Evolution of the Sardinia Channel (Western Mediterranean): new constraints from a diving survey on Cornacya seamount off SE Sardinia

Marine Geology 179 (2001) 179±202

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Evolution of the Sardinia Channel (Western Mediterranean): new constraints from a diving survey on Cornacya seamount off SE Sardinia Georges H. Mascle a,*, Pierre Tricart a, Luigi Torelli b, Jean-Pierre Bouillin a, Franco Rolfo c, Henriette Lapierre a, Patrick Monie d, Stephane Depardon a, Jean Mascle e, Davide Peis b a

Laboratoire de GeÂodynamique des ChaõÃnes Alpines UMR-CNRS 5025/UJF Grenoble/U Savoie ChambeÂry, OSUG, BP 53, 38031 Grenoble Cedex, France b Dipartimento Scienze della Terra, UniversitaÁ di Parma, Parco Area delle Scienze 158, 43100 Parma, Italy c Dipartimento di Scienze Mineralogiche e Petrologiche, Via Valperga Caluso 35, 10125 Torino, Italy d Institut des Sciences de la Terre, de l'Eau et de l'Espace, Place E. Bataillon, 34095 Montpellier Cedex, France e Geosciences Azur, BP 48, 06230 Villefranche sur Mer, France Received 8 March 2000; accepted 20 February 2001

Abstract Sarcya 1 dive explored a previously unknown 12 My old submerged volcano, labelled Cornacya. A well developed fracturation is characterised by the following directions: N 170 to N±S, N 20 to N 40, N 90 to N 120, N 50 to N 70, which corresponds to the fracturation pattern of the Sardinian margin. The sampled lavas exhibit features of shoshonitic suites of intermediate composition and include amphibole-and micabearing lamprophyric xenoliths which are geochemically similar to Ti-poor lamproites. Mica compositions re¯ect chemical exchanges between the lamprophyre and its shoshonitic host rock suggesting their simultaneous emplacement. Nd compositions of the Cornacya K-rich suite indicate that continental crust was largely involved in the genesis of these rocks. The spatial association of the lamprophyre with the shoshonitic rocks is geochemically similar to K-rich and TiO2-poor igneous suites, emplaced in post-collisional settings. Among shoshonitic rocks, sample SAR 1-01 has been dated at 12:6 ^ 0:3 My using the 40Ar/ 39Ar method with a laser microprobe on single grains. The age of the Cornacya shoshonitic suite is similar to that of the Sisco lamprophyre from Corsica, which similarly is located on the western margin of the Tyrrhenian Sea. Thus, the Cornacya shoshonitic rocks and their lamprophyric xenolith and the Sisco lamprophyre could represent post-collisional suites emplaced during the lithospheric extension of the Corsica±Sardinia block, just after its rotation and before the Tyrrhenian sea opening. Drilling on the Sardinia margin (ODP Leg 107) shows that the upper levels of the present day margin (Hole 654) suffered tectonic subsidence before the lower part (Hole 652). The structure of this lower part is interpreted as the result of an eastward migration of the extension during Late Miocene and Early Pliocene times. Data of Cornacya volcano are in good agreement with this model and provide good chronological constraints for the beginning of the phenomenon. q 2001 Elsevier Science B.V. All rights reserved. Keywords: Western Mediterranean; Sardinia Channel; Structure; Shoshonites; Miocene post-collisional evolution

* Corresponding author. E-mail address: [email protected] (G.H. Mascle). 0025-3227/01/$ - see front matter q 2001 Elsevier Science B.V. All rights reserved. PII: S 0025-322 7(01)00220-1

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Fig. 1. Sardinia Channel: general bathmetry and position of the diving sites Sarcya (SAR) and Sartucya (STC). Bathymetric contours after Gennesseaux and Stanley, 1983.

