Journal of Petrology Advance Access originally published online on June 20, 2006
Journal of Petrology 2006 47(10):1943-1971; doi:10.1093/petrology/egl032
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Exhumation History of a Garnet Pyroxenite-bearing Mantle Section from a Continent–Ocean Transition (Northern Apennine Ophiolites, Italy)
1 DIPARTIMENTO DI SCIENZE DELLA TERRA, UNIVERSITÀ DI PARMA PARCO AREA DELLE SCIENZE 157A, 43100 PARMA, ITALY
2 DIPARTIMENTO DI SCIENZE DELLA TERRA, UNIVERSITÀ DI PAVIA I-27100 PAVIA, ITALY
3 IGG-CNR ISTITUTO DI GEOSCIENZE E GEORISORSE PAVIA, I-27100 PAVIA, ITALY
4 DEPARTMENT OF GEOLOGY, ROYAL HOLLOWAY, UNIVERSITY OF LONDON TW20 0EX EGHAM, UK
RECEIVED JUNE 2, 2005; ACCEPTED MAY 17, 2006
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ABSTRACT |
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TOP
ABSTRACT
INTRODUCTIONGarnet clinopyroxenite and garnet websterite layers occur locally within mantle peridotite bodies from the External Liguride Jurassic ophiolites (Northern Apennines, Italy). These ophiolites were derived from an ocean–continent transition similar to the present-day western Iberian margin. The garnet clinopyroxenites are mafic rocks with a primary mineral assemblage of pyrope-rich garnet + sodic Al-augite (Na2O
2·5 wt %, Al2O3 12·5 wt %), with accessory graphite, Fe–Ni sulphides and rutile. Decompression caused Na-rich plagioclase (An50–45) exsolution in clinopyroxene porphyroclasts and extensive development of symplectites composed of secondary orthopyroxene + plagioclase (An85–72) + Al-spinel ± clinopyroxene ± ilmenite at the interface between garnet and primary clinopyroxene. Further decompression is recorded by the development of an olivine + plagioclase-bearing assemblage, locally under syn-kinematic conditions, at the expense of two-pyroxenes + Al-spinel. Mg-rich garnet has been also found in the websterite layers, which are commonly characterized by the occurrence of symplectites made of orthopyroxene + Al-spinel ± clinopyroxene. The enclosing peridotites are Ti-amphibole-bearing lherzolites with a fertile geochemical signature and a widespread plagioclase-facies mylonitic foliation, which preserve in places a spinel tectonite fabric. Lu–Hf and Sm–Nd mineral isochrons (220 ± 13 Ma and 186.0 ± 1·8 Ma, respectively) have been obtained from a garnet clinopyroxenite layer and interpreted as cooling ages. Geothermobarometric estimates for the high-pressure equilibration have yielded T 1100°C and P 2·8 GPa. The early decompression was associated with moderate cooling, corresponding to T 950°, and development of a spinel tectonite fabric in the lherzolites. Further decompression associated with plagioclase–olivine growth in both peridotites and pyroxenites was nearly isothermal. The shallow evolution occurred under a brittle regime and led to the superposition of hornblende to serpentine veining stages. The garnet pyroxenite-bearing mantle from the External Liguride ophiolites represents a rare tectonic sampling of deep levels of subcontinental lithosphere exhumed in an oceanic setting. The exhumation was probably accomplished through a two-step process that started during Late Palaeozoic continental extension. The low-pressure portion of the exhumation path, probably including also the plagioclase mylonitic shear zones, was related to the Mesozoic (Triassic to Jurassic) rifting that led to continental break-up. In Jurassic times, the studied mantle sequence became involved in an extensional detachment process that resulted in sea-floor denudation. KEY WORDS: garnet pyroxenite; ophiolite; non-volcanic margin; mantle exhumation; Sm–Nd and Lu–Hf geochronology
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INTRODUCTION |
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The architecture and tectonic evolution of non-volcanic continental margins have been extensively investigated in the last two decades in the western Iberian margin (e.g. Boillot et al., 1989
; Beslier et al., 1996; Whitmarsh et al., 2001). These studies have shown the existence of a characteristic zone of exhumed continental mantle, between continent and newly formed oceanic crust. The record of Ligurian Tethys ophiolites [here taking ‘ophiolite’ as a descriptive term to denote an association of peridotites with mid-ocean ridge (MOR)-type gabbros, basalt lavas and pelagic sediments] from the Alpine belt has provided fundamental clues about the nature of this continent–ocean transition (e.g. Bertotti et al., 1993; Manatschal & Bernoulli 1999; Desmurs et al., 2001; Manatschal et al., 2001; Müntener & Hermann, 2001).
Recent studies have proposed that extension leading to final continental break-up is accommodated through different fault systems (Whitmarsh et al., 2001; Manatschal, 2004) and showed that exhumation at the sea floor of the subcontinental mantle is a late process accomplished through relatively shallow downward concave faults. Petrological studies demonstrated the effects of the interaction between rising asthenospheric melts and lithospheric mantle rocks, and suggested that these processes played a significant role in modifying the lithospheric rheology, influencing the localization of continental rupture and the onset of sea-floor spreading (Müntener & Piccardo, 2003). However, geochronological constraints on the timing of mantle evolution during rifting are commonly lacking and little is known about the timing of high-temperature structures observed within the exhumed mantle rocks. The involvement of deep levels of subcontinental mantle during the rifting stages that precede the formation of an oceanic basin still remains an open issue.
The opening of the Ligurian Tethys in Middle Jurassic times led to the separation of the continental plates of Europe and Adria and gave rise to uplift and denudation of subcontinental lithospheric mantle, which is at present exposed in the Alpine chain. Some of these mantle bodies were affected by thermochemical erosion and refertilization processes related to asthenospheric upwelling during rifting and oceanization (Müntener & Piccardo, 2003), but rare occurrences of peridotites retaining their subcontinental origin are still preserved (e.g. Rampone et al., 1995; Müntener et al., 2004). New findings of garnet pyroxenite layers within ophiolitic mantle peridotites from the Northern Apennines provide an opportunity to throw light onto the evolution of subcontinental lithospheric mantle affected by extensional rift-related processes. This garnet pyroxenite-bearing mantle section is not contaminated by ascending asthenospheric melts and preserves relics of a deep-seated lithospheric history. Melting of garnet-bearing mafic rocks has been invoked to explain the ‘garnet signature’ of mid-ocean ridge basalt (MORB) (e.g. Hirschmann & Stolper, 1996) and of isotopically enriched MORB compositions (Salters & Dick, 2001). However, garnet pyroxenite layers are characteristic of mantle bodies exhumed in orogenic settings (e.g. Davies et al., 1993; Pearson et al., 1993). Garnet pyroxenites are absent in peridotite exposures on the present-day ocean floors and, to our knowledge, only two occurrences in ophiolitic mantle rocks have been reported (Peters, 1968; Müntener & Hermann, 1996; Blichert-Toft et al., 1999).
To unravel the exhumation history of the garnet pyroxenite-bearing mantle from the Northern Apennines, we combine field, petrological and geochronological investigations. In particular, Sm–Nd and Lu–Hf mineral isochrons for a garnet pyroxenite allow us to constrain the age of the mantle uplift. The determination of the pressure–temperature–time evolution of this mantle section suggests that only the low-pressure portion of the exhumation path, probably including also the development of plagioclase-facies shear zones, is related to rifting and continental break-up leading to the opening of the Ligurian Tethys. We argue that this has implications for the general understanding of the processes leading to the opening of oceanic basins along non-volcanic margins.
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GEOLOGICAL AND PETROLOGICAL FRAMEWORK |
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The Northern Apennine ophiolites are remnants of lithosphere from the Ligurian Tethys ocean, which opened in Middle Jurassic times (Lemoine et al., 1987
; Bill et al., 2001). They occur in two distinct ophiolite-bearing tectonic units, ascribed to different palaeogeographical domains (Fig. 1). The Internal Liguride (IL) ophiolites have lithostratigraphic characteristics that have been related to an intra-oceanic setting, i.e. as on-land analogues of oceanic lithosphere that originated at slow-spreading centres (Barrett & Spooner, 1977; Cortesogno et al., 1987; Tribuzio et al., 2000). In the External Liguride units (EL), ophiolites consist of fertile lherzolites, MOR-type basalts and rare gabbroic rocks occurring as large olistoliths within Cretaceous sedimentary mélanges, together with rare continental crust bodies locally displaying primary relationships with the ophiolites (Marroni et al., 1998). The gabbros have trace element and Nd–Sr isotope characteristics consistent with derivation from N-MORB magmas and their emplacement predated the oldest pelagic sediments of the Ligurian Tethys ophiolites (Tribuzio et al., 2004), thus testifying to the occurrence of a syn-rift MOR-type magmatic activity. The association of continental crust with the ophiolites was attributed to their formation in a fossil ocean–continent transition zone (Marroni et al., 1998), similar to that of modern non-volcanic continental margins. In particular, the External Liguride mantle rocks bear close petrological and geochemical similarities (Rampone & Piccardo, 2000; Tribuzio et al., 2004) to those from the western Iberia margin (e.g. Boillot et al., 1989; Whitmarsh et al., 2001).
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Detailed petrological and geochemical studies were carried out on mantle peridotite slide-blocks located in the southern sector of the External Liguride units (namely Suvero, Mt. Nero and Mt. Ragola massifs, Fig. 1). These peridotites are mainly spinel lherzolites with protogranular to tectonite fabrics, commonly displaying extensive growth of plagioclase (Piccardo, 1976
; Beccaluva et al., 1984; Piccardo et al., 1990). These lherzolites contain disseminated Ti-rich amphibole and sporadic spinel pyroxenite layers. Temperature estimates for the spinel-facies equilibration (1000–1050°C) are compatible with continental geothermal gradients (Rampone et al., 1995). The relatively fertile nature of the southern External Liguride peridotites is shown by high amounts of clinopyroxene (10–15 vol. %), high whole-rock Al2O3 contents and flat to slightly light rare earth element (LREE)-depleted clinopyroxene REE patterns (Rampone et al., 1995). Nd–Sr–Os isotope investigations indicate a Proterozoic age for the accretion of the Suvero mantle body to the subcontinental lithosphere (see also Snow et al., 2000). In addition, recent studies have indicated that the southern External Liguride peridotites were affected by refertilization by asthenospheric melts during the Ligurian Tethys formation (Piccardo et al., 2004a, 2004b).
In the northern sector of the External Ligurides (Fig. 1), ophiolites are scarce and almost exclusively represented by the ultramafic rocks that form the basis of this study (Bernini et al., 1997; Marroni et al., 2002). These mantle rocks form several bodies embedded in a narrow, NW–SE-trending belt of sedimentary mélanges of Late Cretaceous age that crop out along the margin of the Northern Apennine chain close to the Po Plain. The enclosing sediments do not show any significant metamorphic recrystallization related to the Alpine orogenic cycle. The mélange also includes rare MOR-type basalts (Montanini & Tribuzio, 2003) and fragments of continental crust (mafic granulites, amphibolites, granitoids). The ultramafic bodies are of variable size, ranging from only some tens of metres to a few kilometres (Fig. 1).
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FIELD RELATIONS |
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The mantle peridotites commonly show a mylonitic fabric characterized by relatively large (up to 1 cm across) pyroxene porphyroclasts. A thickness of several hundred metres can be conservatively estimated for the studied mylonitic shear zones. The peridotites contain pyroxenite layers concordant with the foliation of the host peridotite, ranging in thickness from a few millimetres to
2 m. The garnet clinopyroxenite layers are, in the central part of the thickest layers, locally isotropic and coarse-grained. In such cases, the external portions of the garnet clinopyroxenite layers display a weak foliation defined by elongated clinopyroxene porphyroclasts. The isotropic rocks are composed of pinkish Al-augite and embayed red–orange garnets with greenish pyroxene–spinel coronas; clusters of tiny graphite flakes occur in places. Garnet-free layers with evident foliation are also present. These layers are fine-grained and dark grey, and show parallel millimetre-scale pyroxene-rich and olivine-rich bands, concordant with the main peridotite foliation.
The most common layers are represented by boudinaged pinkish grey medium- to fine-grained websterites, of 2–10 cm thickness. Websterites may contain thin stretched lenses of olivine. Interlayering of centimetre-scale websterite and olivine-rich bands is also common.
Brittle deformation with multiple stages of veining, hydration and the formation of tectonic breccias is widespread and shows similar patterns in the different mantle bodies. Both the earliest veining events (F1 and F2) yielded orthogonal fracture systems subparallel and normal to the mylonitic foliation. The oldest veins (F1) are up to 2–3 mm thick and filled with dark green amphibole ± chlorite. F2 veins have variable thickness (millimetre- to centimetre-sized) and are filled with white fibrous or massive dark green serpentine (± opaque minerals). Anastomosing F2 veins are also locally present. In strongly foliated serpentinitized peridotites, there are locally millimetre- to centimetre-thick F2 veins of light green massive serpentine or densely spaced subparallel millimetre-sized veinlets of fibrous white serpentine, both subparallel to the mylonitic foliation.
A brittle event (F3) pervasively overprinted the former vein systems, yielding cohesive tectonic breccias (‘jigsaw breccia’). These cataclasites consist of millimetre- to metre-sized angular clasts in a fine-grained, grey–bluish, serpentine matrix with little displacement and no significant rotation of the clasts, as attested to by the coherence of the mylonitic fabric and pyroxenite layer orientation between adjacent clasts. Carbonates are absent in the early vein generations, but they may occur as tiny spherules in the matrix of the jigsaw breccia. Multiple calcite vein generations, commonly associated with Fe-sulphides and massive green serpentine and/or talc, may crosscut the jigsaw breccia and bound the clasts, yielding ophicalcite-like rocks (e.g. Lemoine et al., 1987; Treves & Harper, 1994).
