Melt Migration and Intrusion during Exhumation of ... - Oxford Journals

JOURNAL OF PETROLOGY

VOLUME 51

NUMBERS 1 & 2

PAGES 295^325

2010

doi:10.1093/petrology/egp061

Melt Migration and Intrusion during Exhumation of the Alboran Lithosphere: the Tallante Mantle Xenolith Record (Betic Cordillera, SE Spain) E. RAMPONE1*, R. L. M. VISSERS2, M. POGGIO1, M. SCAMBELLURI1 AND A. ZANETTI3 1

DIPARTIMENTO PER LO STUDIO DEL TERRITORIO E DELLE SUE RISORSE, UNIVERSITY OF GENOVA, CORSO EUROPA

26, I-16132 GENOVA, ITALY 2

DEPARTMENT OF EARTH SCIENCES, FACULTY OF GEOSCIENCES, UTRECHT UNIVERSITY, PO BOX 80.021, 3508 TA

UTRECHT, NETHERLANDS 3

CNR-ISTITUTO DI GEOSCIENZE E GEORISORSE, SEZIONE DI PAVIA, VIA FERRATA 1, I-27100 PAVIA, ITALY

RECEIVED DECEMBER 29, 2008; ACCEPTED AUGUST 12, 2009 ADVANCE ACCESS PUBLICATION SEPTEMBER 22, 2009

Microstructural and in situ mineral chemistry studies on mantle peridotite xenoliths from the Late Neogene alkaline volcanic center of Cabezo Tallante (SE Spain) reveal an exceptional record of a multi-stage history of deformation, recrystallization, melt^rock interaction and melt intrusion tracking the progressive exhumation of this lithospheric mantle sector. Xenoliths include porphyroclastic to equigranular spinel peridotites, impregnated plagioclase peridotites, and composite xenoliths made up of peridotites intruded first by gabbronorite veins and later by amphibole-bearing pyroxenites.The earliest stage involved subsolidus re-equilibration from garnet- to spinel-facies conditions, represented by rounded opx þ spinel cpx clusters indicative of precursor garnet.The spinel-facies equilibration was followed by development of a porphyroclastic fabric, accentuated in many xenoliths by spinel trails, in response to shear deformation that may be related to the early stages of Neogene extension. Porphyroclastic spinel peridotites subsequently underwent multiple episodes of reactive porous melt percolation documented by crystallization of undeformed olivine replacing pyroxene porphyroclasts, and of undeformed poikilitic orthopyroxene at the expense of both pyroxene porphyroclasts and newly crystallized olivines.The porphyroclastic and melt^rock reaction textures are progressively obliterated by an equigranular structure developed as the result of static, possibly melt-assisted, annealing recrystallization. Clinopyroxenes in

equigranular peridotites (i.e. the most equilibrated with the percolating melts) display slight light rare earth element (REE) depletion and almost flat middle to heavy REE spectra (LaN/YbN ¼ 037^ 062; SmN/YbN ¼ 089^123). Computed equilibrium liquids have an enriched tholeiitic affinity, consistent with the sub-alkaline magmatism of the Alboran Domain. Overall, the tectonic and magmatic stages recorded in spinel peridotites from Tallante are remarkably consistent with the evolution documented in the Ronda peridotites of the western Betics. Reactive porous flow and annealing recrystallization were followed by an impregnation event, documented by crystallization of interstitial (plag opx ol) aggregates in porphyroclastic and equigranular xenoliths; this indicates further exhumation to shallower depths. Diffuse melt percolation was followed by intrusion of melts with distinct chemical affinity. The first event is documented by thin gabbronoritic^noritic veins, showing opx reaction rims against the host peridotite. Comparable gabbronorites were previously ascribed to slab-derived melts. The norite veins are crosscut by centimeter-thick dikelets of amphibole pyroxenite. Geobarometric estimates and the observed crystallization order (ol^cpxâmph^plag) point to 07^09 GPa for pyroxenite intrusion. Computed melts in equilibrium with clinopyroxene show alkaline affinity, similar to the host Tallante alkali basalts. Textural and geochemical features in the xenoliths thus indicate that the

*Corresponding author.Telephone: 003910 3538315. Fax: 003910 352169. E-mail: [email protected]

ß The Author 2009. Published by Oxford University Press. All rights reserved. For Permissions, please e-mail: journals.permissions@ oxfordjournals.org


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progressive uplift of the Tallante lithospheric mantle was accompanied by interaction with melts of different sources, reflecting the magmatic evolution of the Alboran Domain in response to lithosphere extension and thinning leading to the formation of the Betic^Rif arc.

KEY WORDS: mantle xenoliths; Alboran Domain; lithosphere extension; reactive porous flow; melt impregnation

I N T RO D U C T I O N Since the latest Oligocene the western Mediterranean region has experienced a complex geodynamic evolution, involving lithosphere extension and development of the Alboran Sea whilst the bounding African and Eurasian plates formed an essentially convergent setting. Extensional thinning of the Alboran lithosphere was most probably accompanied by the ascent of hot asthenosphere and consequent lithosphereâsthenosphere interaction. This tectonic context invites geological and geochemical studies of the pertinent upper mantle rocks, aiming to elucidate the details of these processes and their bearing on the geodynamics of the region. The Betic Cordillera of southern Spain, together with the Rif, Tell and Kabylies chains of Morocco, Algeria and Tunisia, form part of a tight arc-shaped mountain belt making up the westernmost part of the Alpine orogenic system (Fig. 1). The external parts of the belt represent the South Iberian and North African passive margin sequences, which were strongly deformed during Neogene folding and thrusting (Garc|¤ a-Herna¤ndez et al., 1980; Banks & Warburton, 1991). The inner parts of the Betic Cordillera and Rif Mountains are made up of an allochthonous pile of intensely deformed, mostly metamorphic rocks; they are considered as the relics of an early Alpine collisional system formed during Late Mesozoic to Tertiary convergence between Africa and Iberia, subsequently exhumed and strongly dismembered during Neogene late-orogenic extension (Platt & Vissers, 1989; Garc|¤ a-Duen‹as et al., 1992; Lonergan & White, 1997; Comas et al., 1999; Jolivet & Faccenna, 2000). These rocks of the Betic and Rif internal zones, intensely affected by Neogene extension, are often referred to as the Alboran Domain (Garc|¤ a-Duen‹as et al., 1992). The present-day crustal thickness in the internal parts of this domain varies from 20^25 km to less than 10 km, with minimum thicknesses towards the east at the transition from the Alboran Basin to the South Balearic Basin (Torne et al., 2000). In response to Neogene lithospheric extension, the Alboran Sea region has been affected by widespread magmatic activity involving eruption of tholeiitic, calc-alkaline and shoshonitic magmas, followed by Late Neogene alkaline basalts. The region thus records the post-collisional transition from subduction- to intraplate-type magmatism

JANUARY & FEBRUARY 2010

(Turner et al., 1999; Coulon et al., 2002; Duggen et al., 2003, 2004, 2005, 2008). Two major groups of models have been proposed to explain the complex geodynamic and magmatic evolution of the western Mediterranean region: (1) non-subduction models, involving delamination or convective removal of gravitationally unstable, thickened subcontinental lithosphere beneath the Alboran Domain (Platt & Vissers, 1989; Docherty & Banda, 1995; Platt et al., 1996; Comas et al., 1999; Turner et al., 1999; Doblas et al., 2007); (2) subduction models, involving subduction of Tethyan oceanic lithosphere associated with slab roll-back and steepening and/or detachment of the subducted slab (Royden, 1993; Lonergan & White, 1997; Zeck et al., 1998; Hoernle et al., 1999; Wortel & Spakman, 2000; Coulon et al., 2002; Gutscher et al., 2002; Duggen et al., 2003, 2004, 2005, 2008; Gill et al., 2004). This latter group of models is supported by tomographic studies (Wortel & Spakman, 2000; Gutscher et al., 2002; Spakman & Wortel, 2004) providing evidence of an east-dipping slab of cold oceanic lithosphere descending from the Atlantic Domain beneath the Alboran Sea. According to Gutscher et al. (2002) and Spakman & Wortel (2004), Paleogene subduction of the Ligurian ocean in the present-day western part of the Mediterranean region occurred in a northwesterly direction underneath the Balearic Islands, Sardinia and Corsica. During the latest Oligocene to early Miocene, this subduction system became divided into two distinct segments after collision between Iberia and North Africa. Subsequently, the western subduction segment initiated a west-directed roll-back to form the Alboran Basin. A general consensus on the geodynamic evolution of the Alboran Domain does not as yet exist; however, each of the currently competing hypotheses involves large-scale mantle processes, and this feature motivates this study of the structure and petrology of the upper mantle in question. Within the Alboran Domain, this upper mantle is accessible in two different ways; that is, in peridotite massifs exposed in the Betic and Rif internal zones in the westernmost part of the arc, and in the form of mantle xenoliths brought to the surface in southeastern Spain by late Neogene alkaline as well as potassic to ultrapotassic volcanic rocks. Large exposures of upper mantle rocks occur in the peridotite massifs of the Sierra Bermeja, Alpujata and Carratraca near Ronda (western Betics), and in the Beni Bousera massif (internal Rif, north Morocco). Structural and geochemical studies of the Ronda peridotites of the Sierra Bermeja (Van der Wal, 1993; Van der Wal & Vissers, 1993, 1996; Van der Wal & Bodinier, 1996; Garrido & Bodinier, 1999; Lenoir et al., 2001) reveal a structural history marked by early stage exhumation of the peridotites from deep lithospheric levels (Davies et al., 1993), followed by a stage of progressive deformation leading to

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EVOLUTION OF ALBORAN DOMAIN

Fig. 1. Tectonic sketch map of the Betic^Rif arc and Alboran Sea, modified after Meijninger (2006), showing the location of the Cabezo Tallante volcanic centre NW of Cartagena. The peridotite massifs amidst rocks of the Alboran Domain in the western part of the arc near Ronda and in northern Morocco should be noted. Ocean Drilling Program (ODP) Site 976 in the western Alboran Sea is indicated.

porphyroclastic and mylonitic microstructures. These deformational structures in turn became overprinted by a stage of intense annealing recrystallization in the presence of percolating melts, prior to emplacement at crustal levels during the Early Miocene (22 Ma, Priem et al., 1979) facilitated by plagioclase-facies ductile shear zones. In addition, a recent study of the Carratraca massif (Tub|¤ a et al., 2004) has revealed the local preservation of early stage granular spinel-facies peridotites with protogranular orthopyroxene^spinel clusters presumably derived from precursor garnets. The emerging history of early stage exhumation and garnet breakdown, followed by deformation and development of a porphyroclastic microstructure prior to intense melt-assisted recrystallization has been interpreted to reflect Mesozoic rifting and subsequent lithospheric thickening, prior to intense heating related to late orogenic extension (van der Wal & Vissers, 1993, 1996; Vissers et al., 1995; Tub|¤ a et al., 2004).

In this study, we present the results of a microstructural and geochemical study of upper mantle xenoliths from the alkaline lavas of Cabezo Tallante, an eroded cinder cone NW of Cartagena (SE Spain) with Pliocene ArÂr ages ranging from 293 to 229 Ma (Duggen et al., 2005). These xenoliths exceptionally preserve microstructural and geochemical evidence of a multi-stage history of deformation, recrystallization, melt^rock interaction and melt intrusion tracking the progressive uplift of this lithospheric mantle sector in response to lithosphere extension in the Alboran Domain. In terms of tectonic and magmatic processes, the evolution of the Tallante mantle shows marked similarities to that documented in the Ronda peridotites as outlined above. A major aim of this study, therefore, is to provide further insights into the complex geodynamic and magmatic evolution of the Alboran Domain through a mantle perspective. Specific aims concern (1) the origin and chemical affinity of melts migrating through and intruding the peridotites at different lithospheric depths, and (2) the

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nature and conditions of these melt^rock interactions during progressive mantle uplift. The results of this study are discussed in the context of current geodynamic scenarios for the evolution of the western Mediterranean region.

