Journal of Petrology

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Journal of Petrology Volume 43 Number 1 Pages 17-43 2002
© Oxford University Press 2002

Genesis of Pyroxenite-rich Peridotite at Cabo Ortegal (NW Spain): Geochemical and Pb–Sr–Nd Isotope Data

J. F. SANTOS^1,*, U. SCHÄRER², J. I. GIL IBARGUCHI¹ and J. GIRARDEAU³

¹DEPARTAMENTO DE MINERALOGÍA–PETROLOGÍA, UNIVERSIDAD DEL PAÍS VASCO, APTDO. 644, 48080 BILBAO, SPAIN
²LABORATOIRE DE GÉOCHRONOLOGIE, UNIVERSITÉ PARIS 7-IPGP, 2 PLACE JUSSIEU, 75251 PARIS CEDEX 05, FRANCE
³LABORATOIRE DE PLANÉTOLOGIE ET GÉODYNAMIQUE, UNIVERSITÉ DE NANTES, 2 RUE DE LA HOUSSINIÈRE, BP 92208, 44322 NANTES CEDEX 3, FRANCE

Received December 10, 1999; Revised typescript accepted June 25, 2001

	ABSTRACT

TOP ABSTRACT INTRODUCTION GEOLOGICAL SETTING PREVIOUS GEOCHRONOLOGICAL AND... SAMPLE DESCRIPTION MAJOR ELEMENT MINERAL CHEMISTRY WHOLE-ROCK MAJOR AND TRACE... ISOTOPE ANALYTICAL METHODS Pb-Pb AND U-Pb ISOTOPE... Rb-Sr AND Sm-Nd ISOTOPE... DISCUSSION CONCLUSIONS REFERENCES

 
Petrographic and field data indicate the existence of four main rock types within the allochthonous Cabo Ortegal ultramafic units: (1) harzburgites; (2) dunites; (3) massive, occasionally garnet-bearing, pyroxenites; (4) less abundant mafic rocks with variable amounts of garnet-rich pyroxenite. The major and trace element compositions of the analysed ultramafic rocks define well-delimited fields in binary variation diagrams. Normalized trace element patterns, however, exhibit large ion lithophile element (LILE) and light rare earth element (LREE) enrichment that do not correlate with the main rock types distinguished. NiO contents and fo-number of olivine in the harzburgites match those of the mantle olivine array, whereas a fractional crystallization trend is observed from dunites to pyroxenites. Spinel and olivine in the harzburgites have residual characteristics comparable with those of abyssal peridotites or peridotites from arc settings, whereas in most of the dunites and pyroxenites the range of fo-number and Cr/(Cr + Al) ratio suggests crystallization from primitive subduction-related magmas. Whole-rock major and trace element and Pb–Sr–Nd isotope data suggest that regional-scale massive pyroxenites from Cabo Ortegal originated from relatively homogeneous parental melts. Fractional crystallization processes, coeval with intense deformation, might result in the formation of cumulate layers (clinopyroxene, orthopyroxene, olivine, chromite, etc.). Some less abundant mafic rocks and associated pyroxenites are also homogeneous but have different chemical and isotopic signatures suggesting a different parental melt from that of the massive pyroxenites. Although some differences exist in the major element and isotopic composition of the clinopyroxenes, their initial isotopic ratios (²⁰⁶Pb/²⁰⁴Pb = 17·845–18·305, ²⁰⁷Pb/²⁰⁶Pb = 15·433–15·634; ⁸⁷Sr/⁸⁶Sr = 0·70330–0·70476; ¹⁴³Nd/¹⁴⁴Nd = 0·512539–0·512916) suggest involvement of an enriched component in their mantle source, which may be related to the subduction of terrigenous sediments (i.e. EMI). The new data obtained confirm that ultramafic units of Cabo Ortegal experienced a complex tectonothermal history similar to that of other units of the same area and allow us to distinguish at least two different events. Sm–Nd whole-rock–clinopyroxene ages suggest formation of the ultramafic units at 500 Ma, an age similar to that of formation of the protoliths of associated HP/HT units. Internal Sm–Nd isochrons (Cpx–Grt–whole rock) from two pyroxenites indicate ages of 390 Ma for garnet crystallization in these rocks, which is consistent with previous U–Pb dating of HP/HT recrystallization in Cabo Ortegal. Pyroxenite–dunite-rich ultramafic massifs such as Cabo Ortegal might have originated within the lithospheric mantle above a subduction zone, variably modified by fluid or melt infiltration from the subducted oceanic crust.

KEY WORDS: geochemistry; isotopes; Ortegal; pyroxenites; mantle

	INTRODUCTION

TOP ABSTRACT INTRODUCTION GEOLOGICAL SETTING PREVIOUS GEOCHRONOLOGICAL AND... SAMPLE DESCRIPTION MAJOR ELEMENT MINERAL CHEMISTRY WHOLE-ROCK MAJOR AND TRACE... ISOTOPE ANALYTICAL METHODS Pb-Pb AND U-Pb ISOTOPE... Rb-Sr AND Sm-Nd ISOTOPE... DISCUSSION CONCLUSIONS REFERENCES

 
Pyroxenite layers parallel or subparallel to the dominant foliation of the host peridotite generally constitute <10% of the volume of ultramafic massifs (e.g. Fabriès et al., 1991). They are normally centimetres to metres thick but larger pyroxenite bodies within peridotite have been reported from a few localities, e.g. Cabo Ortegal, Kohistan and Tonsina (Girardeau et al., 1989; DeBari & Sleep, 1991; Burg et al., 1998, and references therein). To explain the origin of pyroxenite, either in-folding of subducted oceanic crust through convection into the mantle (Allègre & Turcotte, 1986) or crystallization from a melt derived from a peridotitic mantle source (e.g. Suen & Frey, 1987) has been proposed. Recent work on various pyroxenite-bearing peridotite massifs, however, has left open the question of the origin of the magmas that formed the pyroxenites, and a combination of both processes, including strong interaction with surrounding peridotite, has been considered (Kelemen & Ghiorso, 1986; Downes et al., 1991; Becker, 1996; Zanetti et al., 1999).

In the Cabo Ortegal complex of NW Spain, a concentration of metre-thick websterite layers separated by dunite, locally alternating with garnet-rich mafic rocks, forms a kilometre size pyroxenite body within the Herbeira peridotite massif (Fig. 1). Several explanations exist for this body, including high-pressure crystallization of melts that intruded the solid peridotite at 390 Ma (Girardeau & Gil Ibarguchi, 1991; Santos et al., 1996) or that it is part of a layered complex overlying tectonized mantle harzburgite (Moreno et al., 1999). Controversy exists about the origin of the magma that crystallized to form the pyroxenites. Models have included partial melting of a metasomatized mantle wedge (Gravestock, 1992), of primitive mantle in an oceanic setting (Laribi-Halimi, 1992) or of the mantle wedge underlying an arc root (Moreno et al., 2001). In an attempt to elucidate the origin of the pyroxenites and to investigate mantle heterogeneity on a regional scale, new U–Pb, Pb–Pb, Sm–Nd and Rb–Sr isotopic data have been obtained from mineral separates and whole rocks from selected pyroxenites and peridotites from the Cabo Ortegal complex.

View larger version (52K): [in this window] [in a new window]   Fig. 1. (a) Geological setting of allochthonous complexes in the Palaeozoic orogen of NW Iberia (see Santos et al., 1996). (b) Map of the Cabo Ortegal complex [modified after Abalos et al. (2000)] with sample locations and approximate position of samples along the pyroxenite–dunite section of Herbeira.

 

	GEOLOGICAL SETTING

TOP ABSTRACT INTRODUCTION GEOLOGICAL SETTING PREVIOUS GEOCHRONOLOGICAL AND... SAMPLE DESCRIPTION MAJOR ELEMENT MINERAL CHEMISTRY WHOLE-ROCK MAJOR AND TRACE... ISOTOPE ANALYTICAL METHODS Pb-Pb AND U-Pb ISOTOPE... Rb-Sr AND Sm-Nd ISOTOPE... DISCUSSION CONCLUSIONS REFERENCES

 
The uppermost units of the Cabo Ortegal allochthonous complex comprise ultramafic rocks, granulites, gneisses and eclogites (Fig. 1; Arenas et al., 1986; Gil Ibarguchi et al., 1990). Ultramafic rocks are exposed in the Limo, Herbeira and Uzal massifs and in dispersed outcrops associated with high-pressure granulites. They are mostly harzburgite, pyroxenite and dunite in decreasing order of abundance. Pyroxenites occur as: (1) scattered metre-size layers usually parallel to the main foliation of the host harzburgite; (2) a composite multilayer pyroxenite–dunite unit 3 km long and >500 m thick within the Herbeira peridotite massif. This multilayer sequence may have originated through folding and boudinage of a pyroxenite–dunite-rich fragment of mantle peridotite during shearing related to emplacement of the massif (Girardeau & Gil Ibarguchi, 1991). More recently, the pyroxenites have been viewed as the middle part of an essentially tripartite dunite–pyroxenite–dunite ultramafic layered complex overlying a basement of tectonized harzburgite (Moreno et al., 2001).

New mapping of the Cabo Ortegal complex (Abalos et al., 2000) has revealed that the Herbeira peridotite massif forms an upright asymmetric synform with a thinner gently dipping eastern limb (Fig. 1b). The thickness of ultramafic rocks (normal to the foliation) decreases from 1000–1500 m in the western limb to 750 m in the east. The thickness of the central part of the main pyroxenite–dunite body as seen along the western cliffs of the massif (Fig. 1b) changes from 450 m (S) to 150 m (N). This part is formed essentially of pyroxenite (80–90%) alternating with dunite and less abundant harzburgite. Towards the base and top, this pyroxenite-rich portion changes into a dunite-rich unit in which pyroxenites and harzburgite occur in greatly reduced amounts. The thickness of the lower dunite-rich unit is 100 m, whereas the upper dunite is >400 m thick. Towards the eastern limb of the massif, both dunite-rich sections appear to coalesce in a single outcropping dunite-rich unit 100 m thick (Fig. 1b). The rest of the Herbeira massif is composed of harzburgite up to 500 m in thickness below the basal dunite and 200 m thick above the upper dunite. The wedge shape of the units described is not as spectacular as the thickness variations described seem to imply. The calculated angle between the top and the base of any of the dunite or pyroxenite units is 5°.