1. Introduction The Sardinia Channel is located on a segment of the Apenninic±Maghrebides collision chain. Although this segment is nowadays completely submerged, it has not been broken apart by the opening of the Algerian-Provence and Tyrrhenian oceanic basins. Earlier seismic data have shown that the structure of the Sardinia Channel is due to the superposition of two successive regimes of deformation: a compressional event with crustal thickening followed by extension

and thinning out. The morphology and structural patterns of the Sardinia Channel are still well preserved because the post-orogenic extension was moderate and submergence prevailed important erosions. Thus, the Sardinia Channel represents a key area for the study of the tectonic evolution of a collisional chain. Moreover, the use of the easy handling CYANA submersible allows a detailed structural and morphological survey, and the collection of well located, and sometimes oriented samples. In this paper, we present the results of a structural

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Fig. 2. The Tyrrhenian Sardinia margin at the latitude of SAR 1 dive. Seismic pro®le after Fabretti et al., 1995. LC: Lower Crust; PRU: Post Rift Unconformity.

study of a submerged volcano, discovered during the dive and named Cornacya, and the petrology and geochemistry of the lavas of this volcano, sampled during dive Sarcya 1. The data obtained from this study will help us to precisely understand the geodynamic environment of this part of the southern Tyrrhenian Sea and Sardinia Channel. This paper is mainly devoted to the volcanic rocks 2. Geological setting The Sardinia Channel (Fig. 1) is located on a 400 km long submerged segment of the Apenninic± Maghrebides Alpine collision Chain, where the Algero-ProvencËal oceanic basin joins the Tyrrhenian Sea (Caire, 1973; Amodio-Morelli et al., 1976; Grandjacquet and Mascle, 1978; Bouillin, 1984, 1986). Since the pioneer work of Castany (1955, 1959), the area has been surveyed at different scales by different authors (i.e. Auzende et al., 1974; Colantoni et al., 1981; Mauffret et al., 1981; Catalano et al., 1985; see Mascle and Tricart, 1994; 2001, for a complete bibliography). The structure of the area, evidenced by seismic pro®les (Fig. 2) (Morelli and

Nicolich, 1990; Blundell et al., 1992; Egger, 1992; Giese et al., 1992; Torelli et al., 1992; Tricart and Torelli, 1994), is characterised by the superposition of two successive regimes of deformation. The oldest one corresponds to a period of shortening and crustal thickening, which resulted in the formation of the Apenninic±Maghrebide range during Oligocene± Early Miocene times, contemporaneously to the Corsica±Sardinia rotation and the opening of the Algero-ProvencËal oceanic basin, as a consequence of the Tethyan subduction. During this event, metamorphic recrystallisation occurred, which is documented by 39Ar/ 40Ar datings (Bouillin et al., 1999). A calc-alkaline magmatism, well developed in Sardinia but also present on the Tunisian margin (Galite archipelago and some small outcrops in Northern Tunisia), is representative of the subduction (Fig. 3). The youngest one is associated with moderate extension and crustal thinning related to the Tyrrhenian rifting, which occurred before the Messinian crisis (Bouillin et al., 1998). The latter caused an important erosion in the whole area. It was followed by a period of thermal subsidence in correspondence with the opening of the Tyrrhenian oceanic realm. However, careful survey of the Pliocene-Quaternary sequence show that repeated minor phases of inversion are

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Fig. 3. Oligocene to present day magmatism around the Tyrrhenian basin. After Savelli (1988).

documented in the area, particularly well exposed on the fault planes (Torelli et al., 1992; Tricart and Torelli, 1994; Depardon, 1995; Brocard, 2001). The Sardinia Channel is characterised by a dissymetric WSW±ENE trending depression, the deepest part of which is more than 2000 m deep. This trough, which has been diversely labelled Teulada Valley, Sardinia Valley, Tuniso±Sardinia basin, Deep Sardinia Channel, allows the connection between the North Algerian (Algero-ProvencËal basin, 2600 m) and SW Tyrrhenian (2800 m depth) abyssal plains. The trough has been interpreted as an episutural basin, correlated with the inversion of the Alpine

thrusting of the Sardinian basement upon the CPK units (Tricart et al., 1990, 1991; Torelli et al., 1992; Tricart and Torelli, 1994). The NW Sardinia Channel shows a steep morphology due to the presence of tilted blocks located on the N±S to NNE±SSW striking margin of south-eastern Sardinia. The largest tilted block, named Monte Ichnusa, is less than 200 m deep and its basement belongs to that of southern Sardinia and consists of Paleozoic metasediments and granitoids, covered by post-Paleozoic sedimentary and volcanic rocks (Compagnoni et al., 1987, 1989). According to the Geotraverse pro®le (Morelli and

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Fig. 4. SAR 1 dive course: map and pro®le (geological data have been projected following the 310 average direction).