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PETROGRAPHY AND MICROSTRUCTURES |
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Garnet clinopyroxenites
The garnet clinopyroxenites are commonly affected by a complex retrograde evolution. In particular, retrogression can be divided into two distinct stages characterized, respectively, by two-pyroxene + spinel + plagioclase and olivine + plagioclase assemblages.
The high-pressure protoliths were characterized by a coarse-grained isotropic texture and an anhydrous bimineralic assemblage of garnet + clinopyroxene (Cpx I). The oldest clinopyroxene generation (Cpx I) is preserved only as inclusions in garnet. Accessory graphite, Fe–Ni–Cu sulphides and rutile have been found locally. Graphite occurs as small flakes and stacks of flakes with grain size up to 2–3 mm. No diamond pseudomorphs have been observed. Raman spectra (Montanini et al., 2005) indicate a highly ordered structure similar to that of graphite crystallized at high temperature in peridotite and eclogite xenoliths (Pearson et al., 1993). Sulphides may occur both as tiny (20–50 µm) inclusions in garnet or, locally, as interstitial, lobate larger crystals (0·1–1·0 mm).
Complex pyroxene–spinel–plagioclase symplectites are present at the interface between garnet and Cpx I (Fig. 2a). Fine-grained (µm-scale) radial and fibrous symplectites composed of orthopyroxene (Opx I) + green spinel (Spl I) + plagioclase ± ilmenite (Fig. 3a) occur at the rims of the garnet relics. This inner symplectite is commonly rimmed by coarser symplectitic intergrowths of the same minerals and, locally, by an outer shell of plagioclase + clinopyroxene (Cpx IIa) with vermicular spinel inclusions.
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Coarse clinopyroxene grains (up to 4–5 mm; Cpx IIb) display plagioclase exsolution lamellae and occur as porphyroclasts in medium-grained granoblastic domains composed of clinopyroxene (Cpx III) + orthopyroxene (Opx II) + minor plagioclase and green spinel (Spl II; Fig. 2b). Several garnet clinopyroxenite layers are almost completely retrogressed to foliated rocks characterized by this latter assemblage, which locally preserves pyroxene–spinel symplectitic domains after garnet and clinopyroxene IIb porphyroclasts showing breakdown to secondary clinopyroxene + plagioclase. In these rocks, small (1–2 mm) clinopyroxene and orthopyroxene grains (corresponding to Cpx III and Opx II) are rimmed by fine-grained neoblastic Cpx (IV) + Opx (III) + plagioclase + spinel (Spl III). The garnet textural domain is recognized by orthopyroxene including vermicular grains of spinel (locally with a thin rind of secondary plagioclase) and minor clinopyroxene (Fig. 3b).
An olivine + plagioclase (± Spl IV) assemblage occurs as fine-grained aggregates replacing either the pyroxene–plagioclase–spinel symplectites after garnet, or the garnet relics within the symplectites. The association of olivine + plagioclase is also found as coronas and/or symplectites (Fig. 2c) around pyroxenes at the contact with spinel. In a few layers, the olivine + plagioclase assemblage syn-kinematically overgrows pyroxene + spinel (II) + plagioclase domains, locally associated with trace amounts of Spl (IV) and brown amphibole.
The garnet clinopyroxenites do not record an extensive low-temperature hydration process. Veining and replacement of the pyroxene + spinel + plagioclase assemblage by pale green amphibole + chlorite related to the F1 stage is locally observed. Serpentine F2 veins and serpentine ± chlorite pseudomorphs after olivine occur in places.
Websterites
The websterites are medium-grained rocks consisting of clinopyroxene (40–60%), orthopyroxene (20–30%), spinel (5–10%), plagioclase (0–10%), olivine (0–5%) and brown amphibole (0–3%); Fe–Ni sulphides are common accessory minerals. The most common texture is protomylonitic, characterized by highly stretched, elongated orthopyroxene crystals with high aspect ratios (up to 15; Fig. 2f) and fine-grained neoblastic aggregates of orthopyroxene + clinopyroxene + spinel ± amphibole around pyroxene porphyroclasts. Evidence for a previous garnet-bearing assemblage comes from rounded domains composed of Opx + Cpx + Spl symplectites (Fig. 2d). In one case, we have found tiny garnet relics within these symplectites. Coarsening and progressive annealing of the symplectites after garnet led to the formation of pyroxene–spinel clusters and, locally, to aggregates with allotriomorphic texture (Fig. 2e and f).
The pyroxene + spinel assemblage is locally mantled by fine-grained olivine–plagioclase-rich mylonitic bands (Fig. 2e and f), similar to those occurring in the peridotite (see the next section). Strain localization associated with this deformation is commonly observed at the contact with peridotite. The clinopyroxene porphyroclasts frequently show plagioclase exsolution and rims. Green to greenish brown spinel occurs both as small rounded or vermicular grains within orthopyroxene and as larger allotriomorphic crystals, commonly rimmed by plagioclase. Olivine of probably primary origin occurs as irregularly shaped crystals or thin seams surrounded by pyroxenes similar to the ‘olivine flames’ of Burg et al. (1998). Secondary olivine forms: (1) radial crystals, associated with plagioclase + greenish brown spinel ± ilmenite ± brown amphibole, forming coronas between pyroxenes and green spinel, locally stretched along the main foliation; (2) granoblastic domains composed by the same phases; (3) fine-grained plagioclase-bearing (± brown amphibole) mylonitic bands.
Peridotites
The peridotites are clinopyroxene-rich (10–15 vol. %) spinel–plagioclase lherzolites with accessory brown amphibole. The oldest recognizable texture is preserved in small domains characterized by coarse-grained (up to 5–10 mm across) orthopyroxene and clinopyroxene grains surrounded by smaller polygonal or irregular-shaped neoblasts of olivine + pyroxene + spinel ± brown amphibole (Fig. 2g). These domains thus represent relics of a spinel-facies low-strain tectonite. Mutual pyroxene exsolution may occur in these porphyroclasts.
The spinel-facies assemblage is overprinted by plagioclase-facies recrystallization associated with the development of a mylonitic fabric. The mylonite microstructure is characterized by aligned porphyroclasts of pyroxene + brown spinel + brown amphibole and by an ultrafine-grained (20–50 µm) polyphase aggregate composed of olivine + pyroxenes + plagioclase (± spinel ± accessory brown amphibole) occurring as millimetre-sized bands, lenses and porphyroclast tails (Fig. 2h). Crystals in this polyphase matrix commonly display grain boundary alignment parallel to the mylonitic foliation.
Orthopyroxene porphyroclasts may be rounded, lozenge-shaped or elongate along the mylonitic foliation with high aspect ratio (up to 10). Clinopyroxene is usually more equant than orthopyroxene. Both pyroxenes commonly show evidence for intracrystalline deformation (i.e. undulose extinction, bending and kinking). Plagioclase occurs as (1) thin rims between spinel and pyroxene porphyroclasts, (2) tiny neoblastic grains in the mylonitic matrix and, rarely, (3) exsolution in clinopyroxene porphyroclasts.
The mylonitic foliation is locally crosscut by thin fractures filled with prismatic pale green amphibole, which is commonly associated with minor amounts of chlorite (F1). Amphibole and chlorite also grow adjacent to the fractures, as pseudomorphs after clinopyroxene and brown amphibole, and as coronas around spinel, respectively (Fig. 4a).
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The association of amphibole + chlorite is followed by multiple serpentine veining episodes (F2), associated with serpentinization, to various extents, of the host rock. X-ray diffraction analyses have shown that lizardite is the dominant phase in the serpentinized peridotite. Vein serpentine may be fibrous with antiaxial filling or isotropic, rarely associated with magnetite. Olivine alteration is mainly accomplished through growth of ‘mesh texture’ serpentine (with frequent preservation of olivine cores) and, to a lesser extent, of ‘hourglass’ serpentine. Magnetite abundance in olivine pseudomorphs usually increases with increasing degree of serpentinization. Pyroxenes may be pseudomorphed by light green or bluish fibrous serpentine (‘bastite’). Relics of the former mantle assemblages, mainly represented by sparse brown spinel grains with opaque rims, are only rarely preserved in the serpentinized plagioclase mylonites. Two main types of subparallel veins follow the original plagioclase-bearing foliation: (1) thin ribbon lizardite veins (± magnetite seams) transected by (2) millimetre- to centimetre-thick massive lizardite veins with low birefringence and undulose extinction.
All the structures described above are pervasively disrupted by the F3 stage of brecciation. These breccias show a jigsaw-puzzle pattern at both meso- and micro-scales, and show evidence for polyphase cracking and veining. The breccia matrix (Fig. 4b) is made of fine-grained serpentinite derived from comminution of the larger clasts and by yellow–light brown serpentine in tiny fibrous–spherulitic aggregates. The serpentine matrix also contains clusters of micrometre-scale, curved needles of Ni-sulphide (millerite), and locally abundant calcite spherules, studded with dark fluid inclusions, some tenths of a micrometre across.
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ANALYTICAL METHODS |
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Major element mineral analyses were obtained on samples representative of the different lithologies and different stages of retrograde evolution. Mineral analyses were performed at the Dipartimento di Scienze della Terra of Università di Parma using a JEOL-6400 electron microprobe equipped with LINK-ISIS energy-dispersive microanalytic system. Operating conditions were accelerating voltage of 15 kV and probe current of 0·25 nA; both natural minerals and synthetic compounds were used as standards. Trace amounts of Na2O in garnet were analysed at CNR–Istituto di Geoscienze e Georisorse, Sezione di Firenze, using a JEOL JXA-8600 microprobe. Operating conditions were accelerating voltage of 15 kV, probe current of 100 nA and counting times of 20 s. These conditions provided a detection limit of
40 ppm for Na.
Sm–Nd and Lu–Hf dating techniques were applied to a garnet clinopyroxenite layer (sample AM322). The rock was crushed and sieved to obtain a 150–250 µm fraction. Mineral concentrates were extracted using a Frantz magnetic separator and then purified by careful handpicking under a binocular microscope to obtain clinopyroxene, garnet and plagioclase separates, avoiding inclusion-bearing grains or small impurities derived from symplectitic intergrowths. Sm–Nd and Lu–Hf analyses of the garnet clinopyroxenite layer were carried out on the same mineral and whole-rock splits. Details of sample digestion and ion exchange chromatography were presented by Anczkiewicz et al. (2004). Mass spectrometry procedures follow Thirlwall & Anczkiewicz (2004). Total procedure analytical blanks for Hf and Nd were <30 pg. All elements were analysed in a static, hard extraction mode using the Royal Holloway IsoProbe®. All errors are 2SE and relate to the last significant digits, except for age errors, which are 95% confidence level. All measurements were conducted on a single day to minimize correction for secular variation in the static 176Hf/177Hf of JMC47. 147Sm/144Nd errors are 0·3%. Reproducibility of the Aldrich Nd standard on the day of analyses was 143Nd/144Nd = 0·511388 ± 11 (2SD, n = 4).Daily variations in 143Nd/144Nd ratios were normalized to 143Nd/144Nd = 0·511421. 176Lu/177Hf errors are 0·5%; the JMC475 standard on the day of analysis yielded 0·282173 ± 6 (2SD, n = 3). Daily variations in 176Hf/177Hf ratios were normalized to 176Hf/177Hf = 0·282165. Mass bias corrections were made using 146Nd/144Nd = 0·7219 and 179Hf/177Hf = 0·7325.
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MINERAL CHEMISTRY |
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Garnet pyroxenites
Garnet compositions cover the range Prp53–45Alm29–38Grs11–20, with only minor spessartine component (
1·5 mol %, Table 1, Fig. 5). Cr2O3 contents do not exceed 0·25 wt %. The low TiO2 (0·10–0·20 wt %) and Na2O concentrations (0·02 wt %, confirmed by unpublished LA-ICP-MS analyses that yielded Na 120 ppm) are comparable with those of the diamond-free Group II eclogites defined by McCandless & Gurney (1989). The bulk composition of the symplectitic intergrowths in retrogressed pyroxenites was determined by electron microprobe and is a pyrope-rich garnet composition (Table 1). However, the analyses contain an excess of MgO relative to garnet stoichiometry and high amounts of Na2O, which may be related to involvement of olivine and primary clinopyroxene in the garnet breakdown reaction, respectively.
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The oldest clinopyroxene (Cpx I) has up to 12·6 wt % Al2O3 and 2·7 wt % Na2O. Recalculation of structural formulae gives high proportions of the Ca-Tschermak component (15–20 mol %) and moderate jadeite contents (10–16 mol %). The structural formulae do not display any cation deficiency requiring the presence of a Ca-Eskola component. Similar clinopyroxene compositions were reported for Group II garnet pyroxenites from Beni Bousera (Kornprobst et al., 1990
). The different generations of retrograde clinopyroxene have lower amounts of the Ca-Tschermak and jadeitic component than Cpx I and define a negative correlation in a plot of Na2O–Al2O3 (Fig. 6a). The latest pyroxenes (Cpx IV) have the lowest amounts of the Ca-Tschermak and jadeitic component. The composition of the coarse clinopyroxene porphyroclasts characterized by plagioclase exsolution (Cpx IIb) is highly variable, both for Al2O3 and Na2O contents. In particular, Na2O concentrations vary over a wide range (1·3–0·1 wt %), even within the same crystal, with the lowest values occurring close to the plagioclase lamellae and patches.
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The Mg/(Mg + Fe2+) ratios of clinopyroxene and orthopyroxene from the garnet clinopyroxenites span a wide range (0·73–0·90 and 0·71–0·88, respectively). Cr2O3 varies in the range 0·05–0·30 wt % in both pyroxenes. The lowest Mg/(Mg + Fe2+) values are associated with the lowest Cr contents in both pyroxenes. The forsterite content of olivine (61–85 mol %) exhibits a rough positive correlation with the Mg/(Mg + Fe2+) values of the pyroxenes.