The central pyroxenite-rich section is composed of clinopyroxenite and websterite layers up to several metres thick, separated by dunite and less often by harzburgite. Garnet-bearing pyroxenites occur randomly throughout the pyroxenite–dunite section, whereas olivine websterite, orthopyroxenite and phlogopite-bearing varieties occur in minor amounts. Towards the top of the pyroxenite-rich unit, and also at one locality within the harzburgite (samples GRCP6, GZC7, GZA2; Fig. 1b), decimetre- to metre-thick layers of garnet–zoisite-rich mafic rocks and garnet–rutile-rich pyroxenite occur. Peridotites and pyroxenite layers are locally enriched in amphibole (to >30% modal) and cut by late garnet-rich and pyroxenite veins. Girardeau & Gil Ibarguchi (1991), on the basis of major element mineral data, proposed an affinity of the peridotites with residual oceanic tectonites and suggested that the pyroxenites correspond to the products of high-pressure crystal fractionation in dykes or sills intruded into the peridotite during subduction. The peridotite–pyroxenite association from the equivalent Bragança complex in north Portugal (Fig. 1) has also been interpreted as magma intrusion into a depleted mantle protolith in a supra-subduction zone setting (Bridges et al., 1995). In Cabo Ortegal the pyroxenites, peridotites and mafic rocks have a common tectonothermal evolution with nearby granulites and eclogites (Girardeau & Gil Ibarguchi, 1991; Abalos et al., 1996). They underwent, however, a more complex tectonothermal history involving a first episode of shear deformation at high temperature (>1000°C) and relatively low deviatoric stress (300–600 bars), followed by late- to syn-kinematic development of garnet after spinel in pyroxenites (including olivine websterite) and mafic rocks at 800°C, 1·65 GPa. Amphibole was probably produced during subsequent decompression at relatively high temperature (>750°C). A comprehensive account of the petrography and structure of these rocks has been given by Girardeau & Gil Ibarguchi (1991).

	PREVIOUS GEOCHRONOLOGICAL AND GEOCHEMICAL DATA

TOP ABSTRACT INTRODUCTION GEOLOGICAL SETTING PREVIOUS GEOCHRONOLOGICAL AND... SAMPLE DESCRIPTION MAJOR ELEMENT MINERAL CHEMISTRY WHOLE-ROCK MAJOR AND TRACE... ISOTOPE ANALYTICAL METHODS Pb-Pb AND U-Pb ISOTOPE... Rb-Sr AND Sm-Nd ISOTOPE... DISCUSSION CONCLUSIONS REFERENCES

 
The age of the Cabo Ortegal peridotite unit was poorly constrained at 477 ± 122 Ma by Van Calsteren et al. (1979; Rb–Sr in whole rock). Rutile from pyroxenite and magmatic zircon from pyroxenite and peridotite yielded more precise U–Pb ages of 383–395 Ma (Peucat et al., 1990; Santos et al., 1996; Ordóñez Casado et al., 2001), and a granitic pegmatite intrusive rock in the Uzal peridotite yielded 385–395 Ma (U–Pb monazite and zircon ages and Rb–Sr isochron; Santos et al., 1996). This substantiates a magmatic event around 390 ± 10 Ma nearly coeval with HP/HT metamorphism dated at 382–395 Ma on adjacent eclogites, granulites and HP gneisses (U–Pb zircon, titanite and rutile ages, Santos et al., 1996; Valverde Vaquero & Fernández, 1996; Ordóñez Casado et al., 2001). Phlogopite– and amphibole–whole-rock pairs from harzburgites have been dated at 386 ± 10 and 337 ± 10 Ma, respectively (Rb–Sr age, Van Calsteren et al., 1979).

Previous geochemical data for Cabo Ortegal ultramafic rocks have been reported by Vogel (1967), Maaskant (1970), Van Calsteren et al. (1979), Peucat et al. (1990), Gravestock (1992) and Laribi-Halimi (1992). On the basis of major and limited trace element and Sr isotope data, and the field association with high-temperature metabasites and gneisses, the origin of the ultramafic rocks was considered by Van Calsteren et al. (1979) to be related to Lower Palaeozoic mantle plume activity within an intra-continental setting. For those workers, the peridotites correspond to slightly depleted spinel lherzolites whose partial melting products, represented in part by the pyroxenites, would have crystallized in place during ascent and emplacement of a mantle diapir at the base of the continental crust. In contrast, later results, including Nd isotope data (Peucat et al., 1990) were interpreted in terms of an oceanic mantle setting modified with crustal components for the peridotite–pyroxenite association.

Gravestock (1992) used Pb, Sr and Nd isotope data together with modelling of HFSE and REE compositions to propose that the Cabo Ortegal pyroxenites represent crystal cumulates. However, they could not be related to parent magmas by a simple igneous fractionation process and Gravestock suggested a depleted source in a mantle-wedge environment metasomatized by a low-viscosity carbonatite melt as a likely scenario. Using mainly major and trace element whole-rock data, Laribi-Halimi (1992) proposed variable degrees of partial melting of primitive mantle in relation to mid-ocean ridge basalt (MORB) magma generation for the origin of the peridotites, and fractional crystallization of tholeiitic magmas for the pyroxenites. Finally, on the basis of the platinum-group element (PGE) contents of the ultramafic units, Moreno et al. (2001) considered the existence of two batches of mantle-derived magmas from which the pyroxenite–dunite association crystallized at the crust–mantle interface of an arc-root region.

	SAMPLE DESCRIPTION

TOP ABSTRACT INTRODUCTION GEOLOGICAL SETTING PREVIOUS GEOCHRONOLOGICAL AND... SAMPLE DESCRIPTION MAJOR ELEMENT MINERAL CHEMISTRY WHOLE-ROCK MAJOR AND TRACE... ISOTOPE ANALYTICAL METHODS Pb-Pb AND U-Pb ISOTOPE... Rb-Sr AND Sm-Nd ISOTOPE... DISCUSSION CONCLUSIONS REFERENCES

 
Harzburgite
Intensely deformed and serpentinized (up to >70% serpentine) amphibole-bearing harzburgite (samples 432, H3, Table 1) constitutes most of the ultramafic outcrops. These rocks were previously described as lherzolite and the amphibole was interpreted as resulting from transformation of primary clinopyroxene during retrograde amphibolite-facies metamorphism (Vogel, 1967; Maaskant, 1970; Van Calsteren, 1978). Indeed, they are comparable with mantle lherzolites worldwide, in their high contents of CaO (1·7–4·3%) and Al₂O₃ (2·2–4·1%; Table 1). However, they are clearly harzburgites in terms of their modal composition (Table 1). Harzburgites contain sub-spherical, 5–10 mm, orthopyroxene porphyroclasts or aggregates in a matrix of olivine and amphibole, minor clinopyroxene as isolated crystals or as part of discontinuous microlayers of millimetre thickness, and green to brown spinel. Orthopyroxene accounts for 10–20% of the rock volume and generally contains exsolution lamellae of clinopyroxene parallel to (100). More rarely, orthopyroxene contains exsolution of Al-rich spinel towards the grain margins, which may result in the formation of minute spinel grains along crystal boundaries. Olivine forms 1–3 mm porphyroclasts and <0·5 mm neoblasts. Clinopyroxene (<2% modal) forms anhedral porphyroclasts rich in fine exsolution lamellae of orthopyroxene. Green and brown subidiomorphic spinel up to 5 mm in diameter forms 1–3% of the rock volume. It is usually disseminated and more rarely forms discontinuous aggregates parallel to the main foliation.

View this table: [in this window] [in a new window]   Table 1: Modal composition of Cabo Ortegal ultramafic and related rocks

 

Harzburgite may be rich (10–30 modal %) in subidiomorphic, unzoned, pale green magnesio-hornblende and more rarely pargasitic to edenitic hornblende in textural equilibrium with olivine and pyroxene. Amphibole may contain inclusions of brown–green spinel and often forms aggregates or thin microlayers parallel to the main foliation. Within fine-grained, mylonitized and strongly deformed rocks, amphibole has grown syn- to post-kinematically around or in pressure shadows of orthopyroxene porphyroclasts. More rarely, hornblende has the features of a phase formed from exsolution lamellae in orthopyroxene. Tremolitic amphibole occurs replacing hornblende and orthopyroxene along grain margins. Phlogopite in textural equilibrium with olivine, pyroxene and amphibole is a rare accessory. Rare magnetite may occur associated with Ni and Fe sulphides, and millimetre-size chlorite, often associated with magnetite, is locally present defining a late schistosity.

Dunite
Dunite within the pyroxenite–dunite section of Herbeira is found as (1) centimetre- to decimetre- thick bands alternating with pyroxenite and, less often, with harzburgite bands of similar thickness, in the lower and middle parts of the section, and (2) outcrops of several metres in size that include scattered pyroxenite layers in the upper portions of the section. Less commonly, dunite forms metre- to decametre-size pods with diffuse contacts within harzburgite. Pyroxenites are generally more abundant (up to 30%) in orthopyroxene- and amphibole-bearing varieties of dunite. Massive dunite shows a foliation defined by elongated crystals of spinel and is generally less serpentinized than harzburgite. Massive dunite and dunite bands contain millimetre-size olivine porphyroclasts and <0·5 mm neoblasts as in harzburgite, sporadic, small (1 mm), interstitial orthopyroxene and large (3 mm) interstitial clinopyroxene grains. Subidiomorphic millimetre-size chromian spinel occasionally associated with platinum-group minerals may concentrate in pods or layers >10 cm thick (Monterrubio et al., 1992). In this way, Moreno et al. (2001) have distinguished the PGE-poor lower dunites, which are considered genetically related to the massive pyroxenites, from the PGE-rich upper dunites. Unzoned edenitic- to magnesio-hornblende occurs in minor amounts. Phlogopite is an accessory in some samples. Secondary serpentine, chlorite and magnetite are common. Dunite was not analysed for whole-rock composition but only for mineral composition in this study.

Massive pyroxenites and pyroxenite bands
Pyroxenites, which generally form homogeneous layers with sharp contacts to peridotite, mainly comprise medium-grained, orthopyroxene-rich, websterite and clinopyroxenite in equal amounts. Pegmatoid facies containing large orthopyroxene crystals (>5 cm) occur sporadically, and coarse-grained (2–5 mm) orthopyroxenite locally forms massive bands up to 30 cm thick. Websterites may show a poor layering marked by discontinuous thin orthopyroxenite layers and are occasionally zoned with olivine-rich rims and finer-grained, clinopyroxene-rich central parts. Garnet pyroxenites occur randomly throughout the pyroxenite–dunite section. Samples of orthopyroxenite (OP10), websterite (136, 198a), garnet websterite (GW1, GW7692), olivine websterite (196c), and clinopyroxenite (430a, CP11) have been studied (Table 1).