S A M P L E S A N D A N A LY T I C A L P RO C E D U R E S The investigated xenoliths have decimeter-scale sizes and generally do not exceed 10^15 cm. They comprise (1) clinopyroxene-poor spinel lherzolites (cpx 510 vol. %), (2) plagioclase-bearing spinel peridotites, and (3) composite xenoliths made up of spinel and/or plagioclase peridotites intruded by olivineâmphibole-bearing pyroxenites. The most relevant microstructural relations are shown in Figs 2^5. Xenoliths show little microstructural evidence of interaction with the host basalts. This is confined to rare and very thin (550 mm) glassy fractures mostly occurring along grain boundaries or eventually crosscutting mantle minerals. When in contact with spinel, they cause tiny dark reaction rims. These sites have, however, been carefully avoided, and all the studied minerals were clean grains, with no evidence of interaction with the host basalt. On the other hand, we specifically investigated the interaction between peridotite and intruded amphibolepyroxenite veins in composite xenoliths. Clinopyroxene-poor spinel lherzolites display variable microstructures, mostly represented by porphyroclastic to equigranular types. The microstructure is: (1) dominantly porphyroclastic with large to medium grain sizes in spinel peridotites 92T4, 92T6, 92T19 (Fig. 2a^d), 92T17 (Fig. 3f), and 92T21; (2) porphyroclastic to granular in samples 92T3, 92T7 and T30; (3) fine-grained equigranular in samples 92T18 and 92T20 (Figs 2e and 3b, c) and rather coarse-grained equigranular in sample T18. Plagioclase-bearing peridotites are clinopyroxene-poor spinel lherzolites with variable amounts of plagioclase, mostly crystallized in interstitial aggregates, in association with olivine and/or orthopyroxene (Fig. 4a and b). Like the spinel peridotites, the plagioclase-bearing peridotites exhibit variable microstructure: (1) dominantly porphyroclastic in samples 92T1, 92T9A, T11, T13 and T31A; (2) porphyroclastic to granular in sample T32A; (3) tabular equigranular in samples 92T14 and 92T23. In most samples, the plagioclase modal abundance is lower than 5 vol. %, whereas it occurs in higher amounts (up to 10 vol. %) in samples 92T1 and 92T23. In peridotite 92T14, plagioclase occurs in very low modal amounts (53 vol. %), mostly as thin rims around spinel. Spinel peridotites T18 and 92T20, and plagioclasebearing peridotites T11 and T13 contain thin (51cm thick) gabbronoritic to noritic veins, mostly consisting of orthopyroxene, plagioclase, and subordinate clinopyroxene. They crosscut the peridotite fabric and show a reaction


margin made of fine-grained orthopyroxene towards the host peridotite (Fig. 4c). Similar gabbronoritic lithotypes have been documented in previous studies on the Tallante xenoliths (Arai et al., 2003; Beccaluva et al., 2004; Shimizu et al., 2004, 2008) and ascribed to the intrusion of slabderived melts. Here we describe the main microstructural features of the gabbronorites, which allow us to position this intrusion event in the context of the multi-stage melt interactionîntrusion history recorded by the studied peridotite xenoliths. Detailed chemical investigations of the gabbronorite veins will be subject of a separate paper (Rampone et al., in preparation). Plagioclase-bearing peridotites 92T9A, T31A, T32A and T13 are parts of composite xenoliths, consisting of country peridotite and intruded olivineâmphibole-bearing pyroxenite (samples 92T9B, T31B and T32B, respectively). The pyroxenites display a rather sharp contact against the host peridotites, and cut at high angles across the peridotite foliation defined by spinel trails (Fig. 5a and b). Major and trace element mineral chemistry data have been obtained from a selected number of samples (spinel peridotites 92T20, T30, 92T6, 92T7, 92T19; plagioclasebearing peridotite 92T1, composite xenoliths 92T9A-B, T31A-B), representative of all microstructural types (Tables 1^6). Mineral major element compositions were analysed using: (1) a Philips SEM 515 equipped with an X-ray dispersive analyser (accelerating potential 15 kV, beam current 20 nA), at the Dipartimento per lo Studio del Territorio e delle sue Risorse, University of Genova, and (2) a JEOL JXA 8200 Superprobe equipped with five wavelength-dispersive (WDS) spectrometers, an energydispersive (EDS) spectrometer, and a cathodoluminescence detector (accelerating potential 15 kV, beam current 15 nA), at the Dipartimento di Scienze della Terra, University of Milano. In situ trace element mineral analyses were carried out by laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS) techniques at IGG-CNR in Pavia. Detailed descriptions of these analytical procedures have been given by Tiepolo et al. (2003), Miller et al. (2007) and Rampone et al. (2008a).

M I C RO S T RU C T U R E S Spinel peridotites Porphyroclastic spinel peridotites consist of large (millimeter-size) olivine and orthopyroxene, and smaller clinopyroxene and spinel grains. Accessory amounts (51 vol. %) of phlogopite and amphibole have been observed in a few samples. Pyroxenes often display a preferred orientation of the cleavage planes and exsolution lamellae or blebs, whereas olivines are kinked, thus indicating internal plastic deformation of the primary mineral assemblage (Fig. 2a and b). Spinel occurs both as single grains and in rounded clusters, in association with orthopyroxene and subordinate clinopyroxene and olivine

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Fig. 2. Microstructures in spinel peridotites. (a) Porphyroclastic spinel peridotite 92T6. Large kinked olivine porphyroclast (OLp) (crosspolarized light). (b) Porphyroclastic spinel peridotite 92T19. Large, exolved orthopyroxene porphyroclast partly replaced by new olivine (cross-polarized light). (c) Porphyroclastic spinel peridotite 92T19. Close-up view of (b), showing new underformed olivine grains (OLn) partly replacing, with lobate contacts, pyroxene (OPXp, CPXp) and olivine (OLp) porphyroclasts (cross-polarized light). (d) Porphyroclastic spinel peridotite 92T4. New unstrained poikilitic orthopyroxene (OPXn) partly replacing both new (OLn) and kinked porphyroclast (OLp) olivine grains (cross-polarized light). (e) Equigranular spinel peridotite 92T18. Equigranular olivine, orthopyroxene, clinopyroxene grains showing triple-point junctions (cross-polarized light). (f) Equigranular spinel peridotite 92T14. Orthopyroxene (opx) þ spinel (sp) rounded cluster surrounded by finer-grain equigranular matrix (plane-polarized light).

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Fig. 3. Microstructures in spinel peridotites showing progressive deformation of pyroxene^spinel clusters. (a) Equigranular spinel peridotite 92T14 (plane-polarized light). (b) Equigranular spinel peridotite 92T20 (plane-polarized light). (c) Equigranular spinel peridotite 92T18 (plane-polarized light). (d) Porphyroclastic spinel peridotite 92T17 (plane-polarized light).

(e.g. 92T14; Figs 2f and 3a). In a recent study by Shimizu et al. (2008), similar clusters in the Tallante peridotite xenoliths have been interpreted to result from garnet breakdown. Spinel^pyroxene clusters show variable shapes, ranging from approximately spherical and ellipsoidal in some xenoliths (92T14) to strongly ellipsoidal and elongate (e.g. 92T18); many xenoliths show conspicuous trails of spinel suggesting that the clusters have been deformed to very high strains (e.g. 92T18, 92T17, Fig. 3c and d; 92T3). This interpretation is supported by a distinct preferred orientation of the orthopyroxene [100] cleavage planes in many of the stretched clusters. The highly deformed clusters contribute to defining a foliation in the rock, which is characterized by a spinel lineation in the plane of the foliation and is often accentuated by the elongation of olivine and orthopyroxene. To investigate the deformation associated with the spherical to ellipsoidal clusters, we measured the olivine lattice preferred orientation (LPO) in plagioclase-bearing spinel peridotite 92T14; this is characterized by the occurrence of such clusters and a tabular equigranular microstructure.

The result is shown in Fig. 6 (top panel). The LPO can be classified as a typical A-type pattern (terminology according to Jung & Karato, 2001; Karato et al., 2008), with a point maximum of the [010] axes at high angles to the foliation, a maximum of the [100] axes close to the lineation, and a maximum of the [001] axes in the plane of the foliation normal to the lineation. Similar fabrics have been produced in simple shear deformation experiments (Zhang & Karato, 1995; Katayama et al., 2004) and are commonly interpreted to reflect dominant olivine (010)[100] slip expected for lithospheric conditions of elevated temperatures (around 11008C), low stress (5300 MPa) and low water content (Karato et al., 2008). The evidence in sample 92T14 for simple shear deformation led us to select a set of additional xenoliths showing increasing degrees of deformation of the spinel^pyroxene clusters, to explore if there are any trends in the LPOs and microstructure associated with this deformation. Figure 6 illustrates the microstructure and olivine LPOs in six xenoliths, shown in qualitative order of increasing flattening of the clusters. The upper one (92T14) described

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Fig. 4. Microstructures in plagioclase-bearing peridotites and gabbronorite veins. (a) Porphyroclastic peridotite T11. (ol þ plag þ opx) granoblastic pocket interstitial between kinked mantle olivines (OLp) (cross-polarized light). (b) Porphyroclasticêquigranular peridotite T32. Interstitial orthopyroxene (opx) replacing and lobating triple point junction between equigranular olivines (OLeq) (crosspolarized light). (c) Porphyroclastic peridotite T11: (orthopyroxene^ plagioclase) gabbronoritic vein crosscutting large kinked olivine porphyroclast (OLp). Noteworthy features are the fine-grained orthopyroxene reaction rim towards the host peridotite, and spinel-facies melt^rock interaction textures in the host peridotite [replacement of olivine porphyroclast (OLp) by new orthopyroxene (Opxn)] (crosspolarized light).

Fig. 5. Meso- and micro-structures in composite olâmph-pyroxenite^peridotite xenoliths 92T9 and T31. (a) Decimeter-size xenolith; pyroxenite crosscutting peridotite foliation (marked by spinel trails). (b) Reaction zone at the contact between pyroxenite and host peridotite, characterized by crystallization of new clinopyroxene (plane-polarized light). (c) Euhedral olivine (ol) and poikilitic brown amphibole (amph) in pyroxenite (cross-polarized light). (d) Subhedral clinopyroxene (cpx) and interstitial plagioclase (plag) in pyroxenite (cross-polarized light).

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[100]


[010]

[001]

92T14

92T18

92T21

92T17

T18

92T3

Fig. 6. Microstructures and olivine lattice preferred orientations (LPO) of six selected xenoliths. Microstructure sketches were made from photomicrographs, and show olivine without shading with fine dashed lines indicating subgrain boundaries, orthopyroxene with dashed line pattern, clinopyroxene with dashed-point ornamentation, and spinel in black. Ubiquitous olivine 1208 triple junctions in all samples should be noted. LPO patterns all shown with the foliation defined by stretched spinel aggregates oriented east^west and the spinel stretching fabric horizontal. Samples are shown in qualitative order of increasing strain, from relatively low strain in sample 92T14 to high strain and very intense stretching of the orthopyroxene^spinel aggregates in sample 92T3. The angle between [100] maxima and foliation in the upper four samples should be noted. All diagrams show 100 measurements contoured at 1, 3, 5, 7, etc. times uniform distribution. (For further explanation see text.)

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above contains plagioclase, four xenoliths are spinel lherzolites, and one sample (T18) is transected by a gabbronoritic dikelet. All of these samples were cut perpendicular to the macroscopically visible foliation and parallel to the spinel lineation. As in sample 92T14, most LPO patterns are A-type fabrics, but samples 92T18 and to a lesser extent 92T21 show a tendency of the [010] and [001] axes to form girdles. These two latter LPOs are close to D-type fabrics (Jung & Karato, 2001) and suggest a transition to olivine {0kl}[100] slip, again at elevated (although possibly somewhat lower) temperatures and low water contents, but at higher stresses (4300 MPa; Karato et al., 2008). Close inspection of the LPOs reveals a clear tendency of the [010] axes to be more perpendicular to the plane of the foliation at inferred high strains, whereas the [100] axes at high strains make smaller angles or become parallel to the lineation. It follows that, at higher strains, the foliation defined by the deformed spinel^pyroxene aggregates tends to make decreasing angles with the flow plane, a phenomenon to be expected when sampling along strain gradients in heterogeneous simple shear zones. It should be noted that all LPOs in Fig. 6 are shown with the inferred shear sense dextral. Although the orientations of the LPO patterns seem to systematically vary with strain as inferred from the deformed shapes of the clusters, there is no clear relationship between the increasing strain denoted by the elongation of clusters, LPO patterns and the microstructure of the xenoliths. As shown in Fig. 6, in the six selected samples representing increasing amount of strain, the microstructures vary from porphyroclasticêquigranular (92T14), to equigranular (92T18), to porphyroclasticêquigranular (92T21, 92T17), to again strongly equigranular types (T18, 92T3). The ubiquitous presence in all samples of 1208 olivine triple junctions and relatively straight and slightly curved olivineôlivine grain boundaries indicates that the microstructure predominantly resulted from surface energy driven migration recrystallization. We therefore infer that the microstructure of the xenoliths predominantly reflects a stage of migration recrystallization that to a variable extent obliterated the precursor, presumably porphyroclastic microstructure associated with the hightemperature shearing. There is distinct microstructural evidence that the strong equigranular recrystallization seen in many of the Tallante xenoliths is in some way related to melt^rock interaction processes. The most porphyroclastic types of spinel peridotite exhibit peculiar microstructures indicative of multiple stages of melt^rock interaction: (1) crystallization of undeformed lobate olivine rims partly replacing exolved pyroxenes and kinked olivine porphyroclasts (Fig. 2c) and (2) later diffuse crystallization of unstrained poikilitic orthopyroxene grains that partly

corrode both the porphyroclastic minerals and the new olivine of the previous stage (Fig. 2d). Small unstrained clinopyroxene grains are sometimes associated with the poikilitic orthopyroxene. In peridotites with an intermediate porphyroclastic^granular microstructure, both the porphyroclastic minerals and the melt^rock reaction structures are increasingly replaced by the development of a medium- to fine-grained granoblastic assemblage of olivine þ pyroxene þ spinel grains, typically showing sharp boundaries and triple junctions. The porphyroclastic fabric is almost completely obliterated in the equigranular peridotites, which consist of a recrystallized (olivine þ pyroxene þ spinel) and sometimes tabular, fine-grained equigranular matrix (Fig. 2e). In these peridotites, (pyroxene þ spinel) clusters are still recognizable, although they also consist of equigranular recrystallized grains (Fig. 3b), but the melt^rock interaction textures are only rarely preserved. The equigranular recrystallization thus occurs at the expense of variably deformed porphyroclastic peridotites, as also evidenced by the different degree of cluster deformation and flattening in the equigranular xenoliths. For example, equigranular peridotite 92T20 and tabular equigranular peridotite 92T14 exhibit rounded to ellipsoidal clusters, whereas porphyroclastic to granular peridotites 92T17 and T30 (not shown) are characterized by the occurrence of aligned spinel trails (Fig. 3a^c). We emphasize again that the olivines in the strongly recrystallized equigranular peridotites preserve their LPO related to the development of the spinel trails and partially preserved porphyroclastic microstructure. Microstructural evidence and the LPO results thus indicate that the final stages of equigranular recrystallization occurred subsequent to both the deformation and melt^peridotite interaction events.

