Journal of Petrology Advance Access originally published online on February 25, 2005
Journal of Petrology 2005 46(6):1155-1201; doi:10.1093/petrology/egi013
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Post-Collisional Transition from Subduction- to Intraplate-type Magmatism in the Westernmost Mediterranean: Evidence for Continental-Edge Delamination of Subcontinental Lithosphere
1 GEOMAR RESEARCH CENTRE FOR MARINE GEOSCIENCES, DEPARTMENT OF VOLCANOLOGY AND PETROLOGY, WISCHHOFSTR. 1–3, 24148 KIEL, GERMANY
2 IFM–GEOMAR LEIBNIZ INSTITUTE FOR MARINE SCIENCES, DYNAMICS OF THE OCEAN FLOOR, WISCHHOFSTR. 1–3, 24148 KIEL, GERMANY
3 INSTITUTE FOR GEOSCIENCES, UNIVERSITY OF KIEL, OLSHAUSENSTR. 40, 24118 KIEL, GERMANY
RECEIVED NOVEMBER 12, 2003; ACCEPTED JANUARY 14, 2005
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONREGIONAL GEOLOGYANALYTICAL METHODSRESULTSDISCUSSIONGEODYNAMIC EVOLUTIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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Post-collisional magmatism in the southern Iberian and northwestern African continental margins contains important clues for the understanding of a possible causal connection between movements in the Earth's upper mantle, the uplift of continental lithosphere and the origin of circum-Mediterranean igneous activity. Systematic geochemical and geochronological studies (major and trace element, Sr–Nd–Pb-isotope analysis and laser 40Ar/39Ar-age dating) on igneous rocks provide constraints for understanding the post-collisional history of the southern Iberian and northwestern African continental margins. Two groups of magmatic rocks can be distinguished: (1) an Upper Miocene to Lower Pliocene (8·2–4·8 Ma), Si–K-rich group including high-K (calc-alkaline) and shoshonitic series rocks; (2) an Upper Miocene to Pleistocene (6·3–0·65 Ma), Si-poor, Na-rich group including basanites and alkali basalts to hawaiites and tephrites. Mafic samples from the Si–K-rich group generally show geochemical affinities with volcanic rocks from active subduction zones (e.g. Izu–Bonin and Aeolian island arcs), whereas mafic samples from the Si-poor, Na-rich group are geochemically similar to lavas found in intraplate volcanic settings derived from sub-lithospheric mantle sources (e.g. Canary Islands). The transition from Si-rich (subduction-related) to Si-poor (intraplate-type) magmatism between 6·3 Ma (first alkali basalt) and 4·8 Ma (latest shoshonite) can be observed both on a regional scale and in individual volcanic systems. Si–K-rich and Si-poor igneous rocks from the continental margins of southern Iberia and northwestern Africa are, respectively, proposed to have been derived from metasomatized subcontinental lithosphere and sub-lithospheric mantle that was contaminated with plume material. A three-dimensional geodynamic model for the westernmost Mediterranean is presented in which subduction of oceanic lithosphere is inferred to have caused continental-edge delamination of subcontinental lithosphere associated with upwelling of plume-contaminated sub-lithospheric mantle and lithospheric uplift. This process may operate worldwide in areas where subduction-related and intraplate-type magmatism are spatially and temporally associated.
KEY WORDS: post-collisional magmatism; Mediterranean-style back-arc basins; subduction; delamination; uplift of marine gateways
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONREGIONAL GEOLOGYANALYTICAL METHODSRESULTSDISCUSSIONGEODYNAMIC EVOLUTIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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Subduction-related and intraplate-type igneous rocks occur together worldwide (e.g. Mediterranean area, Colorado Plateau and Andes) (Kay & Mahlburg Kay, 1993
; Arculus & Gust, 1995; Wilson & Bianchini, 1999). Kilometre-scale continental uplift may occur contemporaneously with the magmatic activity and is thought to reflect lithospheric response to rapid changes in upper mantle geometry (Bird, 1979; Turner et al., 1993; Duggen et al., 2003). Large-scale reorganization of the upper mantle may be related to: (1) roll-back and detachment of subducted oceanic lithosphere (Innocenti et al., 1982; Keller, 1982; Davies & Blanckenburg, 1995; Wilson & Bianchini, 1999; Wortel & Spakman, 2000); (2) detachment or convective thinning of subcontinental lithosphere (Pearce et al., 1990; Platt & England, 1993; Turner et al., 1999; López-Ruiz et al., 2002); or (3) delamination (peeling-off) of subcontinental lithosphere (Bird, 1979; Serri et al., 1993; Docherty & Banda, 1995; Gîrbacea & Frisch, 1998; Duggen et al., 2003). Formation of magmas having a subduction-related geochemical signature can result from: (1) subduction of oceanic lithosphere (Hawkesworth et al., 1993; Pearce & Peate, 1995); (2) melting of metasomatically enriched subcontinental lithosphere with an inherited subduction signature (Turner et al., 1996; Benito et al., 1999); or (3) extensive crustal contamination of MORB-like (mid-ocean-ridge basalt) magmas (Turner et al., 1999). Generation of intraplate-type lavas in continental settings is generally related to upwelling sub-lithospheric mantle material (e.g. Western and Central Europe) (Wilson & Downes, 1991; Granet et al., 1995; Hoernle et al., 1995). However, the processes responsible for uplift and the generation of volcanic areas with geochemically diverse rock types such as subduction-related and intraplate-type igneous rocks are still poorly understood.
Combination of systematic geochemical and geochronological data allows reconstruction of the temporal and spatial transition from post-collisional subduction-related to intraplate-type magmatism in the western Mediterranean. The transition occurs regionally in the westernmost Mediterranean realm and locally in individual volcanic systems such as the Gourougou and Guilliz stratovolcanic complexes in northwestern Africa. The new geochemical and geochronological data for igneous rocks from the continental margins of southern Spain and northern Morocco are integrated with published data for rocks from the westernmost Mediterranean sea floor (Duggen et al., 2004) and northwestern Algeria (Louni-Hacini et al., 1995; Coulon et al., 2002) providing the basis for a case study of this transition.
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REGIONAL GEOLOGY |
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TOPABSTRACTINTRODUCTIONREGIONAL GEOLOGYANALYTICAL METHODSRESULTSDISCUSSIONGEODYNAMIC EVOLUTIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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Neogene post-collisional igneous rocks occur along the European and African margins of the western Mediterranean (Wilson & Bianchini, 1999
). In the Alboran region in the westernmost Mediterranean, Middle Miocene to Pleistocene volcanic rocks can be found in two lithospheric domains: (1) the Alboran Basin (hosting the Alboran Sea), which consists of strongly thinned continental crust progressing into Miocene oceanic crust to the east (Lonergan & White, 1997); (2) the southern Iberian and northwestern African continental margins (Fig. 1).
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The Alboran Basin area includes Middle to Late Miocene igneous rocks from the Alboran Sea floor (ridges and seamounts), the Alboran Island and coastal stratovolcanic complexes bordered by major strike-slip faults in southern Spain (the Cabo de Gata and Aguilas block) and northern Morocco (Ras Tarf and Trois Furches) (Duggen et al., 2004
) (Fig. 1). The magmatic evolution of the Alboran Basin area is marked by the eruption of magmas of both the low-K (tholeiitic) series and the medium- to high-K (calc-alkaline) series in the Middle to Late Miocene (Figs 1 and 2). The mafic Alboran Basin lavas have major and trace element and O–Sr–Nd–Pb-isotope compositions characteristic of tholeiitic volcanic front and calc-alkaline rear-arc lavas in active subduction zones such as the Izu–Bonin island arc (Benito et al., 1999; Hoernle et al., 1999; Duggen et al., 2004).
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The Iberian and African continental areas surrounding the Alboran Basin are marked by igneous activity with a different geochemical composition. Upper Miocene to Pleistocene post-collisional magmatism along the southern Iberian and northwestern African continental margins is characterized by Si-rich, K-rich igneous rocks and Si-poor, Na-rich lavas (Fig. 1). Upper Miocene to Lower Pliocene Si–K-rich igneous rocks include latites, shoshonites and ultrapotassic rocks such as lamproites (Fig. 2) (Bellon & Brousse, 1977
; López Ruiz & Rodríguez Badiola, 1980; Venturelli et al., 1984; Hernandez, 1986; Louni-Hacini et al., 1995; El Bakkali et al., 1998; Benito et al., 1999; Coulon et al., 2002). These Si–K-rich igneous rocks are often spatially associated with Si-poor, Na-rich lavas such as alkali basalts and basanites to hawaiites and tephrites. The first Si-poor, Na-rich alkali basalts appeared in the Upper Miocene and continued to erupt into the Quaternary (Figs 1 and 2) (Duggen et al., 2003). The post-collisional transition from Si–K-rich to Si-poor, Na-rich magmatism occurred close to the Miocene–Pliocene boundary that is defined by the Messinian Salinity Crisis and can clearly be seen in the field in the Guilliz volcanic area, where dark alkali basalt lava flows were erupted on top of light-coloured, Si–K-rich, pyroclastic deposits of an eroded stratovolcanic complex.
A large number of models have been proposed to explain the geodynamic evolution of the Alboran region in the westernmost Mediterranean since the Eocene and can be divided into three groups: (1) mantle diapirism (Weijermars, 1985); (2) subduction of oceanic lithosphere involving slab roll-back and slab detachment (Royden, 1993; Zeck, 1996; Lonergan & White, 1997; Hoernle et al., 1999; Coulon et al., 2002; Gutscher et al., 2002; Duggen et al., 2003, 2004); (3) convective removal (detachment) and peeling-off (delamination) of subcontinental lithosphere (Platt & Vissers, 1989; Docherty & Banda, 1995; Houseman, 1996; Seber et al., 1996a; Comas et al., 1999; Turner et al., 1999; López-Ruiz et al., 2002; Platt et al., 2003). In this study, the temporal and spatial geochemical evolution of the magmatism in the westernmost Mediterranean is integrated with geophysical and geological data from the literature to distinguish between these models. We developed a three-dimensional model that describes the mantle geodynamic setting in the westernmost Mediterranean area near the Miocene–Pliocene boundary.
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ANALYTICAL METHODS |
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TOPABSTRACTINTRODUCTIONREGIONAL GEOLOGYANALYTICAL METHODSRESULTSDISCUSSIONGEODYNAMIC EVOLUTIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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40Ar/39Ar isotope ratios were measured on K-bearing phases such as amphibole, biotite and feldspar phenocrysts and glass and matrix chips, at the GEOMAR Tephrochronology Laboratory. The freshest possible glass and matrix chips and phenocrysts were hand-picked from crushed and sieved splits and cleaned using an ultrasonic disintegrator. Plagioclase phenocrysts were additionally etched in hydrofluoric acid (15%). For several samples, fresh matrix chips were separated and analysed together with phenocrysts to examine the reliability of age data measured for matrix chips from aphyric igneous rocks. The samples were neutron-irradiated at the 5 MW reactor of the GKSS Reactor Centre in Geesthacht (Federal Republic of Germany), with crystals and glass and matrix chips in aluminium trays and irradiation cans wrapped in 0·7 mm cadmium foil. Age determinations by laser 40Ar/39Ar analysis were performed by fusing single crystals and glass and matrix chips in a single step or by laser-step heating analysis when necessary. Purified gas samples were then analysed using a MAP 216 series noble gas mass spectrometer. Raw mass spectrometer peaks were corrected for mass discriminations background, and blank values determined every fifth analysis. Neutron flux during irradiation was monitored using TCR sanidine (Taylor Creek Rhyolite, 27·92 Ma) (Dalrymple & Duffield, 1988
; Duffield & Dalrymple, 1990) and an internal standard SAN6165 (0·470 Ma) (van den Bogaard, 1995). Vertical variations in J values were quantified by a cosine function fit. Lateral variations in J were not detected. Optical grade CaF2 and high-purity K2SO4 salt crystals that were irradiated and analysed together with the samples were used for corrections of interfering neutron reactions on Ca and K. For statistical reasons, 3–19 particles of each phase were separately dated to calculate mean apparent ages and isochron ages. For samples that did not produce statistically valid isochron or mean apparent ages, step heating ages on single matrix chips were additionally determined. The age with the highest confidence level was chosen for each sample (Table 1), taking into account mean square weighted deviates (MSWD), plateau portions for step heating analysis, deviation of the data in the isochron plots (see Electronic Appendix 1, which may be downloaded from http://www.petrology.oupjournals.org) and the initial 40Ar/39Ar ratio. Errors on the age data are reported at the 1
confidence level.
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For bulk-rock geochemical analyses, whole-rock sample chips (0·1–1 cm in diameter) were hand-picked using a binocular to minimize the effects of alteration; these were then cleaned ultrasonically with de-ionized water to extract salts, and then dried and pulverized in an agate mill. Bulk-rock analyses of major and selected trace elements (Ba, Co, Cr, Cu, Ga, Rb, Sr, V and Zr) were performed on fused tablets with a Philips X'unique PW 1480 X-ray fluorescence spectrometer (XRF) at the GEOMAR Research Centre. International reference standards JB-2, JB-3 (basalts), JA-2, JA-3 (andesites), JR-2, JR-3 (rhyolites), JG-2, GM (granites) and JF-1 (feldspar) were used to evaluate the precision and accuracy of the measurements. Na contents were normalized to the standard values to account for X-ray tube drift for this element. H2O and CO2 were determined using a Rosemont Infrared Photometer CSA 5003. The new data are presented in Table 2 and Electronic Appendix 2. Values of major elements given in the text and used for plotting and rock classification were calculated on a volatile-free basis following a correction of Ca contents assuming that CO2 comes from CaCO3.
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Additional trace elements (Li, Y, Nb, Mo, Sn, Sb, Cs, REE, Hf, Ta, W, Tl, Pb, Th and U) were determined from mixed acid (HF–aqua regia–HClO4) pressure digests by inductively coupled plasma-mass spectrometry (ICP-MS) at the Institute for Geoscience, University of Kiel, using an upgraded (High-Performance Interface, torch box) VG Plasmaquad PQ1. The precision of analytical results as estimated from duplicate measurements was better than 3% RSD. Details of the analytical procedure have been described in the literature (Garbe-Schönberg, 1993
). Blanks and international standards, BE-N and BHVO-1, were analysed with the samples, to evaluate the precision and accuracy of the measurements (Duggen, 2002). The ICP-MS trace element data are presented in Table 2 and Electronic Appendix 2.
Sr, Nd and Pb isotopic analyses were carried out on whole-rock powders. To minimize the effects of seawater alteration, the powders were leached with hot distilled 6 M HCl for 1 h (except the highly evolved samples). After dissolution with Merck Ultrapur HF, HNO3 and HCl, Sr, Nd and Pb were extracted by conventional ion exchange chromatographic techniques. Isotope ratios of all samples were measured by thermal ionization mass spectrometry (TIMS) in static (Sr, Nd, Pb) and dynamic (Sr, Nd) mode on a Finnigan MAT262 system at GEOMAR. Measured 87Sr/86Sr and 143Nd/144Nd ratios were normalized within-run to 86Sr/88Sr = 0·1194 and 146Nd/144Nd = 0·7219. The Sr and Nd isotope ratios were further normalized to 87Sr/86Sr = 0·71025 (NBS 987 standard) and 143Nd/144Nd = 0·51185 (La Jolla standard) or 143Nd/144Nd = 0·51171 (internal Nd-Spex standard calibrated to La Jolla) to correct for mass spectrometer drift. Replicate analyses of standards NBS 987 (Sr), La Jolla and Nd-Spex (Nd) used for the evaluation of the external precision gave mean values of 87Sr/86Sr = 0·710242 ± 0·000017 (2 SD) (n = 39), 143Nd/144Nd = 0·511841 ± 0·0000010 (2 SD) (n = 13) and 143Nd/144Nd = 0·511706 ± 0·000008 (2 SD) (n = 12). Replicate analyses of the NBS 981 (Pb) standard yielded a mean value of 206Pb/204Pb = 16·894 ± 0·004 (2 SD), 207Pb/204Pb = 15·433 ± 0·005 (2 SD) and 208Pb/204Pb = 36·515 ± 0·017 (2 SD.) (n = 23). Pb-isotope data were corrected for 0·1238
mass discrimination per atomic mass unit (a.m.u.). Total analytical blanks were on average <0·100 ng and are thus negligible given the generally high Pb content of the samples.