Pyroxenites are usually very fresh rocks and exhibit poorly oriented granoblastic textures. Orthopyroxene is often strained with evidence of dynamic recrystallization at grain boundaries. Clinopyroxene appears undeformed, except for local undulose extinction and mechanical twins. Olivine and idiomorphic green spinel in disseminated grains, more rarely forming elongated aggregates, are not abundant; spinel exsolution in pyroxene is exceedingly rare. Optically homogeneous amphibole may constitute >30% by volume in some samples, forming generally large (>1 cm) subidiomorphic crystals without deformation. This amphibole is in textural equilibrium with the pyroxenes although, occasionally, it may have developed after clinopyroxene. Garnet often occurs in coronas around spinel and/or ilmenite and may contain sporadic inclusions of corundum. Garnet may also form aggregates of subidiomorphic crystals with or without spinel relicts. Carbonates, close to pure dolomite and calcite, are rare accessories in apparently textural equilibrium with pyroxene and garnet. Interstitial phlogopite associated with ilmenite is also a rare accessory except in pyroxenites near late tectonic contacts, where it may be abundant. Garnet websterite GW7692 was originally interpreted as a vein cross-cutting the foliation at Uzal (sample 7692, Peucat et al., 1990). Actually, this sample comes from a strongly deformed portion of the Uzal peridotite massif where structural relations are obscured. As it is identical to garnet pyroxenites elsewhere within the pyroxenite–dunite section and pyroxenites parallel to the foliation in harzburgite, and very different from the late pyroxenite veins cross-cutting the foliation (see below), it has been included in the present study as belonging to the common pyroxenite unit.

Mafic rocks and related pyroxenites
Rare mafic rocks with variable amounts of associated garnet-rich pyroxenites occur at two localities near the top of the pyroxenite-rich section lying parallel to the pyroxenite–dunite banding. Garnet- and zoisite-rich metre-thick outcrops are composed of bands of 5–20 cm thickness containing either more clinopyroxene than amphibole (288a, 288b, GZC7), or more amphibole and practically no clinopyroxene (288, 431a, 431b, GZA2; Table 1). These rocks exhibit oriented granoblastic textures with millimetre-size garnet, often with inclusions of Cr-rich spinel and pink corundum, forming up to 70% of the rock volume. Subidiomorphic millimetre- to centimetre-size zoisite in textural equilibrium with pyroxene and amphibole is present in all cases and may constitute up to 40% of the rock volume (GZA2; Table 1). Unzoned amphibole in textural equilibrium is magnesio-hornblende and more rarely tschermakitic hornblende. K-feldspar and plagioclase have been found as accessories in one sample. Kyanite inclusions in subidiomorphic zoisite were also found in one sample. A gradual transition exists between the garnet–zoisite rocks and garnet–rutile clinopyroxenite (GRCP6; Table 1). This is composed essentially of clinopyroxene and garnet in textural equilibrium (60 and 30%, respectively), subidiomorphic, prismatic to rounded rutile and lesser amounts of ilmenite. Amphibole, locally accompanied by clinozoisite, is present in small amounts interstitially and along fractures within clinopyroxene and garnet, whereas rutile is occasionally replaced by titanite.

Other rock types
Late granitic intrusive rocks, pyroxenite dykes and garnet-rich dykes are less abundant rock types. The main granitic intrusive type is a pegmatite granite (OGS23) several metres thick and >30 m long within mylonitic harzburgite of the Uzal massif at the Cartés beach locality (Santos et al., 1996). The granite pegmatite is of the same age as the pyroxenites (see above). It contains decimetre- to centimetre-size blocks of peridotite with rims of chlorite and actinolite. The pegmatite exhibits a porphyroclastic texture with a foliation parallel to that in the host peridotite and is composed of centimetre-size K-feldspar, muscovite and less abundant plagioclase (An₁₅), surrounded by a fine-grained matrix composed of the same minerals, phlogopite and quartz. Allanite, zircon, apatite and monazite are accessories. Decimetre-size granitic pods also occur at several places within the Herbeira peridotite massif. These are formed of plagioclase, quartz and biotite, with reaction rims of Mg-arfvedsonite and wollastonite at the contact with the peridotite; allanite, zircon and apatite are accessories.

Sparse mafic dykes occur cutting through massive pyroxenites and surrounding peridotites. These dykes, generally <10 cm thick, are medium grained (3–5 mm), may display orthopyroxene-rich outer zones and clinopyroxene-rich cores and, when included in harzburgite, dunitic rims. Hornblende, olivine and phlogopite occur in minor amounts. The garnet-rich dykes, sometimes filled only by garnet with spinel cores, are thinner and generally zoned, displaying thin clinopyroxene- or amphibole-rich rims and garnet-rich central parts. Amphibole and chlorite are abundant; carbonates, phlogopite, titanian spinel and ilmenite are rare accessories. Preliminary geochemical data by Gravestock (1992), including mineral isotope compositions, suggest that these rocks are genetically unrelated to the main peridotite–pyroxenite unit.

	MAJOR ELEMENT MINERAL CHEMISTRY

TOP ABSTRACT INTRODUCTION GEOLOGICAL SETTING PREVIOUS GEOCHRONOLOGICAL AND... SAMPLE DESCRIPTION MAJOR ELEMENT MINERAL CHEMISTRY WHOLE-ROCK MAJOR AND TRACE... ISOTOPE ANALYTICAL METHODS Pb-Pb AND U-Pb ISOTOPE... Rb-Sr AND Sm-Nd ISOTOPE... DISCUSSION CONCLUSIONS REFERENCES

 
Mineral compositions were determined by electron microprobe methods using Camebax Micro and SX50 instruments at Clermont-Ferrand and Paris-VII universities, respectively. Mineral analyses with details of analytical procedures, structural formulae and end-member calculations have been presented elsewhere by Girardeau & Gil Ibarguchi (1991). An interpretation of the analytical data in terms of the P–T evolution undergone by the ultramafic rocks, stressing the residual character of the mineralogy of peridotites, was presented in that study. A refined interpretation aimed at the origin of peridotites and pyroxenites is presented here, including new analytical data. Minerals from 46 samples of pyroxenite, eight dunites, one harzburgite and two mafic rocks spanning the whole 500 m thickness of the pyroxenite–dunite section, and 16 harzburgites, eight pyroxenites and one dunite from the Herbeira, Limo and Uzal peridotite massifs have been analysed. The complete dataset is available from the Journal of Petrology Web site at http:www.petrology.oupjournals.org.

Olivines from harzburgite, dunite and pyroxenite are unzoned and have fairly constant low Ca, Al, Cr and Ti contents across the pyroxenite–dunite section of Herbeira. Olivine in pyroxenites has distinctly lower average fo-number and NiO contents (0·85 and 0·19% NiO, respectively) than olivine in dunite and harzburgite. Olivine in the harzburgite has similar fo-number (0·9) but higher Ni content (0·37% NiO, as average) than olivine in dunite (0·29% NiO, as average) and it overlaps the ‘mantle olivine array’ (Fig. 2a). In turn, olivines from dunites and pyroxenites deviate from the ‘mantle olivine array’ and define broader fields than harzburgite olivine, involving a rapid decrease of NiO content and fo-number from dunite to pyroxenite (Fig. 2b).

View larger version (21K): [in this window] [in a new window]   Fig. 2. (a) NiO vs fo-number of olivine for ultramafic rocks of Cabo Ortegal. Detail for dunitic and harzburgitic rocks [see box in (b)]. (b) Data for pyroxenitic and dunitic rocks. Mantle olivine array from Takahashi et al. (1987). I and II are olivine fractionation trends calculated by Varfalvy et al. (1996) and Niida (1997), respectively. ‘Reaction’ is the trend of replacive olivines produced by open-system reaction between harzburgite and infiltrating melts with evolved composition (Niida, 1997). (c) Cr/(Cr + Al) spinel vs fo-number of olivine for ultramafic rocks of Cabo Ortegal. ‘Olivine–spinel mantle array’ (OSMA) and Alpine-type field from Arai (1994); abyssal peridotite field from Arai et al. (1998).

 

Spinel compositions are generally Al-rich in harzburgite, Cr-rich in dunite and range from Al- to Cr-rich in pyroxenites. Limited variations in spinel composition may be due in part to late re-equilibration processes as suggested by textural and colour changes of this mineral (Girardeau & Gil Ibarguchi, 1991). In a plot of Cr/(Cr + Al) of spinel and fo-number of olivine, dunites and harzburgites plot in separate fields and do not follow the ‘olivine–spinel mantle array’ (Fig. 2c). Olivines in the pyroxenites are characterized by lower Fo contents than those in harzburgites and dunites over a wide range in spinel Cr/(Cr + Al) plotting outside the olivine–spinel array (Fig. 2c).

Clinopyroxene from peridotites and pyroxenites is a Cr-, Ti- and Na-poor diopside. Despite an overall similarity, there exist some differences in composition among clinopyroxenes from different rock types. In a plot of Cr vs Na (Fig. 3a), clinopyroxenes from massive pyroxenites and dunites define a broad trend with an increase in Cr and Na towards dunite. Clinopyroxenes from harzburgites have variable Cr and Na contents generally overlapping the lowest values from pyroxenites and dunites, and clinopyroxenes from mafic rocks and related pyroxenites are Cr-poor with variable, generally higher, Na contents. In a plot of Al_Z vs TiO₂, the whole set of clinopyroxene analyses shows a limited compositional range with clinopyroxenes from mafic rocks and related pyroxenites exhibiting generally higher Al_Z contents (Fig. 3b).

View larger version (32K): [in this window] [in a new window]   Fig. 3. (a) Cr x 10 vs Na x 10 (p.f.u.) variation in clinopyroxenes from the pyroxenite–dunite section of Cabo Ortegal. Subcontinental peridotite and Type II ophiolite mantle fields from Kornprobst et al. (1981). Clinopyroxenes from harzburgites are shown as a compositional field for clarity. (b) TiO2 vs AlZ in clinopyroxene. Fields for arc cumulates (Tonsina, Jijal, Peninsular Ranges, Alaska, Lesser Antilles, Aleutians), ophiolite layered cumulates (Bay of Islands, Canyon Mountain and Samail) and Mid-Atlantic Ridge cumulates from Loucks (1990).

 

Amphibole from harzburgites, dunites and pyroxenites is unzoned magnesio-hornblende and, more rarely, tschermakitic-hornblende, with low contents of Ti, Fe and Mn, and moderate contents of K, Na and Cr. Amphiboles from dunites are generally richer in Na (1·6–2·4% Na₂O) and Cr (1–1·6% Cr₂O₃) than those of pyroxenites (1·1–2·1% Na₂O, 0·3–1·6% Cr₂O₃), whereas amphiboles from harzburgites have intermediate Na and Cr contents (1·2–2% Na₂O, 0·5–1·2% Cr₂O₃). All these amphiboles are characterized by high Al^[4]/(Na + K)_A ratios suggesting significant pargasite (edenite + tschermakite) substitution.