Plagioclase-bearing spinel peridotites In the plagioclase-bearing peridotites the spinel-bearing microstructure ranges from porphyroclastic to equigranular, thus covering the whole variability described above. In the less recrystallized porphyroclastic samples, the melt^rock reaction microstructures (olivine and/or orthopyroxene replacement of spinel-facies porphyroclasts) are frequently preserved. The olivine LPOs in plagioclasebearing equigranular peridotites 92T14 (Fig. 6) and 92T23 (not shown) are again high-temperature, low-stress A-type fabrics as seen in the spinel peridotites. Plagioclase occurs in highly variable modal amounts in the samples (from 55% to about 10 vol. %) and mostly crystallizes in small-grained granoblastic aggregates (together with olivine and/or orthopyroxene), interstitial between the spinel-facies minerals. In samples with low plagioclase modal abundance (e.g. T11, 92T14), plagioclase crystallizes both as rims replacing spinel (in association with fine-grained olivine), and in (plagioclase þ olivine)

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Table 1: Major (wt %) element compositions of olivines in peridotites and pyroxenites Sample: 92T20 T30 grain

new

SiO2

4079

FeO

852

MgO

4966

MnO

016

014

CaO

005

005

Total

9918

Mg-no.

912

T30

92T6

92T7

92T7

porph porph new

new

porph new

4112 4095

92T6

92T19 92T19 92T1

92T1

new

porph aggr

porph

new

idiom

small

small

3913

906

4978 4964

4920

4943

016

012

015

013

007

008

008

010

014

015

013

012

014

022

016

003

002

004

002

001

003

002

002

004

001

003

003

003

004

004

007

10010 9978

9973

9943 10025 9959

9986 9994

9963

10024 9945

10026

10041 9961

9984

905

907

914

913

908

908

5047 5033

916

916

874

853

5011 5009

5007

911

872

876

4981 4971

10007 9975 911

911

4121 4074

porph

4075

837

4136 4112

T31A 92T9A 92T9B 92T9B 92T9B

918

822

4091

T31A

4121

828

4127 4101

aggr

900

901

4135 4096

92T7

4078

3976 3927

959

1071

1887 1948

2006

4996 4897

4860

4152 4066

4037

891

909

901

890

797

788

021

782

Grain: olivine in the equigranular recrystallized matrix. Porph: olivine porphyroclast. New: unstrained olivine replacing kinked olivine porphyroclasts. Aggr: olivine grain in interstitial granoblastic aggregates. Idiom: idiomorphic grain. Small: small anhedral crystal.

interstitial aggregates (Fig. 4a). Plagioclase-rich (410 vol. %) peridotites 92T1 and 92T23 exhibit the diffuse occurrence of gabbroic pockets, made of plagioclase þ orthopyroxene olivine, that crystallized interstitially and partly replace porphyroclastic and equigranular spinel-facies minerals. In the equigranular peridotites, thin orthopyroxene and/or plagioclase rims are often observed corroding triple junctions between olivine granoblastic grains (Fig. 4b), indicating that crystallization of the plagioclase-bearing aggregates occurred in an advanced stage of migration recrystallization, or even after the development of the equigranular microstructure.

Intrusive rocks: gabbronorites and olivineâmphibole pyroxenites Thin (51cm) gabbronoritic to noritic veins have been observed both in spinel (sample 92T20) and in plagioclase-bearing (T11, T13, T18) peridotites, indicating that the melt impregnation process causing diffuse plagioclase crystallization in the peridotites and the intrusion of gabbronoritic veins are not directly related events. This is also supported by the distinct geochemical signature of the related melts (see discussion below). The veins show clear intrusive relationships relative to the wall-rock peridotite, and frequently transect large olivine and orthopyroxene porphyroclasts. The inner parts of the veins consist of up to millimeter-size orthopyroxene and plagioclase grains, whereas the vein margins against the host peridotites are made by palisade-type, oriented, fine-grained orthopyroxene (Fig. 4c). Small euhedral apatite crystals are observed as inclusions in both vein-forming plagioclase and orthopyroxene. In plagioclase-bearing peridotite T13, the gabbronoritic vein is cross-cut at a high angle by an olivine^ amphibole-bearing pyroxenite, which indicates that the pyroxenite intrusion was subsequent to the gabbronorite intrusion stage.

Pyroxenites display a clearly magmatic hypidiomorphic texture, defined by large (millimeter-size) euhedral olivine and subhedral clinopyroxene, poikilitic brown amphibole, interstitial plagioclase and minor phlogopite (Fig. 5c and d). Large poikilitic amphibole grains often include both olivine and clinopyroxene. Small euhedral Fe-oxide grains frequently occur as inclusions in clinopyroxene and amphibole. Fine-grained allotriomorphic aggregates of clinopyroxene þ olivine þ amphibole phlogopite plagioclase grains have crystallized between the matrix (olivine, clinopyroxene) minerals. The microstructure therefore indicates the following crystallization order: olivine, Fe-oxides, clinopyroxene, amphibole, phlogopite and plagioclase. In all composite xenoliths, the contact between pyroxenite and the host peridotite is marked by a centimeter-thick reaction zone (Fig. 5b), characterized by an enrichment in clinopyroxene and by the occurrence of small interstitial grains of brown amphibole and phlogopite. The newly formed clinopyroxene is clearly distinguishable from mantle clinopyroxene because it is interstitial, rich in fluid inclusions, and tends to replace mantle minerals (olivine and spinel).

M AJ O R A N D T R AC E E L E M E N T M I N E R A L C H E M I S T RY Despite of the observed microstructural variability, mineral compositions in the studied spinel peridotites are rather homogeneous and consistent with previously published data on Tallante cpx-poor lherzolites (Beccaluva et al., 2004). Olivines show limited chemical heterogeneity (Table 1): Mg-numbers range between 905 and 916, with the lowest values observed in equigranular peridotite 92T20 and porphyroclastic^granular peridotite T30. Within a single sample, no appreciable chemical changes are observed in olivine pertaining to different

304

RAMPONE et al.


microstructures (i.e. porphyroclastic olivine, new olivine rims or recrystallized equigranular olivine grains). Clinopyroxenes show rather narrow Mg-number variation (912^932), consistent with that observed in olivine (Table 2; Fig. 7). They have slightly variable Ti and Al contents (TiO2 ¼ 024^044 wt %; Al2O3 ¼ 542^ 664 wt %), not correlated with the Mg-number; within each sample, the highest Al contents are generally observed in clinopyroxene porphyroclasts, relative to recrystallized grains. In both porphyroclastic and equigranular peridotites, clinopyroxenes exhibit rather homogeneous contents in moderately incompatible trace elements [i.e. rare earth elements (REE), Zr, Hf, Ti, Y], and are characterized by slight light REE (LREE) depletion and almost flat middle to heavy REE (MREE^HREE) spectra (LaN/YbN ¼ 037^062; SmN/YbN ¼ 089^123). On the other hand, they display highly variable Th, U and Nb abundances, and such variations are observed within a single sample (e.g. peridotite 92T20; Table 2, Fig. 8); despite highly variable absolute contents, all clinopyroxenes show U and Th enrichment, and constant negative Nb anomalies. Similar trace element spectra in clinopyroxene, characterized by large U and Th heterogeneities coupled with more constant REE contents, have already been documented in previous studies of the Tallante spinel peridotites (Beccaluva et al., 2004; Shimizu et al., 2008). Orthopyroxenes have Mg-numbers ranging from 901 to 920, consistent with the compositions of olivine and clinopyroxene; the lowest Mg-number values are recorded in samples 92T20 and T30. In the porphyroclastic peridotites, no appreciable chemical differences are observed between orthopyroxene porphyroclasts and new poikilitic grains crystallizing at the expense of olivine, except for slightly lower Al and Cr contents in the latter (Al2O3 ¼ 468^585 wt % and 415^503 wt %, Cr2O3 ¼ 050^063 wt % and 030^048 wt %, in porphyroclastic and poikilitic grains, respectively). Few orthopyroxene porphyroclasts have been analysed for trace elements: they have LREE-depleted spectra (LaN/SmN ¼ 020^ 036), and display U and Th enrichment, similar to clinopyroxene. In plagioclase peridotite 92T1, the major element compositions of primary spinel facies minerals (olivine, orthopyroxene, clinopyroxene), as well as the chemical variations in a specific mineral pertaining to different microstructures (e.g. orthopyroxene porphyroclasts and new poikilitic crystals) are generally similar to those documented in spinel peridotites (see Tables 1, 3 and 4). In terms of minor and trace elements, clinopyroxenes display slightly higher Ti concentrations relative to clinopyroxenes in spinel peridotites, at constant high Mg-numbers (0924^093; Fig. 7). They also exhibit overall enrichment in REE, negative SrN anomalies, sometimes coupled to development of negative EuN, although preserving LREE

depletion (LaN/SmN ¼ 016^028; Fig. 8c). Similar chemical effects in clinopyroxene (i.e. Ti, REE enrichment) are widely documented in plagioclase-bearing impregnated peridotites (Rampone et al., 1997, 2008a; Dijkstra et al., 2003; Piccardo et al. 2007). Clinopyroxenes moreover display variable U and Th enrichment relative to Nb and Ta, although not reaching the high values observed in some spinel peridotites. Plagioclases have anorthite (An) contents ranging from 576 to 58 6, rather high Sr abundances (220^265 ppm) and LREE-enriched REE spectra (LaN/ SmN ¼ 44^54). Minerals in the pyroxenites show rather evolved major element compositions. Olivines have Mg-numbers ranging from 782 to 797, with the lowest values occurring in the fine-grained aggregates (Table 1). Clinopyroxenes also show rather low Mg-number values (80 4^847) and a general Mg-number decrease from large euhedral to fine-grained anhedral crystals. Both clinopyroxene and amphibole have high Ti contents (TiO2 is 073^099 wt % and 270^440 wt % in cpx and amph, respectively; Table 5). According to Leake’s (1997) classification, the amphiboles are pargasites to pargasitic hornblendes, with Mg-numbers of 770^820. Plagioclases have low An contents (358^44 6; Table 6). The trace element compositions of clinopyroxene and amphibole in pyroxenites are shown in Fig. 9. Clinopyroxenes display convex-upward MREE spectra and low HREE abundances (LaN/SmN ¼ 099^ 16; SmN/YbN ¼ 393^705), similar to clinopyroxenes in pyroxenites crystallized from alkaline melts (Fabries et al., 1989; Bodinier et al., 1990; Downes et al., 1991; Downes, 2001). They moreover exhibit marked negative NbN^TaN anomalies, and strong U and Th enrichment [up to 70 Primitive Mantle (PM)]. The REE spectra of amphibole are similar to those of clinopyroxene (LaN/SmN ¼ 088^ 257; SmN/YbN ¼ 627^705), but shifted to higher REE absolute concentrations; in contrast to clinopyroxene, they display NbN^TaN enrichment, and more variable U and Th contents. Higher trace element contents in amphibole are coupled to lower Mg-number values. Phlogopite has high Ba (4865 ppm) and appreciable Sr (61ppm) and Nb (203 ppm) abundances. Clinopyroxene porphyroclasts in the host peridotites T31A and 92T9A have more variable and, on average, lower Mg-number (89 8^926) relative to clinopyroxene in spinel and plagioclase peridotites (see Fig. 7). In terms of trace elements, clinopyroxenes preserve moderate incompatible element spectra similar to those in plagioclase peridotites, but are selectively enriched in LREE (LaN/YbN ¼ 076^081; SmN/YbN ¼ 075^093) and have U and Th abundances similar to those of clinopyroxenes in the pyroxenites. New interstitial clinopyroxenes, crystallized in the centimeter-thick reaction zone at the contact with pyroxenite, have intermediate Mg-number (877^88 6) and TiO2 compositions with respect to clinopyroxenes in

305


VOLUME 51

NUMBERS 1 & 2


Table 2: Major (wt %) and trace (ppm) element compositions of clino- and orthopyroxenes in spinel peridotites 92T20

