex magnetite solid-solution? American Mineralogist 75, 589–600.[Abstract]
Harley, S. L. (1984a). An experimental study of the partitioning of Fe and Mg between garnet and orthopyroxene. Contributions to Mineralogy and Petrology 86, 359–373.[CrossRef][Web of Science]
Harley, S. L. (1984b). The solubility of alumina in orthopyroxene coexisting with garnet in FeO–MgO–Al2O3–SiO2 and CaO–FeO–MgO–Al2O3–SiO2. Journal of Petrology 25, 665–696.
Hauzenberger, C. A., Höller, W. & Hoinkes, G. (1996). Transition of eclogite to amphibolite-facies metamorphism in the Austroalpine Ulten Zone. Mineralogy and Petrology 58, 111–130.[CrossRef][Web of Science]
Hoinkes, G., Thöni, M., Lichem, Ch., Bernhard, F., Kaindl, R., Schweigl, J., Tropper, P. & Cosca, M. (1997). Metagranitoids and associated metasediments as indicators for the pre-Alpine magmatic and metamorphic evolution of the western Ötztal Basement (Kaunertal, Tirol). Schweizerische Mineralogische und Petrographische Mitteilungen 77, 299–314.
Hoinkes, G., Koller, F., Dachs, E., Höck, V., Neubauer, F., Rantitsch, G. & Schuster, R. (1999). Alpine metamorphism in the Eastern Alps. Schweizerische Mineralogische und Petrographische Mitteilungen 79, 155–181.[Web of Science]
Holland, T. J. B. & Powell, R. (1998). An internally consistent thermodynamic data set for phases of petrological interest. Journal of Metamorphic Geology 16, 309–343.[CrossRef][Web of Science]
Johnson, C. A. & Essene, E. J. (1982). The formation of garnet in olivine-bearing metagabbros from the Adirondacks. Contributions to Mineralogy and Petrology 81, 240–251.[CrossRef][Web of Science]
Klötzli-Chowanetz, E., Klötzli, U. & Koller, F. (1997). Lower Ordovician migmatization in the Ötztal Crystalline Basement (Eastern Alps, Austria): linking U–Pb and Pb–Pb dating with zircon morphology. Schweizerische Mineralogische und Petrographische Mitteilungen 77, 315–324.
Koch-Müller, M., Cemic, L. & Langer, K. (1992). Experimental and thermodynamic study of Fe–Mg exchange between olivine and orthopyroxene in the system MgO–FeO–SiO2. European Journal of Mineralogy 4, 115–135.
Kornprobst, J., Piboule, M., Roden, M. & Tabit, A. (1990). Corundum-bearing garnet clinopyroxenites at Beni Bousera (Morocco): original plagioclase-rich gabbros recrystallized at depth within the mantle? Journal of Petrology 31, 717–745.
Koziol, M. & Oberhänsli, R. (1995). Die südliche Eklogitzone im Ötztal. Mitteilungen der Österreichischen Mineralogischen Gesellschaft 140, 95–118.
Kretz, R. (1983). Symbols for rock-forming minerals. American Mineralogist 68, 277–279.[Abstract]
Kushiro, I. & Thompson, R. N. (1972). Origin of some abyssal tholeiites from the Mid-Atlantic Ridge. Carnegie Institution of Washington Yearbook 71, 403–406.
Ladenhauf, Ch., Armstrong, R. A., Konzett, J. & Miller, Ch. (2001). The timing of Pre-Alpine high-pressure metamorphism in the Eastern Alps: constraints from U–Pb SHRIMP dating of eclogite zircons from the Austro-Alpine Silvretta Nappe. Journal of Conference Abstracts 6, 600.
Leake, B. E., Woolley, A. R., Arps, C. E. S., et al. (1997). Nomenclature of amphiboles: report of the Subcommittee on Amphiboles of the International Mineralogical Association, Commission on New Minerals and Mineral Names. American Mineralogist 82, 1019–1037.[Abstract]
Liati, A. & Seidel, E. (1994). Sapphirine and högbomite in overprinted kyanite-eclogites of the central Rhodope, N Greece, first evidence of granulite metamorphism. European Journal of Mineralogy 6, 733–738.