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RESULTS |
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TOPABSTRACTINTRODUCTIONREGIONAL GEOLOGYANALYTICAL METHODSRESULTSDISCUSSIONGEODYNAMIC EVOLUTIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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40Ar/39Ar geochronological data
Laser probe 40Ar/39Ar age data for 45 volcanic rocks from the Iberian and African continental margins in southeastern Spain and northern Morocco are reported in Table 1. Phenocryst ages are identical within error, or in very good agreement with each other, in all samples (e.g. feldspar and amphibole in sample CX110599 [GenBank] -3, biotite and amphibole in sample TA200400-1, biotite and feldspar in GZ280300-1). Matrix chip ages are also identical within error and in accordance with the phenocryst ages (e.g. matrix and feldspar in GG140699 [GenBank] -8, matrix and feldspar and biotite in GZ280300-3b, matrix and feldspar in GZ180699-4, matrix and biotite in OD190699-1a), showing that analyses of carefully prepared matrix chips can provide reliable age data. Two Si–K-rich samples (lamproites) from southeastern Spain yielded phlogopite ages (FT120599-5, BQ170599 [GenBank] -9) with high MSWD values as a result of a large variation in single phenocryst ages. These phlogopite ages are significantly older than the matrix ages and are unlikely to reflect eruption ages. Three other Si–K-rich (lamproite) samples (CT270300 [GenBank] -2c, ZN130599-1, VE290699-10A) provided phlogopite ages having low MSWD and initial 40Ar/36Ar ratios close to the atmospheric value and are thus considered to represent eruption ages. Below we systematically discuss the laser probe 40Ar/39Ar age data for the Si–K-rich igneous rocks followed by the Si-poor lavas from four areas on the Iberian and African continental margins: southeastern Spain and the Gourougou, Guilliz and Oujda volcanic areas in northwestern Africa (Fig. 1).
The southern Iberian continental margin
Laser probe 40Ar/39Ar age data for 13 Si–K-rich rocks (lamproites and their derivatives) from southeastern Spain range from 8·19 to 6·37 Ma (Table 1 and Figs 1 and 2) and are within the range of published K/Ar ages (8·6–6·2 Ma) (Nobel et al., 1981; Bellon et al., 1983). The 10·8 ± 1 Ma K/Ar age for a lamproite from the Vera basin (Bellon & Brousse, 1977) was not confirmed by laser 40Ar/39Ar age dating. The 6·76 ± 0·04 Ma matrix 40Ar/39Ar age from La Celia near Jumilla is identical within error to a published Ar/Ar whole-rock age of 6·8 ± 0·4 Ma (Turner et al., 1999). A systematic relationship between the new Ar/Ar ages and older K/Ar age data was not detected. As inferred from geological maps, the laser 40Ar/39Ar age data are consistent with stratigraphic evidence showing that lamproitic rocks are generally associated with Upper Miocene marine sediments (mostly marls and sandstones) in the Murcia area. Field observations in the Vera basin show that lamproitic lava flows, showing pillow structures, were emplaced and chilled in soft, wet Messinian marine muds (peperites). Two lamproitic peperites were dated at 6·44 Ma and 6·37 Ma and are consistent with the Messinian ages of their host marine sediments. A newly discovered lamproitic dyke (VE290699-10A) yielded a phlogopite age of 7·45 ± 0·08 Ma. It intrudes rhyodacitic breccias cut by the Palomares fault and the dyke itself has been tectonically deformed. Thus the new age data also provide constraints for the minimum age of the breccia and the tectonic history of the Palomares fault zone.
Si-poor, Na-rich volcanic rocks from the Tallante volcanic field (two hawaiites and one trachybasalt), which are spatially associated with the Si–K-rich (lamproitic) rocks, yielded amphibole + biotite, glass and matrix mean apparent ages and step heating ages ranging from 2·93 to 2·29 Ma (Table 1 and Figs 1 and 2). For Tallante, the ages are consistent with field evidence showing that the volcanic units overlie Tortonian, Messinian and Pliocene sediments and are in good agreement with published K/Ar ages ranging from 2·8 to 2·7 Ma (Bellon et al., 1983). The 10·5 Ma Ar/Ar phlogopite age for a Tallante lava reported by Turner et al. (1999) is not consistent with these Pliocene ages. Phlogopite is not a common phenocryst phase in basaltic magmas, as it is stable only at high pressures and becomes unstable upon extrusion (Deer et al., 1996). Ultramafic mantle xenoliths frequently associated with the alkali basalts from Tallante contain phlogopite (Dupuy et al., 1986; Capedri et al., 1989) and, therefore, the phlogopites dated by Turner et al. (1999) may be xenocrysts from disaggregated mantle xenoliths. In conclusion, the transition from Si–K-rich to Si-poor, Na-rich volcanic rocks occurred between 6·4 to 2·9 Ma at the southern Iberian continental margin.
The northwestern African margin
Si–K-rich igneous rocks from the Gourougou volcanic field (basaltic trachyandesites, trachyandesites, andesite and one diorite) yielded matrix mean apparent and isochron ages on biotite and feldspar ranging from 7·58 to 4·8 Ma (Table 1 and Figs 1 and 2). These ages are largely consistent with the range of whole-rock and biotite-separate K/Ar age data (9·0 to 4·6 Ma) (Bellon & Brousse, 1977; Hernandez & Bellon, 1985; El Bakkali et al., 1998) and 40Ar/39Ar age data (6·73 to 6·0 Ma) (Roger et al., 2000) for volcanic tuffs interbedded with Messinian marine sediments, which were deposited on the slope of the Gourougou stratovolcanic complex. Si-poor, Na-rich alkali basalts and trachybasalts were erupted on the flanks of the Gourougou stratovolcano between 6·3 and 3·73 Ma (Table 1). Lava flows directly overlie Messinian marine sediments (e.g. Haidoun flow, dated at 3·73 Ma). Previously determined K/Ar ages generally tend to be younger (e.g. 2·5 Ma for Haidoun) (Bellon & Brousse, 1977) than the laser 40Ar/39Ar age data. In the Gourougou volcanic system, the transition from Si–K-rich to Si-poor, Na-rich igneous activity occurred between 6·3 and 4·8 Ma.
Si–K-rich igneous rocks (trachyandesite to trachyte) from the Jebel Guilliz stratovolcano located in the centre of the Guilliz volcanic field yielded mean apparent ages for feldspar and biotite + feldspar ranging from 6·90 to 6·82 Ma (Table 1 and Figs 1 and 2). Mean apparent ages for matrix and feldspar + matrix from Si-poor, Na-rich alkali basalts in the Guilliz volcanic field range from 6·3 to 0·65 Ma (Table 1 and Figs 1 and 2). Most published K/Ar ages (8·0 to 4·9 Ma) (Hernandez & Bellon, 1985) fall within this range but some K/Ar ages tend to be older. Major discrepancies between the new laser 40Ar/39Ar age data (0·65 Ma and 0·88 Ma) and the K/Ar (1·74 Ma and 2·5 Ma) (Hernandez & Bellon, 1985) data from the literature exist for the Si-poor lavas from the Ain Zora and Jebel El Kehal lava flows in the westernmost Guilliz area. The new age data suggest that volcanism in the westernmost Guilliz volcanic field was active contemperanously with the Quaternary alkali basaltic volcanism of the nearby Middle Atlas mountains (1·8 to 0·5 Ma, K/Ar ages) (Bellon & Brousse, 1977; Harmand & Cantagrel, 1984). In the Guilliz area, the transition from Si–K-rich to Si-poor, Na-rich igneous rocks occurred between 6·8 and 6·3 Ma.
The Oujda–Algeria area includes igneous rocks from northwestern Algeria, the Oujda volcanic field and the Plateau du Rekkam area situated south of the Oujda area. Based on K/Ar age data, Si–K-rich magmatism (high-K calc-alkaline and shoshonitic series) in northwestern Algeria occurred at the Middle–Upper Miocene boundary and ceased in the Upper Miocene (11·7 to 7·5 Ma) (Louni-Hacini et al., 1995; Coulon et al., 2002). Si-poor, Na-rich rocks (basanites, alkali basalts and trachyandesites) from the Oujda volcanic field yielded matrix and biotite mean apparent ages ranging from 3·78 to 3·10 Ma (Table 1). These ages show that the eruption history of the Oujda volcanic field has a much narrower range than hitherto thought based on published whole-rock K/Ar age data (6·2 to 1·5 Ma) (Bellon & Letouzey, 1977; Tisserant et al., 1985; Andries & Bellon, 1989). Two altered lava flows outcropping in the vicinity of the Oued Isly southwest of Oujda were K/Ar dated at 6·2 and 5·6 Ma (Andries & Bellon, 1989). These flows overlie Tortonian blue marls (Marnes bleues) and so may correlate with Messinian marine sediments (Marnes jaunes) in the Oujda area (Jadid et al., 1999). However, redating of the flows in this area with the laser 40Ar/39Ar technique (3·78 to 3·10 Ma) was not able to confirm the Late Miocene K/Ar ages (Table 1). Based on K/Ar age data, Si-poor volcanic rocks from northwestern Algeria show a much wider age range (10·0–7·2 Ma and 4·3–0·8 Ma) (Fig. 2) (Louni-Hacini et al., 1995; Coulon et al., 2002) than Si-poor volcanic rocks from the Oujda volcanic field (3·78–3·10 Ma) (this study). However, additional laser 40Ar/39Ar age dating is necessary to confirm the older K/Ar ages from northwestern Algeria, in particular the onset and termination of Si–K-rich and Si-poor magmatism. Plateau du Rekkam volcanic rocks SE of the Oujda volcanic field probably erupted in the Plio-Quaternary (Saadi et al., 1985). In the Oujda–Algeria area, the transition from Si–K-rich to Si-poor, Na-rich magmatism apparently took place between 10·0 and 7·5 Ma.
In summary, the transition from Si–K-rich to Si-poor volcanism occurred between 6·4 and 2·9 Ma in southern Spain, between 6·3 and 4·8 Ma in the Gourougou area, between 6·8 and 6·3 Ma in the Guilliz area, and possibly between 10·0 and 7·5 Ma in the Oujda–Algeria area.
Major and trace element data
Major and trace element data for 145 igneous rocks from the southern Iberian and northwestern African continental margins are presented in Table 2 and Electronic Appendix 2. The Electronic Appendix includes the full dataset and is available from the Journal of Petrology website at http://www.petrology.oupjournals.org, whereas Table 2 is a reduced version that includes major and trace element data only for samples that were selected for Sr–Nd–Pb-isotope analysis.
The comprehensive dataset allows division of the post-collisional igneous rocks from the westernmost Mediterranean into a Si-rich and a Si-poor group (Fig. 3a). Based on the K2O content, the Si-rich samples can be further subdivided into a relatively K-poor group and a K-rich group (Fig. 3b). The Middle to Upper Miocene (12·1 to 6·1 Ma) K-poor group includes low-K (tholeiitic) and medium- to high-K (calc-alkaline) basalts through rhyolites outcropping in the Alboran Basin area (Hoernle et al., 1999; Duggen et al., 2004; Gill et al., 2004). The occurrence of Upper Miocene to Lower Pliocene (8·2 Ma and 4·8 Ma) K-rich magmatism, including high-K to shoshonitic series basaltic andesites and basaltic trachyandesites to dacites and trachytes and lamproitic rocks is restricted to the southern Iberian and northwestern African continental margins. Upper Miocene to Pliocene (6·3 and 0·65 Ma) Si-poor volcanic rocks are spatially associated with the Si–K-rich group rocks, and range from alkali basalts and basanites to hawaiites and tephrites and tend to have higher Na2O contents for a given K2O concentration than the K-rich group. Below we describe the major and trace element composition of the Si–K-rich and Si-poor group rocks from the continental margins of southern Iberia and northwestern Africa.
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The southern Iberian continental margin
Si–K-rich volcanic rocks from southeastern Spain (lamproites and their derivatives) show a large range in most major and trace element contents. MgO (1·5–12·9 wt %) correlates inversely with SiO2 and Al2O3 and positively with TiO2, Nb, Ta, Nd, Sm, Zr, Eu, Hf, Gd, Sn, Cr and Ni (with r2 for the regression line >0·7, except for Ni, which shows a curved correlation with MgO) (Fig. 4). Most of these systematic geochemical variations can be explained by fractional crystallization of common phenocryst phases in the Spanish lamproites such as olivine (containing Cr-spinel), phlogopite, apatite and sanidine (Venturelli et al., 1988
; Toscani et al., 1995). There is a compositional continuum between high-MgO lamproites, their low-MgO derivatives (some of which contain crustal xenocrysts such as garnet and cordierite) and cordierite–garnet-bearing volcanic rocks from southeastern Spain (e.g. Cerro Hoyazo) (Fig. 4a–d). The Si–K-rich, lamproitic rocks from southern Spain have enriched incompatible element contents and exhibit spiked patterns on primitive mantle normalized multi-element diagrams (Fig. 5a), resulting from an enrichment of fluid-mobile, incompatible elements such as Rb, Ba, (± Th), U, K, Pb (and Cs, Sn, Sb and Li) relative to incompatible elements that are less fluid-mobile or fluid-immobile [e.g. rare earth elements (REE), Nb, Ta]. Strong enrichments of the light and middle REE relative to heavy REE result in high (LREE, MREE)/HREE ratios (e.g. La/Yb = 34–71). A strong LREE enrichment is associated with depletion of Eu relative to Sm and Gd as reflected by negative Eu anomalies (Eu/Eu* = 0·52–0·67). Enrichment of fluid-mobile relative to less fluid-mobile elements and pronounced troughs in Nb and Ta are typical features of subduction zone magmas (e.g. Aeolian Islands and Izu–Bonin arc lavas in Fig. 5c) (Ellam et al., 1988, 1989; Taylor & Nesbitt, 1998).
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Si-poor volcanic rocks from Tallante in southeastern Spain have MgO contents ranging from 3·6 to 8·5 wt %. MgO correlates negatively with Al2O3, Nb, Ta, Zr and Hf and positively with TiO2, Cr and Ni (r2 > 0·7) (Fig. 4a–d), generally consistent with the fractionation of the observed phenocryst assemblages (e.g. olivine, clinopyroxene and magnetite). The Si-poor group samples from the Iberian continental margin generally exhibit contrasting multi-element patterns to the Si–K-rich group showing relative enrichments of fluid-immobile Nb and Ta and relative depletion of fluid-mobile elements K, ± Pb, and have lower Rb, Ba, Th, U, K and Pb, but higher Sr and Ti concentrations than the Si–K-rich volcanic rocks. The Si-poor group lavas have geochemical signatures similar to those of ocean-island basalts (OIB) and lavas from continental intraplate volcanic settings (e.g. Massif Central, St. Helena and Canary Islands in Fig. 5d). The Tallante lavas, however, show deviations from typical intraplate-type, multi-element patterns with positive Th–U anomalies, highly variable K anomalies and Pb anomalies ranging from negative to positive.