	WHOLE-ROCK MAJOR AND TRACE ELEMENT COMPOSITIONS

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Three rock groups appear well characterized in variation diagrams of MgO vs other elements (Fig. 4, Table 1). Harzburgite and dunite have very similar compositions and plot at the Mg-rich end (>30% MgO), massive pyroxenites show limited overlap with peridotites and lower MgO values (down to 16·7%), and mafic rocks and associated pyroxenites have MgO < 16% (Table 1). Compared with pyroxenites from ultramafic massifs elsewhere, e.g. Horoman, Ronda, Beni Boussera, Alps or Pyrenees, massive pyroxenites from Herbeira show very different behaviour with generally higher SiO₂, CaO, Cr and Sc contents and lower Al₂O₃, TiO₂ and Ni contents. Only the D1 subtype of websterites from the Ronda massif are found to be relatively similar, although, as in the other cases, CaO/Al₂O₃ ratios in Cabo Ortegal are significantly higher (Fig. 4).

View larger version (48K): [in this window] [in a new window]   Fig. 4. Selected major and trace element variation in bulk rock vs weight percent MgO for ultramafic and related rocks of Cabo Ortegal (data from this study, and from Vogel, 1967; Maaskant, 1970; Van Calsteren et al., 1979; Peucat et al., 1990; Laribi-Halimi, 1992); vertically and horizontally ruled areas indicate pyroxenites and peridotites, respectively, from Gravestock (1992). Other fields include data from Beni Boussera (Pearson et al., 1993; Kornprobst, 1969), Horoman (Yoshida & Takahashi, 1997; Takazawa et al., 1999, 2000), Pyrenees (Fabriès & Conquéré, 1983; Conquéré & Fabriès, 1984), Alps (Capredi et al., 1977; Garuti et al., 1980; Rivalenti et al., 1980; Pognante et al., 1985) and Ronda (Obata et al., 1980; Garrido & Bodinier, 1999).

 

Normalization to a primitive mantle composition shows that all the analysed rocks have slight LILE enrichment, with values intermediate between those of N-MORB and E-MORB, and heavy REE (HREE) depletion, except GRCP6 (Fig. 5a; Tables 2 and 3). Ti exhibits a variable character, samples with a negative Ti anomaly being richer in La to Sr than other samples and corresponding to clinopyroxene-rich rocks as previously shown by Gravestock (1992), whereas orthopyroxene-rich rocks show the opposite trend with a positive Ti anomaly, in agreement with the results by Rampone et al. (1991) on Ti partition in orthopyroxene and clinopyroxene. Other significant features are the generalized positive anomalies for Zr, Sr and U and negative anomaly for K, the higher contents of U respect to Th, and the enrichment in Nb with respect to mantle values (Fig. 5a).

View larger version (30K): [in this window] [in a new window]   Fig. 5. (a) Primitive mantle-normalized (Sun & McDonough, 1989) multielement diagram for Cabo Ortegal samples. (b) and (c) show chondrite-normalized (Evensen et al., 1978) REE patterns for Cabo Ortegal pyroxenites and peridotites. Field for D1 pyroxenite subtype of Ronda from Garrido & Bodinier (1999). E-MORB and N-MORB compositions from Sun & McDonough (1989).

 

View this table: [in this window] [in a new window]   Table 2: Whole-rock major and trace element compositions of Cabo Ortegal ultramafic and related rocks

 

View this table: [in this window] [in a new window]   Table 3: REE data for Cabo Ortegal ultramafic and related rocks

 

REE contents (Table 3) normalized to chondrite values show generally variable enrichment in LREE [(La/Sm)_n between 1·45 and 20] with small positive Eu anomalies (Eu/Eu^* from 1·08 to 2·26) and nearly flat HREE patterns [(Gd/Lu)_n 0·94–1·49], a feature noted previously by Peucat et al. (1990) and Gravestock (1992). In detail, samples H3, GW1 and GRCP6 have relatively flat patterns with variable Eu anomalies and concentrations (Fig. 5b), whereas samples OP10, CP11 and GZC7 present distinct LREE-enriched patterns and positive Eu anomalies (Fig. 5c). The latter patterns are similar to those of pyroxenites analysed by Gravestock (1992) from the Herbeira massif and to the analysis of harzburgite also from Herbeira reported by Peucat et al. (1990).

Trace element contents of the Cabo Ortegal rocks thus appear unique and different from those of other pyroxenites in ultramafic massifs, including the D1 subtype from Ronda, although locally there may be pyroxenites of composition resembling those from Ortegal (e.g. the low-MgO pyroxenites from Lower Austria; Becker, 1996). Almost all other pyroxenites in ultramafic massifs are LREE depleted and have much higher HREE contents (e.g. Downes et al., 1991). Cabo Ortegal pyroxenites have very low HREE contents, LREE enrichment and higher contents of Zr, Nb and U. Although whole-rock equivalents can hardly be found, a close similarity may be noted with the composition of REE-enriched clinopyroxenes from peridotite xenoliths of Hungary and the French Massif Central (Downes & Dupuy, 1987; Downes et al., 1992).

	ISOTOPE ANALYTICAL METHODS

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Mineral separates were obtained by processing >10 kg samples through crushing, disc-mill, Wilfley table, Frantz magnetic separator and heavy liquids. Each mineral fraction was selected grain by grain, including crack- and inclusion-free grains to the highest transparency possible. Before dissolution all fractions were leached following a two-step process (modified after Polvé, 1983), using first weak HF (2%) at room temperature (5 min), then ultrasound (5 min), HCl (2N) at 100–120°C (30 min), ultrasound (5 min), and finally triple distilled water at 100–120°C (30 min). Samples were dissolved in HF (48%) and, if necessary, the residue was treated in Teflon bombs at 220°C, before processing with HCl. U–Pb and Rb–Sr isotope analyses were performed at the Laboratory of Geochronology of the University of Paris-VII with a Thompson 206 mass spectrometer with ⁸⁷Sr/⁸⁶Sr = 0·71025 ± 4 (2, n = 33) for the Sr standard NBS-987, whereas Sm–Nd analyses were performed at the University of the Basque Country using a Finnigan MAT262 system with ¹⁴³Nd/¹⁴⁴Nd = 0·5118435 ± 13 (2, n = 18) for La Jolla standard.

For U and Pb analysis, isotope dilution methods (²⁰⁸Pb–²³⁵U–²³³U mixed isotope tracer) through aliquot solutions were utilized following a procedure modified after Manhès et al. (1987). Pb and U were loaded together on the same Re filament using Si gel and H₃PO₄ (Schärer & Gower, 1988). Mass discrimination for U and Pb was 0·1% per a.m.u. on both Faraday and secondary electron multiplier systems, and total blanks were 15 pg for Pb and <1 pg for U. Rb and Sr analysis was carried out with an ⁸⁷Rb–⁸⁴Sr mixed isotope tracer in aliquot solutions. Chemical procedures for Rb–Sr extraction and separate loading with a Ta mixture on Re filaments were described by Birck (1986). Isotope ratios were measured with a secondary electron multiplier for Rb and Sr concentrations, and a double Faraday collector for Sr composition. Blanks for Rb and Sr are negligible. Sm and Nd isotope analysis was carried out using a ¹⁵⁰Nd/¹⁴⁹Sm mixed isotope tracer. Chemical procedures for Sm and Nd extraction were described by Pin et al. (1994) and Pin & Santos Zalduegui (1997). Nd and Sm were loaded over Ta filaments with H₃PO₄–HNO₃ (Platzner, 1997) and measured with a Faraday collector in multicollection mode. Blanks for Sm and Nd were negligible. The isochrons were regressed using the Isoplot/Ex program of Ludwig (1999).

	Pb–Pb AND U–Pb ISOTOPE RESULTS

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Three groups can be distinguished on the basis of measured Pb isotopic composition (Table 4, Fig. 6a):

View this table: [in this window] [in a new window]   Table 4: Pb/Pb and U–Pb data for whole rocks and mineral separates of pyroxenite, peridotite and mafic rocks from Cabo Ortegal

 

View larger version (24K): [in this window] [in a new window]   Fig. 6. (a) Measured Pb isotopic compositions of minerals and whole rocks from Cabo Ortegal compared with models of Pb growth (Stacey & Kramers, 1975; Zartman & Doe, 1981). UC, upper crust; AC, average crust; M, mantle. Fields: MORB, OIB, EMII and HIMU (Zindler & Hart, 1986; Hart, 1988); Cabo Ortegal pyroxenites (Gravestock, 1992); feldspars from pegmatitic granite in peridotite of Cabo Ortegal (Santos et al., 1996); Mediterranean ophiolites (Hamelin et al., 1984). Clinopyroxenes from pyroxenite GRCP6 are drawn separately to show apparent age correlation (see text for details). (b) Pbi isotopic compositions of minerals from Cabo Ortegal corrected for in situ decay after 390 Ma. Fields: feldspars from pegmatitic granite in peridotite of Cabo Ortegal and Mediterranean ophiolite group as in (a). Arrows for plume (HIMU), depleted mantle and ancient sediments components from Downes (2001).

 

1. a radiogenic Pb-rich group represented by clinopyroxenes from samples GW7692 and GRCP6, whose relatively high µ values (>15) are in accord with the presence of U-rich phases such as zircon and rutile, respectively, in the rock (Peucat et al., 1990; Santos et al., 1996), and orthopyroxenes from H3. The measured Pb isotopic composition of these minerals is more radiogenic than EMII values and falls within the ocean island basalt (OIB) array. The anomalously high radiogenic Pb isotopic composition of fraction Grt-2 from sample GRCP6 (Table 4) could be due to the presence of clinopyroxene or minute rutile inclusions in garnet.
   
2. Less radiogenic clinopyroxenes, garnet, amphibole and whole-rock data with µ values generally of 2–6·5 (Table 4). These are equivalent to previous results by Gravestock (1992) for analogous rocks of Cabo Ortegal. The data are similar to those of circum-Mediterranean ophiolites (Hamelin et al., 1984) or those of less radiogenic OIB and are intermediate between MORB and EMII values (Fig. 6a).
   
3. Zoisite analyses yielded µ values of 2–5·7 but significantly lower ratios of ²⁰⁷Pb/²⁰⁴Pb than other minerals analysed, having a composition close to present-day MORB (Table 4, Fig. 6a).
   