92T20

92T20

92T20

92T20

T30

T30

T30

92T6

cpx

cpx

cpx

cpx

opx

cpx

cpx

opx

cpx

92T6 cpx

cluster

grain

grain

cluster

cluster

grain

grain

porph

porph

porph

wt % SiO2

5187

5173

5141

5160

5609

5243

5156

5435

5223

5200

TiO2

043

042

040

047

007

044

044

009

032

030

Al2O3

610

629

639

607

324

542

656

585

641

664

Cr2O3

100

108

111

109

031

087

114

054

104

106

FeO

237

238

245

238

617

242

253

570

248

249

MgO

1560

1549

1540

1545

3355

1636

1568

3314

1531

1520

MnO

011

008

010

009

015

013

012

014

008

CaO

2181

2187

2219

2239

066

2218

2158

086

2162

009 214

Na2O

088

091

082

073

003

048

062

000

111

114

Total

10017

10025

10027

10027

10022

10073

10023

10067

10059

10032

921

921

918

920

906

923

917

912

917

916

Mg-no. ppm Sc

52

52

50

52

17

61

53

19

64

56

V

262

260

271

233

84

244

250

97

262

255

Sr

62

58

60

71

047

27

34

023

50

45

Y

134

134

125

152

076

141

145

098

154

127

Zr

19

19

19

24

118

23

20

149

18

14

Nb

064

054

062

073

001

007

012

001

015

018

La

119

123

117

132

001

078

089

001

075

080

Ce

363

367

368

395

007

267

336

004

301

289

Pr

063

067

062

072

001

046

058

001

055

047

Nd

379

363

357

406

005

291

348

007

326

276

Sm

156

162

136

152

002

130

156

003

152

105

Eu

063

066

063

072

001

056

058

001

073

051

Gd

187

201

180

181

005

193

198

005

205

153

Tb

035

034

031

043

001

035

039

001

041

031

Dy

239

255

213

261

011

235

262

012

317

236

Ho

048

048

043

066

003

055

065

004

063

053

Er

159

144

145

142

010

151

141

013

187

161

Tm

019

020

020

024

002

027

023

003

027

024

Yb

136

149

138

158

018

158

161

028

188

158

Lu

019

019

019

020

004

018

027

005

025

021

Hf

068

059

052

077

003

075

062

006

056

037

Ta

0080

0054

0039

0047

50004

0014

0028

0001

0041

0028

002

007

006

0037

0031

0046

0015

0005

018

0082

0041

0024

0002

0070

0028

Pb

5002

004

003

007

Th

057

053

028

0098

U

021

017

012

0044

— 500025

—

Cr

7708

8371

8394

6664

2259

5955

6762

2818

7744

Co

21

23

23

21

53

21

21

51

22

23

2251

2155

2205

2360

540

2570

2274

551

1840

1538

Ti

8445

LaN/SmN

048

048

054

054

035

037

036

016

031

048

SmN/YbN

125

118

107

104

013

089

105

013

088

072

(continued)

306

RAMPONE et al.


Table 2: Continued 92T7

92T7

92T19

92T19

92T19

cpx

cpx

cpx

cpx

cpx

porph

porph

porph

porph

porph

wt % SiO2

5270

5233

5240

5172

5394

TiO2

028

024

029

041

007

Al2O3

577

573

640

620

578

Cr2O3

127

118

099

108

063

FeO

225

215

210

219

547

MgO

1612

1642

1597

1594

3339 015

MnO

009

006

006

004

CaO

2086

2116

2164

2120

078

Na2O

100

136

082

121

004

Total

10034

10063

10067

9999

10025

927

932

931

928

916

Mg-no.

Fig. 7. Variation of Mg-number vs Ti (1000) (atoms per six oxygens) in clinopyroxenes from spinel peridotites, plagioclase-bearing peridotite 92T1, and host peridotites to amphibole pyroxenites. The field refers to the compositions of clinopyroxenes in cpx-poor lherzolites from Beccaluva et al. (2004).

ppm Sc

68

61

50

45

18

V

272

270

260

228

96

Sr

48

57

57

46

106

111

101

102

14

18

17

181

Y Zr

969 16

the host peridotite and pyroxenite (Fig. 10a); their trace element contents are variable, mostly similar to those of clinopyroxenes in pyroxenites (Fig. 10b).

080

Nb

011

052

067

031

002

La

090

101

141

122

001

Ce

301

361

500

415

008

Pr

055

058

084

069

001

Nd

321

384

462

379

006

Sm

118

136

141

130

004

Eu

054

056

069

056

001

Gd

163

167

150

148

005

Tb

028

031

030

027

002

Dy

208

215

220

189

016

Ho

044

047

043

041

004

Er

126

129

120

096

014

Tm

019

018

017

017

002

T R AC K I N G T H E U P L I F T H I S T O RY O F T H E TA L L A N T E MANTLE The Tallante xenoliths record a multi-stage history of deformation, recrystallization, melt migration and melt intrusion. Together with the specific geochemical fingerprints in the various lithotypes, the overprinting relationships between the different microstructural features allow us to reconstruct a relative chronology of tectonic and magmatic events as follows.

Early decompression and deformation events

Yb

123

122

114

114

018

Lu

018

018

015

013

004

Hf

050

059

056

049

005

Ta

0048

0049

0043

0036

0002

Pb

0034

005

002

002

Th

0095

0111

0078

0070

0013

U

0074

0068

0026

0025

Cr Co Ti

9150

5002

11050

8704

7277

0005 3763

20

25

25

20

49

1696

1597

2124

1932

539

LaN/SmN

048

046

062

059

020

SmN/YbN

104

121

135

124

025

Cluster: clinopyroxene grain in (pyroxenes þ spinel) clusters. Grain: clinopyroxene in the equigranular recrystallized matrix. Porph: clinopyroxene porphyroclast.

The earliest microstructures, observed in both the more porphyroclastic and the intensely recrystallized spinel peridotites, are rounded to ellipsoidal clusters made of orthopyroxene þ spinel þ minor clinopyroxene and olivine. A previous study of the Tallante xenoliths (Shimizu et al., 2008) that focused on symplectitic spinel^pyroxene aggregates has shown textural and chemical characteristics similar to the spinel^pyroxene clusters of this study. On the basis of mass-balance calculations, Shimizu et al. (2008) demonstrated that the spinel^pyroxene aggregates were derived from garnet breakdown, according to the subsolidus reaction garnet þ olivine ! spinel þ orthopyroxene þ clinopyroxene. This indicates reequilibration of the Tallante peridotites from garnet- to spinel-facies conditions. In principle, the garnet^spinel

307


VOLUME 51

NUMBERS 1 & 2

(a)

(b)

(c) Fig. 8. Primitive mantle normalized trace element abundances of representative clinopyroxenes, orthopyroxenes and plagioclases in spinel- and plagioclase-bearing peridotites. (a) Porphyroclastic peridotites 92T6, 92T7 and 92T19. The light grey and dark grey fields refer to the compositions of clinopyroxenes in Tallante spinel peridotites from Beccaluva et al. (2004) and Shimizu et al. (2008), respectively. (b) Porphyroclasticêquigranular peridotite T30 and equigranular peridotite 92T20. Compositional fields as in (a). (c) Plagioclase-bearing peridotite 92T1. The dark grey field refers to the compositions of clinopyroxenes in spinel peridotites from this study. Normalizing values from McDonough & Sun (1995).

transition could be related either to isobaric heating or to decompression. However, equilibration temperatures in the porphyroclastic peridotites discussed below do not allow us to make any inference on a possible heating event


related to spinel^pyroxene cluster development. On the other hand, a transition from deep lithospheric (garnetfacies) to spinel-facies conditions has been documented in the Ronda and Beni Bousera massifs, both containing graphite pseudomorphs after diamond (Davies et al., 1993); the inferred garnet breakdown and development of the spinel^pyroxene clusters in the Tallante xenoliths may well reflect this early decompressional evolution. Consistent with previous work by Shimizu et al. (2008), we therefore favour the same interpretation for the Tallante mantle. The spinel^pyroxene clusters in the various xenoliths show variable degrees of flattening, from moderately ellipsoidal to intensely stretched spinel^pyroxene aggregates. Irrespective of the recrystallization microstructure, the olivine LPOs in a selected set of samples with different degrees of flattening of the clusters are mostly of the A-type, common in upper mantle peridotite massifs as well as in basalt-hosted peridotite xenoliths, suggesting high-temperature crystal^plastic ductile flow (e.g. Ave¤ Lallemant & Carter, 1970; Karato et al., 2008, and references therein). In addition, there is a marked trend in the samples studied for the angle between the stretching direction (marked by the elongate spinel^pyroxene clusters) and the flow plane to decrease with increasing strain. This strongly suggests that the strain in the samples and the associated LPOs reflect different degrees of shearing (Zhang & Karato, 1995; Karato et al., 2008). The development of the LPOs was probably related to the simultaneous development of a coarse-grained porphyroclastic microstructure, represented by a spinel-bearing mineral assemblage with evidence of internal plastic deformation such as large kinked olivines and coarse, partially exolved orthopyroxenes. The remnants of this porphyroclastic microstructure are best preserved in some of the spinel peridotites. A critical issue, relevant to the tectonic history of the upper mantle in this area, concerns the timing of the formation of the porphyroclastic microstructure and the development of the LPOs with respect to the inferred garnet breakdown and development of the clusters. It could be argued that intense shear deformation of a precursor garnet-bearing peridotite might have led to strongly stretched garnets, known to occur in garnetperidotite massifs (e.g. van Roermund et al., 2001) which then became transformed to spinel^pyroxene aggregates during later decompression. Alternatively, ductile shear flow may have affected a spinel peridotite containing spinel^pyroxene clusters resulting from a previous stage of garnet breakdown (i.e. ductile flow subsequent to early decompression). This latter interpretation is strongly supported by a distinct preferred orientation of the orthopyroxene [100] cleavage planes in many of the stretched clusters: hence we infer that the development of the porphyroclastic microstructure and related LPOs postdates

308

RAMPONE et al.


Table 3: Major (wt %) and trace element (ppm) compositions of clinopyroxenes in plagioclase peridotites 92T1

92T1

92T1

cpx

cpx

cpx

92T1 cpx

porph

grain

porph

grain

wt % SiO2

5183

5214

5161

5167

TiO2

054

056

068

071

Al2O3

651

569

664

647

Cr2O3

096

105

113

111

FeO

216

223

216

223

MgO

1606

1631

1586

1526

MnO

004

007

007

012

CaO

2176

2166

2186

2224

Na2O

063

067

054

083

K2O

000

000

000

000

10049

10038

10055

10064

930

929

929

924

Total Mg-no. ppm Sc

59

62

75

64

V

275

320

337

273

Sr

52

99

Y

226

214

12 294

245

66

Zr

30

25

31

28

Nb

007

010

048

La

063

064

097

062

Ce

361

383

450

314

Pr

083

078

094

060

Nd

553

488

737

536

Sm

242

204

216

232

Eu

075

072

115

094

Gd

322

310

391

318

Tb

065

053

078

064

Dy

424

402

488

430

Ho

100

082

146

097

Er

250

234

333

288

Tm

032

035

061

048

Yb

230

221

345

252

Lu

029

028

054

038

Hf

099

086

130

Ta

0008

0008

50058

—

5008

—

Pb

002

003

Th

0054

0020

0043

U

0018

0008

50008

Cr Co Ti

6460

7821

7958

—

104

0037 0009 6410

20

23

24

17

3831

3583

3401

3357

LaN/SmN

016

020

028

017

SmN/YbN

114

100

068

100

an earlier phase of decompression and garnet breakdown. At this stage, we note that the development of the spinel clusters probably involved an early stage of static recrystallization leading to grain coarsening before deformation, as breakdown of garnet in mantle rocks in the first instance typically leads to garnets surrounded by symplectites and/or kelyphites (e.g. van der Wal, 1993; Kaeser et al., 2006). In the majority of the spinel peridotites (i.e. in those with a porphyroclastic^granular microstructure) the primary porphyroclastic assemblage is clearly overprinted by largely surface energy driven annealing (migration recrystallization) leading to a commonly finer-grained, sometimes tabular, equigranular microstructure (e.g. samples 92T18 and 92T20). The degree of equigranular recrystallization, however, is highly variable between xenoliths. Strikingly, there is no correlation between the strains, inferred qualitatively from the degree of stretching of the spinel^pyroxene clusters, and the recrystallization microstructure. Highly recrystallized peridotites (e.g. samples 92T14, 92T18 and 92T20) may preserve rounded to ellipsoidal clusters, whereas highly elongate spinel trails occur in porphyroclastic^granular peridotites (e.g. samples 92T17 and T30). We therefore infer that the recrystallized granular microstructure of the xenoliths reflects a stage of migration recrystallization that largely obliterated a precursor porphyroclastic microstructure associated with spinel-facies ductile shearing. This interpretation seems consistent with another aspect of the microstructure, namely that the grain sizes of the xenoliths do not show any systematic change with strain either. An unpublished pilot study (de Boom, 1994) including four xenoliths from our collection (92T18, 92T3, T11 and T18) suggested that both the maximum and average olivine grain sizes depend on the grain sizes and volume fraction of the other phases, hence that the presence of the other phases controlled the recrystallized olivine grain sizes, again consistent with surface energy driven migration recrystallization. The intense recrystallization did not obliterate the LPO patterns, however, a phenomenon already known from structural and microstructural studies in the Ronda massif (van der Wal, 1993; van der Wal & Vissers, 1993, 1996; Vauchez & Garrido, 2001), but also from experimental studies in quartz aggregates (Heilbronner & Tullis, 2002). It can be concluded that the microstructures and LPO data for the spinel peridotites point to the early decompression of the Tallante mantle from garnet- to spinel-facies conditions. This was followed by ductile shear flow in the spinel facies, presumably at lithospheric conditions. This latter notion seems consistent with the occurrence of few LPO patterns showing a transition to D-type fabrics, known from experiments to develop at higher stresses expected at lithospheric temperature conditions

Grain: clinopyroxene in the equigranular recrystallized matrix. Porph: clinopyroxene porphyroclast. Aggr: plagioclase grains in interstitial granoblastic aggregates.