Spinel does not show appreciable compositional variation between the different textural domains. It is highly aluminous and Cr-poor (Cr2O3 1·5 wt %); the Mg/(Mg + Fe2+) values vary between 50 and 75 (Table 1). The brown amphibole occurring in the olivine–plagioclase-bearing domains is kaersutite (Table 1).
The composition of plagioclase (Table 1) is controlled by textural position. Plagioclase from the Opx–Spl–Pl symplectites around garnet (Pl Ia) has high Ca contents (An = 68–85 mol %). Plagioclase lamellae within clinopyroxene (Pl Ib) and plagioclase grains in recrystallized two-pyroxene domains (Pl II) have 45–50 mol % anorthite.
Vein amphibole and amphibole overgrowths on clinopyroxene in the pyroxenitic layers is Mg-hornblende with high Al2O3 (up to 10·6 wt %) and Na2O (1·6–2·2 wt %) contents, and low TiO2 concentrations (0·13–0·25 wt %); significant amounts of Cl (0·15 wt %; see Table 4) were detected. The associated chlorite is clinochlore, with Al2O3 and Mg/(Mg + Fe2+) values of about 19 wt % and 0·88, respectively.
Websterites
Relict garnet has high pyrope (78 mol %, Fig. 5) and Cr2O3 contents (0·7 wt %), similar to the garnets from mantle lherzolites worldwide (Pearson & Nixon, 1996). Clinopyroxene porphyroclasts can have high amounts of Al2O3 (10–11 wt %) and Ca-Tschermak substitution (up to 18 mol %), whereas the jadeite component is low (6 mol %), probably as the result of exsolution of Na-rich plagioclase (An = 58 mol %; Fig. 6b). These clinopyroxenes are compositionally similar to Cpx IIb from the garnet clinopyroxenite, although they have higher Mg-number [Mg/(Mg + Fe2+)] and Cr2O3 concentrations (Table 2). Neoblastic clinopyroxenes, especially those of the fine-grained plagioclase-rich mylonitic bands, have low Al2O3 and Na2O (Table 2, Fig. 6b). The Mg-numbers of both clinopyroxenes and orthopyroxenes (0·82–0·91 and 0·85–0·89, respectively) partially overlap those of the peridotites. Cr2O3 contents of clinopyroxenes range from 0·20 to 0·60 wt %. Amphibole is kaersutite with low K2O (0·3 wt %) and moderate Cr2O3 (1·0–1·3 wt %). Spinel has Cr-number [Cr/(Cr + Al)] and Mg/(Mg + Fe2+) of 0·02–12 and 0·72–0·78, respectively. The green spinel in the symplectites and in the two-pyroxene assemblage has the lowest Cr-number. The highest Cr-numbers are observed in the greenish brown spinel from the olivine-bearing retrograde domains (Fig. 7). Primary and secondary olivines have high forsterite contents (Fo = 89–90 mol %).
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Peridotites
The peridotitic olivine is Fo89. Porphyroclastic spinel-facies clinopyroxene has high Al2O3, Na2O and TiO2 contents (6·7–8·6 wt %, 1·2–2·2 wt % and 0·6–1·0 wt %, respectively; see also Fig. 8a). Its Mg-number ranges from 0·89 to 0·92 and the Cr2O3 concentrations are 0·5–1·1 wt % (Fig. 8a, Table 3). The amounts of Al2O3 and Na2O in porphyroclastic clinopyroxene decrease moderately from core to rim. The lowest Al2O3 and Na2O contents (1·4–5·0 wt % and
0·6 wt %, respectively) are shown by neoblastic plagioclase-facies clinopyroxene. Cr2O3 shows no significant variations between porphyroclastic and neoblastic clinopyroxenes (Fig. 8b).
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Orthopyroxenes have Mg-numbers of 0·89–0·91. The cores of orthopyroxene porphyroclasts (Table 3) are Al2O3-rich (4·7–5·7 wt %) and contain variable amounts of CaO (0·6–1·4 wt %). The CaO and Al2O3 contents are lower in porphyroclast rims (0·5–0·7 wt % and 3·5–4·4 wt %, respectively). Orthopyroxene neoblasts have the lowest concentrations of Al2O3 and CaO (1·3–2·7 wt % and
0·5 wt %, respectively). As a whole, the orthopyroxene analyses show a rough negative correlation between Al2O3 and CaO (Fig. 9).
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Spinel shows (Fig. 7) variable Mg-number and Cr-number. This variability has been also found in the same sample and is probably related to the recrystallization under plagioclase-facies conditions. The lowest Cr-numbers are found in the cores of the rare porphyroclasts not rimmed by plagioclase, whereas increasing values occur towards the plagioclase coronas or in small relict crystals in plagioclase-bearing domains. TiO2 contents in spinel are invariably lower than 0·25 wt %.
The brown amphibole porphyroclasts belonging to the spinel-facies assemblage range in composition from kaersutite to Ti-pargasite, according to the nomenclature of Leake et al. (1997). They have variable TiO2 contents (3·4–5·3 wt %), even at the thin-section scale (Table 3, Fig. 10), and low to moderate K2O concentrations (0·05–0·30 wt %).
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Plagioclase is invariably anorthite-poor. The highest anorthite contents (48–56 mol %) were found in plagioclase coronas around spinel. The neoblastic grains in the mylonitic matrix have about 45 mol % anorthite and the plagioclase exsolution in the clinopyroxene porphyroclasts is albite-rich (An35).
Pale green amphibole from the veins is Mg-hornblende, which spans a wide range of Al2O3 (4·3–10·4 wt %) and Na2O (0·9–2·1 wt %) contents (Fig. 10). Vein amphibole has TiO2 concentrations between 0·14 and 0·75 wt %, and small amounts of Cl (0·18 wt %). Small prismatic tremolite crystals occur locally in the veins, probably as a later Mg-hornblende replacement (Table 4, Fig. 10). The colourless chlorite forming coronas around spinel (Table 4) is penninite.
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The composition of serpentine (Table 4) varies in relation to the microstructural site of nucleation. Serpentine replacing olivine has low Al2O3 and Cr2O3 (<0·1 wt %), and relatively high NiO (0·15–0·38 wt %). On the other hand, fibrous serpentine after orthopyroxene has high Al2O3 (3·8–5·0 wt %) and Cr2O3 (0·17–0·44 wt %) contents, similar to those observed in the primary mineral, thus suggesting a relative immobility of these elements during serpentinization (Bonatti & Hamlyn, 1981
; Agrinier et al., 1988; Mével, 2003). FeO contents are variable in both types of serpentine pseudomorphs (4·9–7·8 wt %). The different veins belonging to the F2 stage have low Al2O3 (1·2–2·8 wt %) and highly variable FeO contents (1·9–8·8 wt %). The lowest FeO values were observed for the serpentine with undulose extinction, which is probably older than the Fe-rich veins. Serpentine from the matrix of the F3 jigsaw breccia has low to moderate Al2O3 (1·7–3·8 wt %), coupled with relatively high FeO. Non-pseudomorphic serpentine from both F2 and F3 stages commonly displays negligible Cr and Ni. Chlorine contents of the different serpentine types vary over a relatively wide range from 0·05 to 0·31 wt %, without a systematic relation with the nucleation site. |
GEOTHERMOBAROMETRY |
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Conventional thermobarometry methods have been applied to minerals occurring in domains formed at different stages and distinguished on the basis of textural relations. Different cation exchange thermometers have been applied to both pyroxenites and enclosing peridotitic rocks to assess their thermal evolution, whereas pressure estimates have been obtained only for the garnet-bearing rocks (Table 5). To avoid the uncertainties related to Fe3+ recalculations from microprobe analyses, all iron in pyroxenes and garnet is assumed to be Fe2+. The computed P–T estimates should, therefore, be considered as maximum values.
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Pyroxenites
The occurrence of clinopyroxene inclusions (Cpx I) within garnets allows us to calculate the equilibration pressure and temperature of the early garnet clinopyroxenite assemblage. We have used the geobarometric method proposed by Simakov & Taylor (2000)
, based on the solubility of the Ca-Tschermaks component in clinopyroxene coexisting with garnet, which, recently, has been successfully applied to both diamond-bearing and diamond-free eclogite xenoliths (Peltonen et al., 2002; Simakov, 2006). Temperature calculations are based on the application of the garnet–clinopyroxene Fe–Mg exchange reaction, according to the calibrations of Ellis & Green (1979), Krogh (1988), Ai (1994) and Krogh Ravna (2000). Pressures and temperatures were calculated iteratively using the clinopyroxene inclusions characterized by the highest Na2O and Al2O3 contents (sample AM322, Table 5). Using the most recent calibration of the garnet–clinopyroxene geothermometer (Krogh Ravna, 2000), we obtained T 1100°C and P 2·8 GPa. The calibrations of Ellis & Green (1979), Krogh (1988) and Ai (1994) yield slightly higher temperature (and pressure) estimates. The high Al2O3 contents of Cpx I are consistent with the geothermobarometric evaluations. Experimental work has shown that clinopyroxenes with high proportions of Ca-Tschermaks molecule form at high-pressure conditions (2·5–3·5 GPa) as segregations from basic melts (Thompson, 1974; Eggins, 1992; Johnson, 1998) and as residual phases after anhydrous partial melting of mafic rocks (Pertermann & Hirschmann, 2003; Yaxley & Brey, 2003). Jadeite contents in these magmatic or residual clinopyroxenes are typically low and comparable with those observed in the present study.
Pyroxene thermometry has been applied to the retrograde pyroxene–spinel-bearing assemblage of the garnet clinopyroxenites and websterites. The two-pyroxene geothermometer of Wells (1977) applied to small, undeformed clinopyroxene-orthopyroxene pairs (Cpx III–Opx II) of the garnet clinopyroxenites yields mean temperatures of 960°C (Table 5). In addition, the neoblastic Cpx IV–Opx III pairs from a garnet clinopyroxenite layer characterized by dynamic recrystallization of pyroxenes + spinel gave T = 920°C. Similar temperature estimates were obtained for the clinopyroxene–orthopyroxene porphyroclasts and neoblasts from the websterites. We have also applied the Ca-in-Opx geothermometer of Brey & Kohler (1990) to the olivine-bearing websterites, which has yielded temperature estimates of 950°C for the orthopyroxene porphyroclasts.
The coexistence of Ti-rich amphibole and plagioclase in the olivine-bearing domains of the retrogressed clinopyroxenites and websterites permits temperature to be estimated on the basis of the Holland & Blundy (1994) method. Temperature values of 960°C have been obtained for both rock types.
Peridotites
Temperatures of the spinel- and plagioclase-facies equilibration in the peridotites have been estimated (Table 5) applying the Ca-in-Opx geothermometer (Brey & Kohler, 1990) and the Ca–Mg exchange geothermometers of Wells (1977) and Brey & Kohler (1990). The cores of spinel-facies orthopyroxene porphyroclasts yield TCa-in-Opx estimates mainly ranging between 1000 and 1100°C, assuming P = 1·5 GPa. Core to rim compositional variations in the orthopyroxene porphyroclasts reflect cooling of 100–150°C. Two-pyroxene porphyroclast temperatures vary over a wide range, shifted to lower values (870–1040°C) than TCa-in-Opx estimates based on the core compositions of orthopyroxene porphyroclasts. The lower estimates resulting from two-pyroxene geothermometers may be a consequence of pyroxene exsolution during slow cooling. The relatively low diffusivity of Ca in orthopyroxene compared with the Ca–Mg diffusivity in clinopyroxene (see Lenoir et al., 2001) may have hampered complete re-equilibration of large orthopyroxene porphyroclasts during cooling. Temperature estimates for the plagioclase-facies equilibration range from 860 to 1030°C, on the basis of Cpx–Opx pairs.
Pressure evolution of the peridotites is difficult to constrain owing to the lack of appropriate geobarometers in garnet-absent systems. The pressure of the garnet–spinel transition in peridotitic systems is mainly controlled by the Cr content of the spinel (O'Neill, 1981; Webb & Wood, 1986; Klemme, 2004). In fertile compositions, characterized by spinel with Cr/(Cr + Al) <0·2, garnet is stabilized towards lower pressure values. A maximum pressure estimate for the spinel-facies equilibration can be determined on the basis of the formulation of Webb & Wood (1986). For this purpose we used the composition of spinel porphyroclast cores not rimmed by plagioclase and obtained pressure values of around 1·8 GPa. This pressure value is consistent with the P–X phase relations recently considered by Klemme (2004) for a simplified system MgO–Cr2O3–Al2O3–SiO2.
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Sm–Nd AND Lu–Hf GEOCHRONOLOGY |
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The summary of the isotopic results is presented in Table 6 and Fig. 11. Age calculations were performed using Isoplot® (Ludwig, 2003
); age errors are given at the 95% confidence level.
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A Sm–Nd isochron age is defined by whole-rock, plagioclase, clinopyroxene and two garnet fractions. The small scatter (MSWD = 0·26) together with the high 147Sm/144Nd values measured in garnets results in a well-constrained age of 186·0 ± 1·8 Ma (initial
Nd = +6·5). Remarkably, both garnet fractions resulted in identical (within error) isotopic ratios. The unusually high 147Sm/144Nd ratio of plagioclase is probably due to its metamorphic origin as a product of garnet + Cpx I breakdown, in a system characterized by an extreme LREE depletion (see also Tribuzio et al., 2005). Although not all analysed minerals grew in equilibrium (e.g. garnet and plagioclase), the geothermometric investigations (Table 5) show that all analysed minerals crystallized significantly above 750–800°C, the closure temperature for the Sm–Nd system in garnet (Ganguly et al., 1988; Van Orman et al., 2002). Hence, the Sm–Nd isochron is interpreted as a cooling age, i.e. the garnet pyroxenite probably passed through the Sm–Nd blocking temperature at 186·0 ± 1·8 Ma.