Loferski, P. J. & Lipin, B. R. (1983). Exsolution in metamorphosed chromite from the Red Lodge district, Montana. American Mineralogist 68, 777–789.[Abstract]
Ludwig, K. R. (1999). Isoplot/Ex Version 2.00: a Geochronological Toolkit for Microsoft Excel. Berkeley Geochronology Center Special Publication 1a, 46 pp.
Ludwig, K. R. (2000). SQUID 1.00, a User's Manual. Berkeley Geochronology Center Special Publication 2, 17 pp.
MacGregor, I. D. (1965). Stability fields of spinel and garnet peridotites in the synthetic system MgO–CaO–Al2O3–SiO2. Carnegie Institution of Washington Yearbook 64, 126–134.
MacGregor, I. D. (1970). The effect of CaO, Cr2O3, Fe2O3 and Al2O3 on the stability of spinel and garnet peridotites. Physics of the Earth and Planetary Interiors 3, 372–377.[CrossRef]
Maggetti, M. & Flisch, M. (1993). Evolution of the Silvretta Nappe. In: von Raumer, J. F. & Neubauer, F. (eds) Pre-Mesozoic Geology in the Alps. Berlin: Springer, pp. 469–484.
Maggetti, M. & Galetti, G. (1984). Chemie und geotektonische Position von Metabasiten aus dem Südosten der Silvretta (Schweiz). Schweizerische Mineralogische und Petrographische Mitteilungen 64, 423–450.
Maggetti, M. & Galetti, G. (1988). Evolution of the Silvretta eclogites: metamorphic and magmatic events. Schweizerische Mineralogische und Petrographische Mitteilungen 68, 467–484.
Maggetti, M., Galetti, G. & Stosch, H.-G. (1987). Eclogites from the Silvretta nappe (Switzerland): geochemical constraints on the nature and geotectonic setting of their protoliths. Chemical Geology 64, 319–334.[CrossRef][Web of Science]
Mancktelow, N. S. (1981). Högbomite of unusual composition from Reedy Creek, South Australia. Mineralogical Magazine 44, 91–94.[Web of Science]
Markl, G., Ferry, J. & Bucher, K. (1998). Formation of saline brines and salt in the lower crust by hydration reactions in partially retrogressed granulites from the Lofoten Islands, Norway. American Journal of Science 298, 705–757.
McCulloch, M. T. & Wasserburg, G. J. (1978). Sm–Nd and Rb–Sr chronology of continental crust formation. Science 200, 1003–1011.
Michard, A. (1989). Rare earth element systematics in hydrothermal fluids. Geochimica et Cosmochimica Acta 53, 745–750.[CrossRef][Web of Science]
Michard, A., Gurriet, P., Soudant, M. & Albarède, F. (1985). Nd isotopes in French Phanerozoic shales: external vs. internal aspects of crustal evolution. Geochimica et Cosmochimica Acta 49, 601–610.[CrossRef][Web of Science]
Miller, Ch. (1974). Reaction rims between olivine and plagioclase in metaperidotites, Ötztal Alps, Austria. Contributions to Mineralogy and Petrology 43, 333–342.[CrossRef][Web of Science]
Miller, Ch. & Thöni, M. (1995). Origin of eclogites from the Austroalpine Ötztal basement (Tirol, Austria) geochemistry and Sm–Nd vs. Rb–Sr isotope systematics. Chemical Geology 122, 199–225.[CrossRef][Web of Science]
Miller, Ch., Thöni, M., Frank, W., Schuster, R., Melcher, F., Meisel, T. & Zanetti, A. (2003). Geochemistry and tectonomagmatic affinity of the Yungbwa ophiolite, SW Tibet. Lithos 66, 155–172.[CrossRef][Web of Science]
Morishita, T. & Arai, S. (2001). Petrogenesis of corundum-bearing mafic rock in the Horoman Peridotite Complex, Japan. Journal of Petrology 42, 1279–1299.
Morishita, T., Arai, S. & Gervilla, F. (2001a). High-pressure aluminous mafic rocks from the Ronda peridotite massif, southern Spain: significance of sapphirine- and corundum-bearing mineral assemblages. Lithos 57, 143–161.[CrossRef][Web of Science]
Morishita, T., Arai, S. & Green, D. H. (2001b). Evolution of low-Al orthopyroxene in the Horoman Peridotite, Japan: an unusual indicator of metasomatizing fluids. Journal of Petrology 44, 1237–1246.