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Table 4: Major element (wt %) compositions of orthopyroxenes in spinel- and plagioclase-bearing peridotites 92T20

92T20

92T20

T30

T30

92T6

92T6

92T7

92T7

92T7

92T19

92T19

92T1

92T9A

92T9A

T31A

T31A

cluster

new

porph

porph

new

porph

new

porph

new

new

porph

new

new

porph

new

porph

new

5505

SiO2

5609

5568

5435

5466

5413

5469

5491

5497

5495

5437

5520

5511

5377

5500

5447

5487

TiO2

007

012

0143

009

008

012

013

011

011

014

010

012

016

012

005

007

011

Al2O3

324

413

495

585

503

538

483

468

451

415

528

403

454

553

404

444

422

Cr2O3

031

049

059

054

030

053

038

050

052

048

055

043

051

056

055

048

052

FeO

617

644

638

570

572

585

605

544

529

533

551

555

566

582

610

573

586

MgO

3355

3311

3242

3314

3404

3310

3353

3383

3412

3386

3321

3422

3287

3328

3334

3394

3405

MnO

015

0139

017

014

016

023

014

005

007

018

015

017

020

013

011

010

017

CaO

066

067

109

086

075

081

086

078

079

073

081

078

086

086

079

076

081

Na2O

003

003

004

000

000

005

004

006

003

003

000

000

002

003

002

002

000

Total

10027

10081

10083

10067

10074

10020

10065

10036

10041

9985

9998

10050

9993

10010

10000

10001

10061

906

902

901

912

914

910

908

917

920

919

915

917

912

911

907

913

912

Mg-no.

Cluster: orthopyroxene grain in (pyroxene þ spinel) clusters. New: unstrained poikilitic orthopyroxene replacing olivine porphyroclasts and new olivine. Porph: orthopyroxene porphyroclast.

(T511008C). The resulting porphyroclastic microstructure became in turn overprinted by variable degrees of migration recrystallization leading to intermediate (porphyroclastic^granular) as well as intensely recrystallized equigranular microstructures. We envisage that this migration recrystallization may well be related to the melt^rock interaction processes discussed below.

Melt migration by reactive porous flow and annealing recrystallization The replacement of porphyroclastic minerals by unstrained olivine rims, and the subsequent crystallization of undeformed poikilitic orthopyroxene at the expense of both porphyroclasts and newly crystallized olivine (see Fig. 2c and d) indicate that the porphyroclastic peridotites were affected by melt migration and melt^rock interaction. Similar microstructures have been described in spinel peridotites from ophiolitic and oceanic settings (Dijkstra et al., 2003; Piccardo & Vissers, 2007; Piccardo et al., 2007; Seyler et al., 2007; Rampone & Borghini, 2008; Rampone et al., 2008a). They have been related to open-system reactive porous flow of olivine-saturated tholeiitic melts that progressively shift towards orthopyroxene saturation during percolation and interaction with the host peridotites. A series of studies on the mechanisms of melt migration in the lithospheric mantle (Quick, 1981; Kelemen, 1990; Kelemen et al., 1992, 1995, 1997; Kelemen & Dick, 1995) have pointed out that melts rising adiabatically are saturated in olivine; they will therefore crystallize olivine and dissolve pyroxenes in the host peridotites until, through continuous ascent and interaction with the host peridotites, they reach pyroxene saturation and start to

crystallize ortho- and clinopyroxene. The diffuse crystallization in the Tallante peridotites of orthopyroxene at the expense of previous olivine-replacement textures may thus indicate various stages of reactive porous flow and interaction with increasingly modified melts. Another remarkable feature in variably deformed porphyroclastic peridotites is the partial obliteration of the porphyroclastic and melt^rock reaction microstructures by the development of a granular texture. This again indicates that the allied migration recrystallization was not induced by the shearing deformation as outlined above, but that the equigranular peridotites developed as the result of extensive static, possibly melt-assisted annealing recrystallization. As textural evidence clearly indicates reactive porous melt flow, we suggest that the annealing recrystallization was largely related to pervasive melt percolation. Inferences on the origin and chemical affinity of the percolating melts at spinel facies conditions can be made considering the trace element composition of clinopyroxenes in the spinel peridotites. Overall, clinopyroxenes show highly variable U and Th enrichment, even within a single thin section. In peridotite 92T20, clinopyroxenes with the highest U and Th contents have been analysed in a (pyroxene^spinel) cluster close to a noritic vein, this latter being significantly U^Th enriched (Rampone et al., 2008b, 2009). Based on this evidence, we believe that heterogeneity in the Th and U contents in clinopyroxenes probably resulted from late-stage percolation of small melt fractions (Bedini et al., 1997; Ionov et al., 2002; Raffone et al., 2009), and we did not consider these elements to constrain the compositions of the migrating melts.

310

RAMPONE et al.


Table 5: Major (wt %) and trace (ppm) element compositions of clinopyroxenes, amphiboles and phlogopites in composite xenoliths 92T9 and T31 Host peridotite

Reaction zone

Pyroxenite

T31A

T31A

92T9A

92T9A

92T9A

92T9A

92T9A

T31B

T31B

92T9B

cpx

cpx

cpx

cpx

cpx

cpx

cpx

cpx

cpx

cpx

92T9B cpx

porph

porph

porph

porph

react

react

react

idiom

idiom

idiom

idiom

wt % SiO2

5212

5104

5245

5230

5166

5314

5361

5104

5200

5166

5060

TiO2

062

043

032

032

076

044

030

073

083

099

098

Al2O3

468

580

442

601

516

321

300

525

430

510

585

Cr2O3

111

108

116

101

102

090

132

027

036

066

066

FeO

326

277

335

226

396

417

377

631

554

476

530

MgO

1602

1559

1670

1585

1586

1679

1650

1450

1559

1568

1474

MnO

012

008

008

007

010

009

008

017

010

009

016

CaO

2155

2183

2133

2196

2128

2142

2106

2097

2135

2113

2177

Na2O

044

083

082

056

056

073

084

111

018

046

021

K2O

003

001

003

000

002

000

002

000

001

006

002

Total

9995

9946

10066

10034

10038

10089

10050

10035

10026

10059

10029

Mg-no.

898

909

899

926

877

878

886

804

834

826

832

ppm Sc

64

63

49

70

54

44

57

57

76

V

257

287

336

241

306

204

338

307

344

Sr

10

18

87

78

116

77

78

94

96

Y

196

164

121

113

111

109

114

102

130

Zr

24

23

67

24

21

70

27

31

53

Nb Ba

017 —

057 —

051

028

—

La

219

172

105

Ce

514

442

276

Pr

079

072

Nd

435

381

11 766 162

349

169

158

631

083 — 636 154 214 102

011

025

01

07

88 278 362 164

704 189 254 127

040 —

068 —

700 200

697 212

268 129

337 176

Sm

181

132

374

167

310

339

342

314

443

Eu

052

055

116

060

092

111

110

100

140

Gd

239

202

332

173

255

313

345

315

447

Tb

046

036

047

029

036

040

045

041

056

Dy

366

295

262

197

237

242

261

231

319

Ho

074

058

043

041

042

036

045

042

055

Er

222

191

113

109

104

092

118

101

126

Tm

031

027

013

015

015

014

016

012

015

Yb

211

190

109

117

091

072

095

082

092

Lu

028

026

012

015

012

012

011

009

014

Hf

069

063

177

064

069

210

111

130

244

Ta

005

006

006

003

006

002

003

004

009

Pb

039

072

33

18

36

21

22

19

18

Th

235

23

55

31

63

24

50

31

33

U

113

082

13

083

15

055

12

079

Cr Co Ti

8514

9159

7436

8231

9627

689

2890

5553

083 5745

19

22

24

22

29

35

37

29

30

2737

2028

8272

1632

4130

1910

6338

5793

8426

LaN/SmN

076

081

175

288

128

163

129

140

099

SmN/YbN

093

075

372

155

372

514

393

415

525

(continued)

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VOLUME 51

NUMBERS 1 & 2

Table 5: Continued Pyroxenite 92T9B

T31B

92T9B

92T9B

cpx

amph

amph

amph

T31B phlog

idiom

poikil

poikil

poikil

idiom

wt % SiO2

5218

4185

4170

4316

TiO2

074

381

439

269

3712 417

Al2O3

436

1297

1303

1314

1514

Cr2O3

060

013

008

011

000

FeO

500

917

932

768

943

MgO

1556

1486

1415

1595

1867 008

MnO

011

016

019

011

CaO

2166

1151

1140

1178

008

Na2O

036

259

240

244

184

K2O Total Mg-no.

001

129

141

131

881

10058

9834

9807

9837

9534

847

780

770

820

780

ppm Sc

63

39

33

40

V

366

428

492

420

296

Sr

84

574

778

514

62

Y

121

219

187

Zr

39

88

55

Nb Ba

061 —

160 136

50

28

72

543

710

534

534

005 149 204 4865

La

61

200

114

266

—

Ce

168

588

400

689

—

Pr Nd

239 121

840 406

639 338

820 325

— —

Sm

340

898

810

648

—

Eu

121

260

243

203

—

Gd

366

748

711

544

—

Tb

048

091

088

063

—

Dy

291

530

460

363

—

Ho

050

085

075

056

—

Er

116

214

185

149

—

Tm

013

026

022

019

—

Yb

093

175

125

112

—

Lu

011

023

016

015

—

Hf

178

309

199

347

034

Ta

008

289

166

358

140

Pb

29

62

94

Th

31

88

064

42

U

059

18

029

13

Cr Co Ti

4145

12

605

231

579

106 008 021 723

33

60

55

55

77

8653

27712

36055

18360

21818

LaN/SmN

112

140

088

257

SmN/YbN

397

557

705

627


Despite variable U and Th enrichment, clinopyroxenes display smooth trace element patterns characterized by slight LREE depletion. Such REE compositions (in whole-rock and clinopyroxene) are widely documented in lithospheric peridotites from both orogenic massifs and xenoliths, and can reflect either a pristine fertile mantle composition or a previously depleted lithospheric mantle modified and refertilized by interaction with percolating melts (Vannucci et al., 1991; Bianchini et al., 2007; LeRoux et al., 2007; Rivalenti et al., 2007; Raffone et al., 2009). In the Tallante xenoliths, significant geochemical variations are not observed between the porphyroclastic and equigranular spinel peridotites. Nevertheless, it is noteworthy that the clinopyroxenes in the equigranular peridotites are slightly more homogeneous in terms of REE contents, and exhibit less pronounced Zr, Hf and Ti anomalies. This suggests, in agreement with microstructural evidence, a larger time-integrated melt^rock ratio in the equigranular peridotites and, in turn, indicates that clinopyroxene chemistry probably reflects refertilization by melt^rock interaction. We therefore used the REE compositions of clinopyroxenes in the equigranular peridotites 92T20 and T30, inferred to represent the peridotites most equilibrated with the percolating melts, to derive the compositions of the equilibrium melts. Two sets of cpx/meltREE distribution coefficients were used, for Si-undersaturated and Si-saturated melts (Vannucci et al., 1998; Ionov et al., 2002), to account for the chemical modification of percolating melts from olivine to orthopyroxene saturation (inferred from microstructural evidence) during ascent and interaction with the host peridotites. The results are shown in Fig. 11a, together with the compositional fields for most primitive basaltic andesites and alkaline magmas of the Alboran region (Duggen et al., 2004, 2005, 2008). Despite variable absolute REE concentrations, depending on the choice of cpx/melt distribution coefficients, computed melts show moderate LREE enrichment and almost flat MREE to HREE patterns, consistent with the sub-alkaline magmatism of the Alboran Domain. We thus infer that the melt percolation events documented in the Tallante spinel peridotites were related to the sub-alkaline magmatism that affected the Alboran regionsince the latest Oligocene, presumably in response to processes in the upper mantle driving late-orogenic extension. It should be noted that a similar scenario has been envisaged for the Ronda peridotites (Van der Wal & Bodinier, 1996).