Initial Pb isotopic compositions calculated at 390 Ma (U–Pb rutile ages from sample GRCP6, Santos et al., 1996) produced significant modifications only in the minerals with higher µ values, which tend to approach a common initial value for all minerals (except zoisite, Table 4, Fig. 6b). In this way, Pb isotopic compositions of garnets might correspond to initial values, as the measured U concentration was below detection limit. However, the difference from the initial Pb isotope ratios of clinopyroxenes (Table 4, Fig. 6b) suggests either that there was some U, or that the garnet initial composition was different from that of clinopyroxene within the same rock, which may be related to a metamorphic origin of the garnet. Orthopyroxenes from harzburgites have the highest initial Pb isotopic ratios (²⁰⁶Pb/²⁰⁴Pb = 18·286–18·426), clinopyroxene and amphibole of mafic rocks and associated pyroxenites show more scatter and have the lowest initial Pb isotopic ratios (²⁰⁶Pb/²⁰⁴Pb = 17·91–18·17), whereas clinopyroxene and amphibole of massive pyroxenites have intermediate values (²⁰⁶Pb/²⁰⁴Pb = 18·05–18·3) (Fig. 6b).

An attempt to derive geochronological information from the Pb–Pb and U–Pb isotope data provides two types of results. Clinopyroxenes from GRCP6, excluding the strongly deviating Cpx-6 fraction of slightly turbid grains, are the only samples to define a trend on a ²⁰⁶Pb/²⁰⁴Pb vs ²⁰⁷Pb/²⁰⁴Pb diagram (Fig. 6a). The trend corresponds to an age range of 366 Ma (Cpx1 to Cpx5) to 497 Ma (Cpx1 to Cpx4) Ma; however, errors in excess of 150 my render the calculated age meaningless. On the other hand, regression of data with known µ values yields considerably more consistent results in the ²⁰⁶Pb/²⁰⁴Pb vs ²³⁸U/²⁰⁴Pb diagram. A best-fit line through all the data points (20 mineral analyses) results in a calculated age of 400 ± 26 Ma [mean square weighted deviation (MSWD) = 3·4, Fig. 7a]. This result is barely modified if the three zoisites from mafic rocks, which show some evidence of Pb isotope disequilibrium, or the two values for clinopyroxenes of the Uzal massif, are excluded (394 ± 29 Ma). These ages are consistent with previous U–Pb data for the study area (Peucat et al., 1990; Santos et al., 1996; Ordóñez Casado et al., 2001) and the Sm–Nd ages obtained on the same samples (see next section).

View larger version (15K): [in this window] [in a new window]   Fig. 7. (a) U–Pb diagram and best-fit line for minerals from ultramafic and related mafic rocks from Cabo Ortegal. (b) Rb–Sr diagram for whole rocks and mineral separates from peridotites, pyroxenites and mafic rocks of Cabo Ortegal (Table 5), including data from Van Calsteren et al. (1979) and Gravestock (1992).

 

View this table: [in this window] [in a new window]   Table 5: Sm–Nd and Rb–Sr data for whole rocks and separated minerals from the Cabo Ortegal Complex

 

	Rb–Sr AND Sm–Nd ISOTOPE RESULTS

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Rb–Sr and Sm–Nd elemental and isotopic compositions from whole rock and separated minerals are presented in Table 5. The Sr content of clinopyroxene from massive pyroxenites (except GW1 with the lowest values) is different from that of clinopyroxenes from the mafic rocks and associated pyroxenites (142–177 and 67–81 ppm, respectively). After correction to 390 Ma, the initial ⁸⁷Sr/⁸⁶Sr ratio is also, as in the case of Pb isotopic data, different for both types of rock: 0·7044–0·7048 and 0·7040, respectively. Nevertheless, in both cases ⁸⁷Sr/⁸⁶Sr_i ratios are higher than those of present-day N-MORB (0·7020–0·7030) suggesting the participation of a component with higher ⁸⁷Sr/⁸⁶Sr ratio compared with mantle values in the genesis of these rocks.

_Nd values for whole rocks and clinopyroxene from clinopyroxenites, age corrected to 390 Ma, have a limited range of +2·4 to +4·9. Sm-poor orthopyroxenite OP10 yields an even lower value of -1·7, very different from the maximum value obtained of +6·8, corresponding to the harzburgite H3. A plot of _Nd vs ⁸⁷Sr/⁸⁶Sr ratios for whole rocks and clinopyroxenes at 390 Ma shows a tendency towards higher ⁸⁷Sr/⁸⁶Sr ratios with respect to the ‘palaeo-mantle array’ (Fig. 8). These results are consistent with previous data for the ultramafic rocks of Cabo Ortegal (Peucat et al., 1990; Gravestock, 1992) represented in Fig. 8. A difference with respect to associated eclogites (interpreted as ancient N-MORB rocks) and a tendency towards EM (enriched mantle) characteristics, as inferred also from the Pb compositions (Fig. 6a), is also observed. Other minerals analysed exhibit even higher ⁸⁷Sr/⁸⁶Sr compositions (garnets of sample GRCP6) or lower _Nd (amphibole and zoisite) compared with clinopyroxene and whole-rock values (Fig. 8). The new Rb–Sr results do not allow for calculation of a whole-rock isochron either for the pyroxenites or peridotites, even if previously published data are included. A best-fit line, highly dependent upon one Rb-rich phlogopite composition (Van Calsteren et al., 1979), is close to 390 Ma (Fig. 7b).

View larger version (23K): [in this window] [in a new window]   Fig. 8. Nd vs 87Sr/86Sr at 390 Ma for Cabo Ortegal ultramafic and related mafic whole-rock samples and separated minerals. Fields for published data from Cabo Ortegal rocks include ultramafics, eclogites and granulites from Gravestock (1992) and Peucat et al. (1990), and gneisses from Peucat et al. (1990). * Ultramafics with Rb <1 ppm of Gravestock (1992; XRF analyses) have been recalculated for 0·8 ppm of Rb taking into account the results of this study. Present-day Nd member mantle components (DM, HIMU, EMI and EMII) are from Hart (1988). Dashed line represents the palaeo-mantle array at 390 Ma.

 

In the case of Sm–Nd, selected whole-rock data for pyroxenites excluding orthopyroxenite OP10 with the lowest Sm content, and adding data for similar rocks from Gravestock (1992), except one websterite with the greatest errors, result in an errorchron at 493 ± 130 Ma (_Nd = +4, Fig. 9a). Sm–Nd isotope data for clinopyroxene from pyroxenites provide an equivalent age of 506 ± 91 Ma (_Nd = +4·1), whereas a better result is obtained by combining whole-rock and clinopyroxene data yielding 506 ± 67 Ma (_Nd = +4, Fig. 9b).

View larger version (26K): [in this window] [in a new window]   Fig. 9. Sm–Nd isotopic correlation diagrams for minerals and whole rocks of Cabo Ortegal. (a) Whole-rock data. (b) The same as (a) including clinopyroxene data from this work. (c) Internal isochron (whole-rock–clinopyroxene–garnet) for sample GRCP6. (d) The same for sample GZC7. (e) Two-point line fits for sample GZA2. (f) Regression line for the three garnet fractions.

 

These results contrast with the more precise internal isochron (whole-rock–garnet–clinopyroxene) of 394 ± 22 Ma (_Nd = +4·4, Fig. 9c) for clinopyroxenite sample GRCP6, which is in good agreement with the previous U–Pb age of 383 ± 1 Ma for rutile from the same sample (Santos et al., 1996). The spatially associated sample GZC7 provides a similar though less precise internal (whole-rock–garnet–clinopyroxene) age of 392 ± 69 Ma (_Nd = +2·6, Fig. 9d). Amphibole-rich sample GZA2 gives 384 Ma for the garnet–amphibole pair and a distinctly younger age of 345 Ma for garnet and zoisite, in agreement with the previous suggestion of a late origin for zoisite (Fig. 9e). Finally, the three garnet fractions analysed yielded an errorchron at 413 ± 130 Ma with slightly lower _Nd values of +1·8 (Fig. 9f).

	DISCUSSION

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Mineral data
The new major element mineral chemical data suggest an arc-related origin for ultramafic units in Cabo Ortegal, reveal their similarity with Bragança (Bridges et al., 1995) and exclude a purely oceanic setting (Figs 2–4). Moreover, the harzburgites and dunite–pyroxenite units appear to have different origins.

In a plot of fo-number vs NiO, olivines from both dunites and pyroxenites deviate from the ‘mantle olivine array’ and from the trend for replacive olivines, while defining a trend that is compatible with that produced by olivine fractionation in accord with a cumulus origin for the dunites (Fig. 2). In contrast, olivines from harzburgites exhibit higher NiO contents than those of dunites and overlap the ‘mantle olivine array’ pointing towards a residual origin for these rocks (Fig. 2a). Additionally, although the fo-number is slightly lower than in oceanic mantle peridotites, the harzburgite olivine–spinel pairs plot within a portion of the ‘olivine–spinel mantle array’ that may correspond to refractory peridotite, either abyssal or not (Fig. 2). Actually, the fo-number of olivine in the harzburgites is slightly lower than in typical oceanic peridotites, overlapping the composition of harzburgite–lherzolite rocks from arc complexes elsewhere (e.g. Japan arc and Horoman; Takahashi, 1991; Arai et al., 1998). The cr-number of spinel and the NiO content of coexisting olivine in harzburgite vary with the fo-number of olivine in a way that suggests a residual origin for these rocks after various degrees of partial melting [7–10% of primitive mantle according to calculations by Laribi-Halimi (1992)]. The chemical compositions of olivine and spinel pairs reflect, therefore, a residual origin for the harzburgites.

In contrast, olivine compositions from the dunites clearly deviate from the mantle array and rarely plot in the high-Fo region of refractory rocks on a spinel peridotite restite trend, mostly having Fo values lower than or similar to those of harzburgites (Fig. 2c). This excludes either a residual or a replacive origin by selective consumption of pyroxenes (Arai, 1990; Kelemen, 1990). Most dunite samples contain relatively Fo-poor olivine and Cr-rich spinel similar to those in arc-related high-Mg and high-Si magmas, which suggest an origin as cumulates from relatively primitive magmas (Fig. 2c; Arai, 1992, 1994). Olivine–spinel pairs from pyroxenites define a distinct area characterized by a wide range in Cr/(Cr + Al) ratios and systematically lower Fo contents than harzburgite and dunite. An evolution from more primitive compositions in websterite towards more evolved compositions in mafic rocks and related pyroxenites, spanning the range of volcanic rocks from MORB to island arc tholeiites, is observed (Fig. 2c; see Arai, 1992, 1994).