The northwestern African continental margin
The Si–K-rich igneous rocks from the northwestern African margin also show a large variation in major and trace element composition. MgO ranges from 0·1 to 6·6 wt % and correlates positively with CaO, TiO2, FeO, CaO/Al2O3, Co (Gourougou) and MnO, Co, V (Guilliz) and negatively with SiO2, K2O, total alkalis, Rb (Gourougou) and K2O and Li (Guilliz) (r2 > 0·7) (Fig. 4e–h). In multi-element diagrams, Si–K-rich group rocks from Gourougou at the northwestern African margin show spiked patterns as a result of enrichment of fluid-mobile elements Rb, U and K relative to less fluid-mobile or fluid-immobile elements (e.g. REE, Nb and Ta) (Fig. 6a). However, the behaviour of Pb is more complex, showing both positive and negative anomalies, and the troughs in Nb and Ta of some samples are only moderately pronounced. The more evolved Si–K-rich samples from Guilliz have smooth patterns and show neither positive nor negative anomalies for the fluid-mobile elements Rb, Th, U, K and Pb and fluid-immobile elements Nb and Ta. The only negative anomalies are for Ba, Sr and Ti, and there is a slight positive anomaly for Zr (Fig. 6c). The Si–K-rich sample from the Oujda–Algeria area (Ahfir sample; open circle, Fig. 6e) shows enrichment of fluid-mobile Rb, Ba, Th and U relative to fluid-immobile Nb and Ta but lacks peaks for fluid-mobile elements K and Pb. The multi-element pattern of the Ahfir sample is largely subparallel to that of both the mafic and evolved Si–K-rich samples from northwestern Algeria (Fig. 6e) (Louni-Hacini et al., 1995; Coulon et al., 2002). Volcanic rocks having distinct troughs in Nb and Ta but only slight positive peaks for fluid-mobile elements can be found in the active Aeolian subduction zone (Fig. 5c).
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For the Si-poor, Na-rich lavas, MgO (1·8–12·6 wt %) correlates positively with Cr, Ni, Ga (Gourougou), Cr, Ni (Guilliz) and CaO, CaO/Al2O3, Cr, Ni (Oujda area) and inversely with Al2O3, K2O, Th (Gourougou), Al2O3 (Guilliz) and Al2O3, MnO, Cs, Zr, Hf, Eu, Yb, Lu and Y (Oujda) (r2
0·7, except for Cr and Ni, which show curved correlations with MgO) (Fig. 4e–h). The Oujda samples also show a strong decrease in TiO2, FeO, Co and V as MgO decreases below 3 wt %. These systematic correlations between major and trace elements point to the fractional crystallization of the observed phenocryst assemblages (olivine, clinopyroxene and magnetite). In multi-element diagrams (Fig. 6b, d, f), the Si-poor lavas from northwestern Africa generally show contrasting patterns to the Si–K-rich group as a result of a relative depletion in fluid-mobile elements Rb, Th, U, K and Pb, compared with less fluid-mobile elements such as the REE, Nb and Ta. These patterns are similar to those of lavas from oceanic intraplate-type volcanic settings (e.g. St. Helena and Canary Islands in Fig. 5d) (Chaffey et al., 1989; Thirlwall, 1997). Several Si-poor samples from Gourougou and Guilliz, however, show only weak Nb and Ta anomalies, and a distinct negative K anomaly is lacking. It should be emphasized that several samples from both the Si–K-rich group and the Si-poor group show deviations from typical subduction-related or intraplate-type geochemical fingerprints and have hybrid compositions between the two end-member geochemical signatures (Figs 5 and 6). Some of the Si–K-rich samples lack relative enrichment and even show slight relative depletion in some fluid-mobile elements (in particular Th, U, K and Pb for Guilliz and Pb for Gourougou lavas). These samples also display large ranges in fluid-mobile to fluid-immobile element ratios, which have values between those found in OIB and subduction zone lavas. The Si–K-rich volcanic rocks from Gourougou, for example, have U/Nb = 0·09–0·33 and K/La = 732–1143, whereas similar rocks from Guilliz have U/Nb = 0·03–0·07 and K/La = 461–582. Some mafic Si-poor samples exhibit less pronounced depletion or even slight enrichment in fluid-mobile elements (e.g. Th, U, K and Pb for Guilliz and Gourougou lavas and Th, U, K and Pb for Tallante lavas). Selective addition of strongly fluid-mobile elements relative to much less fluid-mobile elements, which have almost identical partition coefficients (such as Pb relative to Ce) cannot be significantly controlled by differences in the degree of partial melting or by fractional crystallization processes in mafic magmas but could reflect binary mixing of magmas from different sources, derivation from a subduction-modified mantle source or crustal assimilation.
Sr–Nd–Pb-isotope data
Sr–Nd–Pb-isotope data for 52 igneous rocks from the southern Iberian and northwestern African continental margins are reported in Table 3. Figures 7, 8, 10 and 11 display age-corrected Sr–Nd–Pb-isotope ratios.
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The southern Iberian continental margin
Si–K-rich, lamproitic rocks have high initial 87Sr/86Sr ratios (0·7161–0·7230) and low initial 143Nd/144Nd ratios (0·51215–0·51223) (Fig. 7a, c, d). The Sr-isotope ratios reported here slightly extend the range of published data (0·7166–0·7207) (Nelson, 1992
; Turner et al., 1999) but the new Nd-isotope ratios are significantly higher and show much less scatter than published data (143Nd/144Nd = 0·51197–0·51214, normalized to 146Nd/144Nd = 0·7219 and age corrected to 7 Ma) (Nelson, 1992; Turner et al., 1999). The Si–K-rich, lamproitic rocks show a relatively narrow range in initial 206Pb/204Pb (18·71–18·85), 207Pb/204Pb (15·66–15·69) and 208Pb/204Pb ratios (38·90–39·13) and have high 
7/4 (14·3–16·5) and 8/4 (61·4–76·8) (Fig. 7b) [7/4 and 8/4 notations denote the vertical deviation from the Northern Hemisphere Reference Line (NHRL), in the 206Pb/204Pb vs 207Pb/204Pb and 208Pb/204Pb diagrams, respectively] (Hart, 1984). It should be noted that the Si–K-rich, lamproitic rocks, however, have Nd and Pb isotopic compositions very similar to those of cordierite-bearing rocks from southeastern Spain (Fig. 7a, b, d).
Si-poor, Na-rich lavas from southeastern Iberia have relatively low initial 87Sr/86Sr ratios (0·7037–0·7046) and relatively high initial 143Nd/144Nd ratios (0·51283–0·51298) (Fig. 7a, c, d). Despite higher initial 206Pb/204Pb (18·88–18·98) compared with the Si–K-rich, lamproitic rocks, the Si-poor lavas from Tallante have similar 207Pb/204Pb and 208Pb/204Pb ratios (15·65–15·67 and 38·9–39·0) to the Si–K-rich rocks from the same area and therefore plot well above the NHRL, also having relatively high 7/4 and 8/4 (10·2–13·3 and 38·6–54·7; Fig. 7b).
The northwestern African continental margin
Si–K-rich igneous rocks from the African continental margin show a wide range of initial 87Sr/86Sr ratios (0·7040–0·7100) and initial 143Nd/144Nd ratios (0·51236–0·51284), similar to the range reported for the Oujda–Algeria area (87Sr/86Srinitial = 0·7043–0·7096 and 143Nd/144Ndinitial = 0·51226–0·51254) (Coulon et al., 2002) (Fig. 7a, c, d). On Pb-isotope diagrams, they show a limited range for initial 206Pb/204Pb (18·7–19·0) and 207Pb/204Pb (15·63–15·70) ratios and plot above the NHRL (7/4 = 7·1–12·9, 8/4 = 30·7–56·9; Fig 7b).
Most Si-poor, Na-rich lavas from the northwestern African continental margin have lower initial 87Sr/86Sr (0·7029–0·7045) and higher 143Nd/144Nd (0·51274–0·51303) than the spatially associated Si–K-rich rocks and overlap with the range reported for Si-poor rocks from northern Algeria (87Sr/86Srinitial = 0·7032–0·7060 and 143Nd/144Ndinitial = 0·51255–0·51292) (Coulon et al., 2002) (Fig. 7a, c, d). The African Si-poor rocks exhibit a large range in initial Pb-isotope ratios (206Pb/204Pb = 18·85–20·75, 207Pb/204Pb = 15·65–15·67, 208Pb/204Pb = 38·9–39·0), although the range within the individual volcanic areas is much more limited, with the exception of lavas from the Plateau du Rekkam volcanic field in the Oujda–Algeria area (Fig. 7b). Si-poor samples from Gourougou and Guilliz plot well above the NHRL (except one sample) (7/4 = –0·1–9·2, 8/4 = –2·2–43·8) and the data fields overlap those of the spatially associated Si–K-rich rocks but extend to more radiogenic Pb isotopic composition. Most Si-poor volcanic rocks from the Oujda–Algeria area plot below the NHRL (7/4 = –7·7–5·8, 8/4 = –42·1–34·8) and fall within the field for Ahaggar volcanic rocks (southern Algeria) (Allègre et al., 1981) and two samples have 206Pb/204Pb ratios as high as St. Helena OIB (Chaffey et al., 1989). The variation in 206Pb/204Pb ratio for Si-poor lavas from the northwestern African margin almost covers the field for the modelled composition of young HIMU plotting below the NHRL, reflecting the composition of MORB with a 238U/204Pb of 10 and a 0·5–1·5 Gyr recycling time (Thirlwall, 1997) (Fig. 7b). In contrast to the young HIMU field, most Si-poor lavas plot closer to or lie above the NHRL; Si-poor Oujda lavas, however, overlap with the low-velocity component (LVC) (Hoernle et al., 1995).
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DISCUSSION |
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TOPABSTRACTINTRODUCTIONREGIONAL GEOLOGYANALYTICAL METHODSRESULTSDISCUSSIONGEODYNAMIC EVOLUTIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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Temporal evolution of volcanism at the Iberian and African margins
Dating of the westernmost Mediterranean igneous rocks suggests that the transition from Si–K-rich to Si-poor, Na-rich compositions occurred between 6·4 and 2·9 Ma in Spain along the southern Iberian continental margin, between 6·8 and 4·8 Ma in Morocco and possibly between 10·0 and 7·5 Ma in Algeria within the northwestern continental African margin (Fig. 8a–h) (Louni-Hacini et al., 1995
; El Bakkali et al., 1998; Coulon et al., 2002).
The transition also occurred in individual volcanic systems such as the Gourougou and Guilliz stratovolcanic complexes over shorter timescales (Figs 1, 2 and 8a–h). At the Gourougou stratovolcanic complex, Si–K-rich igneous activity occurred from the Late Miocene to Early Pliocene (7·58 to 4·8 Ma) partially overlapping the Latest Miocene to Early Pliocene Si-poor volcanism (6·3 to 3·73 Ma), which erupted from localized vents on the flanks of the stratovolcano. In the Guilliz volcanic field, a short interval of c. 0·4 Myr separates Late Miocene Si-rich explosive and intrusive volcanism (6·90 to 6·69 Ma) from the eruption of lava flows from scoria cones with Si-poor compositions in the Latest Miocene to Pleistocene (6·3 to 0·65 Ma). The Si–K-rich Gourougou igneous rocks show a systematic change in their geochemical characteristics with decreasing age over a time span of about 3 Myr (7·58 to 4·8 Ma). This is observed in their major and trace element and Nd isotope ratios (e.g. K2O/Na2O, K2O/Al2O3, Ba/Sr, Th/Nb, K/Nb, K/La, Pb/Ce, 87Sr/86Sr, 7/4Pb decrease and Nb/La, Nb/Zr and 143Nd/144Nd increase) (Fig. 8a–c, f, h). A similar geochemical evolution is observed in the Si-poor Gourougou lavas, which show a systematic decrease of incompatible element ratios (e.g. Ba/Sr, K/La, K/Nb, Pb/Ce) with decreasing age (Fig. 8a, f). The systematic temporal variations of both isotope and incompatible element ratios with very similar partition coefficients cannot be explained by differences in the degree of partial melting or by fractional crystallization, but instead must reflect crustal assimilation or binary mixing between different components from a heterogeneous mantle source.
Mantle source identification
As illustrated in the Sr–Nd–Pb-isotope diagrams (Fig. 7a–d), a number of distinct mantle and/or source components appear to play a role in the petrogenesis of the Si–K-rich and Si-poor groups: (1) lithospheric components such as the continental crust and subduction-modified lithospheric mantle (e.g. Enriched Mantle 2, EM2) having high 87Sr/86Sr, 7/4 and 8/4, moderate 206Pb/204Pb ratios and generally low 143Nd/144Nd ratios, and (2) sub-lithospheric components such as young and old HIMU mantle with low 87Sr/86Sr, high 143Nd/144Nd and moderate 206Pb/204Pb ratios combined with negative 7/4 (e.g. Canary Islands and Ahaggar) or high 206Pb/204Pb ratios combined with positive 7/4 (e.g. St. Helena) (Thirlwall, 1997). Below we explore the role of the continental crust, the subcontinental lithospheric mantle, sub-lithospheric mantle reservoirs and the possible depths of melting in the origin of Si–K-rich and Si-poor group magmas from the southeastern Iberian and northwestern African continental margins.
The Si–K-rich group from the southern Iberian margin (8·2–4·8 Ma)
The Si–K-rich volcanic rocks (lamproites) from the southeastern Iberian continental margin form a compositional continuum with cordierite–garnet-bearing lavas from southern Spain, which are thought to result from crustal anatexis of a metapelitic protolith (Fig. 4a–d) (Zeck, 1970; Munksgaard, 1984; Cesare et al., 1997; Cesare & Maineri, 1999). The geochemical similarity of the Spanish low-MgO lamproites (some of which contain cordierite and garnet, e.g. the lamproitic rocks from Mazarron based on petrographic observations) and the cordierite–garnet-bearing lavas from Cerro Hoyazo suggests that the cordierite–garnet-bearing igneous rocks themselves may be highly crustally contaminated derivatives of mafic lamproitic melts rather than pure crustal anatexites. This is consistent with the results of geochemical modelling by Benito et al. (1999), who proposed a mantle-derived origin for the parental melts of the Cerro Hoyazo lavas. However, as inferred from the previously mentioned compositional continuum between high-MgO lamproites, low-MgO lamproites and cordierite–garnet-bearing lavas, the Si–K-rich, lamproitic lavas with MgO below 6–8 wt % appear to have apparently assimilated significant amounts of continental crust. Although the high-MgO lamproites (MgO >8 wt %) seem to be unaffected by crustal contamination, they have isotopic compositions very similar to crustal materials such as local continental crust and marine sediments (Figs 5a and 7a–d). High-MgO lamproites also show extreme enrichments of most incompatible elements (e.g. K2O = 3·6–10·2 wt %) and have high Cr (740–1230 ppm) and Ni (100–340 ppm) contents (Fig. 5a and Table 2). These more mafic magmas appear, therefore, to represent primitive mantle melts, despite their crustal trace element and isotopic geochemical signatures. This suggests that they are derived from enriched subcontinental mantle lithosphere modified by fluids or melts probably released from subducted marine sediments, resulting in extremely high 87Sr/86Sr and low 143Nd/144Nd ratios as well as a large positive deviation from the NHRL (high positive 7/4 and 8/4) (Venturelli et al., 1984, 1988; Nelson et al., 1986; Benito et al., 1999; Turner et al., 1999).