Mottana, A., Carswell, D. A., Chopin, C. & Oberhänsli, R. (1990). Eclogite facies mineral parageneses. In: Carswell, D. A. (ed.) Eclogite Facies Rocks. Glasgow: Blackie, pp. 14–52.
Müntener, O. (1997). The Malenco peridotites (Alps): petrology and geochemistry of subcontinental mantle and Jurassic exhumation during rifting. Ph.D. thesis, Swiss Federal Institute of Technology, Zürich, 205 pp.
Müntener, O. & Hermann, J. (1996). The Malenco lower crust–upper mantle complex and its field relations (Italian Alps). Schweizerische Mineralogische und Petrographische Mitteilungen 76, 475–500.[Web of Science]
Müntener, O., Hermann, J. & Trommsdorff, V. (2000). Cooling history and exhumation of lower-crustal granulite and upper mantle (Malenco, Eastern Central Alps). Journal of Petrology 41, 175–200.
Niida, K. & Green, D. H. (1999). Stability and composition of pargasitic amphibole in MORB pyrolite under upper mantle conditions. Contributions to Mineralogy and Petrology 135, 18–40.[CrossRef][Web of Science]
Obata, M. (1980). The Ronda Peridotite: garnet-, spinel- and plagioclase lherzolite facies and the P–T trajectories of a high-temperature mantle intrusion. Journal of Petrology 21, 533–572.
Obata, M. & Morten, L. (1987). Transformation of spinel lherzolite to garnet lherzolite in ultramafic lenses of the Austridic crystalline complex, northern Italy. Journal of Petrology 28, 599–623.
Obata, M. & Thompson, A. B. (1981). Amphibole and chlorite in mafic and ultramafic rocks in the lower crust and upper mantle. Contributions to Mineralogy and Petrology 77, 74–81.[CrossRef][Web of Science]
O'Neill, H. St. C. (1981). The transition between spinel lherzolite and garnet lherzolite, and its use as a geobarometer. Contributions to Mineralogy and Petrology 77, 185–194.[CrossRef][Web of Science]
O'Neill, H. St. C. & Wood, B. J. (1979). An experimental study of Fe–Mg partitioning between garnet and olivine and its calibration as a geothermometer. Contributions to Mineralogy and Petrology 70, 59–70.[CrossRef][Web of Science]
Paces, J. B. & Miller, J. D. (1993). Precise U–Pb ages of Duluth Complex and related mafic intrusions, northeastern Minnesota: geochronological insights to physical, petrogenetic, paleomagnetic and tectonomagnetic processes associated with the 1·1 Ga Midcontinent Rift System. Journal of Geophysical Research 98B, 13997–14013.
Pearce, J. A. (1983). A ‘user's guide’ to basalt discrimination diagrams. Unpublished report, The Open University, Milton Keynes, 37 pp.
Petersen, E. U., Essene, E. J., Peacor, D. R. & Marcotty, L. A. (1989). The occurrence of högbomite in high-grade metamorphic rocks. Contributions to Mineralogy and Petrology 101, 350–360.[CrossRef][Web of Science]
Rammlmair, D., Mogessie, A., Purtscheller, F. & Tessadri, R. (1988). Högbomite from the Vumba schist belt, Botswana. American Mineralogist 73, 651–656.[Abstract]
Roeder, P. L. & Emslie, R. F. (1970). Olivine–liquid equilibrium. Contributions to Mineralogy and Petrology 29, 275–289.[CrossRef][Web of Science]
Rubatto, D. & Hermann, J. (2003). Zircon formation during fluid circulation in eclogites (Monviso, Western Alps): implications for Zr and Hf budget in subduction zones. Geochimica et Cosmochimica Acta 67, 2173–2187.[CrossRef][Web of Science]
Schweinehage, R. & Massonne, H. J. (1999). Geochemistry and metamorphic evolution of metabasites from the Silvretta Nappe, Eastern Alps. In: Gosso, G., Jadoul, F., Sella, M. & Spalla, M. I. (eds) 3rd Workshop on Alpine Geological Studies. Memorie di Scienze Geologiche 51, 191–203.