Clinopyroxenes from harzburgites generally have the lowest Cr and Na contents, overlapping the lowest values from pyroxenite and dunite (Fig. 3a). These are the clinopyroxenes most similar to those of ophiolitic ultramafic rocks (Kornprobst et al., 1981). Clinopyroxenes from the massive pyroxenites and dunites have Cr and Na contents clearly different from those of subcontinental peridotite and comparable with those of Type II ophiolitic mantle (Fig. 3a). Type II ophiolites are interpreted as generated in active continental margin systems where the depleted suboceanic mantle may rise as diapirs (Kornprobst et al., 1981). In a plot of Al_Z vs TiO₂ (Fig. 3b), clinopyroxene data are very different from those of Mid-Atlantic Ridge ultramafic rocks or from those of tholeiitic and alkalic suites generated in rift-related tectonic settings and define a trend similar to that of ultramafic and mafic cumulates from subduction-generated arc magmas (Loucks, 1990).

The tectonic discrimination diagram of Koloskov & Zharinov (1993), based on a multivariate statistical comparison of the chemical composition of clinopyroxene in ultramafic and mafic xenoliths in volcanic rocks worldwide, discriminates linear zones corresponding to clinopyroxene compositions in equilibrium with garnet (fields I and II), spinel (III and IV) and plagioclase (V, VI and VII). Within these zones, fields are defined for pyroxenes in mafic and ultramafic rocks representing different temperature conditions. In the original P₁–P₂ diagram of Koloskov & Zharinov (1993), the data for island arcs and mid-ocean ridges plot within the same field and partially overlap the field of Kamchatka ultramafic intrusions (Fig. 10a). The statistical treatment of the database by Koloskov & Zharinov (1993) has been discussed by Buccianti & Vaselli (1996), who showed significant errors in the position of some samples and suggested that the P₁–P₂ diagram may be strongly biased. To test the classification proposed by Koloskov & Zharinov (1993), we have plotted data for ultramafic rocks of the MARK area and for corundum-bearing and graphite-bearing pyroxenites from the Ronda massif that bear evidence of spinel to garnet reaction (Fig. 10a). Both plot in distinct fields, with most data from Ronda in the spinel pyroxenites field and those from the MARK area in the field for plagioclase-bearing ultramafic rocks and island arc pyroxenites. The data from Kohistan, Izu–Bonin–Mariana and Aleutians–Tonsina are correctly classified as arc clinopyroxenes (see also Fig. 3b) although overlapping with mid-ocean ridge data (Fig. 10b and c). These overlappings would suggest that the P₁–P₂ diagram allows for a discrimination of clinopyroxenes from different origins, although the precise setting is not well enough established.

View larger version (21K): [in this window] [in a new window]   Fig. 10. Tectonic discrimination diagram for clinopyroxenes from various rock types from Cabo Ortegal. Fields I–VII and parameters P1 and P2 are as defined by Koloskov & Zharinov (1993); P1 = -0·02SiO2 + 0·09TiO2 + 0·64Al2O3 + 0·59Cr2O3 + 1·35FeO + 9·65MnO - 0·50MgO + 0·29CaO - 2·26Na2O + 8·0; P2 = 0·93TiO2 + 0·07Al2O3 + 1·23Cr2O3 - 0·46FeO + 1·74MnO + 0·36MgO + 0·10CaO - 1·66Na2O + 8·0. Fields: I, peridotite xenoliths in kimberlites; II, eclogites and pyrope-bearing pyroxenites in kimberlites; III, spinel peridotites; IV, spinel pyroxenites; V, island arc peridotites; VI, island arc pyroxenites; VII, plagioclase-bearing xenoliths in volcanic rocks from island arcs. (a) Data sources: island arc and mid-ocean ridges (Koloskov & Zharinov, 1993); Izu–Bonin–Mariana forearc (Parkinson & Pearce, 1998); Ronda pyroxenites (Obata, 1977; J. F. Santos et al., unpublished data, available on request); Mid-Atlantic Ridge Kane Fracture Zone (MARK area, Gaggero & Cortesogno, 1997). In (b) and (c), vertically dashed area indicates Kamtchatkan alpine-type ultramafic intrusions (Koloskov & Zharinov, 1993): Aleutians–Tonsina (Conrad & Kay, 1984; DeBari et al., 1986; DeBari & Coleman, 1989) and Kohistan (Qasin Jan & Windley, 1990).

 

Our clinopyroxene data for Cabo Ortegal harzburgites plot in the region of island arc peridotites (which includes websterites according to Koloskov & Zharinov) in this diagram (Fig. 10a). The data from the dunites overlap those of harzburgites, whereas clinopyroxenes from pyroxenites show a larger compositional range and extend into the field of pyroxenites from island arc environments (Fig. 10b). Most clinopyroxenes from mafic rocks and related pyroxenites have parameter values similar to those of the massive pyroxenites and plot on the same geotectonic discrimination field of the diagram (Fig. 10b). P₁ and P₂ parameters of clinopyroxenes show thus nearly constant values throughout that might indicate a common geotectonic setting, that of an arc region, for all rock types. The data of Cabo Ortegal, which in part overlap the fields for Kohistan and less so Aleutians–Tonsina, or Kamchatka intrusive rocks and peridotites of Izu–Bonin–Mariana forearc (Fig. 10b and c), would correspond to Mg-rich compositions and elevated temperatures of formation for clinopyroxenes (Koloskov & Zharinov, 1993).

Amphiboles have low Na_(M4) and Al^[6] contents, which suggests crystallization at relatively low pressure [<1·3 GPa, according to experimental results by Niida & Green (1999)]. The overall composition of analysed amphibole is similar to that of Ti-poor amphibole from alpine-type peridotite massifs attributed to near-isochemical hydration reactions involving clinopyroxene and spinel (Niida & Green, 1999, and references therein), whereas it markedly differs from the Ti- and K-rich compositions of amphibole related to mantle metasomatism (Wilkinson & Le Maitre, 1987; Zanetti et al., 1999, and references therein). This suggests extensive recrystallization of the ultramafic rocks under relatively high-temperature, H₂O-deficient conditions to produce widespread homogeneous, subidiomorphic amphibole in textural equilibrium in all rocks analysed.

Whole-rock geochemical data
On the basis of petrography and field data, four main rock types can be distinguished within the ultramafic massifs of Cabo Ortegal: (1) harzburgites; (2) dunites, mostly associated with pyroxenites in a composite pyroxenite–dunite-rich section of the Herbeira massif; (3) massive, occasionally garnet-bearing, pyroxenites ranging in composition from olivine websterite to orthopyroxenite; (4) rare mafic rocks with variable amounts of associated garnet-rich pyroxenite. Variations in CaO/Al₂O₃ ratios in the peridotites can be related to the presence of fine layers of pyroxenite in the rocks analysed. The geochemical data plotted in the binary diagrams (Fig. 4), however, show a distribution that reflects well the mineralogy of the rocks and allows for a distinction between three main groups: (1) harzburgites and dunites; (2) massive pyroxenites; (3) mafic rocks and associated pyroxenites. In these diagrams, a lack of linear correlation between these rock groups is observed, which may suggest that the pyroxenites and mafic rocks do not represent melts extracted from the peridotites. In addition, the chemical composition of the pyroxenites is very different from that of similar rocks in ultramafic massifs elsewhere (e.g. Ronda, Beni Boussera, Horoman, Alps, Pyrenees) except for a restricted group of pyroxenites from the Ronda peridotite (Fig. 5).

The limited amount of data does not allow us to differentiate the above groups in multi-element diagrams where, nevertheless, normalized patterns are slightly enriched in Rb, Ba, U, Th, K and Sr with respect to primitive mantle and differ substantially from those of E-MORB and N-MORB (Fig. 5a). Similar diagrams for mafic rocks and pyroxenites from orogenic peridotites elsewhere show a characteristic depletion in incompatible elements (e.g. Takazawa et al., 1999; Garrido & Bodinier, 1999). The absence of this depletion in the Cabo Ortegal rocks suggests that the source area of the pyroxenite parental melts was originally enriched in incompatible elements or underwent LILE enrichment by reaction with external fluids or melts. The last hypothesis is preferred here, as it would also account for the observed LILE enrichment in the peridotites [sample H3 and Gravestock (1992)].

The observed enrichment in LILE is associated with a generalized although variable LREE enrichment of peridotites and pyroxenites (Fig. 5b; see also Peucat et al., 1990; Gravestock, 1992). Such generalized LREE enrichment of whole rocks and clinopyroxenes from Cabo Ortegal is noteworthy. Flat to LREE-enriched patterns have been described also from the Horoman, Alpine and Ronda massifs but their occurrence there is limited and their origin is not well constrained (e.g. Takazawa et al., 1992; Garrido & Bodinier, 1999; Zanetti et al., 1999). Recently, Gruau et al. (1999) have proposed various mechanisms for selective LREE enrichment in ophiolite peridotites. In the present case, and by comparison with the models proposed by Gruau et al. (1999), the REE and isotopic data (see below) point to high-temperature contamination of the ultramafic rocks through interaction with LREE-rich melts or fluids derived from subducted continental materials in a supra-subduction setting.

As previously mentioned, some mafic rocks from the Ronda peridotite massif, the D1 subtype, have petrographic (clinopyroxenite, websterite and olivine websterite modes) and major element compositions (Fig. 4) similar to the rocks studied from Cabo Ortegal. According to Garrido & Bodinier (1999), the origin of the majority of mafic rocks from Ronda could be related to a magmatic event caused by melting of the base of thinned subcontinental lithosphere by upwelling asthenosphere. In the waning stages of this magmatic event, D1 subtype was formed by replacement of older mafic rocks by small-volume melt fractions that percolated pervasively through the massif and had a refractory, calc-alkaline character. In the case of Cabo Ortegal, a similar process is not excluded; however, the higher LILE and LREE values (and also isotope data; see below) suggest melting of subcontinental lithosphere with participation of subducted crustal components. Limited evidence of late melt percolation might be represented in some harzburgites and dunites by large interstitial clinopyroxene crystals at the base of the pyroxenite–dunite-rich section.

Finally, the presence of anomalous and volumetrically restricted rock types, such as the pyroxenites GCZ7 and GRCP6, suggest other petrogenetic processes of limited extent. In GCZ7, the weak positive Eu anomaly might be related to the occurrence of significant amounts of zoisite in this rock (traces of feldspar also occur in similar pyroxenites). Zoisite might have contributed to the Eu enrichment through Eu–Ca substitution in plagioclase or by substitutions between zoisite and allanite. As for the garnet–rutile pyroxenite GRCP6, the HREE and HFSE enrichment could be related to accumulation processes through high-pressure crystal segregation of those minerals. The presence of Ti minerals suggests that the equilibrium melts were saturated in Ti (see Becker, 1996).