Based on results from experimental studies, the primary Spanish lamproite melts are thought to originate from depths of about 30–50 km within the uppermost portion of the subcontinental mantle lithosphere (Foley, 1993), i.e. within the stability field of spinel lherzolite. Derivation of the southern Iberian primary Si–K-rich melts from a spinel lherzolite mantle source, even if it contains K-rich amphibole, is not, however, consistent with their high Dy/Yb ratios (Fig. 9a). High Dy/Yb ratios require residual garnet in the source of the magmas, which preferentially incorporates the HREE. Very low degrees of partial melting of a garnet–phlogopite lherzolite with K contents ranging from 3000 to 9000 ppm could explain the variation of K/Yb and Dy/Yb (Fig. 9a). According to Foley, however, petrogenetic models for lamproites based on partial melting of phlogopite- (and amphibole-) bearing lherzolitic or harzburgitic mantle sources belie a more complex melting history including phlogopite–amphibole-rich pyroxenitic veins and vein–melt–wall-rock interaction (Foley, 1992, 1993). Based on geochemical modelling, 1–30% partial melts from a vein having the composition Orthopyroxene0·05–Clinopyroxene0·20–Garnet0·05–Phlogopite0·40–Amphibole0·25–Apatite0·03–Rutile0·02 [as estimated from the diagrams of Foley (1992)] ought to have relatively low Dy/Yb and high K/Yb ratios. The partial melting curve for such vein material plots far outside the diagram in Fig. 9a, as the modelled melts have extremely high K/Yb (130 000–140 000) combined with relatively low Dy/Yb ratios (1·30–1·35) (i.e. to the right of the legend box in Fig. 9a). The position of the modelled curve, however, is relatively robust to moderate changes in the orthopyroxene, clinopyroxene, phlogopite, amphibole and rutile contents in the vein but very sensitive to minor changes in the amount of garnet and apatite. For a wide range of 0–10% garnet and apatite contents, however, K/Yb and Dy/Yb ratios stay above 100 000 and below 2·5, respectively. Therefore vein melting alone seems to be inadequate to explain the high Dy/Yb ratios observed for the lamproites (2·8–3·8). Foley (1992) proposed an attractive model of interaction of vein melts with a garnet peridotite wall rock. Significant interaction of lamproitic vein melts with garnet lherzolite or melts therefrom could increase the Dy/Yb and lower the K/Yb ratios to the range observed in the lamproites. High Dy/Yb, however, indicates interaction of lamproite melts with garnet lherzolite and thus an origin from within the garnet peridotite stability field at depths 70–80 km.
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, 0·007, 0·0015; 
, 0·022, 0·042; 
, 0·33, 0·28; 
, 1·06, 4·01; 
, 0·01, 0·01; 
, 0·029, 0·03; 
, 0·78, 0·59, respectively (these values were selected from the GERM website, at The Si-poor group from the southern Iberian margin (2·93–2·29 Ma)
Si-poor lavas from the Tallante volcanic field at the southeastern Iberian continental margin show correlations between Sr–Nd–Pb-isotope ratios and major and trace element concentrations and inter-element ratios. These correlations provide important information about the possible role of crustal and mantle components in their petrogenesis. 87Sr/86Sr correlates positively with SiO2, Th, U, Pb, (Th, U)/Nb, Nb/La and Pb/Nd, and negatively with CaO, 143Nd/144Nd and 206Pb/204Pb (see examples in Figs 10 and 11). 143Nd/144Nd shows inverse correlations with U, Th, Pb, (U, Th)/Nb and Pb/Nb. 206Pb/204Pb displays inverse correlations with CaO, U/Nb and Pb/Nd (with r2 > 0·7). Whereas the most mafic samples are likely to reflect the composition of their mantle source, the more evolved samples are likely to provide information about the composition of assimilants during fractional crystallization (AFC). On this basis, it can be inferred that the contaminant has higher 87Sr/86Sr, (Th, U)/Nb and Pb/Nd but lower 143Nd/144Nd and 206Pb/204Pb than the mafic Si-poor Tallante magmas and their mantle sources.
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Such geochemical variations may result from interaction of primitive Si-poor Tallante magmas with the continental lithosphere, e.g. the continental crust or metasomatized subcontinental mantle. Volcanic rocks from the Tallante volcanic field frequently contain lower crustal xenoliths and mantle xenoliths (Vielzeuf, 1983
; Dupuy et al., 1986). Although the presence of crustal xenoliths indicates the possibility for assimilation of upper crustal material, the ubiquity of mantle xenoliths in the Tallante lavas argues for rapid magma ascent and against significant residence time in the continental crust, which would allow for AFC processes to occur. Therefore, the increase of 87Sr/86Sr, (Th, U)/Nb and Pb/Nd coupled with a decrease of 143Nd/144Nd and 206Pb/204Pb in the Si-poor Tallante melts is most likely to reflect interaction with metasomatized sub-continental mantle lithosphere. Interaction with lamproitic melts derived from the lithospheric mantle serves as an alternative explanation. Interaction with metasomatized sub-continental mantle or melts therefrom is consistent with the occurrence of phlogopite in the mantle xenoliths found in the Tallante lavas (Dupuy et al., 1986; Capedri et al., 1989). As Si–K-rich lamproitic melts both have higher 87Sr/86Sr and SiO2, and lower CaO contents than the Si-poor Tallante magmas (Table 2) (Venturelli et al., 1988; Benito et al., 1999; Turner et al., 1999), interaction with lamproitic melts could also partly explain: (1) the observed decrease in CaO and increase in SiO2 coupled with increasing 87Sr/86Sr; (2) their potentially high 7/4 and 8/4 well above the NHRL (Fig. 7b); (3) the hybrid geochemical characteristics between intraplate-type and subduction-related geochemical signatures in multi-element diagrams (e.g. positive Th–U anomalies, variable Nb, Ta, K and Pb anomalies) (Fig. 5b). However, the geochemical similarity to lavas from intraplate volcanic areas such as the Canary Islands in terms of major and most trace elements and Sr–Nd-isotope ratios points to derivation of the primary Si-poor melts from sub-lithospheric mantle sources and has already been pointed out in the literature (Wilson & Bianchini, 1999; López-Ruiz et al., 2002; Duggen et al., 2003).
The FeO and SiO2 contents of mafic partial melts in intraplate volcanic settings can vary as a function of melting depths (Hirose & Kushiro, 1993). As illustrated in Fig. 12a, the FeO and SiO2 contents of the more mafic Tallante lavas combined with experimental data for partial melting of dry peridotite at high pressures (Hirose & Kushiro, 1993) point to a derivation from a peridotitic source (Mg-number between 85 and 89) at depths between c. 40 and 80 km, i.e. presumably primarily within the spinel stability field. It should be emphasized that neither SiO2 nor FeO in the mafic Si-poor, Na-rich Tallante lavas (MgO >6 wt %) correlates with MgO. Therefore the variations in SiO2 and FeO observed for the mafic Tallante lavas are unlikely to result from fractional crystallization of magnesian phenocrysts and thus more probably reflect mantle processes such as partial melting and/or mixing with high-Si–K, low-Fe lamproitic melts. However, the moderately high Dy/Yb ratios at around three in the most mafic Si-poor Tallante lavas (>6 wt % MgO) require the presence of residual garnet in the source of the Tallante melts.
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In the K/Yb vs Dy/Yb diagram (Fig. 9a), the Si-poor samples from Tallante plot close to partial melting curves modelled for three different scenarios: (1) partial melting of a garnet lherzolite source; (2) a spinel–garnet lherzolite source; (3) a garnet-facies amphibole lherzolite source with low K contents (250 ppm, i.e. a Na-rich rather than a K-rich amphibole). Garnet in peridotite is stable at depths >70–80 km, which is the approximate depth for the spinel–garnet transition zone in the upper mantle. Amphibole is unstable at depths exceeding 70–100 km (Green, 1973
; Wyllie & Wolf, 1993) and therefore garnet-facies amphibole peridotite can exist only in a narrow zone roughly between 70 and 100 km depth, probably within the mechanical boundary layer of the subcontinental lithosphere (Pearce et al., 1990). Because the subcontinental lithosphere beneath southeastern Iberia at this depth appears to be highly metasomatized, providing the source for the Si–K-rich lamproitic melts, it seems unlikely that garnet-facies amphibole lherzolite constitutes the source of the Si-poor Tallante melts. This is consistent with their geochemical signature, pointing to derivation from a sub-lithospheric mantle source (Figs 5b and 7a–d). Alternatively, the Si-poor Tallante melts could have been formed through partial melting of an amphibole-free source: the mafic Tallante lavas plot close to the model curves for 6–8% partial melting of a garnet lherzolite and 1–2% partial melting of lherzolite containing both spinel and garnet. Melting in the spinel–garnet transition zone is able to explain slightly elevated K/Yb as observed for one Tallante sample without involving K-bearing phases in the mantle source. Melting of a garnet lherzolite, however, requires an additional process that is capable of enhancing K/Yb, such as, for example, interaction of Si-poor magmas with K-rich, metasomatized mantle lithosphere or lamproitic melts derived from such lithosphere. Interaction with MgO-rich lamproitic melt components may also explain why both SiO2 and K/Yb correlate positively with MgO in the three most mafic lavas with MgO between 7·7 and 8·5 wt %, which is the opposite of what is expected for fractional crystallization combined with crustal assimilation; lamproitic melts have higher SiO2, higher K/Yb (and lower FeO) than the Tallante lavas (Figs 3 and 9a, and Table 2).
In conclusion, the Si-poor Tallante lavas from the southeastern Iberian continental margin probably represent low-degree partial melts of garnet-bearing sub-lithospheric mantle material such as garnet lherzolite or lherzolite from the garnet–spinel transition zone at depths 70–80 km. Variable interaction of these Si-poor melts with Si–K-rich components or melts from the metasomatized subcontinental mantle lithosphere is interpreted to have partially modified the intraplate geochemical fingerprint of the sub-lithospheric melts towards subduction-related signatures. How sub-lithospheric mantle material may undergo partial melting at depths between c. 70 and 100 km to form the Si-poor Tallante lavas in an area in which the normal sub-continental lithosphere thickness is c. 150 km will be addressed subsequently in the discussion of the geodynamic evolution of this region.
The Si–K-rich group from the northwestern African margin (7·58–4·82 Ma)
Si–K-rich igneous rocks from Gourougou and Guilliz show linear correlations between radiogenic isotope ratios and major and trace elements. 87Sr/86Sr correlates positively with La and Nd (Gourougou) and negatively with Na2O (Guilliz), 143Nd/144Nd increases with increasing Sr (Gourougou and Guilliz) and CaO (Guilliz) and 206Pb/204Pb correlates positively with Na2O (Guilliz) (with r2 > 0·7). Hyperbolic correlations exist between Sr-isotope ratios and incompatible trace element ratios such as U/Nb and Nb/La (Fig. 11a). Such correlations, in which several samples achieve elevated 87Sr/86Sr up to 0·710, indicate that AFC processes affected the geochemical composition of some Si–K-rich magmas in the Gourougou and Guilliz stratovolcanic complexes. As illustrated in the SiO2 and 1/Sr vs 87Sr/86Sr diagrams (Figs 10a and 11c), three Gourougou igneous rocks, one Guilliz sample and the most Si–K-rich igneous rocks from the Oujda–Algeria area have elevated 87Sr/86Sr coupled with high SiO2 and 1/Sr (Coulon et al., 2002). However, several Si–K-rich samples from Gourougou and Guilliz (and one Si–K-rich sample from the Oujda–Algeria area) have low Sr-isotope ratios despite a significant variation in SiO2 and Sr contents, which argues against significant crustal contamination. The Sr and Nd isotopic composition of the Si–K-rich igneous rocks from Gourougou and Guilliz with low Sr-isotope ratios overlaps the field for the Si-poor lavas from the same area despite different Pb isotopic compositions. This is not consistent with a co-genetic derivation of the Si–K-rich lavas from the Si-poor ones by AFC (Fig. 11) and, therefore, we infer that the primary magmas of the Si–K-rich and Si-poor volcanic rocks at the northwestern African continental margin were derived from different mantle sources. Major and trace element and Sr–Nd–Pb radiogenic isotope ratios of the Si–K-rich igneous rocks, therefore, suggest derivation from a K-enriched, metasomatized mantle source with a subduction-related geochemical signature, possibly inherited from an older subduction zone. Because of its vicinity to the Alboran Basin (Fig. 1), the source of the Gourougou magmas may also have been affected by hydrous fluids and melts from the Miocene Alboran subduction zone (Hoernle et al., 1999; Duggen et al., 2004; Gill et al., 2004).
The K/Yb vs Dy/Yb diagram in Fig. 9b provides constraints on the depths of melting for the Si–K-rich Gourougou magmas. Two of the more mafic Gourougou samples with MgO > 4 wt % and low 87Sr/86Sr are plotted. The Guilliz Si–K-rich trachyandesites are not included in Fig. 9b because of uncertainties arising from their evolved nature such as fractionation of REE through fractional crystallization of apatite, which is evident from a negative correlation of SiO2 and P2O5 (r2 = 0·9) for Si–K-rich samples with >57 wt % silica. The more mafic Si–K-rich Gourougou lavas (SiO2 <55·5 wt %) (and one mafic Si–K-rich sample from Algeria) have Dy/Yb ratios slightly below two and elevated K/Yb ratios. Elevated Dy/Yb ratios require residual garnet in the source of the melts but they are too low to result from partial melting of garnet lherzolite alone (Fig. 9b). Different scenarios can be considered: (1) mixing between 1–10% partial melt of a garnet-facies phlogopite lherzolite with 0·5–10% partial melt of a spinel lherzolite or 1–15% melt from a spinel-facies amphibole lherzolite; (2) mixing of 1–15% garnet-facies amphibole lherzolite melts with spinel lherzolite melts or spinel-facies amphibole lherzolite melts; or (3) mixing of garnet lherzolite melts or spinel–garnet transition zone melts with spinel-facies amphibole lherzolite melts. Each scenario requires residual garnet in the source region and the presence of K-bearing phases. A plausible explanation would be a melting column with the lower end located in the upper garnet-stability field and the upper end situated in the spinel lherzolite field. In conclusion, Si–K-rich melts from the northwestern African continental margin were probably derived from a K-enriched, metasomatized mantle source from depths between 50 and 100 km, i.e within the mechanical boundary layer of the subcontinental lithosphere. K enrichment may have been associated either with fluids or hydrous melts introduced from the adjacent Miocene subduction zone in the Alboran Basin (e.g. Gourougou) or with an earlier subduction event (e.g. Guilliz) or both.