Melt impregnation: emplacement at shallow lithospheric depth

Porph: mantle porphyroclast. React: new clinopyroxene crystallized in the reaction zone. Idiom: idiomorphic grain. Poikil: poikilitic grain.

Exhumation of spheric levels is clase-bearing microstructures porphyroclastic

312

the Tallante mantle to shallower lithodocumented by the occurrence of plagiospinel peridotites showing peculiar (e.g. gabbroic pockets interstitial between and equigranular grains) and anomalous

RAMPONE et al.


Table 6: Major (wt %) and trace (ppm) element compositions of plagioclases T31A

92T9B

T31B

585

92T1

92T1

92T1

wt % SiO2

5269

5700

5462

5463

TiO2

000

007

012

002

000

5437 001

Al2O3

3055

2749

2606

2873

2876

2896

FeO

034

022

023

006

007

009

MgO

000

000

000

003

003

005

MnO

000

000

000

000

001

002

CaO

1193

899

700

1190

1196

1193

Na2O

465

618

694

484

481

483

K2O

015

023

053

007

010

010

10031

10018

9938

10028

10037

10035

586

446

358

576

579

577

Total An

Fig. 9. Primitive mantle normalized trace element abundances of clinopyroxenes and amphiboles in pyroxenites T31B and 92T9B. Normalizing values as in Fig. 8.

ppm Sr

189

223

256

265

Y

011

012

012

041

La

137

070

113

099

Ce

164

144

183

170

Pr

011

012

018

022

Nd

038

044

074

062

Sm

0051

009

013

014

Eu

017

060

083

084

Gd

0045

035

012

014

Ti

51

108

157

(a)

142

(410 vol. %) plagioclase modal enrichment. Similar microstructures have been widely documented in ophiolitic and oceanic peridotites, and are commonly ascribed to entrapment and crystallization of migrating melts (Dick, 1989; Cannat et al., 1990; Rampone et al., 1997, 2008a; Dijkstra et al., 2001; Tartarotti et al., 2002; Piccardo et al., 2004, 2007; Chazot et al., 2005). In plagioclase-bearing peridotite 92T1, clinopyroxenes exhibit significant Sr depletion and overall enrichment in REE, Ti and Zr for a similar LREE depletion as compared with clinopyroxenes in the spinel peridotites (Figs 7 and 8c). Comparable chemical features, specifically the REE increase, have previously been documented in clinopyroxenes from impregnated peridotites (Rampone et al., 1997, 2008a; Barth et al., 2003; Dijkstra et al., 2003; Piccardo et al., 2004, 2007; Rampone & Borghini, 2008), and ascribed to one (or a variable combination) of the following effects: (1) entrapment within the peridotites of small (53%) melt fractions; (2) melt^rock reaction at decreasing melt mass; (3) increase of the solid/liquid trace element partition coefficients as a result of the chemical evolution

(b)

Fig. 10. (a) Variation of Mg-number vs Ti (1000) (atoms per six oxygens) in clinopyroxenes from pyroxenites, the reaction zone at the host peridotite^pyroxenite contact, and host peridotites. The dark grey field refers to the compositions of clinopyroxenes in the spinel peridotites (this study). (b) Primitive mantle normalized trace element abundances in clinopyroxenes as in (a); the dark and light grey fields refer, respectively, to the compositions of clinopyroxenes in spinel and plagioclase peridotites (this study). Normalizing values as in Fig. 8.

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VOLUME 51

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Specifically, the impregnated plagioclase-bearing peridotites would correspond to the stage when upward migrating melts reached rather shallow and colder lithospheric environments where conductive cooling caused melt crystallization and possibly entrapment. In this context it is important to note the occurrence, in a few peridotites (e.g. 92T14, T11), of fine-grained rims of plagioclase þ olivine around spinel, indicating that some plagioclase formation can be related to subsolidus recrystallization from spinel- to plagioclase-facies conditions according to the reaction orthopyroxene þ spinel ! plagioclase þ olivine þ new pyroxenes. According to a recent experimental study of the spinel^plagioclase transition (Borghini et al., 2009), crystallization in fertile and depleted lherzolites of such a plagioclase-bearing assemblage indicates re-equilibration at P 5 09 GPa at about 1000^11008C. In addition, the microstructures of the plagioclase-bearing peridotites clearly indicate that plagioclase crystallization occurred in mantle rocks that previously experienced the deformation, melt^rock interaction and annealing recrystallization history documented in the spinel peridotites. This constitutes strong evidence that the plagioclase-bearing xenoliths are not simply samples from the shallower levels of the Tallante lithosphere, but that they record the progressive exhumation of the Tallante mantle, presumably in response to progressive extension in the Alboran region.

(a)

(b)

Fig. 11. (a) Primitive mantle normalized trace element abundances in computed melts in equilibrium with average clinopyroxene from equigranular peridotite 92T20 and porphyroclastic^granular peridotite T30. Symbols refer to different sets of partition coefficients used for the calculation (filled triangle, Ionov et al., 1992; filled square, Vannucci et al., 1998). The grey fields refer to the compositions of the most primitive basaltic andesites from the Alboran Domain (light grey; data from Duggen et al., 2004, 2008) and alkali basalts from Cabezo Tallante (dark grey; data from Duggen et al., 2005). (b) Primitive mantle normalized trace element abundances of computed melts in equilibrium with averaged clinopyroxenes from pyroxenites T31B and 92T9B. Dark grey field as in (a).

of the percolating melts. On the other hand, the similarity of LREE fractionation in clinopyroxenes from spinel and plagioclase peridotites suggests that the impregnating melts probably had a similar (sub-alkaline) chemical affinity to that of the percolating melts of the previous reactive porous flow stages. This is also consistent with the diffuse occurrence of orthopyroxene (olivine) in the plagioclasebearing gabbroic pockets, indicative of an opx-saturated signature of the impregnating melts. As inferred in ophiolitic peridotites from the AlpsÂpennine system and the Othris Massif (Dijkstra et al., 2003; Piccardo et al., 2007; Rampone & Borghini, 2008; Rampone et al., 2008a), open-system reactive porous flow at spinel-facies conditions and melt impregnation at plagioclase-facies conditions could represent different stages of the same melt percolation event at different lithospheric levels.

The melt intrusion events: from diffused to focused melt migration Veins and dykes of different compositions (gabbronorites and olivineâmphibole pyroxenites) with clear crosscutting relationships intrude both spinel and plagioclasebearing peridotites. Veins and dykes provide further evidence for the exhumation of the Tallante mantle towards a shallower and colder lithospheric environment, reflected by the transition from porous melt flow to magma emplacement in fractures, presumably related to a change in the rheology of the lithospheric mantle during extension-related exhumation and cooling. Gabbronoritic xenoliths and veins comparable with those described in this study have been reported in previous studies of the Tallante xenoliths (Arai et al., 2003; Shimizu et al., 2004, 2008; Beccaluva et al., 2007). On the basis of their Si-oversaturated composition, trace element (Th, U, REE, Sr) enrichment and radiogenic Sr isotope compositions, it has been inferred that they represent subduction-related melts, either approaching Miocene Betic lamproites in composition (Beccaluva et al., 2007) or adakitic melts produced by partial melting of a sinking slab beneath the Alboran Domain (Shimizu et al., 2004). Detailed microstructural and geochemical investigations of the gabbronorites will be the subject of a separate paper (Rampone et al., in preparation), because they constitute one of the few documented occurrences of Si-oversaturated

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large ion lithophile element enriched melts intruded and crystallized within the shallow lithospheric mantle (Bali et al., 2008; Mazzucchelli et al., 2004), which justifies specific attention. Preliminary geochemical investigations (Rampone et al., 2009) have confirmed that the parental melts to the gabbronorites are significantly enriched in LREE, Th, U and volatile (Cl) components: this provides further support for their slab-derived origin, and points to a continental crust (or sediment) source component, rather than oceanic crust. In the context of this study, which aims to explore the whole geodynamic evolution recorded in the Tallante mantle, the gabbronorites represent an important stage, as they provide striking evidence for the existence of a subducting slab beneath the region. The gabbronorite^norite veins are crosscut by olivine^ amphibole pyroxenite dykes, which represent the latest magmatic event affecting the Tallante mantle. According to phase relations in the H2O-bearing picrobasaltic system (Helz, 1973; Allen & Boettcher, 1983; Ulmer, 1989; Grove et al., 2003), the observed crystallization order in pyroxenites (olivine^clinopyoxeneâmphibole^ plagioclase) is indicative of intrusion pressures above 05 GPa, probably in the range 06^09 GPa. This is consistent with the evidence that the pyroxenites are intruded within plagioclase-bearing peridotites (equilibrated at P508 GPa). The inferred pressures are further supported by the 07^09 GPa estimates achieved using the Nimis & Ulmer (1998) single-clinopyroxene geobarometer. Major element mineral compositions in the amphibole pyroxenites indicate that they crystallized from rather evolved melts (e.g. Mg-number in olivine and clinopyroxene 782^797 and 80 4^847, respectively). Nevertheless, inferences on the chemical affinity of the parental melts can be made using the trace element compositions of clinopyroxene, which is an early crystallizing phase and thus better reflects the chemical composition of the migrating melt. In terms of REE patterns, clinopyroxenes in the pyroxenites closely resemble clinopyroxenes crystallized from alkaline melts (Fabries et al., 1989; Bodinier et al., 1990; Downes et al., 1991; Downes, 2001; see Fig. 9). This is confirmed by the REE compositions of computed equilibrium melts (using cpx/meltREE distribution coefficients from Ionov et al., 2002; Fig. 11b); they are similar to the compositions of the host alkali basalts (Duggen et al., 2005) in terms of LREE/HREE fractionation, although shifted to higher absolute concentrations because of their more evolved chemical signature. Parental melts to the amphibole pyroxenites were therefore alkaline magmas similar to the Tallante host alkali basalts. It is noteworthy that the clinopyroxenes exhibit more pronounced depletion in Nb, and to a minor extent Ta, relative to clinopyroxenes in equilibrium with alkaline melts. In principle, this could be an effect of the early precipitation of Fe-oxides (similar to those occurring as inclusions

in clinopyroxene and amphibole), which can fractionate these elements relative to the REE (Bodinier et al., 1996; Gregoire et al., 2000; Rivalenti et al., 2004). At the bulkrock scale, the low Nb and Ta contents in clinopyroxene are then primarily compensated by high Nb and Ta concentrations in pargasitic amphibole, as well as moderate Nb abundances in phlogopite. Both clinopyroxene and amphibole in the pyroxenites also show remarkable Th and U enrichment, at least one order of magnitude higher than expected on the basis of experimentally determined mineral/melt partition coefficients (Hauri et al., 1994; Lundstrom et al., 1994; Tiepolo et al., 2007) and ocean island basalt (OIB) compositions. Large Th and U concentrations can result from late-stage entrapment of small volumes of alkaline melts (Raffone et al., 2009). This could be especially true for poikilitic minerals such as amphibole, as suggested by its larger Th and U variability (see Fig. 9). On the other hand, new clinopyroxenes crystallized in the wall^peridotite reaction zone are similar to clinopyroxenes in pyroxenites in terms of U, Th and LREE (Fig. 10b): this argues against a simple trapped melt effect, and rather suggests that the observed Th and U enrichment in clinopyroxene represents the chemical signature of the parental melts. Chromatographic percolation of small, highly incompatible element enriched melt volumes through the host peridotites is then documented by the compositions of clinopyroxene porphyroclasts in the host peridotites (i.e. outside the reaction zone), these latter being selectively enriched in Th, U and LREE relative to clinopyroxenes in porphyroclastic spinel peridotites (see the comparison field in Fig. 10b). This process has been widely documented in metasomatized peridotites (Bedini et al., 1997; Ionov et al., 2002; Bodinier & Godard, 2003; Bodinier et al., 2004; Rivalenti et al., 2007; Raffone et al., 2009, and references therein), particularly in the host peridotites of amphibolebearing veins crystallized from alkaline melts (Bodinier et al., 1990, 2004; Zanetti et al., 1996). In summary, we infer that the observed Th and U enrichment in pyroxenite minerals possibly reflects the chemical signature of parental alkaline melts that could have been acquired by interaction with lithospheric mantle material enriched in these elements during previous U^Th-rich metasomatic stages (such as gabbronorite^ norite intrusion). In spite of this, the petrographic and goechemical characteristics of the amphibole pyroxenites clearly indicate that their parental melts had alkaline affinity, consistent with the Tallante host alkali basalts.