Lu–Hf dating of the same whole-rock and mineral fractions (excluding plagioclase) shows only 0·06 spread in 176Lu/177Hf ratios, which does not give an age with high precision. The regression line with MSWD = 0·23 defines an age of 220 ± 13 Ma (initial Hf = +4·5). Interpretation of the Lu–Hf age is more problematic, as little is known about the closure temperature for this system. It has been proposed that the closure temperature for Lu–Hf in garnet is equal to, or slightly higher (100°C) than for Sm–Nd (Scherer et al., 2001). Although few coupled Lu–Hf and Sm–Nd age determinations have been carried out on mantle garnet-bearing rocks, the available data corroborate this interpretation. Lu–Hf cooling ages older than Sm–Nd ages have been reported for garnet pyroxenite layers from Beni-Bousera (Blichert-Toft et al., 1999; Pearson & Nowell, 2004) and for a garnet pyroxenite sample in a Mirdita (Albania) ophiolite (Blichert-Toft et al., 1999), in agreement with this notion. In particular, the large difference between the Lu–Hf and Sm–Nd mineral isochron ages (203 ± 3·4 Ma and 166 ± 2 Ma, respectively) obtained for the Mirdita ophiolite sample is similar to that found for the External Liguride pyroxenite in this study. In addition, recent work on eclogite xenoliths in kimberlites (Bedini et al., 2004) yielded Lu–Hf clinopyroxene–garnet ages generally older than the corresponding Sm–Nd ages. This suggests that radiogenic Hf diffuses out of garnet more slowly than radiogenic Nd, giving further support to the notion that Tc(Lu–Hf) 
Tc(Sm–Nd). Because the estimated temperature for garnet equilibration in the studied rock is 1100°C (i.e. above any reasonable closure temperature for the Lu–Hf system; see also Bedini et al., 2004), it seems plausible to interpret the Lu–Hf age in terms of cooling.
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DISCUSSION |
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Decompression-related evolution of pyroxenite layers
Textural observations indicate that the retrograde transformations within the garnet clinopyroxenites started with the breakdown of the garnet + Al-rich clinopyroxene assemblage to a garnet-free two-pyroxene + spinel + plagioclase assemblage. The bulk mineralogical composition (Opx + Spl + Pl + Cpx) of the coronas around garnet in the garnet clinopyroxenites indicates that breakdown of garnet and clinopyroxene was coupled with element diffusion across the boundary between the two phases, suggesting a general reaction
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) and peridotites (Godard & Martin, 2000) exhumed from great depths. The breakdown of primary clinopyroxene to a secondary clinopyroxene (Cpx IIb) with lower Ca-Tschermaks and jadeite contents + plagioclase (Pl Ib) is commonly observed for clinopyroxenes from retrogressed eclogites (e.g. Boland & Van Roermund, 1983; Joanny et al., 1991).
In the Mg-rich garnet clinopyroxenites and websterites, the breakdown products of garnet include spinel and secondary orthopyroxene and clinopyroxene, in the absence of plagioclase. The relatively high Na contents of both secondary clinopyroxene and bulk symplectite (Tables 1 and 2) indicate that the symplectites are not simply the result of garnet decomposition, requiring the involvement of an earlier Na-bearing phase in the reaction, probably a primary clinopyroxene that has not been preserved. In addition, the bulk symplectite compositions have excess Mg, suggesting a contribution from primary olivine (Morishita & Arai, 2003) or orthopyroxene. The following reaction is therefore proposed:
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; Gasparik, 1984). This is also consistent with the remarkable grain-size decrease towards the inner portion of the coronas between garnet and primary clinopyroxene (e.g. Messiga & Bettini, 1990). Geothermometry results for the pyroxenites actually show that the evolution from garnet–to pyroxene–spinel assemblage involved significant cooling (Table 5).
The garnet stability field for mafic to ultramafic systems is known to expand with decreasing whole-rock Mg and Cr, and increasing Al (O'Neill, 1981; Nickel, 1986; O'Hara et al., 1971; Klemme, 2004), so that garnet may be stable at lower pressure in pyroxenite than in peridotite. The pressure of garnet breakdown can be roughly constrained considering the experimental curves obtained for mafic systems (e.g. Irving, 1974; Adam et al., 1992). The pressure of garnet (+ Cpx I ± Ol ± Opx I) breakdown to Cpx II + Opx II + Spl I ± Pl in garnet clinopyroxenites and websterites approximately lies in the range 1·2–1·5 GPa, assuming T = 1000°C. This decompression stage was accompanied by significant cooling (Table 5).
A later decompression stage is recorded by the local formation of olivine + plagioclase after two-pyroxenes + spinel in different microstructural sites. Similar decompression-related subsolidus formation of olivine + plagioclase coronas has been reported for pyroxenite layers within other peridotite massifs (e.g. Van der Wal & Vissers, 1996) and within the peridotites drilled along the western Iberian margin (Cornen et al., 1996). The increase of Cr-number commonly observed in the spinel from the olivine-bearing domains (in the websterites) may be accounted for by the breakdown of relatively Cr-rich clinopyroxene according to the reaction Cpx + Opx + Al-Spl = Ol + Pl + Cr-Spl. The olivine + plagioclase formation is commonly associated with the involvement of the pyroxenite layers in the mylonitic deformation that affected the host peridotite under plagioclase-facies conditions. In garnet clinopyroxenites, in addition, olivine locally developed as a late breakdown product of garnet relics, together with plagioclase and spinel. The reaction Grt = Ol + Pl + Spl is not frequently observed in garnet pyroxenites or eclogites, but has been reported as a LP–HT retrograde stage in retrogressed eclogites from the Bohemian Massif (O'Brien, 1997). The conversion of the pyroxene + spinel to olivine + plagioclase may be approximately constrained at a pressure of 0·7–0·8 GPa (Green & Ringwood, 1967; Irving, 1974; Adam et al., 1992).
The temperatures calculated for the two-pyroxene porphyroclastic and neoblastic garnet- and olivine-free assemblages of both garnet clinopyroxenites and garnet websterites overlap significantly and are not distinguishable from those of the olivine-bearing stage inferred from amphibole in equilibrium with plagioclase (Table 5). This indicates that the late stage of the decompression path was nearly isothermal.
Decompression-related evolution of the peridotites and origin of the plagioclase
No mineralogical evidence has been preserved for an early garnet-bearing assemblage in the peridotites of the present study. However, the pressure estimate of 2·8 GPa for the garnet clinopyroxenite assemblage [even considering the large uncertainty reported for the Simakov & Taylor (2000) geobarometer], suggests that the investigated mantle section underwent a previous stage of equilibration under garnet-facies conditions. The relatively low pressure required for the spinel–garnet transition in a peridotitic system characterized by Cr-poor spinel further supports this view. The early occurrence of garnet was also hypothesized by Rampone et al. (1995) for the southernmost External Liguride peridotite bodies on the basis of clinopyroxene trace element compositions. In addition, the presence of rounded orthopyroxene–spinel clusters indicating a former garnet-bearing precursor has been reported for other subcontinental peridotite bodies of the Alpine–Apennine system exhumed during the same geodynamic event, i.e. Erro–Tobbio (Ligurian Alps, Hoogerduijn Strating et al., 1993) and Lanzo (Western Alps, Piccardo et al., 2004a). In the External Liguride peridotites, therefore, the original garnet-bearing assemblage was probably erased by re-equilibration under spinel-facies conditions.
The spinel peridotite assemblage in the studied peridotites is preserved in relict domains within the plagioclase mylonite. These domains display a porphyroclastic texture testifying to the development of a former spinel-facies tectonite. The TCa-in-Opx registered by the cores of some of the orthopyroxene porphyroclasts, consistent with those obtained for the garnet–pyroxene assemblage of the pyroxenite layers, may preserve a record of the former garnet-facies equilibration. On the other hand, two-pyroxene and plagioclase–amphibole geothermometry indicates that the spinel- and plagioclase-facies recrystallizations occurred under similar temperature ranges; that is, the related decompression did not involve significant cooling.
The thermal history of the investigated mantle section, the fertile nature of the peridotites and the preservation of basic pyroxenite layers, which are characterized by lower melting temperatures than the host peridotite (Hirschmann & Stolper, 1996), rule out partial melting as an origin for the plagioclase. Plagioclase growth in the External Liguride peridotites has been ascribed to subsolidus decompression in Middle to Late Jurassic times (Rampone et al., 1993, 1995). Recent studies, however, have provided new textural and geochemical evidence indicating that some of the External Liguride peridotites experienced extensive melt–rock reaction, impregnation and heating by ascending asthenospheric melts during the opening of the Ligurian Tethys (Müntener & Piccardo, 2003; Müntener et al., 2004). However, the peridotites of the present study lack typical reaction or impregnation microtextures, even in the portions unaffected by the plagioclase-facies mylonitic deformation. In addition, chemical features indicating infiltration and reaction with basaltic melts (e.g. Ti enrichment of spinel and/or clinopyroxene) are lacking.
The studied peridotites have remarkably similar parageneses (i.e. occurrence of Na–Al-rich clinopyroxene porphyroclasts coexisting with Cr–Ti-poor spinels) to other Ligurian Tethys peridotites interpreted as old lithospheric mantle domains that had escaped impregnation processes (Rampone et al., 1995; Müntener et al., 2004; Piccardo et al., 2004a). In addition, there is no geothermometric indication of local heating to near-asthenospheric conditions, as recorded, for example, in other mantle bodies from the External Liguride domain (Piccardo et al., 2004a), which could result from melt infiltration. We conclude that plagioclase formed at subsolidus conditions during exhumation of this mantle section. Such a process is also recorded by clinopyroxene compositions, which show a decrease of Al2O3 and Na2O from porphyroclast cores to rims, and to neoblasts (Fig. 8a; see also Rampone et al., 1993, 1995). The similar Al2O3 decrease between orthopyroxene porphyroclasts and neoblasts (Fig. 9) is consistent with the inferred process of plagioclase-facies re-equilibration. The lack of a concomitant Cr2O3 increase in clinopyroxene and orthopyroxene, as expected in spinel-free recrystallized assemblages (Rampone et al., 1993), can probably be accounted for by Cr redistribution in the relict spinel, which is significantly enriched in Cr with respect to the porphyroclastic spinel (Table 3).
Maximum pressure estimates for the plagioclase-facies equilibration have been obtained using the reaction orthopyroxene + clinopyroxene + spinel = olivine + plagioclase. The intersection between the temperature calculated for the neoblastic Cpx–Opx pairs in the plagioclase-bearing assemblage and the curve obtained by interpolating the curves of Newman et al. (1998) for the albite-rich plagioclase compositions of the analysed peridotites constrain the pressure of the spinel to plagioclase transition at 0·8–0·9 GPa.
The fine-grained polyphase matrix of the plagioclase mylonites, characterized by grain boundary alignment parallel to the mylonitic foliation, is analogous to that described in other anhydrous plagioclase mylonites, where grain-size sensitive creep (including diffusion creep and grain boundary sliding) acts as the dominant deformation mechanism (e.g. Newman et al., 1998; Furusho & Kanagawa, 1999; Dijkstra et al., 2002). The plagioclase mylonite assemblage locally includes trace amounts of Ti-rich amphibole, suggesting that deformation took place in an H2O-deficient environment. Notably, an external fluid input is not required at this stage, as the small amphibole amounts found in the plagioclase domains may form at the expense of the Ti-rich amphibole occurring in the spinel-bearing assemblage. The formation of 100-m-scale shear zones made up of plagioclase mylonites is consistent with the concept that subsolidus plagioclase formation under decompression, possibly associated with grain-size reduction, may lead to reduction of the lithospheric mantle strength, promoting strain localization and development of large shear zones (‘reaction-enhanced ductility’; see Newman et al., 1998; Furusho & Kanagawa, 1999; Handy & Stünitz, 2002; de Ronde et al., 2004).
P–T conditions of retrograde hydration stages
The anhydrous high-temperature ductile evolution is overprinted by episodes of deformation (F1 to F3 stages) occurring in a brittle regime. In particular, the formation of Mg-hornblende ± chlorite in veins and as pseudomorphs after clinopyoxene and brown amphibole in both pyroxenites and peridotites marks the onset of hydration of the studied mantle section. Unfortunately, no clear textural evidence of an equilibrium assemblage (e.g. coexistence of hornblende + chlorite with olivine and a spinel phase in peridotites) is preserved, thus precluding the evaluation of accurate P–T conditions for the onset of hydration. Nevertheless, the early formation of an aluminous amphibole instead of tremolite indicates amphibolite-facies conditions (Evans, 1982), thus yielding a conservative minimum temperature of 500°C for the early hydration event. An upper temperature limit is provided by chlorite stability in ultramafic systems at low-pressure conditions (700°C, Schmidt & Poli, 1998; Ulmer & Trommsdorff, 1999).
Pressure estimates for the amphibolite-facies event are provided by the absence of epidote in the Mg-hornblende + chlorite assemblage within mafic pyroxenite layers, which indicates P 0·3 GPa (Liou et al., 1974; Apted & Liou, 1983). The appreciable Cl contents of the Mg-hornblende imply an external fluid source whose origin is difficult to constrain. Hornblende veining was observed in mantle rocks from the western Iberia margin and attributed to interaction with seawater-derived fluids (Kimball & Evans, 1988). Alternatively, continental crust derivation of Cl-rich fluids migrating from thinned overlying crust could be hypothesized, as proposed by Muntener et al. (2000) for the hydration of the lower crustal–upper mantle sequence of Val Malenco (Central Alps).