The Si-poor group from the northwestern African margin (6·25–0·65 Ma)
Several samples from the Si-poor group from the northwestern African continental margin plot significantly below the NHRL and therefore have negative 7/4 values up to –7·7 (e.g. igneous rocks from Plateau du Rekkam, the Oujda volcanic field except for one sample, and one sample from Guilliz) (Fig. 7b). On the African plate, basaltic lavas plotting on or below the NHRL can be found in the adjacent intraplate volcanic areas of the Canary Island Archipelago (eastern North Atlantic) and Ahaggar (southern Algeria); these have geochemical characteristics suggesting derivation from sub-lithospheric mantle sources (Allègre et al., 1981; Hoernle & Schmincke, 1993a, 1993b). It is very unlikely that lavas with negative 7/4 values have been significantly affected by continental crustal contamination, as the continental crust plots well above the NHRL and has a high Pb content (Taylor & McLennan, 1985; Thirlwall, 1997). Therefore, such lavas provide direct geochemical information about their mantle sources, in terms of radiogenic isotope ratios and ratios of elements having very similar partition coefficients during partial melting and fractional crystallization in basaltic melts (e.g. the pairs U–Nb, K–La and Pb–Nd). It can, thus, be inferred that the source of the Si-poor melts with negative 7/4 had low 87Sr/86Sr <0·7039, high 143Nd/144Nd >0·5127, 206Pb/204Pb ranging from 19·3 to 20·8, and U/Nb = 0·0172–0·0230, K/La = 80·2–411·1 and Pb/Nd = 0·0326–0·0751. In terms of Sr–Nd-isotope and inter-element ratios this source is very similar to that of lavas from the Canary Islands derived from sub-lithospheric mantle sources but it tends to have higher 206Pb/204Pb ratios (e.g. Gran Canaria 87Sr/86Sr <0·7034, 143Nd/144Nd >0·5128, 206Pb/204Pb = 19·5–20·0, U/Nb = 0·0146–0·0229, K/La = 98·8–418·7, Pb/Nd = 0·0369–0·0557) (Hoernle et al., 1991; Thirlwall, 1997). Interestingly, both Oujda and Plateau du Rekkam volcanic rocks with negative 7/4 yield regression lines subparallel to the Ahaggar array in the uranogenic Pb-isotope diagram (Fig. 7b). If interpreted as Pb–Pb isochrons, these regression lines correspond to c. 840 Ma and 975 Ma ages, respectively. Similar Pb–Pb isochron ages were found for the neighbouring Madeira hot spot system in the eastern North Atlantic and were interpreted to reflect short-term recycling of oceanic lithosphere in the Earth's mantle (Geldmacher & Hoernle, 2000; Geldmacher et al., 2000). Many Oujda lavas, however, also plot close to the low-velocity component (LVC) (Fig. 7b–d), which is observed in intraplate-type lavas from Central and Western Europe, the Central Mediterranean and the Eastern North Atlantic (Hoernle et al., 1995). This suggests that most Si-poor igneous rocks from the Oujda–Algeria area were derived from a sub-lithospheric mantle source with plume-like characteristics.
The Si-poor igneous rocks from the Gourougou and Guilliz volcanic fields extend from the NHRL on Pb isotope diagrams to the field for Si–K-rich rocks with high 7/4 values (e.g. Fig. 7b). For Si-poor Gourougou lavas 7/4 correlates negatively with Nb/Zr (r2 = 0·92) and Nb/Yb (r2 = 0·85), whereas for the Si-poor Guilliz samples 87Sr/86Sr correlates negatively with Eu anomaly (Eu/Eu* ranging from 1·07–0·95, r2 = 0·92) (Fig. 10d). These correlations provide information about the high 7/4 end-member or contaminant of the sublithospheric melts, which has 87Sr/86Sr >0·7045, 143Nd/144Nd <0·5127, 7/4 >8·6, 8/4 >44, which can also be inferred from the Sr–Nd–Pb isotope diagrams (Fig. 7a–d). Components with such a geochemical fingerprint can be found in the continental crust or metasomatized subcontinental mantle lithosphere or melts therefrom, for example, the Si–K-rich melts in the Gourougou and Guilliz volcanic centres (Figs 7, 10 and 11). As no clear correlations exist between MgO or SiO2 with Nb/Zr, Nb/Yb, Eu anomaly or Sr–Nd–Pb-isotope ratios, it seems unlikely that fractional crystallization of magnesian phenocryst phases was associated with an increase of 87Sr/86Sr, 7/4 and 8/4 and a decrease of 143Nd/144Nd, Nb/Zr and Nb/Yb. These observations argue against significant crustal contamination of the Si-poor Gourougou and Guilliz magmas by AFC requiring significant fractional crystallization. Instead, the geochemical systematics point to an interaction of the Si-poor melts with the metasomatized subcontinental mantle lithosphere or melts therefrom. Unfortunately, AFC cannot conclusively be excluded for some of the samples having elevated 7/4 and 8/4. However, the aforementioned geochemical correlations and the composition of the more mafic samples (MgO >6–8 wt %) provide information about the composition of the mantle source of the primary Si-poor magmas. It can be inferred that the source of the Si-poor Gourougou lavas had 87Sr/86Sr <0·7033, 143Nd/144Nd <0·5130, 7/4 <4·3 and 8/4 <9·0, and that of the Guilliz melts had 87Sr/86Sr <0·7029, 143Nd/144Nd <0·5130, negative 7/4 and 8/4 < –2·2. This isotopic composition along with the intraplate-type multi-element patterns (Figs 6b, d, f, and 7a–d) is consistent with derivation of the primary Si-poor Gourougou and Guilliz lavas from a sub-lithospheric mantle source. As can be inferred from the uranogenic Pb-isotope diagram (Fig. 7b), this source was probably contaminated by plume material having a young-HIMU geochemical fingerprint similar to that of the Canary Islands (Hoernle & Tilton, 1991; Hoernle et al., 1991; Thirlwall, 1997).
The major and trace element compositions of the mafic Si-poor igneous rocks from the northwestern African continental margin can provide information about the melting regime of their mantle source. As indicated by the SiO2 vs FeO diagram in Fig. 12b, Si-poor melts from the Plateau du Rekkam volcanic field could have been generated by partial melting of a peridotite with decreased fertility at depths exceeding 100 km, i.e. outside the amphibole stability field. Many Oujda samples plot at the end of the experimental field for the melting of a peridotite with Mg-number of 
89, consistent with derivation from depths exceeding 100 km. Several Oujda samples, however, lie close to the experimental field for partial melts of a more fertile peridotite with Mg-number = 85, pointing to derivation at depths exceeding 70–80 km depth from a more fertile source. In contrast, most of the Si-poor lavas from Algeria have SiO2 and FeO contents pointing to magma generation at shallower depths (c. 40–100 km) than the Plateau du Rekkam and Oujda lavas. These estimates are consistent with geochemical modelling results in the K/Yb vs Dy/Yb diagram in Fig. 9b. Lavas from the Plateau du Rekkam area have very high Dy/Yb ratios (>3·5) combined with low K/Yb ratios and therefore plot very close to a garnet lherzolite partial melting curve (4–5% partial melting). Very high Dy/Yb and relatively low K/Yb argue for the absence of spinel, phlogopite and K-bearing amphibole in the source of Plateau du Rekkam lavas. The situation is slightly different for the remaining samples from the Oujda–Algeria area. They have high Dy/Yb but show a large variation in K/Yb ratios. The Oujda–Algeria lavas therefore plot close to the garnet lherzolite, the garnet-facies amphibole lherzolite and the spinel–garnet transition zone model partial melting curves. Several Oujda samples with elevated K/Yb, shown in Fig. 9b, plot close to the high-pressure tip of the experimental field for partial melting of a peridotite with Mg-number = 89 indicating derivation from >100 km, i.e. outside the stability fields for amphibole and spinel. Those Oujda lavas plotting close to the Mg-number = 85 field could have been derived from 70–100 km depths where spinel and amphibole may have been stable. However, the mafic samples from Oujda do not plot along one of the partial melting curves. Instead, they define an array with a positive correlation of K/Yb with Dy/Yb, which suggests binary mixing of two melt end-members (Fig. 9b). Therefore, a more plausible model to explain the composition of the Si-poor lavas from the Oujda volcanic field is mixing of low-degree (<1%) partial melts from a garnet lherzolite source with moderate-degree (5–10%) partial melts from the spinel–garnet transition zone.
The Si-poor Gourougou and Guilliz magmas were apparently generated from a more fertile peridotite source than the Si-poor lavas from the Oujda–Algeria area. As neither SiO2 nor FeO correlates with MgO for the mafic Si-poor Gourougou and Guilliz samples, elevated SiO2 and lowered FeO are unlikely to result from fractional crystallization. Therefore, SiO2 and FeO contents indicate shallower depths of melting on the order of 40–80 km. The Dy/Yb ratios of mafic Gourougou and Guilliz lavas are relatively high (>2·0) and, therefore, require the involvement of melts from the garnet stability field. One mafic Guilliz sample (GZ160699-5, MgO 11·2 wt %) has higher Dy/Yb ratios than the other samples in Fig. 9b and also has the lowest SiO2 (46·4 wt %) and the highest FeO content (10·7 wt %) of the mafic samples shown in Fig. 12b, which is consistent with a derivation from the spinel–garnet transition zone at 70–80 km depth. The other Guilliz samples plot below the spinel–garnet transition zone partial melting curve in Fig. 9b. Their higher SiO2 and lower FeO contents are not consistent with a significant melt component from a garnet lherzolite source (80 km) and point to involvement of a spinel lherzolite source (Fig. 12b). Therefore we propose derivation by mixing of low- to moderate-degree partial melts from the spinel–garnet transition zone with low- to moderate-degree partial melts from a spinel lherzolite (Fig. 9b). Interaction of Si-poor Guilliz melts with Si–K-rich material in the metasomatized subcontinental mantle lithosphere may also be involved, and could result in a decrease of Dy/Yb, FeO and increase of K/Yb and SiO2. Such a process is more clearly indicated for the mafic Si-poor melts from Gourougou as they form an array with their mafic Si–K-rich associates in Fig. 9b. Relatively low FeO contents and elevated SiO2 and K/Yb in the mafic Si-poor Gourougou melts may result from interaction with Si–K-rich melts in the subcontinental lithosphere (Figs 9b and 12); lavas with the lowest K/Yb would be derived from the deepest sources in this scenario. Taking into account both K/Yb and Dy/Yb ratios and SiO2 and FeO contents, we infer that the Si-poor Gourougou melts were derived by low to moderate degrees of partial melting in the garnet–spinel transition zone at c. 70–80 km depth, and mixed with partial melts (Si–K-rich) from the (amphibole-bearing) spinel lherzolite field in the subcontinental lithosphere.
Summary of the inferred mantle sources and the partial melting regime
Based on major and trace element and Sr–Nd–Pb-isotope ratios, the mantle sources of Upper Miocene to Lower Pliocene Si–K-rich igneous rocks from the Iberian and African continental margins appear to have been modified by Miocene or more ancient subduction events. The mafic Si–K-rich igneous rocks require garnet and K-bearing phases such as phlogopite and/or amphibole in their mantle source, which are different for the Iberian and African Si–K-rich rocks. The Spanish lamproites are inferred to have been generated by interaction of melts from phlogopite–amphibole–pyroxene–garnet veins with garnet peridotite wall-rock (Foley, 1993) and were generated deeper than 70 km based on the geochemical modelling results in this study. For the Si–K-rich rocks from the northwestern African margin, partial melting probably involved a melting column ranging from uppermost garnet lherzolite to well within the spinel lherzolite stability field, i.e. from depths between 50 and 100 km. K-bearing phases such as phlogopite or amphibole could have been stable within the full range of the melting column. The mantle sources of the Si–K-rich igneous rocks from the Iberian and African continental margins, however, were probably located in the mechanical boundary layer of the subcontinental lithosphere.
Based on intraplate-like geochemical signatures, the Late Miocene to Pleistocene mafic Si-poor igneous rocks were probably derived from sub-lithospheric mantle sources contaminated with mantle plume material such as young HIMU mantle (Thirlwall, 1997) or the low-velocity component (LVC; Hoernle et al., 1995). Mafic Si-poor melts with negative 7/4 (e.g. lavas from the Oujda–Algeria area and Guilliz) have probably not interacted with the continental lithosphere. Si-poor melts with positive 7/4 (e.g. lavas from southeastern Spain, Gourougou and Guilliz) show interaction with components in, or melts from, the metasomatized subcontinental lithospheric mantle. Depths of partial melting range from >100 km (garnet lherzolite stability-field) to about 40 km (spinel lherzolite stability-field). Partial melting and generation of Si-poor melts in several volcanic areas obviously involved melting columns ranging from garnet- to spinel-facies depths.
In summary, the depths of partial melting of the younger Si-poor melts (c. 40 to >100 km) overlap with the depths inferred for the spatially associated but generally older Si–K-rich igneous rocks (c. 50–100 km) despite fundamental differences in their major and trace element and Sr–Nd–Pb-isotope compositions. A discussion of the westernmost Mediterranean geodynamic evolution may help to solve the puzzle of why two fundamentally different mantle sources existed at the same depths but at different times.
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TOPABSTRACTINTRODUCTIONREGIONAL GEOLOGYANALYTICAL METHODSRESULTSDISCUSSIONGEODYNAMIC EVOLUTIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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The geochemical transition from subduction-related to intraplate-type geochemical signatures in mantle-derived rocks strongly points to major changes in mantle geometry beneath the westernmost Mediterranean close to the Miocene–Pliocene boundary (El Bakkali et al., 1998
; Hoernle et al., 1999; Coulon et al., 2002; Duggen et al., 2003). These large-scale mantle processes were associated with large-scale uplift and closure of marine gateways connecting the Mediterranean Sea with the Atlantic Ocean, which resulted in the isolation and desiccation of the Mediterranean Sea and major faunal changes that define the Miocene–Pliocene boundary (Hsü et al., 1973; Duggen et al., 2003). Several mantle sources with fundamentally distinct geochemical compositions were partially melted in the course of the large-scale reorganization of the western Mediterranean upper mantle: (1) a shallow mantle wedge (consisting of depleted asthenosphere and attenuated continental lithospheric mantle) enriched by fluids or melts from subducted oceanic lithosphere beneath the Alboran Basin in the Middle to Late Miocene (Hoernle et al., 1999; Duggen et al., 2004; Gill et al., 2004); (2) metasomatically enriched subcontinental mantle lithosphere beneath the southeastern Iberian and northwestern African continental margins modified by fluids or melts from the adjacent Miocene Alboran Basin subduction zone or an older subduction event; (3) plume-contaminated, ambient sub-lithospheric mantle upwelling beneath the continental margins of southeastern Iberia and northwestern Africa. Below we explore the possible role of these mantle reservoirs in the Miocene to Quaternary geodynamic evolution of the westernmost Mediterranean by integrating geochronological, geochemical, geological and geophysical data.
The Alboran Basin
Several models have been proposed for the formation of the Alboran Basin that invoke (1) subduction of oceanic lithosphere (Royden, 1993; Lonergan & White, 1997; Hoernle et al., 1999; Gutscher et al., 2002; Duggen et al., 2003, 2004; Gill et al., 2004); (2) detachment of near-vertical subducted oceanic lithosphere (Blanco & Spakman, 1993; Zeck, 1996, 1999; Hoernle et al., 1999); (3) detachment or delamination of subcontinental mantle lithosphere (Docherty & Banda, 1995; Comas et al., 1999; Turner et al., 1999; Calvert et al., 2000; López-Ruiz et al., 2002; Platt et al., 2003). Subduction of oceanic lithosphere and detachment or delamination of subcontinental lithosphere should produce volcanism with characteristic geochemical compositions. Therefore, we can use the geochemistry of the Alboran Basin volcanism to distinguish between these fundamentally different tectonic processes for the formation of the Alboran ‘Mediterranean-style back-arc basin’ (Horvath & Berckhemer, 1982).
Volcanic rocks from the Alboran Basin (Fig. 1) belong to the low-K (tholeiitic) and medium- to high-K (calc-alkaline) series with major and trace element and O–Sr–Nd–Pb-isotopic compositions very similar to those of volcanic front and rear-arc lavas found in active subduction zones (e.g. Izu–Bonin) (Hoernle et al., 1999; Duggen et al., 2004; Gill et al., 2004). Their geochemical characteristics provide evidence for magma generation involving high degrees of partial melting of depleted (harzburgitic) upper mantle material triggered by hydrous fluids and melts released from subducted oceanic lithosphere in the Middle to Late Miocene (c. 12–6 Ma) (Hoernle et al., 1999; Duggen et al., 2004). This is consistent with the boninitic affinity of mafic lavas from Alboran Island and the Alboran sea floor (Hoernle et al., 1999; Gill et al., 2004). The geochemical characteristics of the Alboran Basin lavas such as mantle-like 
18O as low as 5·3 (Duggen et al., 2003, 2004; Duggen et al. unpublished data) undoubtedly rule out an origin exclusively by anatexis of continental crust as proposed in the literature (Zeck et al., 1998, 1999).