Isotopic characterization
The isotopic compositions of clinopyroxene mineral separates are assumed to be representative of the pristine whole rocks (Reisberg & Zindler, 1986) and thus of the magmas from which the pyroxenites originated. Relatively homogeneous initial Pb isotopic composition of amphiboles suggests equilibrium with clinopyroxene. Pb isotopic disequilibrium observed between the pyroxenites related to mafic rocks and the massive pyroxenites (Fig. 6b) is similar to intra-complex isotopic heterogenities in ophiolite massifs, suggesting a variety of magma sources (Encarnación et al., 1999).

The similarity of ²⁰⁶Pb/²⁰⁴Pb initial ratios in all analysed minerals (except for zoisite) with those of present-day MORBs contrasts with their higher ²⁰⁷Pb/²⁰⁴Pb ratios. Pb isotopic compositions are intermediate between those of MORB and EMII (Zindler & Hart, 1986), suggesting the incorporation of external material into the mantle, probably of an ancient component with low µ values. Comparison with the Pb isotopic composition of oceanic sediments and pyroxenites elsewhere (e.g. Beni Boussera, Pearson et al., 1993) indicates that the magmas that formed the pyroxenites at Cabo Ortegal might involve a significant contribution from recycled crustal materials. Also, the initial Pb isotopic ratios are similar to those of Mediterranean ophiolites (Fig. 6b), which are higher than those of present-day accretion zones (MORB) and have been interpreted as evidence of an island arc origin involving a significant continental component (Hamelin et al., 1984). Such initial values are similar to average values reported from typical island arc regions such as NE Japan and the Philippines (Tatsumoto & Nakamura, 1991; Castillo et al., 1999).

Although there exists some variability in Sr contents and ⁸⁷Sr/⁸⁶Sr_i isotopic composition that might be related to various processes during the evolution of the pyroxenites, a modification of Sr contents through post-crystallization alteration by crustal-derived fluids (Polvé & Allègre, 1980; Zindler et al., 1983) can be excluded, as the pyroxenites are fresh rocks devoid of chlorite, serpentine, talc or other secondary products. These petrographic features, together with the different ranges of Sr and Pb isotopic initial ratios (Fig. 7) of pyroxenites related to mafic rocks with respect to massive pyroxenites and harzburgites, suggest primary Sr isotope compositions.

After age correction to 390 Ma, _Nd and ⁸⁷Sr/⁸⁶ initial ratios of the analysed rocks deviate greatly from DM (depleted mantle) values, at the same time suggesting source enrichment. Taking into account previous data by Peucat et al. (1990) and Gravestock (1992) on similar rocks, this could correspond to an EMI-like composition. If we consider other major rock types from the Cabo Ortegal complex, there appears to be a mixing trend from eclogites (equivalent to DM) to gneisses (equivalent to upper crust), with the mafic to intermediate granulites and ultramafic rocks approaching the composition of the eclogites (Fig. 9). If the mantle source were similar isotopically to that of the eclogites, the observed variability in _Nd in the rocks studied (+2 to +4; Table 5) and the difference with respect to the Nd and Sr isotopic compositions of the associated eclogites would imply the participation of some external material in the melts from which the pyroxenites crystallized. Alternatively, a different scenario might be envisaged; for instance, a mantle portion (mantle wedge?) whose isotopic composition would be different from that of DM. In this respect, and although Sr and Nd isotopic compositions comparable with those of Cabo Ortegal are recorded in different geodynamic settings, a close similarity may be noted with those of arc regions elsewhere, e.g. NE Japan and Aeolian Arc areas (Tatsumoto & Nakamura, 1991; De Astis et al., 2000).

Despite the textural relationships that suggest equilibrium of clinopyroxene with garnet and zoisite, variations in the isotopic composition of clinopyroxenes and, particularly, of garnet (high Sr isotope ratios) or zoisite (very low Pb isotopic compositions) do not appear to correspond to a simple event, either magmatic or metamorphic. For zoisite, the Sm–Nd age of the Grt–Zo pair suggests a younger re-equilibration of feldspar-bearing rocks, as has been suggested above. Garnet and zoisite both require a more detailed study to allow a definite conclusion on this subject.

Interpretation of ages
Van Calsteren et al. (1979) reported a Rb–Sr errorchron age of 480 Ma for the harzburgites of Cabo Ortegal. The ages around 500 Ma obtained in this study for the pyroxenites using the Sm–Nd whole-rock and clinopyroxene data (Fig. 8) are comparable with those previously reported. Such an age is also similar to known ages for the mafic protoliths of associated eclogites and mafic granulites within the same complex (U–Pb in zircons, Peucat et al., 1990; Ordóñez Casado et al., 2001). Despite the lack of precision, an age around 490 Ma for the ultramafic rocks appears therefore realistic and suggests that the mafic and ultramafic protoliths from the uppermost units of the complex originated during the same period. However, ages around 390 Ma are more frequently found in high-grade rocks within those units (Ordóñez Casado et al., 2001, and references therein).

The younger ages are corroborated by the new Sm–Nd internal isochrons for pyroxenites GRCP6 and GZC7 (Grt–Cpx–whole-rock) at 390 Ma (Fig. 8), the three garnet fractions errorchron at 413 Ma, and the Grt–Amp age of 384 Ma from sample GZA2. All these are, furthermore, consistent with previous age determinations on the same rocks, e.g. 383 Ma in rutile from GRCP6 (U–Pb, Santos et al., 1996) or from the nearby Uzal peridotite massif; U–Pb dating of zircon and monazite and Rb–Sr whole-rock–mineral dating of pyroxenite and intrusive pegmatite (Peucat et al., 1990; Santos et al., 1996) yielding values close to 390 Ma. Other ages around 380 Ma for the peridotites of the Herbeira massif were obtained also by the K–Ar and Rb–Sr methods on mineral–whole-rock pairs by Van Calsteren et al. (1979).

The fact that the Sm–Nd garnet data alone give a similar or slightly older age and slightly lower initial _Nd(T) values than the internal isochrons has been interpreted in high-pressure terranes elsewhere as evidence that the mafic protoliths intruded shortly before the metamorphism (Chavagnac & Jahn, 1996), which would be in agreement with the results presented above. Because the inferred closure temperature for Nd isotope systematics in garnet is at least 700–750°C (Hensen & Zhou, 1995), the present results would suggest that the ages around 390 Ma are representative of a regional high-temperature (and high-pressure) event, including partial melting of crustal protoliths (pegmatite and associated HP migmatitic gneiss units of the complex). This event would be relatively short in time as deduced from the slightly younger ages for the closure temperature of U–Pb in rutile (450°C) in pyroxenites and HP gneisses at 380 Ma (Santos et al., 1996; Valverde Vaquero & Fernández, 1996).

The younger Sm–Nd age of 345 Ma obtained for the Grt–Zo pair of GZC7 is in the range of the oldest ages obtained for the D1 tectonothermal event in the para-autochthon of the complex (Fig. 1a, Dallmeyer et al., 1997). This corresponds to a rather pervasive low-grade retrogressive episode in the HP/HT rocks of the complex (Abalos et al., 1994). As such, the age obtained might reflect re-equilibration of zoisite during a younger tectonothermal episode in accord with the previous suggestion that zoisite is a late mineral, possibly after feldspar, with a different Pb isotopic composition.

The Pb–Pb, Pb–U and Rb–Sr isotope data obtained do not allow us to provide better constraints on the age of the protoliths of the peridotites or of the massive pyroxenites (Figs 7a and 8). Clinopyroxenes from GRCP6 yielded ²⁰⁶Pb/²⁰⁴Pb vs ²⁰⁷Pb/²⁰⁴Pb regression lines from 366 Ma (five points) to 497 Ma (four points, Fig. 6a) whereby the 490 Ma event is again observed although more analyses are needed to confirm this age. In contrast, the age of 400 ± 26 Ma (Fig. 7a) obtained through regression of all data with known µ values (20 mineral analyses) appears consistent with previous U–Pb dating of the high-grade recrystallization events in the study area (Peucat et al., 1990; Santos et al., 1996; Valverde Vaquero & Fernández, 1996; Ordóñez Casado et al., 2001) and with the new Sm–Nd results of the present study.

Geodynamic scenario
By comparison with similar peridotite massifs elsewhere (e.g. Ronda, Horoman), studied ultramafic rocks from Cabo Ortegal exhibit generalized, although variable, trace element enrichment that suggests widespread interaction with a LREE- and, to some extent, LILE-enriched component. LREE enrichment in ophiolite peridotites has been attributed to high-temperature metasomatism through interaction with LREE-rich melts or fluids derived from subducted continental material in a supra-subduction setting (Gruau et al., 1999). Also, Gamble et al. (1996) and Plank & Langmuir (1998), among others, have proposed tectonic settings related to mantle-wedge areas at the origin of fluid or melts enriched in LILE. This allows us to consider that the various rock types, including harzburgites, pyroxenites, dunites and mafic rocks, from Cabo Ortegal might correspond to part of a mantle wedge modified by infiltration of fluids or partial melts derived from subducting oceanic lithosphere and sediments.

Variations in modal mineralogy and mineral composition of the massive pyroxenites–dunites at Cabo Ortegal may be related to fractional crystallization and accumulation processes, whereas variations resulting from interaction of the pyroxenite parental melts with the host peridotite are less clear. Some features, such as the depletion in HREE of the garnet-bearing massive pyroxenites, favour the interpretation that the parental melts were already strongly depleted in HREE. The origin of minor, but generalized, positive Eu anomalies and variable LREE and middle REE (MREE) enrichment is less clear, and these may be interpreted as inherited features or the result of subsequent interaction with fluids or melts. Detailed, small-scale studies of pyroxenite–dunite-rich sections are necessary to understand the relationships and/or interactions between pyroxenites or mafic layers and the host peridotites (e.g. the role of diffusion-controlled Fe–Mg exchange).

The mineral composition of the harzburgites reflects a residual character, which might result from island arc magma(s) extraction and not from an oceanic origin as previously suggested (Laribi-Halimi, 1992). Simultaneously, harzburgites within a mantle-wedge environment would be affected by fluid or melt infiltration as signalled by moderate LREE enrichment. As for the scarce mafic rocks and related pyroxenites, the positive anomalies of Eu, Sr and Ba, and the relatively high Na, Al and Ca contents, suggest the former presence of plagioclase that has been subsequently recrystallized to form other Ca-rich silicates (e.g. zoisite) under metamorphic conditions.