The amphibole compositional variations, i.e. decreasing amount of edenite and Ti-tschermakite substitution (Fig. 10), indicate a metamorphic evolution characterized by decreasing temperature conditions. The P–T conditions of the subsequent retrograde evolution during the F2 stage, dominated by intense brittle fracturing and serpentinization, are difficult to constrain. No definitive consensus on lizardite thermal stability exists. It should form at low temperature (<300°C) during low-pressure retrograde metamorphism (Evans et al., 1976; O'Hanley, 1996; Evans, 2004), but departure from the ideal formula by incorporation of Fe and Al enlarges its thermal stability up to 400–500°C (e.g. Caruso & Chernosky, 1979; O'Hanley et al., 1989). However, it has been experimentally shown that under relatively high-temperature conditions (T = 400°C) orthopyroxene and clinopyroxene are altered into talc and tremolite, respectively, in ultramafic systems, whereas olivine remains stable (Allen & Seyfried, 2003). The absence of talc pseudomorphs after orthopyroxene in the studied peridotites suggests that the main hydration event started at low-grade conditions, unless talc was later replaced by serpentine.
Several features of the jigsaw breccia formed during the F3 stage, such as the extremely wide range of clast size, the matrix made of comminuted serpentinite fragments and the absence of significant clast displacement, indicate a tectonic origin. Similar cataclastic rocks have been observed in the upper levels of the mantle exhumed along the western Iberia margin (Comas et al., 1996). The jigsaw-puzzle pattern of fractures indicates expansion in all directions and is considered diagnostic of fluid-induced, fault-related implosion brecciation (Sibson, 1987). The occurrence of polyphase cracking and veining suggests a complex interplay of extensional stresses leading to intermittent rupturing of the partially serpentinized peridotite, fluid infiltration and precipitation of new phases. In particular, the formation of spherulitic serpentine ± spheroidal calcite in the cataclasite matrix requires rapid growth from nucleation sites in a hydrothermal fluid (Sibson, 1987; Skelton & Valley, 2000).
The precipitation of Cl-bearing serpentine + sulphides + calcite in the cataclasite matrix indicates the involvement of Cl-, CO2- and S-bearing fluids. In particular, the pyrite + millerite + calcite assemblage indicates high 
conditions and high water:rock ratios (Frost, 1985; Palandri & Reed, 2004); the late calcite + pyrite veins, which crosscut all the earlier brittle structures, could be ascribed to cold seawater-related alteration (T < 50°C; Hopkinson et al., 2004). Textural and mineral analogies between the F3 cataclasites and the uppermost levels of the serpentinized mantle exhumed along the western Iberia margin (Comas et al., 1996; Alt & Shanks, 1998) support the notion that the investigated mantle section was emplaced at shallow levels, in proximity to the ocean floor.
Pressure–temperature–time evolution of the External Liguride mantle section
Combining petrographic and mineral chemistry data with thermobarometric estimates for both pyroxenites and peridotites, the following P–T evolution (Fig. 12) has been reconstructed for the External Liguride mantle section: (1) equilibration occurred under garnet-facies conditions (P 2·8 GPa, T 1100°C), compatible with a relatively hot subcontinental geotherm (Fig. 12), which rules out the possibility that this early high-pressure stage was related to involvement of these mantle rocks in the Variscan subduction event, in contrast to other mantle bodies now exposed in the Eastern Alps (Tumiati et al., 2003); (2) an early decompression stage, shown by garnet breakdown in the pyroxenites and associated with the formation of the spinel tectonite fabric in the peridotites, occurred under moderately decreasing temperature (P 1·8 GPa, T 950°C); (3) further, nearly isothermal, decompression to P 0·8 GPa led to the development of plagioclase-bearing peridotite mylonites and growth of the olivine + plagioclase association in the pyroxenites; (4) inception of hydration occurred under amphibolite (Mg-hornblende + chlorite)-facies conditions at P 0·3 GPa and T = 500–700°C, coupled with brittle fracturing; (5) shallow-level polyphase brittle deformation occurred, associated with peridotite serpentinization, which culminated in sea-floor exposure.
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As the closure temperature of the Sm–Nd system is higher than the temperature estimated for Mg-hornblende vein formation, the Sm–Nd isochron age of 186·0 ± 1·5 Ma represents a maximum time estimate for the onset of mantle hydration. In addition, the Lu–Hf isochron indicates that the mantle section cooled below 750–900°C (i.e. the inferred closure temperature of the Lu–Hf system) at 220 ± 13 Ma, which may be a minimum age for the development of plagioclase-facies mylonites. The ages obtained by the Sm–Nd and Lu–Hf methods thus record a retrograde history of slow cooling (
5°C/Ma) in the plagioclase stability field between the Late Triassic and the Early Jurassic. The retrograde evolution from plagioclase- to amphibolite-facies conditions was probably associated with the intrusion of syn-rift MOR-type gabbros in other sectors of the External Liguride mantle section, which have given Sm–Nd clinopyroxene–plagioclase–whole-rock isochron ages of 179 ± 9 Ma (Tribuzio et al., 2004).
Two peridotites from the southernmost Mt. Nero body of the External Liguride units yielded Sm–Nd clinopyroxene–plagioclase isochron ages of 163 ± 20 Ma and 165 ± 20 Ma (Rampone et al., 1995). These ages were interpreted to date the exhumation of the External Liguride mantle section to plagioclase-facies conditions, thus contrasting with the Lu–Hf age reported in this study, unless the mantle bodies from the External Liguride domain record diachronous exhumation. The Mt. Nero mantle body retains a fertile geochemical signature but is pyroxenite-poor, characterized by the occurrence of impregnated plagioclase-rich peridotites and locally intruded by basaltic dykes and small gabbroic bodies (Terranova & Zanzucchi, 1984; Piccardo et al., 2004b; Poggi et al., 2005). In addition, the Mt. Nero peridotites display typical MORB-like Nd and Sr isotope compositions (Rampone et al., 1995). It seems plausible, therefore, that the age reported by Rampone et al. (1995) records the effects on the previously exhumed lithospheric mantle of MORB–rock reaction and MORB impregnation associated with plagioclase formation (see also Müntener & Piccardo, 2003).
In the External Liguride units, lower crustal granulites derived from gabbroic protoliths of Late Palaeozoic age yielded an 40Ar/39Ar amphibole cooling age of 228 ± 2 Ma (Meli et al., 1996). This radiometric date is within error of the Lu–Hf cooling age obtained from the garnet pyroxenites in the present study. When the mantle section was under temperature conditions of 750–900°C, therefore, the gabbro-derived granulites cooled below a temperature of about 500°C (see also Marroni et al., 1998). The subsequent sequence of low-temperature brittle deformations recorded by the lower crustal rocks (see also Montanini, 1997) was related to their exhumation in conjunction with the rifting event that led to oceanization in the Middle Jurassic.
Timing of the high-temperature decompression
The relationships between high-temperature structures and rifting processes in the mantle exhumed along modern and fossil non-volcanic margins are puzzling (e.g. Whitmarsh et al., 2001; Manatschal, 2004). In the case of the External Liguride mantle section, constraints on the timing of high-temperature exhumation from garnet- to plagioclase-facies conditions are provided by new geochronological, petrological and regional geological data.
The Mesozoic rifting that led to the opening of the Ligurian Tethys was preceded by Late Palaeozoic (Late Carboniferous to Early Permian) extensional processes associated with extensive tholeiitic magmatism (Voshage et al., 1990; Tribuzio et al., 1999; Hermann et al., 2001; Montanini & Tribuzio, 2001). The Late Palaeozoic extension was related to asthenospheric upwelling, possibly following delamination of a lithospheric root (Schott & Schmeling, 1998; Arnold et al., 2001) at the end of the Variscan orogeny. Whether the Late Palaeozoic extension and Mesozoic rifting reflected a continuous evolution that led to continental rifting (Piccardo et al., 1990; Lardeaux & Spalla, 1991; Dal Piaz, 1993) or were related to independent geodynamic processes (Handy & Zingg, 1991; Müntener et al., 2000; Müntener & Hermann, 2001) is still a matter of debate.
In this tectonic framework, the exhumation of the External Liguride mantle section from garnet- to plagioclase-facies conditions may have occurred in a unique stage starting in the Late Palaeozoic to Triassic, similar to that proposed by Vissers et al. (1991) for the the Ligurian Alps peridotites. However, the inferred change of slope in the decompressional P–T path (Fig. 12) is difficult to reconcile with this evolution, suggesting that exhumation may have been accomplished through two distinct decompression stages, probably related to separate geodynamic events.
Two main scenarios may thus be envisaged for the timing of the high-temperature decompression recognized for the External Liguride mantle section. In the first, the nearly isothermal decompression from spinel- to plagioclase-facies conditions occurred in relation to Late Palaeozoic lithospheric thinning, and the previous exhumation from the garnet to spinel stability field was related to an ancient geodynamic event. Such a scenario is supported by the comparison with the mantle sequence from Erro Tobbio (Western Alps), which records a similar high-temperature decompression (Vissers et al., 1991; Hoogerduijn Strating et al., 1993). On the basis of Sm–Nd clinopyroxene–plagioclase–whole-rock isochrons yielding ages of 313 ± 13 Ma and 273 ± 16 Ma, Rampone et al. (2005) recently ascribed the decompression from spinel- to plagioclase-facies conditions of the Erro-Tobbio peridotite to Late Palaeozoic lithospheric extension. The second scenario implies that the decompression from garnet- to spinel-facies conditions was related to the Late Palaeozoic evolution, and that the subsequent exhumation associated with plagioclase-facies shear zone formation was due to the onset of Mesozoic rifting (i.e. it occurred in Triassic times).
In the Northern Apennines, evidence for Middle Triassic rifting phases is provided by the lower crustal granulites from the External Liguride units (Marroni et al., 1998; see preceding section) and by the occurrence of an extensional basin sequence consisting of Middle Triassic marine deposits (Martini et al., 1986) associated with alkali basaltic lava flows (Ricci & Serri, 1975). Evidence for rift-related Middle Triassic faults (e.g. Bertotti et al., 1993, 1999) and coeval involvement of mantle-derived melts with alkaline affinity (e.g. Stähle et al., 1990, 2001) is also reported for the South Alpine domain of the Western Alps. In addition, preliminary geochemical and geochronological investigations of the Finero gabbroic complex (South Alpine domain of Western Alps) have shown a MOR-type affinity for the parental melts and an intrusion age of about 230 Ma, thus testifying to asthenospheric ascent in Middle to Late Triassic times (Peressini et al., 2004).
The comparison with the P–T–t path of the fosssil mantle–lower crust section cropping out at Malenco, in the central Alps (Müntener et al., 2000; Villa et al., 2000; Müntener & Hermann, 2001), supports the idea that the near-isothermal decompression recorded by the External Liguride mantle section was of Triassic age. The spinel-facies mantle peridotites from Malenco equilibrated at the base of the continental crust (P = 1·0–1·2 GPa) in Late Palaeozoic times, at relatively low-temperature conditions (T 800°C). The subsequent evolution of the Malenco peridotites was characterized by a Late Palaeozoic isobaric cooling stage that was followed by near-isothermal decompression (from 0·8 to 0·4 GPa, at 550°C) associated with hydration and development of olivine + tremolite + chlorite mylonites (Müntener et al., 2000; Villa et al., 2000; Müntener & Hermann, 2001). The near-isothermal decompression started at about 225 Ma, on the basis of 40Ar/39Ar amphibole age determinations, and has been attributed to the onset of the rifting event that led to continental break-up. The 40Ar/39Ar amphibole age of the Malenco section is close to the time at which the External Liguride mantle passed through the Lu–Hf blocking temperature (750–900°C), under plagioclase-facies conditions. The Malenco and External Liguride mantle sections were, therefore, at different thermal conditions in Middle to Late Triassic times. The higher temperature recorded by the External Liguride mantle section may reflect its rapid uplift from deeper, spinel-facies mantle levels. If the spinel- to plagioclase-facies decompression was of Late Palaeozoic age, as envisaged for the first scenario, the External Liguride mantle section must have remained in the plagioclase stability field for 90–130 Myr. Maintaining these high-temperature–low-pressure conditions for such a long time span is difficult to reconcile with a realistic geodynamic model.
Although the new geochronological data are not conclusive, we propose a Middle to Late Triassic age for the near-isothermal spinel- to plagioclase-facies decompression recorded by the External Liguride mantle section. This decompression probably occured in conjunction with the onset of the rifting that led to the opening of the Ligurian Tethys in the Middle Jurassic. The subsequent slow cooling in the plagioclase stability field, as indicated by Sm–Nd and Lu–Hf mineral isochrons for the garnet pyroxenites, is consistent with a rifting situation in which conductive heat loss was partially counteracted by the underlying ascending asthenosphere. Such asthenospheric ascent during the Late Triassic to Early Jurassic evolution could be related to the melt impregnation processes recorded by the southernmost mantle bodies of the External Liguride units (Müntener & Piccardo, 2003). The age of such melt–peridotite interactions is not known, but could precede the intrusion of the syn-rift gabbros, which represent an advanced stage of the process of upward migration of asthenospheric melts (Piccardo et al., 2004a, 2004b), and have been dated at 179 ± 9 Ma (Tribuzio et al., 2004).