Detachment or delamination models advanced for the formation of the Alboran Basin are difficult to envisage for several reasons. Detachment of a near-vertical slab beneath the centre of the Alboran Basin in the Early Miocene following northward subduction of Tethys oceanic lithosphere under Iberia (Zeck, 1996) requires subduction prior to detachment and should involve typical subduction-zone volcanic activity before the slab detached. There is, however, no evidence for large-volume Late Oligocene–Early Miocene continental subduction-zone volcanism on the southern Iberian margin. Additionally, slab detachment should cause a rapid cessation of tholeiitic and calc-alkaline arc volcanism (Hoernle et al., 1999). As volcanism with a clear subduction-related geochemical fingerprint was active between 12·1 and 6·1 Ma in the Alboran Basin (Hoernle et al., 1999; Duggen et al., 2004), slab detachment in the Early Miocene seems to be very unlikely.
Models involving detachment or delamination of subcontinental lithosphere beneath the Alboran Basin at the Oligocene–Miocene boundary are unable to explain a number of key geological, geochemical and geophysical observations in the westernmost Mediterranean. These are: (1) the oceanic subduction zone geochemical characteristics of Middle to Late Miocene Alboran Sea lavas (Hoernle et al., 1999; Duggen et al., 2004; Gill et al., 2004); (2) the transition from thinned continental to Neogene oceanic crust in the easternmost Alboran Basin (Dewey et al., 1989; Comas et al., 1999) (Fig. 1); (3) directed rather than radial Miocene crustal nappe emplacement onto the southern Iberian and northwestern African continental margins associated with vertical axis-rotations of crustal nappes anti-clockwise in the Betics and clockwise in the Rif mountains (Lonergan & White, 1997); (4) an east-dipping positive seismic anomaly beneath the Alboran Basin that shows westward continuity with the Atlantic oceanic lithosphere (Gutscher et al., 2002; Gutscher et al., 2003); (5) a concentration of seismicity along a north–south-directed line at 60–120 km depth (as is commonly observed within subducting slabs where the slab curvature rapidly increases) beneath the westernmost Alboran Basin (Gutscher et al., 2002); (6) geophysical evidence for an active accretionary wedge due west of Gibraltar (Gutscher et al., 2002). These important observations can convincingly be explained with westward slab roll-back and steepening of subducted, east-dipping oceanic Tethys lithosphere since the Oligocene (Lonergan & White, 1997; Wilson & Bianchini, 1999; Gutscher et al., 2002; Duggen et al., 2004). A slab roll-back model for the Alboran Basin fits well into the widely accepted view that the western and central Mediterranean geodynamic evolution was dominated by slab roll-back of old subducting Tethys oceanic lithosphere (Dewey et al., 1989; Lonergan & White, 1997; Carminati et al., 1998; Wilson & Bianchini, 1999; Jolivet & Faccenna, 2000; Wortel & Spakman, 2000; Gutscher et al., 2002; Duggen et al., 2004). However, below we argue that the situation for the continental margins of southern Iberia and northwestern Africa is different.
The southern Iberian and northwestern African continental margins
During the Late Miocene to Pleistocene, igneous activity along the southern Iberian and the Moroccan part of the northwestern African continental margin was marked by the occurrence of Si–K-rich igneous rocks (8·2–4·8 Ma) with subduction-like geochemical signatures and Si-poor lavas (6·3–0·65 Ma) with intraplate-type signatures (Figs 1 and 2). Available K/Ar age data indicate that the onset of both Si–K-rich and Si-poor igneous activity was earlier in Algeria than in Spain and Morocco (Louni-Hacini et al., 1995; Coulon et al., 2002) (Fig. 2). As inferred from the major and trace element data, geochemical modelling and interpretation of experimental partial melting of spinel and garnet peridotites presented above (e.g. Figs 9 and 12), the parental Si–K-rich and Si-poor magmas are considered to be derived by partial melting of geochemically different mantle sources that existed largely at similar depths in a given area—metasomatized, subcontinental lithosphere and plume-contaminated sub-lithospheric mantle. New Ar/Ar age data suggest that the transition from Si–K-rich to Si-poor magmatism at the continental margins of Iberia and Africa (Fig. 8) reflects the progressive replacement of subcontinental lithosphere by sub-lithospheric mantle.
We propose that delamination of subcontinental lithosphere occurred beneath the continental margins of southern Iberia and northwestern Africa. It could be argued that slab detachment was responsible for the transition from Si–K-rich, subduction-related to Si-poor, intraplate-type igneous activity. Any geodynamic model, however, has to explain why the transition from Si–K-rich to Si-poor magmatism occurred both along the southern Iberian and northwestern African continental margins. If slab detachment occurred in the Alboran region, then double slab detachment, following north- and south-directed subduction beneath the Iberian and African margins, is required. There is no evidence in the available geophysical data that indicates such a process. Finally, as mentioned above, slab detachment is not consistent with recent results from seismic tomography and the observed temporal and spatial geochemical variation of the westernmost Mediterranean volcanism.
Detachment of subcontinental lithosphere has been proposed before based on geochemical and geochronological data for igneous rocks from southern Spain, by Turner et al. (1999). Those workers extended their detachment model to the entire Alboran region including the Alboran Basin area. As discussed above, this model is not supported by geophysical evidence or by the geochemical composition of Miocene volcanic rocks from the Alboran Basin area. Instead, we propose that the removal of the sub-continental lithosphere beneath the continental margins of southern Iberia and northwestern Africa was progressive (delamination) rather than convective (detachment).
Both delamination and detachment of subcontinental lithosphere could be associated with the generation of Si–K-rich, subduction-related igneous rocks and Si-poor, intraplate-type lavas. Delamination of subcontinental lithosphere in conjunction with subduction of oceanic lithosphere has been proposed for the petrogenesis of the Neogene to Quaternary K-rich magmatism of central Italy (Serri et al., 1993) and the association of Si–K-rich and Si-poor magmatism of the Andean Puna Altiplano Plateau (Kay & Mahlburg Kay, 1993). Detachment of subcontinental lithosphere following continental thickening is thought to be responsible for the generation of K-rich melts in the Tibetan Plateau (Turner et al., 1996) and the calc-alkaline to K-rich igneous rocks of Eastern Anatolia (Turkey) (Pearce et al., 1990). Based on mantle mineral stabilities, mantle solidi and geothermal gradients, Pearce et al. (1990) showed that continental thickening can also bring metasomatized subcontinental lithosphere to greater depths, causing partial melting.
Delamination of subcontinental mantle lithosphere beneath the continental margins of southern Iberia and northwestern Africa is also consistent with geophysical data such as the location of earthquake hypocentres and seismic-wave velocities pointing to the presence of a thin wedge-shaped layer (c. 20–40 km) of (possibly hotter) asthenospheric material between the Moho and a slab of lithospheric mantle down-bending beneath southern Spain and northern Morocco near Gibraltar in the western part of the Alboran Basin (Seber et al., 1996b). This process can explain why Si-poor, intraplate-type magmas were formed at largely similar depths in the same areas where Si–K-rich magmatism had ceased, as the upwelling sub-lithospheric mantle that replaced the delaminating subcontinental lithosphere at a given depth was probably the source for the primary Si-poor magmas. Such sub-lithospheric mantle could either be depleted convecting asthenosphere capable of producing MORB melts at sufficiently high degrees of melting or enriched mantle plume material. The major and trace element and Sr–Nd–Pb-isotope composition of most of the Si-poor lavas is consistent with derivation from a HIMU-type mantle source rather than depleted MORB-source mantle (Fig. 7b–d). The HIMU-type mantle could be plume material or could reflect concentrations of pyroxenitic and eclogitic material within the upper mantle. A few Si-poor samples, however, have 87Sr/86Sr, 143Nd/144Nd and 206Pb/204Pb similar to enriched MORB that might indicate a marginal involvement of a depleted MORB mantle (DMM) component (Fig. 7c and d), which, however, does not seem to be supported by the Pb isotope data (Fig. 7b). An overlap with enriched MORB composition in the 206Pb/204Pb vs 87Sr/86Sr and 143Nd/144Nd diagrams is more likely to result from interaction of sub-lithospheric melts with the subcontinental mantle or melts derived therefrom (indicated by arrow 2 in Fig. 7c and d). Involvement of a MORB source therefore appears to be unnecessary.
Low magma volumes and the absence of an age progression in the western Mediterranean Si-poor rocks (Fig. 1) and lack of evidence from seismic tomography for low-velocity cylindrical structures beneath the Alboran region (Seber et al., 1996b; Calvert et al., 2000; Gutscher et al., 2002) makes derivation of the Si-poor lavas directly from mantle plumes such as found below the Canary Islands unlikely. The similarity in geochemistry of the western Mediterranean Si-poor group lavas with intraplate-type volcanic rocks from Central Europe, the eastern North Atlantic and NW Africa provides evidence for derivation from a common source beneath these regions (Allègre et al., 1981; Hoernle et al., 1991, 1995; Wilson & Downes, 1991; Hoernle & Schmincke, 1993a, 1993b; Granet et al., 1995; Thirlwall, 1997; Geldmacher & Hoernle, 2000). Global seismic tomography data suggest that a large low-velocity structure extends from >2000 km depth to the base of the upper mantle (c. 660 km) and spreads out in the upper mantle beneath Europe and the Mediterranean (Hoernle et al., 1995; Goes et al., 1999) thereby probably contaminating the convecting asthenosphere with plume material. Subduction of oceanic lithosphere may trigger sub-lithospheric mantle material to well up around slab ‘graveyards’ localized at the base of the upper mantle or slab structures present in the upper mantle (Goes et al., 1999). We propose that further upwelling of such plume-contaminated, sub-lithospheric mantle material may be triggered by delaminating or detaching subcontinental mantle lithosphere. Melting of upwelling plume-contaminated asthenosphere replacing delaminated subcontinental lithosphere can explain the occurrence of Si-poor lavas on the continental margins of southern Iberia and northwestern Africa.
Reconciling the Miocene to Pliocene geodynamic evolution of the Alboran Basin with that of the continental margins of southern Iberia and northwestern Africa
In Figure 13, we present a three-dimensional geodynamic model that illustrates the connection between the changes in mantle geometry beneath the westernmost Mediterranean area and the transition from subduction-related to intraplate-type magmatism occurring at the Miocene–Pliocene boundary. As the Alboran Block was forced to move to the SW in the Miocene it overrode subducting east-dipping Tethys oceanic lithosphere (Gutscher et al., 2002; Duggen et al., 2004). Hydrous fluids and melts were released from the subducted oceanic lithosphere into the mantle wedge beneath the Alboran Basin, which consisted of asthenospheric mantle and attenuated continental lithosphere. These hydrous fluids–melts caused high degrees of partial melting in the mantle wedge, resulting in the eruption of low-K (tholeiitic) and medium-K (calc-alkaline) series lavas in Alboran Basin area. These Middle to Late Miocene lavas have typical subduction-related geochemical signatures similar to those from modern subduction zones (e.g. Izu–Bonin arc) (Hoernle et al., 1999; Duggen et al., 2004; Gill et al., 2004). Roll-back and steepening of the subducted Tethys oceanic lithosphere could have had several effects, as follows.
- It could cause peeling off (delamination) of bands of subcontinental mantle lithosphere beneath the southern Iberian and northwestern African continental margins (continental-edge delamination). We propose that the lithospheric bands were attached to the cold and dense subducting oceanic lithosphere through mechanical coupling, which would be an effective mechanism for the removal of even buoyant subcontinental lithosphere.
- It could cause asthenosphere contaminated with plume material to well up from depth around the margins of the delaminating subcontinental lithosphere.
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Perturbation of the metasomatically enriched subcontinental lithosphere resulted in the generation of Si–K-rich, subduction-type melts in the Late Miocene to Early Pliocene. Decompressional partial melting of the upwelling asthenosphere contaminated with plume material led to the formation of Si-poor, intraplate-type alkali basalts in the Late Miocene to Pleistocene. Interaction of Si-poor melts with the metasomatized subcontinental mantle or Si–K-rich melts derived from it could have generated magmas with hybrid geochemical characteristics between subduction- and intraplate-type. Westward slab roll-back and steepening coupled with delamination of subcontinental lithospheric mantle is consistent with the earlier onset of the subduction- to intraplate-type magmatic transition in Algeria (10·0 to 7·5 Ma) (Louni-Hacini et al., 1995
; Coulon et al., 2002) than in southeastern Iberia (6·4 to 2·9 Ma) and northern Morocco (6·8 to 4·8 Ma) (El Bakkali et al., 1998). Thermo-mechanical modelling shows that the delamination of bands of subcontinental lithosphere could also have caused large-scale uplift of the order of 1 km (continental-edge effect) sufficient to close the Late Miocene marine gateways at the southern Iberian and northwestern African continental margins, causing the Mediterranean to dry up in the Messinian (Duggen et al., 2003). |
CONCLUSIONS |
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TOPABSTRACTINTRODUCTIONREGIONAL GEOLOGYANALYTICAL METHODSRESULTSDISCUSSIONGEODYNAMIC EVOLUTIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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The transition from subduction-related, Si–K-rich to intraplate-type, Si-poor magmatism in the westernmost Mediterranean at the Miocene–Pliocene boundary can be observed both on a regional scale in the entire Alboran region and on a local scale in individual volcanic centres. Whereas the Alboran Basin area is marked by the eruption of low-K (tholeiitic) series and medium- to high-K calc-alkaline series volcanic rocks related to Middle to Late Miocene subduction (Duggen et al., 2004
), the southern Iberian and northwestern African continental margins are characterized by Upper Miocene to Lower Pliocene Si–K-rich, subduction-type and Upper Miocene to Pleistocene Si-poor, intraplate-type magmatism. The mafic Si–K-rich group magmas at the Iberian and African continental margins resulted from low degrees of partial melting of subcontinental mantle lithosphere metasomatically enriched by fluids or melts from earlier subduction processes. Mafic intraplate-type, Si-poor magmas were generated by low-degree decompressional melting of upwelling asthenospheric mantle contaminated with enriched plume material. The hybrid geochemical characteristics of the mafic magmas generated close to the Miocene–Pliocene boundary reflect interaction between subduction-type, Si–K-rich magmas or their source regions and intraplate-type, Si-poor magmas of sub-lithospheric origin.
At the Miocene–Pliocene boundary, the primary source of the magmatism shifted from subduction-modified subcontinental lithosphere to plume-contaminated sub-lithospheric sources not previously affected by subduction processes. The evolution of the Gourougou stratovolcanic complex in Morocco exhibits the progressively decreasing influence of the subduction-type component and increase of the intraplate-type component through time.
The transition from subduction-related to intraplate-type magmatism points to major changes in the geometry of the westernmost Mediterranean mantle. Roll-back and steepening of an east-dipping, subducting slab of old Tethys oceanic lithosphere in the Miocene triggered (1) the peeling off (delamination) of bands of the subcontinental lithosphere at the southern Iberian and northwestern African continental margins and (2) the upwelling of plume-contaminated sub-lithospheric material beneath southern Iberia and northwestern Africa, replacing the delaminating subcontinental lithospheric mantle. Continental-edge delamination associated with slab roll-back and steepening provides a plausible explanation for the complex spatial and temporal geochemical evolution of magmatism in the westernmost Mediterranean since the Oligocene, and is consistent with the Late Miocene uplift and closure of marine gateways connecting the Mediterranean Sea to the Atlantic Ocean, causing the desiccation of the Mediterranean and the Messinian Salinity Crisis.