Geothermobarometry data and their significance Our microstructural and geochemical study of the Tallante xenoliths allows us to identify a multi-stage history of recrystallization, deformation, melt migration and intrusion as outlined above. The P^T evolution of the Tallante

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mantle associated with this history may help to elucidate its geodynamic significance; however, the thermobarometric data summarized below and their interpretation are surrounded with uncertainties, mostly related to latestage cooling of the studied xenoliths at shallow lithospheric depths. Early uplift of the Tallante mantle rocks and development of spinel^pyroxene clusters (qualitatively shown in Fig. 12 as field Ta1) was followed by ductile shear flow at spinel-facies conditions leading to a porphyroclastic microstructure. Geothermometric estimates on the porphyroclastic minerals using different methods (Brey & Ko«hler, 1990; Taylor, 1998; Witt-Eickschen & O’Neill, 2005) have yielded variable results, as has been frequently observed in the case of significantly exolved pyroxene grains. More consistent results were obtained with the Brey & Ko«hler (1990) Ca-in-Opx and the Witt-Eickschen & O’Neill (2005) geothermometers. Most temperature estimates are in the range 960^10108C, with a few values up to 1050^ 11008C (shown as field Ta2 in Fig. 12). In a few orthopyroxene porphyroclasts we analysed core^rim traverses; Ca profiles are generally smooth, and the resulting geothermometric estimates, using the Brey & Ko«hler (1990) method, mostly range from 950 to 10008C, with a few higher temperatures (to 11008C) rarely preserved in the cores. SimilarTestimates (960^10208C) were also obtained in spinel-facies porphyroclastic minerals from plagioclasebearing peridotites. Overall this indicates that any early high-Tequilibration (predating exolutions in pyroxenes) is poorly preserved. The data are consistent, however, with the development of the porphyroclastic microstructure in a lithospheric environment. These early microstructures are variably overprinted by recrystallization, which in all likelihood occurred in the presence of percolating melts, as suggested by the replacement of porphyroclastic minerals by unstrained olivine rims and the subsequent crystallization of undeformed poikilitic orthopyroxene at the expense of porphyroclasts and newly crystallized olivines. Temperature estimates on these poikilitic, replacive orthopyroxenes using the Brey & Ko«hler (1990) Ca-in-Opx method are in the range 960^ 10508C, similar to the T estimates on exsolved porphyroclasts; but pressures (within the spinel stability field) are poorly constrained (field Ta3 in Fig. 12). In several spinel peridotites recrystallization led to the development of a granular microstructure. Despite the microstructural evidence for high-temperature annealing, geothermometric estimates on these granular assemblages again yield rather low equilibration temperatures of 950^10008C. Progressive melt^rock interaction and migration (annealing) recrystallization was probably accompanied by exhumation of the mantle rocks to shallower levels, where melts began to crystallize interstitial plagioclase and gabbroic pockets, causing the anomalous


(410 vol. %) plagioclase modal enrichment observed in the plagioclase-bearing peridotite xenoliths. The latest stages of the Tallante mantle history are marked by the intrusion in spinel- as well as plagioclase-bearing peridotites of gabbronorite veins followed by olivineâmphibole pyroxenites, which occurred at 07^09 GPa (tentatively shown as field Ta4 in Fig. 12). The sequence of P^T conditions obtained from the Tallante xenoliths is characterized by markedly low temperatures, too low to account for the observed reactive porous flow and melt impregnation events. In all documented stages, Testimates vary in a narrow range, despite different microstructural sites and mineral assemblages. This suggests that all of the geothermometric results may largely reflect the effects of late-stage cooling before the xenoliths were sampled by the ascending host magmas. Such late-stage cooling in the Tallante mantle seems consistent with documented geochronological results from the crustal metamorphic rocks of the Alboran Domain, in particular in the western Betics, in which U^Pb ion microprobe dating of zircon, Ar/Ar dating of hornblende, Ar/Ar laserprobe dating of muscovite and biotite, and Essiontrack analysis of zircon and apatite all reveal that cooling was extremely rapid in the interval 212^204 Ma (e.g. Platt & Whitehouse, 1999; Platt et al., 2003); that is, well before xenolith entrainment in the uprising host lavas, during the Pliocene. Unlike the uncertainty on the geothermometry results, the reconstructed multi-stage history of deformation, recrystallization, melt migration and intrusion in the Tallante xenoliths clearly points to about 30 km of exhumation of the Tallante mantle, from P42 GPa (as indicated by the orthopyroxene^spinel clusters after garnet), to 07^09 GPa (marked by partial peridotite re-equilibration at plagioclase-facies conditions and intrusion of alkaline pyroxenites). This implies exhumation to a shallow level in response to Neogene lithosphere extension. In this context we emphasize that equilibration at 07^09 GPa of the Tallante xenoliths is entirely consistent with current estimates of the present-day crustal thickness of 15^20 km in the Alboran Sea and about 22 km in the Cartagena area (Torne et al., 2000).

T H E TA L L A N T E M A N T L E I N T H E A LBOR A N CONTEXT Comparison with the Ronda massif The Alboran Domain is an exceptional region for upper mantle studies, in that both mantle xenoliths (Tallante) and orogenic peridotites (Ronda) are clearly associated in space and time with one and the same large-scale orogenic process. This invites comparison of our xenolith results with the Ronda history as inferred by other studies (van der Wal, 1993; van der Wal & Vissers, 1993, 1996; van der

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Fig. 12. P^Tdata and inferred P^T paths for the Tallante mantle and Ronda peridotites. Peridotite solidus, garnet^spinel (g-s) reaction curve and ariegite^seiland subfacies boundary (a-s) redrawn after van der Wal & Vissers (1993); spinel^plagioclase peridotite boundary for fertile lherzolite (s-p) and depleted lherzolite (s-p’) according to Borghini et al. (2009). GRT, SPL and PLAG denote garnet, spinel and plagioclase peridotite facies, respectively. Fields Ta1^Ta4 show qualitative estimates (circled fields) and thermobarometry results (rectangles) for the Tallante mantle, with Ta1 garnet breakdown to spinel^pyroxene clusters, Ta2 geothermometry results obtained in porphyroclastic microstructures, Ta3 results for poikilitic replacive orthopyroxenes formed by melt^rock interaction and recrystallization, Ta4 development of plagioclase-bearing assemblages, fine granoblastic recrystallization, and intrusion of gabbronoritic and olivineâmphibole-pyroxenite veins. A (dashed black line), plausible P^T trajectory for the Tallante mantle inferred from extensive evidence for melt^rock interaction, as well as from comparison with Ronda peridotites. (Note inferred cooling stage towards field Ta4 at shallow levels.) B (light grey), P^T path for the Ronda peridotites according to Van der Wal & Vissers (1993), with R1 early spinel-facies equilibration, R2 development of spinel tectonites, R3 garnet^spinel mylonites, R4 conditions in spinel tectonites close to recrystallization front, R5 granular spinel peridotites, R6 emplacementrelated plagioclase-facies shear zones. C (dark grey), P^T path inferred by Platt et al. (2003) for rocks at the Ronda recrystallization front. Shaded field near R2 and R3 represents equilibrium conditions during development of spinel tectonites and garnet^spinel mylonites, shaded field labelled R5L shows conditions for the granular peridotites as estimated by Lenoir et al. (2001), shaded field at the s-p boundary between wet and dry solidus represents plausible conditions for the plagioclase-facies shear zones. Depth scale assumes average crustal density of 2800 kg/m3. (For further explanation see text.)

Wal & Bodinier, 1996; Garrido & Bodinier, 1999; Lenoir et al., 2001); with this aim we briefly review the main characteristics of that history. The Ronda peridotites contain evidence of an early stage uplift from the diamond stability field to the deeper part of the spinel-facies field (Davies et al., 1993). This was followed by intense deformation and development of

porphyroclastic spinel-facies tectonites, representing the oldest deformational structure preserved (e.g. van der Wal, 1993). Along the outer margin of the Ronda body, adjacent to the high-grade metamorphic crustal envelope, these tectonites pass into garnet^spinel mylonites in a mylonitic shear zone of a few hundred meters width (see also Precigout et al., 2007). Towards the core of the massif,

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the porphyroclastic microstructures, as well as occasional narrow mylonite zones, are overprinted by intense recrystallization and development of granular peridotites. The boundary between the spinel tectonites and the granular rocks, currently known as the Ronda recrystallization front (van der Wal & Vissers, 1993, 1996) has been shown to result from extensive melt^rock interaction and meltassisted recrystallization (van der Wal & Bodinier, 1996) affecting the foliated spinel tectonites and occasional mylonites at spinel-facies conditions. In a more recent study, Lenoir et al. (2001) have interpreted the Ronda recrystallization front as a lithospheric melting front. The melting and melt^rock interaction processes were followed by peridotite emplacement into the crust along high-temperature plagioclase-facies shear zones developed in the deeper parts of the spinel-facies granular peridotites. Van der Wal & Vissers (1993) ascribed the early uplift of the Ronda peridotites from diamond facies across the garnet^spinel facies boundary (stage R1 in Fig. 12) to the Jurassic phase of extension and breakup in the Neotethys (Dercourt et al., 1986). They attributed the subsequent deformation in the porphyroclastic spinel peridotites (spinel tectonites) and garnet^spinel mylonites to progressive ductile shearing in a subduction-zone hanging-wall environment during Cretaceous^Paleogene collision. A subduction-zone setting for the spinel tectonites and garnet^spinel mylonites was principally suggested on the basis of the higher inferred pressures associated with the garnet^spinel mylonites in combination with low temperatures obtained from geothermometry on syntectonic assemblages (810^9008C and 830^8808C for the spinel tectonites and garnet^spinel mylonites, respectively, denoted in Fig. 12 as stages R2 and R3). A subduction zone setting (i.e. with the Ronda mantle in the subduction zone hanging wall) was also proposed by Davies et al. (1993) on the basis of the geochemistry of the garnet pyroxenites. Uplift of the Ronda rocks in the margin of the body proceeded under relatively cool conditions between 800 and 9008C (shown as a dashed trajectory in Fig. 12), whereas the deeper part of the body became heated and extensively affected by melting and melt^rock interaction processes (shown in Fig. 12 as stage R5), probably in response to convective removal of overthickened lithosphere or detachment of a subducting slab, causing ascent of asthenosphere and extensional thinning of the remaining overlying lithosphere (Van der Wal & Vissers, 1993; Van der Wal & Bodinier et al., 1996; Garrido & Bodinier, 1999). This led to extensional exhumation of the peridotites along plagioclase-facies ductile shear zones (stage R6 in Fig. 12) estimated at 22 Ma (Priem et al., 1979). Prior to any comparison of the Tallante xenolith results with the Ronda peridotite history, three aspects of that history need to be considered that have been discussed in recent studies. First, as emphasized by Platt et al. (2003),


the P^T path suggested by van der Wal & Vissers (1993) is in fact a composite P^T trajectory for the Ronda peridotite as a whole, rather than for single material points within the massif. Second, van der Wal & Vissers (1993) obtained rather low temperatures (770^8808C) in the spinel tectonites close to the recrystallization front (labelled in Fig. 12 as R4), but renewed thermobarometry study by Lenoir et al. (2001) of similar samples has yielded much higher temperatures of around 1050^11008C. In addition, Lenoir et al. (2001) concluded that the development of the recrystallization front occurred at temperatures in the range 1180^12258C and pressures near those of the ariegite^seiland subfacies boundary (field R5L in Fig. 12); that is, at higher temperatures but also higher pressures than inferred by van der Wal & Vissers (1993). Third, with the aim of reconciling the structural and geothermometric data from the peridotites with those seen in the crustal envelope, Platt et al. (2003) have reconsidered the significance of the spinel tectonites and in particular the garnet^spinel mylonites. Those workers have elegantly shown that, instead of (Cretaceous to Paleogene) subduction-related lithosphere thickening, the inferred low-temperature conditions during development of the garnet^spinel mylonites may equally be consistent with the onset of lithospheric extension, and that the mylonites may represent an extensional ductile shear zone that deformed at relatively low temperatures as a result of continuous cooling against a hanging wall of previously thickened crust. Recent structural work (Precigout et al., 2007) in addition suggests that the spinel tectonites and garnet^ spinel mylonites may form one heterogeneous shear zone system, such that the spinel tectonites may equally represent extensional deformation in the upper mantle. On the basis of this alternative interpretation, and using the van der Wal & Vissers (1993) estimates for the garnet^spinel mylonites in combination with the new Lenoir et al. (2001) thermobarometric results for the granular peridotites, Platt et al. (2003) proposed an alternative P^T trajectory for the Ronda peridotites, also shown in Fig. 12. Consistent with van der Wal & Vissers (1993, 1996) and Vissers et al. (1995), the mantle uplift associated with this evolution largely results from continuous extension and thinning during the early Miocene of the overlying crust, initially thickened to values of up to 55 km as a result of Alpine collision, and thinned in response to a late Oligocene delamination of gravitationally unstable lithosphere and consequent ascent of asthenosphere and thermal erosion of the remaining lithospheric mantle. The sequence documented in the Tallante xenoliths of early uplift, followed by ductile flow in porphyrocastic spinel peridotites, in turn overprinted by extensive meltassisted recrystallization, clearly recalls the earlier part of the Ronda history. The early uplift in the Tallante mantle evidenced by the spinel^pyroxene clusters after former