Sölva, H., Thöni, M., Grasemann, B. & Linner, M. (2001). Emplacement of eo-Alpine high-pressure rocks in the Austroalpine Ötztal Complex (Texel Group, Italy/Austria). Geodinamica Acta 14, 345–360.[CrossRef][Web of Science]
Teale, G. S. (1980). The occurrence of högbomite and taafeite in a spinel–phlogopite schist from the Mount Painter Province of South Australia. Mineralogical Magazine 43, 575–577.[Web of Science]
Thöni, M. (1999). A review of geochronological data from the Eastern Alps. Schweizerische Mineralogische und Petrographische Mitteilungen 79, 209–230.[Web of Science]
Tiepolo, M., Bottazzi, P., Palenzona, M. & Vannucci, R. (2003). A laser probe coupled with ICP-double-focusing sector-field mass spectrometer for in situ analysis of geological samples and U–Pb dating of zircon. Canadian Mineralogist 41, 259–272.[CrossRef][Web of Science]
Tribuzio, R., Riccardi, M. P. & Ottolini, L. (1995). Trace element redistribution in high temperature deformed gabbros from East Ligurian ophiolites (northern Appenines, Italy): constraints on the origin of syndeformation fluids. Journal of Metamorphic Geology 13, 367–377.[Web of Science]
Tribuzio, R., Tiepolo, M., Vannucci, R. & Bottazzi, P. (1999). Trace element distribution within olivine-bearing gabbros from the Northern Apennine ophiolites (Italy): evidence for post-cumulus crystallization in MOR-type gabbroic rocks. Contributions to Mineralogy and Petrology 134, 123–133.[CrossRef][Web of Science]
Trommsdorff, V. (1983). Metamorphose magnesiumreicher Gesteine: Kritischer Vergleich von Natur, Experiment und thermodynamischer Datenbasis. Fortschritte der Mineralogie 61, 283–308.[Web of Science]
Trommsdorff, V. & Evans, B. W. (1969). The stable association enstatite–forsterite–chlorite in amphibolite facies ultramafics of the Lepontine Alps. Schweizerische Mineralogische und Petrographische Mitteilungen 49, 325–332.
Tumiati, S., Thöni, M., Nimis, P., Martin, S. & Mair, V. (2003). Mantle–crust interactions during Variscan subduction in the Eastern Alps (Nonsberg–Ulten Zone); geochronology and new petrological constraints. Earth and Planetary Science Letters 210, 509–526.[CrossRef][Web of Science]
Ulmer, P. (1989). The dependence of Fe2+–Mg cation partitioning between olivine and basaltic liquid on pressure, temperature and composition. An experimental study to 30 kbars. Contributions to Mineralogy and Petrology 101, 261–273.[CrossRef][Web of Science]
Ulmer, P. & Trommsdorff, V. (1995). Serpentine stability to mantle depths and subduction-related magmatism. Science 268, 858–861.
von Seckendorff, V. & O'Neill, H. St. C. (1993). An experimental study of Fe–Mg partitioning between olivine and orthopyroxene at 1173, 1273 and 1423 K and 1·6 GPa. Contributions to Mineralogy and Petrology 113, 196–207.[CrossRef][Web of Science]
Witt-Eickschen, G. & Seck, H. A. (1991). Solubility of Ca and Al in orthopyroxene from spinel peridotite: an improved version of an empirical geothermometer. Contributions to Mineralogy and Petrology 106, 431–439.[CrossRef][Web of Science]
Yalçin, Ü., Schreyer, W. & Medenbach, O. (1993). Zn-rich högbomite formed from gahnite in the metabauxites of the Menders Massif, SW Turkey. Contributions to Mineralogy and Petrology 113, 314–324.[CrossRef][Web of Science]
Zakrzewski, M. A. (1977). Högbomite from the Fe–Ti deposit of Liganga (Tanzania). Neues Jahrbuch für Mineralogie, Monatshefte 8, 373–380.
Zimmermann, R. & Franz, G. (1988). Die Eklogite der Unteren Schieferhülle; Frosnitztal/Südvenediger (Tauern, Österreich). Mitteilungen der Österreichischen Mineralogischen Gesellschaft 81, 167–188.
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