Wider implications are that subcontinental lithosphere can effectively be delaminated when attached to dense subducting oceanic lithosphere. This process may be responsible for continental uplift and the generation of post-collisional magmatism in the Mediterranean area and elsewhere in the world.
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SUPPLEMENTARY DATA |
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TOPABSTRACTINTRODUCTIONREGIONAL GEOLOGYANALYTICAL METHODSRESULTSDISCUSSIONGEODYNAMIC EVOLUTIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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Supplementary data for this paper are available on Journal of Petrology online.
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ACKNOWLEDGEMENTS |
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We express our gratitude to K. Abdelmalki, M. Bouabdellah, Y. Bouabdellah, M. Chaieb, M. Jadid, A. Milhi and A. Moukadiri (from the Universities of Oujda and Fez and the Centre Régional Géologique d'Oujda) for their valuable help with our fieldwork in Morocco. We thank T. Arpe, D. Garbe-Schönberg, F. Hauff, L. Hoffmann, S. Krolikowska, S. Laukat, S. Plagmann, D. Rau, J. Sticklus, S. Vetter, K. Wolff and E. Zuleger for their help with the analytical work in Kiel. Martin Menzies and Matthew Thirlwall are gratefully acknowledged for comments on the manuscript prior to submission. We are grateful to José-Maria Cebriá, Szabolcs Harangi, Julian Pearce and Marjorie Wilson for their thorough reviews and valuable comments on the manuscript. The project was supported by the Deutsche Forschungsgemeinschaft (German Research Council) (HO1833/5).
* Corresponding author. Present address: Geological Institute, University of Copenhagen, Øster Voldgade 10, 1350 Copenhagen K, Denmark. E-mail: sduggen{at}geol.ku.dk
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REFERENCES |
|---|
|
TOPABSTRACTINTRODUCTIONREGIONAL GEOLOGYANALYTICAL METHODSRESULTSDISCUSSIONGEODYNAMIC EVOLUTIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
|---|
Ait-Brahim, L. & Chotin, P. (1990). Oriental Moroccan Neogene volcanism and strike-slip faulting. Journal of African Earth Sciences 11, 273–280.[CrossRef]
Allègre, C.-J., Dupre, B., Lambret, B. & Richard, P. (1981). The subcontinental vs suboceanic debate; I, Lead–neodymium–strontium isotopes in primary alkali basalts from a shield area; the Ahaggar volcanic suite. Earth and Planetary Science Letters 52, 85–92.[CrossRef][Web of Science]
Andries, D. & Bellon, H. (1989). Ages isotopiques 40K–40Ar du volcanisme alcalin néogène d'Oujda (Maroc oriental) et implications tectoniques. Sciences Géologiques, Mémoire 84, 107–116.
Arculus, R. J. & Gust, D. A. (1995). Regional petrology of the San Francisco Volcanic Field, Arizona, USA. Journal of Petrology 36, 827–861.
Bellon, H. & Brousse, R. (1977). Le magmatisme périméditerranén occidental. Essai de synthèse. Bulletin de la Société Géologique de France 7, 469–480.
Bellon, H. & Letouzey, J. (1977). Volcanism related to plate-tectonics in the western and eastern Mediterranean. In: Biju-Duval, B. & Montadert, L. (eds) Structural History of the Mediterranean Basins. Paris: Technip, pp. 165–184.
Bellon, H., Bordet, P. & Montenat, C. (1983). Chronologie du magmatisme néogène des Cordillères bètiques (Espagne méridionale). Bulletin de la Société Géologique de France 7, 205–217.
Benito, R., López-Ruiz, J., Cebriá, J. M., Hertogen, J., Doblas, M., Oyarzun, R. & Demaiffe, D. (1999). Sr and O isotope constraints on source and crustal contamination in the high-K calc-alkaline and shoshonitic Neogene volcanic rocks of SE Spain. Lithos 46, 733–802.
Bird, P. (1979). Continental delamination and the Colorado plateau. Journal of Geophysical Research 84, 7561–7571.
Blanco, M. J. & Spakman, W. (1993). The P-wave velocity structure of the mantle below the Iberian Peninsula: evidence for subducted lithosphere below southern Spain. Tectonophysics 221, 13–34.[CrossRef][Web of Science]
Calvert, A., Sanvol, E., Seber, D., Barazangi, M., Roecker, S., Mourabit, T., Vidal, F., Alguacil, G. & Jabour, N. (2000). Geodynamic evolution of the lithosphere and upper mantle beneath the Alboran region of the western Mediterranean: constraints from travel time tomography. Journal of Geophysical Research 105, 10871–10898.[CrossRef]
Capedri, S., Venturelli, G., Salvioli-Mariani, E., Crawford, A. J. & Barbieri, M. (1989). Upper-mantle xenoliths and megacrysts in an alkali basalt from Tallante, south-eastern Spain. European Journal of Mineralogy 1, 685–699.
Carminati, E., Wortel, M. J. R., Spakman, W. & Sabadini, R. (1998). The role of slab detachment processes in the opening of the western–central Mediterranean basins: some geological and geophysical evidence. Earth and Planetary Science Letters 160, 651–665.[CrossRef][Web of Science]
Cebriá, J.-M. & López-Ruiz, J. (1995). Alkali basalts and leucitites in an extensional intracontinental plate setting: the late Cenozoic Calatrava Volcanic Province (central Spain). Lithos 35, 27–46.[CrossRef][Web of Science]
Cesare, B. & Maineri, C. (1999). Fluid-present anatexis of metapelites at El Joyazo (SE Spain): constraints from Raman spectroscopy of graphite. Contributions to Mineralogy and Petrology 135, 41–52.[CrossRef][Web of Science]
Cesare, B., Salvioli Mariani, E. & Venturelli, G. (1997). Crustal anatexis and melt extraction during deformation in the restitic xenoliths at El Joyazo (SE Spain). Mineralogical Magazine 61, 15–27.[Abstract]
Chaffey, D. J., Cliff, R. A. & Wilson, B. M. (1989). Characterization of the St Helena magma source. In: Saunders, A. D. & Norry, M. J. (eds) Magmatism in the Ocean Basins. Geological Society, London, Publications 42, 257–276.
Comas, M. C., Platt, J. P., Soto, J. I. & Watts, A. B. (1999). The origin and tectonic history of the Alboran Basin: insights from Leg 161 results. In: Zahn, R., Comas, M. C. & Klaus, A. (eds) Proceedings of the Ocean Drilling Program, Scientific Results, 161. College Station, TX: Ocean Drilling Program, pp. 555–580.
Coppier, G., Griveaud, P., de Larouziere, F. D., Montenat, C. & d'Estevou, O. (1989). Example of Neogene tectonic indentation in the Eastern Betic Cordilleras: the Arc of Aguilas (Southeastern Spain). Geodinamica Acta 3, 37–51.[Web of Science]
Coulon, C., Megartsi, M., Fourcade, S., Maury, R. C., Bellon, H., Louni-Hacini, A., Cotten, J., Coutelle, A. & Hermitte, D. (2002). Post-collisional transition from calc-alkaline to alkaline volcanism during the Neogene in Oranie (Algeria): magmatic expression of a slab breakoff. Lithos 62, 87–110.[CrossRef][Web of Science]
Dalrymple, G. B. & Duffield, W. A. (1988). High precision 40Ar/39Ar dating of Oligocene tephra from the Mogollon–Datil volcanic field using a continuous laser system. Geophysical Research Letters 15, 463–466.[Web of Science]
Davies, J. H. & Blanckenburg, F. v. (1995). Slab breakoff: a model of lithosphere detachment and its test in the magmatism and deformation of collisional orogens. Earth and Planetary Science Letters 129, 85–102.[CrossRef][Web of Science]
Deer, W. A., Howie, R. A. & Zussman, J. (1996). An Introduction to the Rock-forming Minerals. Harlow: Addison Wesley Longman, 549 pp.
Dewey, J. F., Helman, M. L., Hutton, D. H. W. & Knott, S. D. (1989). Kinematics of the western Mediterranean. In: Coward, M. P. & Park, R. G. (eds) Alpine Tectonics. Geological Society, London, Special Publications 45, 265–283.
Docherty, C. & Banda, E. (1995). Evidence for the eastward migration of the Alboran Sea based on regional subsidence analysis: a case for basin formation by delamination of the subcrustal lithosphere? Tectonics 14, 804–818.[CrossRef][Web of Science]
Dosso, L., Hanan, B. B., Bougault, H., Schilling, J. G. & Joron, J. L. (1991). Sr–Nd–Pb geochemical morphology between 10 degrees and 17 degrees N on the Mid-Atlantic Ridge: a new MORB isotope signature. Earth and Planetary Science Letters 106, 29–43.[CrossRef][Web of Science]
Duffield, W. A. & Dalrymple, G. B. (1990). The Taylor Creek Rhyolite of New Mexico: a rapidly emplaced field of lava domes and flows. Bulletin of Volcanology 52, 475–487.[CrossRef][Web of Science]
Duggen, S. (2002). Spatial and temporal geochemical evolution of igneous rocks in the Alborán Region (westernmost Mediterreanean): implications for the origin of Mediterranean-style back-arc basins and the Messinian Salinity Crisis. Ph.D. thesis, University of Kiel.
Duggen, S., Hoernle, K., van den Bogaard, P., Rüpke, L. & Phipps Morgan, J. (2003). Deep roots of the Messinian salinity crisis. Nature 422, 602–606.[CrossRef][Medline]
Duggen, S., Hoernle, K., van den Bogaard, P. & Harris, C. (2004). Magmatic evolution of the Alboran Region: the role of subduction in forming the western Mediterranean and causing the Messinian Salinity Crisis. Earth and Planetary Science Letters 218, 91–108.[CrossRef][Web of Science]
Dupuy, C., Dostal, J. & Boivin, P. A. (1986). Geochemistry of ultramafic xenoliths and their host alkali basalts from Tallante, southern Spain. Mineralogical Magazine 50, 231–239.[CrossRef][Web of Science]
Ellam, R. M., Menzies, M. A., Hawkesworth, C. J., Leeman, W. P., Rosi, M. & Serri, G. (1988). The transition from calc-alkaline to potassic orogenic magmatism in the Aeolian Islands, Southern Italy. Bulletin of Volcanology 50, 386–398.[CrossRef][Web of Science]
Ellam, R. M., Hawkesworth, C. J., Menzies, M. A. & Rogers, N. W. (1989). The volcanism of Southern Italy: role of subduction and the relationship between potassic and sodic alkaline magmatism. Journal of Geophysical Research 94, 4589–4601.
El Bakkali, S., Gourgaud, A., Bourdier, J.-L., Bellon, H. & Gundogdu, N. (1998). Post-collision Neogene volcanism of the Eastern Rif (Morocco): magmatic evolution through time. Lithos 45, 523–543.[CrossRef][Web of Science]
Esteban, M., Braga, J. C., Martín, J. & de Santisteban, C. (1996). Western Mediterranean Reef Complexes. In: Franseen, E. K., Esteban, M., Ward, W. C. & Rouchy, J.-M. (eds) Models for Carbonate Stratigraphy. From Miocene Reef Complexes of Mediterranean Regions. Tulsa, OK: SEPM (Society for Sedimentary Geology), pp. 55–72.
Foley, S. (1992). Vein-plus-wall-rock melting mechanisms in the lithosphere and the origin of potassic alkaline magmas. Lithos 28, 435–453.[CrossRef][Web of Science]
Foley, S. F. (1993). An experimental study of olivine lamproite: first results from the diamond stability field. Geochimica et Cosmochimica Acta 57, 483–489.[CrossRef][Web of Science]
Garbe-Schönberg, C. D. (1993). Simultaneous determination of thirty-seven trace elements in twenty-eight international rock standards by ICP-MS. Geostandards Newsletter 17, 81–97.[Web of Science]
Geldmacher, J. & Hoernle, K. (2000). The 72 Ma geochemical evolution of the Madeira hotspot (eastern North Atlantic): recycling of Paleozoic (
500 Ma) oceanic lithosphere. Earth and Planetary Science Letters 183, 73–92.[CrossRef][Web of Science]
Geldmacher, J., van den Bogaard, P., Hoernle, K. & Schmincke, H.-U. (2000). The 40Ar/39Ar age dating of the Madeira Archipelago and hotspot track (eastern North Atlantic). Geochemistry, Geophysics, Geosystems 1, paper number 1999GC000018.
Gill, R. C. O., El Aparicio, A., El Azzouzi, M., Hernandez, J., Thirlwall, M. F., Bourgois, J. & Marriner, G. F. (2004). Depleted arc volcanism in the Alboran Sea and shoshonitic volcanism in Morocco: geochemical and isotopic constraints on Neogene tectonic processes. Lithos 78, 363–388.[CrossRef][Web of Science]
Gîrbacea, R. & Frisch, W. (1998). Slab in the wrong place: lower lithospheric mantle delamination in the last stage of the Eastern Carpathian subduction retreat. Geology 26, 611–614.
Goes, S., Spakman, W. & Bijward, H. (1999). A lower mantle source for Central European volcanism. Science 286, 1928–1931.
Granet, M., Wilson, M. & Achauer, U. (1995). Imaging a mantle plume beneath the French Massif Central. Earth and Planetary Science Letters 136, 281–296.[CrossRef][Web of Science]
Green, D. H. (1973). Experimental melting studies on a model upper mantle composition at high pressure under water-saturated and water-undersaturated conditions. Earth and Planetary Science Letters 19, 37–53.[CrossRef][Web of Science]
Gutscher, M.-A., Malod, J., Rehault, J.-P., Contrucci, I., Klingelhoefer, F., Mendes-Victor, L. & Spakman, W. (2002). Evidence for active subduction beneath Gibraltar. Geology 30, 1071–1074.
Gutscher, M.-A., Malod, J., Raehault, J.-P., Construcci, I., Klinelhoefer, F., Mendes-Victor, L. & Spakman, W. (2003). Evidence for active subduction beneath Gibraltar: Reply. Geology: Online Forum, e23.
Harmand, C. & Cantagrel, J. M. (1984). Le volcanisme alcalin tertiaire et quaternaire du Moyen Atlas (Maroc): chronologie K/Ar et cadre géodynamique. Journal of African Earth Sciences 2, 51–55.[CrossRef]
Hart, S. R. (1984). A large-scale isotope anomaly in the Southern Hemisphere mantle. Nature 309, 753–757.[CrossRef]
Hawkesworth, C. J., Gallagher, K., Hergt, J. M. & McDermott, F. (1993). Mantle and slab contributions in arc magmas. Annual Review of Earth and Planetary Sciences 21, 175–204.[Web of Science]
Hernandez, J. (1986). Pétrologie du massif volcanique du Guilliz (Maroc oriental). Cristallisation fractionée, mélanges de magmas et transferts de fluides dans une série shoshonitique. Journal of African Earth Sciences 5, 381–399.[CrossRef]
Hernandez, J. & Bellon, H. (1985). Chronologie K–Ar du volcanisme miocéne du Rif oriental (Maroc): implications tectoniques et magmatique. Revue de Géologie et de Géographie Physique 26, 85–94.
Hirose, K. & Kushiro, I. (1993). Partial melting of dry peridotites at high pressures: determination of compositions of melts segregated from peridotite using aggregates of diamond. Earth and Planetary Science Letters 114, 477–489.[CrossRef][Web of Science]
Hoernle, K. (1998). Geochemistry of Jurassic oceanic crust beneath Gran Canaria (Canary): implications for crustal recycling and assimilation. Journal of Petrology 39, 859–880.[CrossRef][Web of Science]
Hoernle, K. & Schmincke, H. U. (1993a). The petrology of the tholeiites through melilite nephelinites on Gran Canaria, Canary Islands: crystal fractionation, accumulation, and depths of melting. Journal of Petrology 34, 573–597.