Implications for continental rifting and ocean opening
The age determinations reported in this study permit mantle exhumation rates during rift evolution to be estimated. The age of the oldest pelagic sediments associated with the Ligurian Tethys ophiolites (164 Ma; Chiari et al., 2000; Bill et al., 2001) can be assumed to indicate the timing of mantle exhumation at the sea floor. The pressure at which the studied mantle section cooled below the closure temperature of the Sm–Nd system cannot be constrained precisely. An upper pressure limit of 0·8 GPa is provided by the olivine + plagioclase assemblage in the basic pyroxenites, and a lower pressure limit of 0·4 GPa is presumed, in agreement with the onset of amphibolite-facies metamorphism at P 0·3 GPa. These pressure limits suggest uplift of 12–25 km (assuming a rock density value of 3·3 g/cm3) and integrated exhumation rates between 1·2 and 0·6 mm/year.
Recent studies of modern rifted continental margins have shown that the top of the exhumed mantle consists of a brittle fault zone (e.g. Manatschal, 2004). Continental break-up is inferred to develop through extensional detachment structures that root into the lithospheric mantle. These structures develop at an advanced stage of rift evolution and lead to exhumation of progressively deeper mantle levels (see also Whitmarsh et al., 2001). The ubiquitous brittle structures observed in the mantle bodies of the present study, superimposed on a plagioclase-facies shear zone, are consistent with a shallow detachment fault event (Fig. 13).
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The plagioclase mylonites from the External Liguride mantle section could have originated along the same extensional detachment that gave rise to the brittle structures. The occurrence of such fine-grained mylonites is known to induce strain softening and reduction of the bulk strength of the lithospheric mantle (Vissers et al., 1995
; Frederiksen & Braun, 2001) and to promote further strain localization (Dijkstra et al., 2004). This implies that the Ligurian Tethys rifting was a simple-shear dominated process, which resulted in asymmetric extension (Lemoine et al., 1987; Vissers et al., 1991). However, most numerical and analogue models of the formation of non-volcanic rifted margins (Brun & Beslier, 1996; Bowling & Harry, 2001; Michon & Merle, 2003), as well as those based on observations from modern and fossil margins (e.g. Whitmarsh et al., 2001), predict an early stage of pure shear-controlled lithospheric necking, in relation to the ascent of the underlying asthenosphere (Whitmarsh et al., 2001). In addition, slow cooling in the plagioclase stability field and the relatively large time interval between plagioclase mylonite formation and the onset of brittle deformation–hydration argues against continuous evolution along a detachment structure. We thus propose that the External Liguride plagioclase mylonites, formed by an extensional shearing event related to a lithospheric necking stage, mark the onset of the rift evolution. Alternatively, the plagioclase-facies shear zone may have marked the boundary between a relatively cold mantle section (e.g. the garnet pyroxenite-bearing peridotites of this study), and the underlying hotter lithospheric mantle heated by asthenospheric melt ascent and infiltration (e.g. the southernmost External Liguride massifs). This hypothesis implies that the mantle impregnation process was already active in Middle to Late Triassic times. |
CONCLUSIONS |
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The fertile mantle section from the External Liguride units (Northern Apennines) was exhumed at an ocean–continent transition. The mantle bodies considered in the present study escaped melt contamination and intrusion and preserve relics of a deep-seated lithospheric history. In particular, garnet-bearing pyroxenite layers record an early stage of equilibration in the subcontinental lithosphere under high P–T conditions (
2·8 GPa and 1100°C).
Pyroxenites and enclosing peridotites preserve evidence for a two-stage high-temperature decompression. Garnet-bearing parageneses in the pyroxenites were initially replaced by two pyroxenes + spinel and then by olivine + plagioclase assemblages. The peridotites are characterized by relics of a spinel tectonite fabric extensively overprinted by plagioclase-facies mylonitic deformation. The early decompression, from the garnet to spinel stability field, was associated with cooling to T 950°C and presumably occurred in relation to Late Variscan (Late Carboniferous to Early Permian) continental extension. The subsequent near-isothermal decompression, including the plagioclase-facies shearing event, was probably related to the onset of the Mesozoic rifting event (i.e. in the Middle to Late Triassic) that led to oceanization in the Middle Jurassic.
The high-temperature decompression was overprinted by a widespread polyphase brittle deformation under decreasing temperature conditions, coupled with hydration. In particular, the brittle evolution includes an early amphibolite-facies stage, followed by low-temperature hydrothermal alteration and serpentinization associated with polyphase cataclasis. The brittle evolution stage is attributed to involvement of the studied mantle section in the footwall of an extensional detachment system. This probably reactivated rheological weaknesses developed under high-temperature, ductile, conditions (i.e. the plagioclase-facies shear zones) and led to continental break-up and mantle exhumation at the sea floor.
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ACKNOWLEDGEMENTS |
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We wish to thank Matteo Beneventi for helping with mineral separation and field assistance. We are also grateful to Gloria Vaggelli for EMP technical assistance, Giuliana De Grandis, Luca Barchi and Andrea Comelli for sample preparation, and Edvige Masini for drawing the maps. Dr Sergei Simakov is acknowledged for helping with P–T calculations. Othmar Müntener and Dante Canil provided constructive comments on the manuscript. Editorial handling and remarks by Colin Devey are also appreciated. This work was supported by funds from the Ministero Italiano dell'Università e della Ricerca (Progetti di Ricerca d'Interesse Nazionale).
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FOOTNOTES |
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Present address: Polish Academy of Sciences, Krakow Research Centre, Krakow, Poland. 
*Corresponding author. E-mail: alessandra.montanini{at}unipr.it.
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REFERENCES |
|---|
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]
Agrinier P., Mevel C., Girardeau J. (1988) Hydrothermal alteration of the peridotites cored at the ocean/continent boundary of the Iberian margin: petrologic and stable isotope evidence. In Boillot G. and Winterer E. L. (Eds.), et al. Ocean Drilling Program Scientific Results, 103.(Ocean Drilling Program, College Station, TX) pp. 225–233.
Allen D. E. and Seyfried W. E. Jr. (2003) Compositional controls on vent fluids from ultramafic-hosted hydrothermal systems at mid-ocean ridges: an experimental study at 400°C, 500 bars. Geochimica et Cosmochimica Acta 67:1531–1542.[CrossRef][Web of Science]
Ai Y. (1994) A revision of the garnet–clinopyroxene Fe2+–Mg exchange geothermometer. Contributions to Mineralogy and Petrology 115:467–473.[CrossRef]
Alt J. C. and Shanks W. C. (1998) Sulphur in serpentinized oceanic peridotites: serpentinization processes and microbial sulfate reduction. Journal of Geophysical Research 103:9917–9929.[CrossRef]
Anczkiewicz R., Platt P., Thirlwall M. F., Wakabayashi J. (2004) Franciscan subduction off to a slow start: evidence from high precision Lu–Hf garnet ages on high grade blocks. Earth and Planetary Science Letters 225:147–161.[CrossRef][Web of Science]
Apted M. J. and Liou J. G. (1983) Phase relations among greenschist–epidote amphibolite and amphibolite in a basaltic system. American Journal of Science 283:328–354.
Arnold J., Jacoby W. R., Schmeling H., Schott B. (2001) Continental collision and the dynamic and thermal evolution of the Variscan orogenic crustal root—numerical models. Journal of Geodynamics 31:273–291.[CrossRef][Web of Science]
Barrett J. J. and Spooner E. F. C. (1977) Ophiolitic breccias associated with allochthonous oceanic crustal rocks in the East Ligurian Apennines, Italy—a comparison with observations from rifted oceanic ridges. Earth and Planetary Science Letters 35:79–91.[CrossRef][Web of Science]
Beccaluva L., Macciotta G., Piccardo G. B., Zeda O. (1984) Petrology of lherzolitic rocks from the Northern Apennine ophiolites. Lithos 17:299–316.[CrossRef][Web of Science]
Bedini R. M., Blichert-Toft J., Boyet M., Albarède F. (2004) Isotopic constraints on the cooling of the continental lithosphere. Earth and Planetary Science Letters 223:99–111.[CrossRef][Web of Science]
Bernini M., Vescovi P., Zanzucchi G. (1997) Schema strutturale dell'Appennino settentrionale. Ateneo Parmense Acta Naturalia 33:43–54.
Bertotti G. (1991) Early Mesozoic extension and Alpine shortening in the western Southern Alps: the geology of the area between Lugano and Menaggio (Lombardy, Northern Italy). Memorie della Società Geologica di Padova 43:17–123.
Bertotti G., Picotti V., Bernoulli D., Castellarin A. (1993) From rifting to drifting: tectonic evolution of the South-Alpine upper crust from the Triassic to the Early Cretaceous. Sedimentary Geology 86:53–76.[CrossRef][Web of Science]
Bertotti G., Seward D., Wijbrans J., Ter Voorde M., Urford A. J. (1999) Crustal thermal regime prior to, during, and after rifting: a geochronological and modeling study of the Mesozoic South Alpine rifted margin. Tectonics 18:185–200.[CrossRef][Web of Science]
Beslier M. O., Cornen G., Girardeau J. (1996) Tectonometamorphic evolution of peridotites from the ocean/continent transition of the Iberia Abyssal Plain. In Sawyer D. S., Whitmarsh R. B., Klaus A. (Eds.), et al. Ocean Drilling Program Scientific Results, 149.(Ocean Drilling Program, College Station, TX) pp. 397–411.
Bill M., O'Dogherty L., Guex J., Baumgartner P. O., Masson H. (2001) Radiolarite ages in Alpine–Mediterranean ophiolites: constraints on the oceanic spreading and the Tethys–Atlantic connection. Geological Society of America Bulletin 113:129–143.
Blichert-Toft J. and Albarède F. (1997) The Lu-Hf isotope geochemistry of chondrites and the evolution of the mantle-crust system. Earth Planetary Scence Letters 148:243–258.
Blichert-Toft J., Albarède F., Kornprobst J. (1999) Lu–Hf isotope systematics of garnet pyroxenites from Beni Bousera, Morocco: implications for basalt origin. Science 283:1303–1306.
Boillot G., Feraud G., Recq M., Girardeau J. (1989) Undercrusting by serpentinite beneath rifted margins. Nature 6242:523–525.
Boland J. N. and Van Roermund H. L. M. (1983) Mechanisms of exsolution in omphacites from high-temperature type B eclogites. Physics and Chemistry of Minerals 9:30–37.[CrossRef][Web of Science]
Bonatti E. and Hamlyn P.R. (1981) Oceanic ultramafic rocks. In Emiliani C. (Ed.). The Sea(John Wiley, New York) pp. 241–283.
Bowling J. C. and Harry D. L. (2001) Geodynamic models of continental extension and the formation of non-volcanic rifted continental margins. In Wilson R. C. L., Withmarsh R. B., Taylor B., Froitzheim N. (Eds.). Non-Volcanic Rifting of Continental Margins: a Comparison of Evidence from Land and Sea. Geological Society, London, Special Publications 187: pp. 511–536.
Brey G. P. and Kohler T. (1990) Geothermometry in four-phase lherzolites II. New thermobarometers, and practical assesment of existing thermobarometers. Journal of Petrology 31:1353–1378.
Brun J. P. and Beslier M. O. (1996) Mantle exhumation at passive margins. Earth and Planetary Science Letters 142:161–173.[CrossRef][Web of Science]
Burg J. P., Bodinier J. L., Chaudry S., Hussain S., Dawood H. (1998) Infra-arc mantle–crust transition and intra-arc mantle diapirs in the Kohistan Complex (Pakistani Himalaya): petro-structural evidence. Terra Nova 10:74–80.[CrossRef][Web of Science]
Caruso L. and Chernosky J. V. Jr. (1979) The stability of lizardite. Canadian Mineralogist 17:757–769.
Chiari M., Marcucci M., Principi G. (2000) The age of the radiolarian cherts associated with the ophiolites in the Apennines (Italy) and Corsica (France): a revision. Ofioliti 25:141–146.[Web of Science]
Comas M. C., Sànchez-Gòmez G., Cornen E., de Kaenel, E. (1996) Serpentinized peridotite breccia and olistostromes on basement highs from the Iberia Abyssal Plain: implications for tectonic margin evolution. In Sawyer D. S., Whitmarsh R. B., Klaus A. (Eds.), et al. Ocean Drilling Program Scientific Results, 149.(Ocean Drilling Program, College Station, TX) pp. 577–591.
Cornen G., Beslier M., Girardeau J. (1996) Petrologic characteristics of the ultramafic rocks from the ocean/continent transition in the Iberia Abyssal Plain. In Sawyer D. S., Whitmarsh R. B., Klaus A. (Eds.), et al. Ocean Drilling Program Scientific Results, 149.(Ocean Drilling Program, College Station, TX) pp. 377–395.
Cortesogno L., Galbiati B., Principi G. (1987) Note alla ‘Carta geologica delle ofioliti del Bracco’ e ricostruzione della paleogeografia Giurassico-Cretacica. Ofioliti 12:261–342.
Dalmasso J., Barci-Funel G., Ardison G. J. (1992) Reinvestigation of the decay of the long-lived odd–odd 176Lu nucleus. Applied Radiation and Isotopes 43:69–76.[Web of Science]
Dal Piaz G. V. (1993) Evolution of Austro-alpine and Upper Penninic basement in the North-Western Alps from Variscan convergence to post-Variscan extension. In von Raumer J. F. and Neubauer F. (Eds.). Pre-Mesozoic Geology in the Alps.(Springer-Verlag, New York) pp. 327–344.
Davies G. R., Nixon P. H., Pearson D. G., Obata M. (1993) Tectonic implications of graphitized diamonds from the Ronda peridotite massif, soutrhern Spain. Geology 21:471–474.
de Ronde A. A., Heilbronner R., Stuenitz H., Tullis J. (2004) Spatial correlation of deformation and mineral reaction in experimentally deformed plagioclase–olivine aggregates. Tectonophysics 27:93–109.[CrossRef]
Desmurs L., Manatschal G., Bernoulli D. (2001) In Wilson R. C. L., Withmarsh R. B., Taylor B., Froitzheim N. (Eds.). The Steinmann trinity revisited: exhumation and magmatism along an ocean–continent transition: the Platta nappe, eastern Switzerland. Non-Volcanic Rifting of Continental Margins: a Comparison of Evidence from Land and Sea. Geological Society, London, Special Publications 187:235–266.