Likewise, isotopic compositions plotted in Pb–Pb (Fig. 6), Sr–Nd (Fig. 9) or Pb–Nd diagrams suggest for the parental melts of the pyroxenites a mantle source that does not correspond to a depleted mantle or MORB source. The relatively high ⁸⁷Sr/⁸⁶Sr and ²⁰⁷Pb/²⁰⁴Pb ratios and low _Nd values with respect to DM compositions point to the involvement of a source component with EM characteristics, which may be recycled terrigenous sediment or subducted continental crust. Such components are normally related to subduction settings—in the case of EMII either to subducted continental material, principally terrigenous sediments (e.g. Hauri et al., 1993), whereas for EMI, sediment recycling of pelagic ocean-floor sediments (e.g. Weaver, 1991) or subduction of subcontinental lithosphere from craton margins (e.g. Tatsumoto & Nakamura, 1991) has been envisaged. Notwithstanding, the absence of coherent isochron relationships and decoupling between the Sr–Pb isotope compositions of minerals from any given sample point to a multistage evolution of massive pyroxenites in the Herbeira massif.

¹⁴³Nd/¹⁴⁴Nd and ⁸⁷Sr/⁸⁶Sr ratios of clinopyroxenes match the trend proposed for clinopyroxenes from chemically and mineralogically heterogeneous garnet-bearing peridotites and pyroxenites formed from spinel-bearing precursors bearing evidence of high primary temperatures, metasomatic processes and, significantly, protolith ages close to those of the HP recrystallization (<250 my older; Fig. 11; Brueckner & Medaris, 1998). This is a different type from that represented by LILE-poor, garnet-bearing peridotites and pyroxenites formed from cold and >250 my older protoliths (e.g. those from the Western Gneiss Region; Fig. 11). Accordingly, the observed features in the Cabo Ortegal ultramafic rocks suggest an evolution similar to that experienced by (spinel)–garnet-bearing ultramafic rocks elsewhere within the Variscan fold belt (e.g. Bohemian Massif, Brueckner & Medaris, 1998, 2000). The specific geodynamic context for the Cabo Ortegal rocks might correspond to a convecting depleted mantle wedge associated with the collision between Gondwana (para-autochthon) and an outboard arc system (mafic–felsic granulites and gneisses). This mantle section could be metasomatized by fluid or melt migration during the subduction of the intervening oceanic lithosphere, as represented by associated N-MORB eclogites, terrigenous sediments or continental crust.

View larger version (29K): [in this window] [in a new window]   Fig. 11. Sr–Nd isotopic ratio covariance diagram for clinopyroxenes from Cabo Ortegal. Trends for two types of garnet-bearing peridotites and pyroxenites, e.g. the Bohemian massif and Western Gneiss Region, representing different mantle settings and tectonothermal evolution, are from Brueckner & Medaris (1998). Ancient depletion, ancient enrichment and hydrous fluid components from Downes (2001). Shaded area indicates Sr–Nd composition of ultramafic massifs in western Europe (Downes, 2001).

 

The proposed geodynamic scenario integrates previous suggestions for the origin of the Cabo Ortegal ultramafic rocks and similar units in the northwestern Iberian Massif. Gravestock (1992), for instance, proposed a mantle-wedge environment metasomatized by a low-viscosity carbonatite melt, whereas Bridges et al. (1995) suggested a supra-subduction zone setting for the nearby Bragança complex, and Moreno et al. (2001) proposed an arc-root setting to accommodate the crystallization of pyroxenite–dunite units at the crust–mantle interface region. As for the participation of different mantle sources or melts, this has been discussed previously by Laribi-Halimi (1992) in relation to partial melting of primitive mantle for the peridotites and crystallization of tholeiitic magmas for the pyroxenites, whereas Moreno et al. (2001) considered the existence of two batches of mantle-derived magmas from which the pyroxenite–dunite association crystallized.

Within such a tectonic context, the processes that operated to produce the geochemical characteristics of the pyroxenites are not well established and are almost certainly complex and multi-stage. For example, high CaO/Al₂O₃ ratios (>5 in the Ortegal pyroxenites) have been related to interaction between carbonate magmas and subcontinental lithospheric mantle (Zangana et al., 1997). This process could also result in the appearance of primary mantle carbonates, U-shaped, LREE-enriched patterns, low Ti/Eu ratios or Zr depletion in whole rocks (Ionov et al., 1993; Lenoir et al., 1997; Yaxley et al., 1998). Although carbonates do occur in the Cabo Ortegal pyroxenites (Girardeau & Gil Ibarguchi, 1991), the observed LREE enrichment does not result in U-shaped REE patterns. Also Ti/Eu ratios (2000–12 000) are not very different from those in other western and central European mantle-derived rocks, where carbonatite metasomatism apparently had a limited effect (Downes, 2001). The high Zr and Nb contents of Cabo Ortegal rocks also do not suggest the passage of carbonate magmas. In this respect, although Zr and Nb depletions are most typical of arc regions, their absence is by no means unknown in such environments (e.g. Shinjo et al., 2000).

However, as mentioned above, the LILE enrichment may be due to reaction with subduction-related hydrous fluids. As such fluids would largely be derived from seawater-altered basalts and sea-floor sediments, this could affect the Sr isotopic compositions and concentrations and, eventually, if sea-floor sediments were derived from old crustal rocks, result in higher ²⁰⁷Pb/²⁰⁴Pb ratios for a given ²⁰⁶Pb/²⁰⁴Pb value. In Cabo Ortegal, LILE contents and Sr and Pb isotopic compositions are more radiogenic than those of MORB, favouring the participation of a subducted component.

	CONCLUSIONS

TOP ABSTRACT INTRODUCTION GEOLOGICAL SETTING PREVIOUS GEOCHRONOLOGICAL AND... SAMPLE DESCRIPTION MAJOR ELEMENT MINERAL CHEMISTRY WHOLE-ROCK MAJOR AND TRACE... ISOTOPE ANALYTICAL METHODS Pb-Pb AND U-Pb ISOTOPE... Rb-Sr AND Sm-Nd ISOTOPE... DISCUSSION CONCLUSIONS REFERENCES

 
Mineral, chemical and whole-rock major and trace element and Pb–Sr–Nd isotope data suggest that regional-scale massive pyroxenites from Cabo Ortegal originated from relatively homogeneous parental melts. Fractional crystallization processes, coeval with intense deformation, might result in the formation of cumulate layers (clinopyroxene, orthopyroxene, olivine, chromite, etc.). Some less abundant mafic rocks and associated pyroxenites are also homogeneous but have different chemical and isotopic signatures, which suggest a different parental melt from that of the massive pyroxenites. LILE and LREE enrichment with respect to similar rocks from other ultramafic massifs and clinopyroxene Pb, Nd and Sr initial isotopic ratios suggest an enriched mantle source. The enriched component may have been introduced into the mantle source during subduction.

Previous studies of the Cabo Ortegal ultramafic units have concluded that it experienced a complex tectonothermal history similar to that of other units in the same area (e.g. Abalos et al., 1996). The new data presented here confirm this hypothesis and identify at least two different events, as follows:

1. Sm–Nd whole-rock–clinopyroxene ages for harzburgites, massive pyroxenites and rare mafic rocks suggest formation of the ultramafic units at 500 Ma. This age is similar to that of formation of mafic to felsic protoliths, whether N-MORB, arc-related or continental in origin, of the associated HP/HT units (Peucat et al., 1990; Ordóñez Casado et al., 2001).
   
2. Sm–Nd internal isochrons for two pyroxenites are consistent with U–Pb dating of rutile (Santos et al., 1996) and indicate subduction at 390 Ma. This event involved recrystallization under HT/HP conditions with the formation of eclogites and HP granulites derived from N-MORB and arc-related rocks, respectively, production of garnet-bearing assemblages in spinel-bearing ultramafic rocks, and limited partial melting as evidenced by rare cross-cutting pyroxenite veins (Peucat et al., 1990) and intrusive granite bodies (Santos et al., 1996).
   

The younger event might be responsible for the introduction from a subducting slab of LILE- and LREE-rich melts or fluids into the overlying mantle wedge in a sub-arc setting. These fluids or melts might possibly induce some partial melting to produce restricted granitic liquids and late, occasionally garnet-rich, pyroxenite veins, as well as the metasomatism of peridotite and massive pyroxenite–dunite rocks to produce their LILE- and LREE-enriched signatures. However, it cannot be precluded that the anomalous geochemical characteristics of the ultramafic rocks reflect the mantle composition at the time of their formation at 500 Ma.

On a regional scale, the formation of abundant island-arc mafic magmas at 395 Ma, and, therefore, the likely presence of substantial astenospheric component within a mantle wedge overlying a deep subducting slab at that time, has been documented in a number of ophiolite units from the allochthonous complexes of the northwestern Iberian massif (Díaz García et al., 1999; Pin et al., 2001). This might indicate a relationship between the origin of those magmas and the subduction recorded within the at present structurally overlying ultramafic–mafic units such as those studied here.

The great amount of pyroxenite in the Cabo Ortegal massif and the seemingly complex processes involved in the formation of the ultramafic–mafic association require further detailed investigations. This should include detailed mapping and systematic classification of the various pyroxenite subtypes together with in-depth studies of the interactions between pyroxenites or mafic layers and the host peridotites.

	ACKNOWLEDGEMENTS

 
Financial support by Spanish DGES grants PB97/617 and PB98/143 is acknowledged. The authors are grateful to Drs B. Abalos and L. A. Ortega for helpful discussions on the structural and petrological features of the study area. Thorough and thoughtful reviews by H. Downes, M. Wilson, S. DeBari, O. Müntener and P. Kelemen are greatly appreciated. L. Gaggero provided useful mineral data on the MARK area.

	FOOTNOTES

 
Extended dataset can be found at http://www.petrology.oupjournals.org

^*Corresponding author. Telephone: +34-94-601-2641. Fax: +34-94-601-3078. E-mail: nppsazaf{at}lg.ehu.es

	REFERENCES

TOP ABSTRACT INTRODUCTION GEOLOGICAL SETTING PREVIOUS GEOCHRONOLOGICAL AND... SAMPLE DESCRIPTION MAJOR ELEMENT MINERAL CHEMISTRY WHOLE-ROCK MAJOR AND TRACE... ISOTOPE ANALYTICAL METHODS Pb-Pb AND U-Pb ISOTOPE... Rb-Sr AND Sm-Nd ISOTOPE... DISCUSSION CONCLUSIONS REFERENCES

 
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