Hoernle, K. & Schmincke, H. U. (1993b). The role of partial melting in the 15-Ma geochemical evolution of Gran Canaria: a blob model for the Canary hotspot. Journal of Petrology 34, 599–626.
Hoernle, K. A. & Tilton, G. R. (1991). Sr–Nd–Pb isotope data for Fuerteventura (Canary Islands) basal complex and subaerial volcanics: applications to magma genesis and evolution. Schweizerische Mineralogische und Petrographische Mitteilungen 71, 3–18.
Hoernle, K., Tilton, G. & Schmincke, H. U. (1991). Sr–Nd–Pb isotopic evolution of Gran Canaria: evidence for shallow enriched mantle beneath the Canary Islands. Earth and Planetary Science Letters 106, 44–63.[CrossRef][Web of Science]
Hoernle, K., Zhang, Y. S. & Graham, D. (1995). Seismic and geochemical evidence for large-scale mantle upwelling beneath the eastern Atlantic and western central Europe. Nature 374, 34–39.[CrossRef]
Hoernle, K., van den Bogaard, P., Duggen, S., Mocek, B. & Garbe-Schönberg, D. (1999). Evidence for Miocene subduction beneath the Alboran Sea: 40Ar/39Ar dating and geochemistry of volcanic rocks from Holes 977A and 978A. In: Zahn, R., Comas, M. C. & Klaus, A. (eds) Proceedings of the Ocean Drilling Program, Scientific Results, 161. College Station, TX: Ocean Drilling Program, pp. 357–373.
Hoernle, K., Duggen, S., Geldmacher, J., Klügel, A., et al. (2003). METEOR Cruise No. 51, Leg 1, Vulkosa: Vulkanismus Ostatlantik–Alboran in Meteor Berichte 03-1 Ostatlantik–Mittelmeer–Schwarzes Meer, Cruise No. 51, 12 September to 28 December 2001. Kiel: GEOMAR Research Centre.
Hofmann, A. W. (1988). Chemical differentiation of the Earth. The relationship between mantle, continental crust, and oceanic crust. Earth and Planetary Science Letters 90, 297–314.[CrossRef][Web of Science]
Hofmann, A. W. (1997). Mantle geochemistry: the message from oceanic volcanism. Nature 385, 219–229.[CrossRef]
Horvath, F. & Berckhemer, H. (1982). Mediterranean backarc basins. In: Berckhemer, H. & Hsü, K. (eds) Alpine–Mediterranean Geodynamics. Washington, DC: American Geophysical Union; Boulder, CO: Geological Society of America, pp. 141–173.
Houseman, G. (1996). From mountains to basins. Nature 379, 771–772.[CrossRef]
Hsü, K. J., Ryan, W. B. F. & Cita, M. B. (1973). Late Miocene desiccation of the Mediterranean. Nature 242, 240–244.[CrossRef]
Innocenti, F., Mazzuoli, R., Pasquare, G., Radicati di Brozolo, F. & Villari, L. (1982). Tertiary and Quaternary volcanism of the Erzurum–Kars area (Eastern Turkey): geochronological data and geodynamic evolution. Journal of Volcanology and Geothermal Research 13, 223–240.[CrossRef][Web of Science]
Ito, E., White, W. M. & Gopel, C. (1987). The O, Sr, Nd and Pb isotope geochemistry of MORB. Chemical Geology 62, 157–176.[CrossRef][Web of Science]
Jadid, M., Wernli, R. & Rezqi, H. (1999). Biostratigraphie et paléogéographie du Mio-Pliocène de la region d'Oujda (Maroc nord-oriental). Revue de Paléobiologie, Genève 18, 253–258.
Jolivet, L. & Faccenna, C. (2000). Mediterranean extension and the African–Eurasia collision. Tectonics 19, 1095–1106.[CrossRef][Web of Science]
Kay, R. W. & Mahlburg Kay, S. (1993). Delamination and delamination magmatism. Tectonophysics 219, 177–189.[CrossRef][Web of Science]
Keller, J. (1982). Mediterranean island arcs. In: Thorpe, R. S. (ed.) Andesites: Orogenic Andesites and Related Rocks. Chichester: John Wiley, pp. 307–325.
Krijgsman, W., Hilgen, F. J., Raffi, I., Sierro, F. J. & Wilson, D. S. (1999). Chronology, causes and progression of the Messinian salinity crisis. Nature 400, 652–655.[CrossRef]
Lonergan, L. & White, N. (1997). Origin of the Betic–Rif mountain belt. Tectonics 16, 504–522.[CrossRef][Web of Science]
López Ruiz, J. & Rodríguez Badiola, E. (1980). La region volcanica neogena del sureste de España. Estudios Geologicos 36, 5–63.
López-Ruiz, J., Cebria, J. M. & Doblas, M. (2002). Cenozoic magmatism I: The Iberian peninsula. In: Gibbons, W. & Moreno, M. T. (eds) The Geology of Spain. London: Geological Society, pp. 417–438.
Louni-Hacini, A., Bellon, H., Maury, R. C., Megartsi, M., Coulon, C., Semroud, B., Cotten, J. & Coutelle, A. (1995). Datation 40K–40Ar de la transition du volcanisme calco-alcalin au volcanisme alcalin en Oranie au Miocène supérieur. Comptes Rendus de l'Académie des Sciences 321, 975–982.
Montenat, C. & Ott d'Estevou, P. (1995). Late Neogene basins evolving in the Eastern Betic transcurrent fault zone: an illustrated review. In: Friend, P. F. & Dabrio, C. J. (eds) Tertiary Basins of Spain. Cambridge: Cambridge University Press, pp. 372–386.
Munksgaard, N. C. (1984). High 18O and possible pre-eruptional Rb–Sr isochrons in cordierite-bearing Neogene volcanics from SE Spain. Contributions to Mineralogy and Petrology 87, 351–358.[CrossRef][Web of Science]
Nelson, D. R. (1992). Isotopic characteristics of potassic rocks: evidence for the involvement of subducted sediments in magma genesis. Lithos 28, 403–420.[CrossRef][Web of Science]
Nelson, D. R., McCulloch, M. T. & Sun, S. S. (1986). The origins of ultrapotassic rocks as inferred from Sr, Nd and Pb isotopes. Geochimica et Cosmochimica Acta 50, 231–245.[CrossRef][Web of Science]
Nobel, F. A., Andriessen, P. A. M., Hebeda, E. H., Priem, H. N. A. & Rondeel, H. E. (1981). Isotopic dating of the post-alpine Neogene volcanism in the Betic Cordilleras, Southern Spain. Geologie en Mijnbouw 60, 209–214.
Pearce, J. A. & Peate, D. W. (1995). Tectonic implications of the composition of volcanic arc magmas. Annual Review of Earth and Planetary Sciences 23, 251–285.[CrossRef][Web of Science]
Pearce, J. A., Bender, J. F., De Long, S. E., Kidd, W. S. F., Low, P. J., Güner, Y., Saroglu, F., Yilmaz, Y., Moorbath, S. & Mitchell, J. G. (1990). Genesis of collision volcanism in Eastern Anatolia, Turkey. Journal of Volcanology and Geothermal Research 44, 189–229.[CrossRef][Web of Science]
Platt, J. P. & England, P. C. (1993). Convective removal of lithosphere beneath mountain belts: thermal and mechanical consequences. American Journal of Science 293, 307–336.
Platt, J. P. & Vissers, R. L. M. (1989). Extensional collapse of thickened continental lithosphere: a working hypothesis for the Alboran Sea and Gibraltar Arc. Geology 17, 540–543.
Platt, J. P., Whitehouse, M. J., Kelley, S. P., Carter, A. & Hollick, L. (2003). Simultaneous extensional exhumation across the Alboran Basin: implications for the causes of late orogenic extension. Geology 31, 251–254.
Rickwood, P. C. (1989). Boundary lines within petrological diagrams which use oxides of major and minor elements. Lithos 22, 247–263.[CrossRef][Web of Science]
Roger, S., Münch, P., Cornee, J. J., Saint Martin, J. P., Feraud, G., Pestrea, S., Conesa, G. & Ben Moussa, A. (2000). 40Ar/39Ar dating of the pre-evaporitic Messinian marine sequences of the Melilla basin (Morocco): a proposal for some biosedimentary events as isochrons around the Alboran Sea. Earth and Planetary Science Letters 179, 101–113.[CrossRef][Web of Science]
Royden, L. H. (1993). Evolution of retreating subduction boundaries formed during continental collision. Tectonics 12, 629–638.[Web of Science]
Saadi, M., Hilali, E. A., Bensaid, M., Boudda, A. & Dahmani, M. (1985). Carte géologique du Maroc. Paris: Ministère de l'énergie et des mines, direction de la géologie.
Seber, D., Barazangi, M., Ibenbrahim, A. & Demnati, A. (1996a). Geophysical evidence for lithospheric delamination beneath the Alboran Sea and Rif–Betic mountains. Nature 379, 785–790.[CrossRef]
Seber, D., Barazangi, M., Tadili, B. A., Ramdani, M., Ibenbrahim, A. & Sari, D. B. (1996b). Three-dimensional upper mantle structure beneath the intraplate Atlas and interplate Rif mountains of Morocco. Journal of Geophysical Research 101, 3125–3138.[CrossRef]
Serri, G., Innocenti, F. & Manetti, P. (1993). Geochemical and petrological evidence of the subduction of delaminated Adriatic continental lithosphere in the genesis of the Neogene–Quaternary magmatism of central Italy. Tectonophysics 223, 117–147.[CrossRef][Web of Science]
Sun, S.-s. & McDonough, W. F. (1989). Chemical and isotopic systematics of oceanic basalts: implications for mantle composition and processes. In: Saunders, A. D. & Norry, M. J. (eds) Magmatism in the Ocean Basins. Geological Society, London, Special Publications 42, 313–345.
Taylor, R. N. & Nesbitt, R. W. (1998). Isotopic characteristics of subduction fluids in an intra-oceanic setting, Izu–Bonin Arc, Japan. Earth and Planetary Science Letters 164, 79–98.[CrossRef][Web of Science]
Taylor, S. R. & McLennan, S. M. (1985). The Continental Crust: its Composition and Evolution. Oxford: Blackwell Scientific, 312 pp.
Thirlwall, M. F. (1997). Pb isotopic and elemental evidence for OIB derivation from young HIMU mantle. Chemical Geology 139, 51–74.[CrossRef][Web of Science]
Tisserant, D., Laouina, A. & Clauer, N. (1985). Datation K–Ar des basaltes de la plaine des Angad au sud d'Oujda (Maroc-oriental). Bulletin de la Société Géologique de France 38, 141–145.
Toscani, L., Contini, S. & Ferrarini, M. (1995). Lamproitic rocks from Cabezo Negro de Zeneta: brown micas as a record of magma mixing. Mineralogy and Petrology 55, 281–292.[CrossRef][Web of Science]
Turner, S., Hawkesworth, C., Liu, J., Rogers, N., Kelley, S. & van Calsteren, P. (1993). Timing of Tibetan uplift constrained by analysis of volcanic rocks. Nature 364, 50–54.[CrossRef]
Turner, S., Arnaud, N., Liu, J., Rogers, N., Hawkesworth, C., Harris, N., Kelley, S., van Calsteren, P. & Deng, W. (1996). Post-collision, shoshonitic volcanism on the Tibetan Plateau: implications for convective thinning of the lithosphere and the source of ocean island basalts. Journal of Petrology 37, 45–71.
Turner, S. P., Platt, J. P., George, R. M. M., Kelley, S. P., Pearson, D. G. & Nowell, G. M. (1999). Magmatism associated with orogenic collapse of the Betic–Alboran Domain, SE Spain. Journal of Petrology 40, 1011–1036.[CrossRef][Web of Science]
Van den Bogaard, P. (1995). 40Ar/39Ar ages of sanidine phenocrysts from Laacher See tephra (12,900 yr BP): chronostratigraphic and petrological significance. Earth and Planetary Science Letters 133, 163–174.[CrossRef][Web of Science]
Venturelli, G., Capedri, S., Di Battistini, G., Crawford, A., Kogarko, L. N. & Celestini, S. (1984). The ultrapotassic rocks from southeastern Spain. Lithos 17, 37–54.[CrossRef][Web of Science]
Venturelli, G., Salvioli-Mariani, E., Foley, S. F., Capedri, S. & Crawford, A. J. (1988). Petrogenesis and conditions of crystallization of Spanish lamproitic rocks. Canadian Mineralogist 26, 67–79.[Web of Science]
Vielzeuf, D. (1983). The spinel and quartz associations in high grade xenoliths from Tallante (S.E. Spain) an their potential use in geothrmometry and barometry. Contributions to Mineralogy and Petrology 82, 301–311.[CrossRef][Web of Science]
Weijermars, R. (1985). Uplift and subsidence history of the Alboran Basin and a profile of the Alboran diapir (W-Mediterranean). Geologie en Mijnbouw 64, 349–356.
Wilson, M. & Bianchini, G. (1999). Tertiary–Quaternary magmatism within the Mediterranean and surrounding regions. In: Durand, B., Jolivet, L., Horváth, F. & Sérrane, M. (eds) The Mediterranean Basins: Tertiary Extension within the Alpine Orogen. Geological Society, London, Special Publications 156, 141–168.
Wilson, M. & Downes, H. (1991). Tertiary–Quaternary extension-related alkaline magmatism in Western and Central Europe. Journal of Petrology 32, 811–849.
Wortel, M. J. R. & Spakman, W. (2000). Subduction and slab detachment in the Mediterranean–Carpathian region. Science 290, 1910–1917.
Wyllie, P. J. & Wolf, M. B. (1993). Amphibolite dehydration-melting: sorting out the solidus. In: Alabaster, P. H. M., Harris, N. B. W. & Neary, C. R. (eds) Magmatic Processes and Plate Tectonics. Geological Society, London, Special Publications 76, 405–416.
Zeck, H. P. (1970). An erupted migmatite from Cerro del Hoyazo, SE Spain. Contributions to Mineralogy and Petrology 26, 225–246.[Web of Science]
Zeck, H. P. (1996). Betic–Rif orogeny: subduction of Mesozoic Tethys lithosphere under eastward drifting Iberia, slab detachment shortly before 22 Ma, and subsequent uplift and extensional tectonics. Tectonophysics 254, 1–16.[CrossRef][Web of Science]
Zeck, H. P. (1999). Alpine plate kinematics in the western Mediterranean: a westward-directed subduction regime followed by slab roll-back and slab detachment. In: Durand, B., Jolivet, L., Horváth, F. & Sérrane, M. (eds) The Mediterranean Basins: Tertiary Extension within the Alpine Orogen. Geological Society, London, Special Publications 156, 109–120.
Zeck, H. P., Kristensen, A. B. & Williams, I. S. (1998). Post-collisional volcanism in a sinking slab setting—crustal anatectic origin of pyroxene-andesite magma, Caldear volcanic province, southeastern Spain. Lithos 45, 499–522.[CrossRef][Web of Science]
Zeck, H. P., Kristensen, A. B. & Nakamura, E. (1999). Inherited Palaeozoic and Mesozoic Rb–Sr isotopic signatures in Neogene calc-alkaline volcanics, Alboran Volcanic Province, SE Spain. Journal of Petrology 40, 511–524.[CrossRef][Web of Science]
Zindler, A. & Hart, S. (1986). Chemical geodynamics. Annual Review of Earth and Planetary Sciences 14, 493–571.[CrossRef][Web of Science]
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