Dijkstra A.H., Drury M.R., Vissers R.L.M., Newman J., Van Roermund H.L.M. (2004) In Alsop G.I. and Holdsworth R.E. (Eds.). Shear zones in the upper mantle: Evidence from alpine- and ophiolite-type peridotite massifs. Flow processes in Faults and Shear zones 224:11–24 Geological Society London Special Publication.
Eggins S. M. (1992) Petrogenesis of Hawaiian tholeiites: 1, phase equilibria constraints. Contributions to Mineraogy and Petrology 110:387–397.
Ellis D. J. and Green D. H. (1979) An experimental study of the effect of Ca upon garnet–clinopyroxene Fe–Mg exchange equilibria. Contributions to Mineralogy and Petrology 71:13–22.[CrossRef][Web of Science]
Evans B. W., Johannes W., Otterodoom H., Tromsdorff V. (1976) Stability of chrysotile and antigorite in the serpentine multisystem. Schweizerische Mineralogische und Petrographische Mitteilungen 56:9–93.
Evans B.W. (1982) Phase relations of metamorphic amphiboles; natural occurrence and theory; Amphiboles in metamorphosed ultramafic rocks. Reviews in Mineralogy and Geochemistry 9B:98–113.
Evans B.W. (2004) The serpentinite multisystem revisited: chrysotile is metastable. International Geology Review 46:479–506.[Web of Science]
Frederiksen S. and Braun J. (2001) Numerical modelling of strain localisation during extension of the continental lithosphere. Earth and Planetary Science Letters 188:241–251.[CrossRef][Web of Science]
Frost B. R. (1985) On the stability of sulfides, oxides and native metals in serpentine. Journal of Petrology 26:31–63.
Furusho M. and Kanagawa K. (1999) Transformation-induced strain localization in a lherzolite mylonite from the Hidaka metamorphic belt of central Hokkaido, Japan. Tectonophysics 313:411–432.[CrossRef][Web of Science]
Ganguly J., Tirone M., Hervig R. L. (1988) Diffusion kinetics of samarium and neodymium in garnet, and a method for determining cooling rates of rocks. Science 281:805–807.
Gasparik T. (1984) Two-pyroxene thermobarometry with new experimental data in the system CaO–MgO–Al2O3–SiO2. Contributions to Mineralogy and Petrology 87:87–97.[CrossRef][Web of Science]
Godard G. and Martin S. (2000) Petrogenesis of kelyphites in garnet peridotites: a case study from the Ulten zone, Italian Alps. Journal of Geodynamics 30:117–145.[CrossRef][Web of Science]
Green D. H. and Ringwood A. E. (1967) An experimental investigation of the gabbro to eclogite transformation and its petrological applications. Geochimica et Cosmochimica Acta 31:767–833.[Web of Science]
Handy M. R. and Stünitz H. (2002) In de Meer S., Drury M. R., de Bresser J. H. P., Pennock G. M. (Eds.). Strain localization by fracturing and reaction weakening—a mechanism for initiating exhumation of subcontinental mantle beneath rifted margins. Deformation Mechanisms, Rheology and Tectonics: Current Status and Future Perspectives. Geological Society, London, Special Publications 200:387–407.
Handy M. R. and Zingg A. (1991) The tectonic and rheological evolution of an attenuated cross section of the continental crust: Ivrea crustal section, southern Alps, northwestern Italy and southern Switzerland. Geological Society of America Bulletin 103:236–253.
Hermann J., Müntener O., Günther D. (2001) Differentiation of mafic magma in a continental crust-to-mantle transition zone. Journal of Petrology 42:189–206.
Hirschmann M. M. and Stolper E. M. (1996) A possible role for garnet pyroxenite in the origin of the garnet signature in MORB. Contributions to Mineralogy and Petrology 124:185–208.[CrossRef][Web of Science]
Holland T. and Blundy J. D. (1994) Non-ideal interactions in calcic amphiboles and bearing on amphibole–plagioclase thermometry. Contributions to Mineralogy and Petrology 116:433–447.
Hoogerduijn Strating E. H., Rampone E., Piccardo G. B., Drury M., Vissers R. L. M. (1993) Subsolidus emplacement of mantle peridotite during incipient oceanic rifting and opening of the Mesozoic Tethys (Voltri Massif, NW Italy). Journal of Petrology 34:901–927.
Hopkinson L., Beard J. S., Boulter C. A. (2004) The hydrothermal plumbing of a serpentinite-hosted detachment: evidence from the West Iberia non-volcanic rifted continental margin. Marine Geology 204:1–15.[CrossRef][Web of Science]
Irving A. J. (1974) Geochemical and high pressure experimental studies on garnet pyroxenite and pyroxene granulite xenoliths from the Delegate basaltic pipes. Australia. Journal of Petrology 15:1–40.
Jacobsen S. B. and Wasserburg G. J. (1980) Sm–Nd isotopic evolution of chondrites. Earth and Planetary Science Letters 50:139–155.[CrossRef][Web of Science]
Joanny V., Van Roermund H. L. M., Lardeaux J. M. (1991) The clinopyroxene/plagioclase symplectite in retrograde eclogites: a potential geothermobarometer. Geologische Rundschau 80:303–320.
Johnson K. T. M. (1998) Experimental determination of partition coefficients for rare earth and high-field-strength elements between clinopyroxene, garnet and basaltic melt at high pressures. Contributions to Mineralogy and Petrology 133:60–68.[CrossRef][Web of Science]
Kennedy C. S. and Kennedy G. C. (1976) The equilibrium boundary between graphite and diamond. Journal of Geophysical Research 81:2467–2470.
Kimball K.L. and Evans C.A. (1988) Hydrothermal alteration of peridotite from the Galicia Margin, Iberian Peninsula. In Boillot G. and Winterer E. L. (Eds.), et al. Ocean Drilling Program Scientific Results, 103(Ocean Drilling Program, College Station, TX) pp. 241–251.
Klemme S. (2004) The influence of Cr on the garnet–spinel transition in the Earth's mantle: experiments in the system MgO–Cr2O3–SiO2 and thermodynamic modelling. Lithos 77:639–646.[CrossRef][Web of Science]
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.
Krogh E. J. (1988) The garnet–clinopyroxene Fe–Mg geothermometer—a reinterpretation of existing experimental data. Contributions to Mineralogy and Petrology 99:44–48.[CrossRef][Web of Science]
Krogh Ravna E. (2000) The garnet–clinopyroxene Fe2+–Mg geothermometer: an updated calibration. Journal of Metamorphic Geology 18:211–219.[CrossRef][Web of Science]
Lardeaux J. M. and Spalla M. I. (1991) From granulites to eclogites in the Sesia zone (Italian Western Alps): a record of opening and closure of the Piedmont ocean. Journal of Metamorphic Geology 9:35–58.[Web of Science]
Leake B.E., Woolley A.R., Arps C.E.S., Birch W.D., Gilbert M.C., Grice J.D., Hawthorne F.C., Kato A., Kisch H.J., Krivovichev V.G., Linthout K., Laird J., Mandarino J.A., Maresch W.V., Nickel E.H., Rock N.M.S., Schumacher J.C., Smith D.C., Stephenson N.C.N., Ungaretti L., Whittaker E.J.W., Guo Y.Z. (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]
Lemoine M., Tricart P., Boillot G. (1987) Ultramafic and gabbroic ocean floor of the Ligurian Tethys (Alps, Corsica, Apennine): in search of a genetic model. Geology 15:622–625.
Lenoir X., Garrido C. J., Bodinier J. L., Dautria J. M., Gervilla F. (2001) The recrystallization front of the Ronda peridotite: evidence for melting and thermal erosion of subcontinental lithospheric mantle beneath the Alboran Basin. Journal of Petrology 42:141–158.
Liou J. G., Kuniyoshi S., Ito K. (1974) Experimental studies of the phase relations between greenschist and amphibolite in a basaltic system. American Journal of Science 274:613–632.[Abstract]
Isoplot/Ex version 3.0. A Geochronological Toolkit for Microsoft Excel. Berkeley Geochronology Center Special Publication 2003. Ludwig K. R. (2003) 1A, 47.
Manatschal G. (2004) New models for evolution of magma-poor rifted margins based on a review of data and concepts from West Iberia and Alps. International Journal of Earth Sciences 93:432–466.[Web of Science]
Manatschal G. and Bernoulli D. (1999) Architecture and tectonic evolution of nonvolvcanic margins: present-day Galicia and ancient Adria. Tectonics 18:1099–1119.[CrossRef][Web of Science]
Manatschal G., Froitzheim N., Rubenach M., Turrin B. D. (2001) The role of detcahment faulting in the formation of an ocean–continent transition: insights from the Iberia Abyssal plain. In Wilson R. C. L., Withmarsh R. B., Taylor B., Froitzheim N. (Eds.). Non-Volcanic Rifting of Continental Margins: a Comparison of Evidence from Land and Sea. Geological Society, London, Special Publications 187: pp. 405–428.
Marroni M., Molli G., Montanini A., Tribuzio R. (1998) The association of continental crust rocks with ophiolites in the Northern Apennines (Italy): implications for the continent–ocean transition in the Western Tethys. Tectonophysics 292:43–66.[CrossRef][Web of Science]
Marroni M., Molli G., Montanini A., Ottria G., Pandolfi L., Tribuzio R. (2002) The External Ligurian units (Northern Apennine, Italy): from rifting to convergence of a fossil ocean–continent transition zone. Ofioliti 27:119–131.[Web of Science]
Martini I.P., Rau A., Tongiorgi M. (1986) Syntectonic sedimentation in a Middle triassic rift, Northern Apennines, Italy. Sedimentary Geology 47:191–219.[CrossRef][Web of Science]
McCandless T. E. and Gurney J. J. (1989) Sodium in garnet and potassium in clinopyroxene: criteria for classifying mantle eclogites. In Ross J. (Ed.). Kimberlites and Related Rocks. Special Publication of the Geologiacal Society of Australia 2: pp. 827–832.
Meli S., Montanini A., Thoni M., Frank W. (1996) Age of mafic granulite blocks from the External Liguride Units (northern Apennine, Italy). Memorie di Scienze Geologiche di Padova 48:65–72.
Messiga B. and Bettini E. (1990) Reactions behaviour during kelyphite and symplectite formation: a key study of mafic granulites and eclogites from the Bohemian Massif. European Journal of Mineralogy 2:125–144.
Mével C. (2003) Serpentinization of abyssal peridotites at mid-ocean ridges. Comptes Rendus de l'Académie des Sciences 335:825–852.
Michon L. and Merle O. (2003) Mode of lithospheric extension: conceptual models from analogue modelling. Tectonics 22:1028 doi 10.1029/2002TC001435.[CrossRef]
Montanini A. (1997) Mafic granulites in the Cretaceous sedimentary mélanges from the northern Apennine (Italy): petrology and tectonic implications. Schweizerische Mineralogische und Petrographische Mitteilungen 77:43–64.
Montanini A. and Tribuzio R. (2001) Gabbro-derived and felsic granulites from the Northern Apennines (Italy): evidence for emplacement of tholeiitic basalts in the post-Variscan lower crust. Journal of Petrology 42:2259–2277.
Montanini A. and Tribuzio R. (2003) Igneous to metamorphic evolution of mafic rocks from a fossil ocean–continent transition (External Liguride ophiolites, Italy). Geophysical Research Abstracts 5: EAE03-A-06549.
Montanini A., Tribuzio R., Beneventi M., Bersani D. (2005) Insights into the origin of mantle graphite from a new occurence in garnet pyroxenites from the External Liguride peridotites (Northern Apennine, Italy). Peridotite Workshop 200527–30 SeptemberLanzo (Torino)Ofioliti 30: pp. 200.
Morishita T. and Arai S. (2003) Evolution of spinel–pyroxene symplectite in spinel-lherzolites from the Horoman Complex, Japan. Contributions to Mineralogy and Petrology 144:509–522.[Web of Science]
Müntener O. and Hermann J. (1996) The Val Malenco lower crust–upper mantle complex and its field relations (Italian Alps). Schweizerische Mineralogische und Petrographische Mitteilungen 763:475–500.
Müntener O. and Hermann J. (2001) The role of lower crust and continental upper mantle during formation of nonvolcanic passive margins: evidence from the Alps. In Wilson R. C. L., Withmarsh R. B., Taylor B., Froitzheim N. (Eds.). Non-Volcanic Rifting of Continental Margins: a Comparison of Evidence from Land and Sea. Geological Society, London, Special Publications 187: pp. 267–288.
Müntener O. and Piccardo G. B. (2003) In Dilek Y. and Robinson P. T. (Eds.). Melt migration in ophiolites: the message from Alpine–Apennine peridotites and implications for embryonic ocean basins. Ophiolites in the Earth History. Geological Society, London, Special Publications 218:69–89.[CrossRef]
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.
Müntener O., Pettke T., Desmurs L., Meier M., Schaltegger U. (2004) Refertilization of mantle peridotite in embryonic ocean basins: trace element and Nd-isotopic evidence and implications for crust–mantle relationships. Earth and Planetary Science Letters 221:293–308.[CrossRef][Web of Science]
Newman J., Lamb W. M., Drury M. R., Vissers R. L. M. (1998) Deformation processes in a peridotite shear zone: reaction-softening by an H2O-deficient, continuous net transfer reaction. Tectonophysics 303:193–222.[Web of Science]
Nickel K. G. (1986) Phase equilibria in the system SiO2–MgO–Al2O3–CaO–Cr2O3 (SMACCR) and their bearing on spinel/garnet lherzolite relationships. Neues Jahrbuch für Mineralogie, Abhandlungen 155:259–287.[Web of Science]
Niida K. and Green D. H. (1999) Stability and chemical 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. (1994) Material transfer and local equilibria in a zoned kelyphite from a garnet pyroxenite, Ronda, Spain. Journal of Petrology 35:271–287.