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garnet could be the equivalent of the early stage uplift of the Ronda peridotite body. The porphyroclastic microstructures and associated LPOs in the Tallante xenoliths may represent a same stage of upper mantle deformation as the Ronda spinel tectonites and garnet^spinel mylonites. It is plausible that these ductile flow structures mark the onset of extensional deformation at upper mantle levels as outlined above. The subsequent melt^rock interaction processes and associated annealing recrystallization in the Tallante xenoliths and the melting and melt^rock interaction processes documented in the Ronda massif probably reflect the same orogen-scale heating event. Accommodated by plagioclase-facies shear zones, the Ronda massif became emplaced in a rapidly thinning pile of HP^LTcrustal rocks at 22 Ma (Priem et al.,1979), shortly after heating since around 27 (Platt & Vissers, 1989) or 25 Ma (Platt et al., 2003). During the stages of melt^rock interaction and melt impregnation, the Tallante mantle rocks record a similar uplift, from pressures around 2 GPa to those of the plagioclase stability field (07^09 GPa), and we suggest that this uplift likewise reflects rapid crustal thinning. However, in contrast to the Ronda peridotites, the Tallante mantle rocks continued to reside in the upper mantle until they were sampled by ascending alkaline melts at 293^229 Ma (Duggen et al., 2005); that is, for a time span of almost 20 Myr during which a significant amount of the transient heat was probably lost by conductive cooling. This notion may have important consequences for comparisons between the two suites of mantle rocks, and for the inferred P^T evolution of the Tallante mantle. Although the similarity of the tectonic sequences in the Tallante xenoliths and Ronda peridotites is obvious, correlation between rock microstructure and associated P^T conditions in the two mantle domains is surrounded with uncertainties. In particular, there are significant microstructural differences between the coarse-granular peridotites of Ronda and the much finer-grained equigranular peridotites of Tallante. In addition, the Ronda coarsegranular peridotites differ geochemically from the spinel tectonites and have been interpreted as the solid residues after partial melting of a spinel tectonite protolith (Lenoir et al., 2001). In Tallante, the equigranular peridotites do not show a marked geochemical difference from the porphyroclastic xenoliths. As all recent studies converge on the idea that the Ronda recrystallization^melting front was thermally controlled and associated with a significant thermal gradient through the spinel tectonites, one would expect ‘Tallante-type’ equigranular peridotites ahead of the front (i.e. in the spinel tectonite domain). Such finegranular, recystallized microstructures, however, are not observed in Ronda. Following a suggestion by J. L. Bodinier (personal communication, 2009) we hypothesize that the equigranular microstructures did not develop in

Ronda because of the rapid cooling of the massif, which effectively ‘quenched’ the deformation and melt^rock reaction microstructures. If our inferences on the analogy of the Tallante microstructural record and the Ronda structural and thermal history are essentially correct, it is likely that the Tallante mantle underwent re-equilibration during Neogene cooling over a time span of up to 20 Myr. This cooling stage explains, in our view, the lack of a geothermometric record in the Tallante xenoliths of elevated temperatures associated with melt percolation and annealing recrystallization, as any mineral equilibria reached during meltassisted high-temperature annealing may well have been reset by diffusion. In addition, and as a consequence of this younger part of the thermal history, it is also impossible to ascertain whether the Tallante mantle ever experienced cooling before the development of (equi)granular peridotites; that is, before pervasive porous melt percolation, such as documented in the Ronda rocks for the transition from spinel tectonites to garnet^spinel mylonites (with temperatures as low as 8508C). Early mineral equilibria attained at such relatively low temperatures can be expected to have been affected by re-equilibration during subsequent melt^rock interaction, and the analytical results from the porphyroclastic microstructures may simply reflect the cumulative effects of reheating and subsequent late-stage cooling.

P^T history of the Tallante mantle In view of the above considerations on the thermobarometric results, inferences on the P^T evolution of the Tallante mantle are necessarily qualitative. However, given the marked similarity to the Ronda peridotite record, from the early uplift to spinel-facies conditions, followed by ductile flow, in turn overprinted by melt-assisted recrystallization and exhumation from pressures around 2 GPa to those of the plagioclase field, we suggest that the Tallante mantle must have gone through an analogous P^T evolution, but probably at lower temperatures than recorded in Ronda, and ending in a distinct cooling stage at plagioclase-facies conditions (Fig. 12, black dashed path) prior to sampling in Pliocene times. Lower climax temperatures during the stage of melt^rock interaction and melt impregnation in the Tallante xenoliths are largely based on the lack of evidence for partial melting related to the development of equigranular peridotites, whereas throughout the granular domain the Ronda massif was affected by a partial melting event implying that super-solidus (at least some ‘lithospheric’ hydrous solidus) temperatures were reached. More depleted, harzburgitic peridotites have been documented at Tallante by Beccaluva et al. (2004) but they were ascribed to prePaleozoic melting events. In addition, we recall that the Ronda P^T path as proposed by Platt et al. (2003) and

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shown in Fig. 12 is in part based on thermobarometric estimates for the coarse-granular rocks at the recrystallization front by Lenoir et al. (2001), who noted that these temperature estimates should be considered as minimum values. On the other hand, microstructural evidence for melt percolation and melt^rock interaction in the Tallante xenoliths does indicate that ambient temperatures should have been high enough for such processes to occur, and it seems likely that temperatures exceeded 11008C during that stage (Fig. 12).

Implications for west Mediterranean geodynamics There are three aspects of the Tallante mantle record with important consequences for the geodynamic evolution of the Alboran Domain. First, spinel-facies shearing was followed by a stage of reactive porous melt flow, melt impregnation and melt intrusion, and despite the lack of reliable geothermometric evidence this calls for elevated temperatures during that stage. Second, the Tallante mantle underwent some 30 km of exhumation, from P 2 GPa to 07^ 09 GPa, indicating an equivalent thinning of the overlying crustal and mantle rocks. Third, reactive porous flow and melt impregnation in the Tallante spinel- and plagioclasebearing peridotites were probably related to migration of sub-alkaline melts, and this was followed by intrusion of melts of alkaline affinity. Consistent with interpretations of the upper mantle structure in the Alboran region by Gutscher et al. (2002), Duggen et al. (2005) inferred that west-directed roll-back and steepening of a subducting oceanic plate induced extension and thinning in the Betic^Rif belt, leading to the formation of the Alboran Basin. We emphasize, however, that slab roll-back by itself cannot account for the both extreme and rapid heating at shallow depths documented in the metamorphic crustal envelope of the Ronda peridotites (Platt et al., 2003) and in the metamorphic rocks at Site 976 in the Alboran Sea (Comas et al., 1999), nor for the high temperatures implied by the melt^ rock interaction and melt impregnation processes in the Tallante mantle and similar melt^rock interaction and development of a lithospheric melting front in the Ronda massif. Both the extreme heating and rapid crustal thinning and mantle exhumation lend support to some form of removal of overthickened lithosphere and ascent of hot asthenosphere (Platt & Vissers, 1989; Vissers et al., 1995). In this context we recall that Paleogene subduction of the Ligurian ocean in the western Mediterranean region occurred in a northwesterly direction underneath the Balearic Islands, Sardinia and Corsica, leading to a collisional ridge between Iberia and North Africa (Fig. 13). The roll-back process is inferred to have started during the latest Oligocene to early Miocene (i.e. around 25 Ma), when this subduction^collision system became


divided into two segments by some form of lithosphere tearing (Spakman & Wortel, 2004) leading to separation of the Alboran and Calabrian parts of the collisional domain and subsequent west-directed roll-back of the Alboran segment to form the Alboran Basin (Fig. 13; see also Gutscher et al., 2002; Booth-Rea et al., 2007). Late Oligocene to early Miocene extensional collapse of the collisional ridge as envisaged by Platt & Vissers (1989) would have coincided with this process of lithosphere rupture. We therefore hypothesize that the retreating subduction system, now reflected by a curved slab observed in seismic tomography studies at the western end of the Betic^Rif arc (Fig. 13, shaded domain), evolved from Late Oligocene removal of gravitationally unstable lithosphere and ascent of hot asthenosphere, triggering extension and rapid transient heating of the overlying remaining lithospheric column. As a result, this remaining lithosphere deprived of its lithospheric root started to extend and spread over a Ligurian oceanic remnant located WSW of the collisional domain, which effectively led to subduction and roll-back of that oceanic lithosphere underneath the extending and thinning collisional ridge. This scenario is consistent with and supported by the notion that the Flysch Trough units in the western Betics were most probably floored by oceanic crust (e.g. Booth-Rea et al., 2007, and references therein). The interaction with melts of different sources in the Tallante spinel and plagioclase-bearing peridotites lends strong support to the above scenario. Uplift of the Tallante mantle was accompanied by reactive porous flow and melt impregnation probably related to migration of subalkaline melts, and this is consistent with the sub-alkaline magmatism affecting the Alboran region since the Oligocene. This magmatism has been largely ascribed to melting of mantle sources contaminated by hydrous fluids or melts derived from dehydration and/or melting of subducting oceanic lithosphere, induced by slab roll-back and steepening (Duggen et al., 2004, 2005, 2008; see, however, Turner et al., 1999). Duggen et al. (2005) also proposed a geodynamic scenario to explain the Neogene transition from subductionrelated to intraplate-type alkaline magmatism in the Alboran region, which can be observed both on a regional scale and in single volcanic systems. Consistent with the interpretations of the upper mantle structure in the Alboran region by Gutscher et al. (2002) they inferred that, close to the Miocene^Pliocene boundary, continuing slab roll-back triggered delamination of bands of subcontinental lithosphere (continental edge delamination) at the edges of the subducting plate (i.e. at the southern Iberia and northern African plate margins), and that this caused upwelling of plume-contaminated sub-lithospheric mantle, which generated the alkaline magmatism as seen in the Tallante volcanic center. Turner et al. (1999) instead

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ds

aric

IBERIA Early Miocene lithosphere tearing

n Isla

le Ba

23 Ma

ne goce ain i l O Late nal dom io collis

Late Miocene slab detachment

Tallante Ronda

Alboran segment

Calabrian segment

23 Ma slab detachment in Middle Miocene

23 Ma lithosphere tearing

AFRICA

500 km

Fig. 13. Kinematics of slab roll-back in the Betic^RifÂlboran region, slightly modified after Spakman & Wortel (2004). Grey shaded area indicates location of the BeticÂlboran slab at a depth of 200 km as observed in seismic tomography. Present-day coastlines shown as continuous lines; dashed lines indicate location of African margin and Balearic Islands 23 Myr ago. Dashed north^south-trending boundary in the east indicates Late Oligocene separation of the Calabrian and Alboran segments. Set of curved dashed lines illustrates westward migration of the subduction trench with time from near the Balearic Islands, and ending at the time of slab detachment under the Betics (Late Miocene). Initially, the trench retreated in a south to southwesterly direction while slab bending progressed. This was accommodated by Early Miocene WSW-directed lithosphere tearing along the Balearic margin, and simultaneous west-directed detachment evolving into lithosphere tearing along the African margin. Along the Balearic margin, lithosphere tearing came to a halt during the Miocene when the trench became roughly parallel to the Iberian margin, and laterally propagating slab detachment allowed the last stages of west-directed roll-back. (Note position of Cabezo Tallante above zone of Late Miocene slab detachment.)

explained the underlying process in terms of an advanced stage of convective removal leading to the loss of the remaining thick lithosphere underneath the Iberian (and also African) margin, whereas Spakman & Wortel (2004) described this stage of the upper mantle history in a different way; namely, as a laterally propagating detachment of the slab. In this context we note that a diachronous change from marine to continental conditions in the Neogene basins (Meijninger, 2006) of the eastern Betics, from Late Tortonian in the east to Early Messinian further west, lends independent support to this latter hypothesis. The consequences in terms of mantle upwelling, however, should be the same in each of these interpretations, and we emphasize that Cabezo Tallante is indeed located roughly above the slab edge as envisaged by Duggen et al. (2005); that is, above the zone of slab detachment as observed in seismic tomography images by Spakman & Wortel (2004). Remarkably, the melt migration and intrusion stages recognized in the Tallante peridotite xenoliths record the

same transition from sub-alkaline to alkaline magmatism, and this magmatic evolution occurred during progressive exhumation probably accommodated by large-scale extension-related crustal thinning as outlined above. We therefore propose that lithosphere extension, initiated by some detachment process in a position near the Balearic Islands and followed by west-directed slab roll-back, led to uplift and migration of lithospheric mantle sectors (such as the mantle sampled at present at Tallante) from an inner part of the mantle wedge, where they experienced deep (spinel-facies) pervasive porous melt flow, towards a position above a westward propagating slab edge or slab detachment zone developing along the southern Iberian continental margin since the early Tortonian. This allowed upwelling of plume-contaminated sub-lithospheric mantle, which generated the alkaline magmatism reflected in the shallow intrusion of olivine^ amphibole pyroxenites at 07^09 GPa, followed by further ascent of alkaline melts and sampling of the Tallante mantle domain.

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AC K N O W L E D G E M E N T S We acknowledge Laura Negretti (Genova) and Andrea Risplendente (Milano) for assistance with the EDS and WDS analyses. We are indebted for a helpful review by Tomoaki Morishita, for challenging and helpful comments by Arjan Dijkstra, and for a well-thought out and extremely constructive review by Jean-Louis Bodinier. Their reviews have been of great help to better clarify the results of our work and to considerably improve the paper, not only with respect to our study of the Cabezo Tallante xenoliths but also regarding our comparison with a state-ofthe-art structural and geochemical interpretation of the Ronda massif.

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