O'Brien P. J. (1997) Garnet zoning and reaction textures in overprinted eclogites, Bohemian Massif, European Variscides: a record of their thermal history during exhumation. Lithos 41:119–133.[CrossRef][Web of Science]
O'Hanley D. S. (1996) Serpentinites(Oxford University Press, New York) pp. 277.
O'Hanley D. S., Chernovsky J. V. Jr, Wicks F. J. (1989) The stability of lizardite and chrysotile. Canadian Mineralogist 27:483–493.[Web of Science]
O'Hara M. J., Richardson S. W., Wilson G. (1971) Garnet-peridotite stability and occurrence in crust and mantle. Contributions to Mineralogy and Petrology 32:48–68.[CrossRef][Web of Science]
O'Neill H. S. 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]
Palandri J. L. and Reed M. H. (2004) Geochemical models of metasomatism in ultramafic systems: serpentinization, rodingitization and sea floor carbonate chimney precipitation. Geochimica et Cosmochimica Acta 68:1115–1133.[CrossRef][Web of Science]
Pearson D. G. and Nixon P. H. (1996) Diamonds in young orogenic belts: graphitised diamonds from Beni Bousera, N. Morocco, a comparison with kimberlite-derived diamond occurrences and implications for diamond genesis and exploration. African Geosciences Review 3:295–316.
Pearson D. G. and Nowell G. M. (2004) Re–Os and Lu–Hf isotope constraints on the origin and age of pyroxenites from the Beni Bousera Peridotite Massif: implications for mixed peridotite–pyroxenite mantle sources. Journal of Petrology 45:439–455.
Pearson D. G., Davies G. R., Nixon P. H. (1993) Geochemical constraints on the petrogenesis of diamond facies pyroxenites from the Beni Bousera peridotite massif, North Morocco. Journal of Petrology 34:125–172.
Peltonen P., Kinnunen K. A., Huhma H. (2002) Petrology of two diamondiferous eclogite xenoliths from the Lahtojoki kimberlite pipe, eastern Finland. Lithos 63:151–164.[CrossRef][Web of Science]
Peressini G., Mazzucchelli M., Rivalenti G., Hoffmann A. W. (2004) Triassic emplacement of the External Gabbro Unit of the Finero mafic complex: U–Pb SHRIMP zircon ages and their implication for the Ivrea–Verbano zone, Western Italian Alps. Geophysical Research Abstracts 6:05072.
Pertermann M. and Hirschmann M. M. (2003) Anhydrous partial melting experiments on MORB-like eclogite: phase relations, phase compositions and mineral–melt partitioning of major elements at 2–3 GPa. Journal of Petrology 44:2173–2201.
Peters T. (1968) Distribution of Mg, Fe, Al, Ca and Na in coexisting olivine, orthopyroxene and clinopyroxene in the Totalp serpentinite (Davos, Switzerland) and in the Alpine metamorphosed Malenco serpentinite (N. Italy). Contributions to Mineralogy and Petrology 18:65–75.[CrossRef]
Piccardo G. B. (1976) Petrologia del massiccio lherzolitico di Suvero (La Spezia). Ofioliti 1:279–317.
Piccardo G. B., Rampone E., Vannucci R. (1990) Upper mantle evolution during continental rifting and ocean formation: evidence from peridotite bodies of the Western Alpine–Northern Apennine system. Mémoires de la Société Géologique de France 156:326–333.
Piccardo G. B., Müntener O., Zanetti A., Pettke T. (2004a) Ophiolitic peridotites of the Alpine–Apennine system: mantle processes and geodynamic relevance. International Geological Review 46:1119–1159.[CrossRef]
Piccardo G. B., Müntener O., Zanetti A. (2004b) Alpine–Apennine ophiolitic peridotites: new concepts on their composition and evolution. Ofioliti 26:193–200.
Poggi E., Piccardo G. B., Zanetti A. (2005) Melt percolation and impregnation in peridotites from a fossil rifted margin: evidence from the Mt. Nero peridotite (External Ligurides, Northern Apennine, Italy). Geophysical Research Abstracts 7:00464.
Rampone E. and Piccardo G. B. (2000) In Dilek J., Moores E., Elthon D., Nicolas A. (Eds.). The ophiolite–oceanic lithosphere analogue: new insights from the Northern Apennines (Italy). Ophiolites and Oceanic Crust: New Insights from Field Studies and Ocean Drilling Program. Geological Society of America Special Papers 349:21–34.
Rampone E., Piccardo G. B., Vannucci R., Bottazzi P., Ottolini L. (1993) Subsolidus reactions monitored by trace element partitioning: the spinel- to plagioclase-facies transition in mantle peridotites. Contributions to Mineralogy and Petrology 115:1–17.[CrossRef][Web of Science]
Rampone E., Hoffmann A. W., Piccardo G. B., Vannucci R., Bottazzi P., Ottolini L. (1995) Petrology, mineral and isotope geochemistry of the External Liguride peridotites (northern Apennines, Italy). Journal of Petrology 36:81–105.
Rampone E., Romairone A., Abouchami W., Piccardo G. B., Hoffmann A. W. (2005) Chronology, petrology and isotope geochemistry of the Erro–Tobbio peridotites (Ligurian Alps, Italy): records of Late Palaeozoic lithospheric extension. Journal of Petrology 46:799–827.
Ricci C. A. and Serri G. (1975) Evidenza geochimica sulla diversa affinità petrogenetica delle rocce basiche comprese nella serie a facies toscana. Bollettino della Società Geologica Italiana 94:1187–1198.
Röhm A. H. E., Snieder R., Goes S., Trampert J. (2000) Thermal structure of continental upper mantle inferred from S-wave velocity and surface heat flow. Earth and Planetary Science Letters 181:395–407.[CrossRef][Web of Science]
Salters V. J. M. and Dick H. J. B. (2001) Mineralogy of the mid-ocean-ridge basalt source from neodymium isotopic composition of abyssal peridotites. Nature 418:68–72.
Scherer E. E., Cameron K. L., Blichert-Toft J. (2001) Lu–Hf garnet geochronology: closure temperature relative to the Sm–Nd system and the effects of trace mineral inclusions. Geochimica et Cosmochimica Acta 64:3414–3432.
Schmidt M.W. and Poli S. (1998) Experimentally based water budgets for dehydrating slabs and consequences for arc magma generation. Earth and Planetary Science Letters 163:361–379.[CrossRef][Web of Science]
Schott B. and Schmeling H. (1998) Delamination and detachment of a litospheric root. Tectonophysics 296:225–247.[CrossRef][Web of Science]
Sibson R. H. (1987) Earthquake rupturing as a mineralizing agent in hydrothermal systems. Geology 15:701–704.
Simakov S. (2006) Redox state of eclogites and peridotites from sub-cratonic upper mantle and a connection with diamond genesis. Contributions to Mineralogy and Mineralogy 151:282–296.
Simakov S. and Taylor L. A. (2000) Geobarometry for mantle eclogites: solubility of Ca-Tschermak in clinopyroxene. International Geological Review 42:534–544.
Skelton A. D. L. and Valley J. W. (2000) The relative timing of serpentinization and mantle exhumation at the ocean–continent transition, Iberia: constraints from oxygen isotopes. Earth and Planetary Science Letters 178:127–138.
Snow J. E., Schmidt G., Rampone E. (2000) Os isotopes and highly siderophile elements (HSE) in the Ligurian ophiolites, Italy. Earth and Planetary Science Letters 175:119–132.[CrossRef][Web of Science]
Stähle V., Frenzel G., Michard A., Schneider W. (1990) Zircon syenite pegmatites in the Finero peridotite (Ivrea zone): evidence for a synite from a mantle source. Earth Planet. Sci. Lett 101:196–205.
Stähle V., Frenzel G., Hess J.C., Saupé F., Schmidt S.Th., Schneider W. (2001) Permian metabasalts and Triassic alkaline dykes in the northern Ivrea Zone: clues to the post-Variscan geodynamic evolution of the Southern Alps. Schweizerische Mineralogische und Petrographische Mitteilungen 81:1–21.
Taylor L. A. and Neal C. R. (1989) Eclogites with oceanic crustal and eclogites, Udachnaya kimberlite pipe, Yakutia, Siberia: evidence of mantle signatures from the Bellsbank kimberlite, South Africa, Part I: Mineralogy, petrography and whole-rock chemistry. Journal of Geology 118:91–100.
Terranova R. and Zanzucchi G. (1984) Relationships between ophiolites and flysch sequences in the Mt. Penna–Mt. Ragola area. Ofioliti 6:287–292.
Thirlwall M. F. and Anczkiewicz R. (2004) Multidynamic isotope ratio analysis using MC–ICP–MS and the causes of secular drift in Hf, Nd and Pb isotope ratios. International Journal of Mass Spectroscopy 235:9–81.
Thompson R. N. (1974) Some high-pressure pyroxenes. Mineralogical Magazine 39:768–787.[Web of Science]
Treves B. E. and Harper G. D. (1994) Exposure of serpentinites on the ocean floor; sequence of faulting and hydrofracturing in the Northern Apennine ophicalcites. In Principi G. (Ed.). Sedimentary Cover of Ophiolitic and Oceanic Sequences. Ofioliti 19: pp. 435–466.
Tribuzio R., Thirlwall M. F., Messiga B. (1999) Petrology, mineral and isotope geochemistry of the Sondalo gabbroic complex (Central Alps, Northern Italy): implications for the origin of post-Variscan magmatism. Contributions to Mineralogy and Petrology 136:48–62.[CrossRef][Web of Science]
Tribuzio R., Tiepolo M., Vannucci R. (2000) Evolution of gabbroic rocks from the Northern Apennine ophiolites (Italy): comparison with the lower oceanic crust from modern slow-spreading ridges. In Dilek J., Moores E., Elthon D., Nicolas A. (Eds.). Ophiolites and Oceanic Crust: New Insights from Field Studies and Ocean Drilling Program. Geological Society of America, Special Papers 349: pp. 129–138.
Tribuzio R., Thirlwall M., Vannucci R. (2004) Origin of the gabbro–peridotite association form the Northern Apennine ophiolites (Italy). Journal of Petrology 45:1109–1124.
Tribuzio R., Montanini A., Thirlwall M. F. (2005) Geochemistry of garnet pyroxenite layers from the External Liguride mantle peridotites (Northern Apennine, Italy): evidence for recycling of crustal material in a ophiolitic mantle sequence. Ofioliti 30:222 Peridotite Workshop 2005, Lanzo (Torino), 27–30 September.
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 Planetary Science Letters 210:509–526.
Ulmer P. and Trommsdorff V. (1999) Phase relations of hydrous mantle subducting to 300 km. In Fei Y., Bertka C.M., Mysen B.O. (Eds.). Mantle Petrology: Field Observations and High Pressure Experimentation: A Tribute to Francis R. Boyd, Geochemical Society Special Publication 6: pp. 259–281.
Van der Wal V. and Vissers R. L. M. (1996) Structural petrology of the Ronda peridotite, SW Spain: deformation history. Journal of Petrology 37:23–43.
Van Orman J. A., Grove T. L., Shimizu N., Layne G. D. (2002) Rare earth element diffusion in a natural pyrope single crystal at 2·8 GPa. Contributions to Mineralogy and Petrology 142:416–424.[Web of Science]
Villa I. M., Hermann J., Müntener O., Trommsdorff V. (2000) 39Ar–40Ar dating of multiply zoned amphibole generations (Malenco, Italian Alps). Contributions to Mineralogy and Petrology 140:363–381.[CrossRef][Web of Science]
Vissers R. L. M., Drury M. R., Hoogerduijn Strating E. H., Van der Wal D. (1991) Shear zones in the upper mantle: a case of study in an Alpine lherzolite massif. Geology 19:990–993.
Vissers R. L. M., Drury M. R., Hoogerduijn S, trating E. H., Spiers C. J., van der Wal D. (1995) Mantle shear zones and their effect on lithosphere strength during continental breakup. Tectonophysics 249:155–171.[CrossRef][Web of Science]
Voshage H., Hofmann A. W., Mazzucchelli M., Rivalenti G., Sinigoi S., Raczek I., Demarchi G. (1990) Isotopic evidence from the Ivrea Zone for a hybrid lower crust formed by magmatic underplating. Nature 347:731–736.[CrossRef]
Webb S. A. C. and Wood B. J. (1986) Spinel–pyroxene–garnet relationship and their dependence on Cr/Al ratio. Contributions to Mineralogy and Petrology 92:471–480.[CrossRef][Web of Science]
Wells P. R. A. (1977) Pyroxene thermometry in simple and complex systems. Contributions to Mineralogy and Petrology 62:129–139.[CrossRef][Web of Science]
White A. J. R. (1964) Clinopyroxenes from eclogites and basic granulites. American Mineralogist 49:883–888.[Web of Science]
Whitmarsh R. B., Manatschal G., Minshull T. A. (2001) Evolution of magma-poor continental margins from rifting to seafloor spreading. Nature 413:150–154.[CrossRef]
Yaxley G. M. and Brey G. P. (2003) Phase relations of carbonate-bearing eclogite assemblages from 2·5 to 5·5 GPa: implications for petrogenesis of carbonatites. Contributions to Mineralogy and Petrology 146:606–619.
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