Nicolich, 1990; Blundell et al., 1992; Egger, 1992; Giese et al., 1992; Research Group for Lithospheric Structure in Tunisia, 1992), the thickness of the continental crust is less than 20 km. This estimated thickness is in good agreement with the heat ¯ow (more than 120 mW/m 2) (Battici et al., 1983) and gravimetric data (Barbieri et al., 1984). 3. Dive data and morphostructural observations The dive SARCYA 1 was completed on 16th

September 1994 (Georges Mascle Observer, Yves Potier Pilot, Serge Richard Co-pilot) on the SE Sardinia scarp off Capo Carbonara, by 39803 0 80 N, 10818 0 50 E, between 2475 and 1697 m; its total duration was of 7 h 58 min, with 5 h 12 min of observation on the bottom. This dive is located on the lower part of the Sardinian scarp which separates the two subbasins of Ichnusa and Cornaglia, precisely at the location where the characteristic NE±SW directions of the Sardinia Channel are replaced by the N±S direction characteristic of the Tyrrhenian margin of Sardinia, and just south of the submarine alkaline Pliocene

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Table 1 Description and location of the rock samples from the SAR1 dive Sample

Depth

Lat. N

Long. E

Description: dimensions, colour, form composition

SAR 1-01 SAR 1-02

2381 2208

39803 0 67 39803 0 67

10818 0 20 10817 0 90

SAR 1-03

2207

39803 0 67

10817 0 80

SAR 1-04

2039

39804 0 10

10817 0 90

SAR 1-05 SAR 1-06

1839 1823

39804 0 48 39804 0 54

10817 0 40 10817 0 30

SAR 1-07

1824

39804 0 65

10817 0 20

SAR 1-08

1707

39804 0 65

10817 0 00

SN

1707

39804 0 65

10817 0 00

10 £ 8 £ 7 cm; yellow-green ball, thin black crust, serpulians, andesite 5 £ 4 £ 2 cm; yellow reddish blunt plate, thin black crust, polyps, andesite 9 £ 4 £ 3 cm; light grey concretioned block, black in super®cy, carbonatic concretion 6 £ 5 £ 4 cm; grey greenish plate, black concretioned crust, polyps, andesite 6 £ 5 £ 5 cm; yellow greenish ball, thin black crust, andesite 7 £ 5; 5 £ 4 cm; yellow greenish blunt tip, thin black crust, andesite with oriented biotites 11 £ 7 £ 6 cm; yellow greenish blunt pyramide, thin black crust, serpulians, fractures, trachy-andesite with rare sanidine and quartz 6 £ 5 £ 3 cm; grey concretioned plate with many perforations, carbonatic concretion Orange brown ooze

Quirra volcano (Fig. 1). A distance of 3.7 km was covered on the bottom following an average direction of N 310 (Fig. 4). The exposed rocks were detected using the sonar STRASA because of important ooze blanket. As a consequence, the road following the outcrops was more or less sinuous with an average direction N 310 up to the depth of 2400 m; then, it followed a N 260 track up to 2320 m, succeeded by a N±S route up to 2170 m. Finally, the last route direction was N 300 (Fig. 4). The bathymetric pro®le was projected following the N 310 average direction (Fig. 4). Above a depth of 2385 m, the outcrops form a succession of subvertical cliffs, separated by steep dipping talus, blanketed by ooze. Eight rock samples were collected (Table 1) in four main zones of outcrops: (1) SAR 1-01 was sampled between 2385 and 2270 m depth, (2) SAR 1-02 and SAR 1-03 were sampled between 2220 and 2150 m depth, (3) SAR 104 was sampled between 2080 and 2000 m depth, and (4) SAR 1-05, SAR 1-06, SAR 1-07 and SAR 1-08 were collected between 1870 m and the end of the dive). The sampling in situ was very dif®cult because of the abundant carbonate concretions. Each zone of exposure consists of a succession of small scarps, between 1.50 and 12 m high, with an average of 3± 4 m, their tops blanketed by ooze. The subvertical cliffs are often covered by a more or less thick crust made of worm tubes (serpulids) and dead ahermatypic polyps, both covered by a ®lm of ferro-manganesiferous oxides.

A very dense network of fractures control both the general and detailed morphology. Four main directions have been recognized: N 20/N 40, N 50/N 70, N 90/N 120, N 170/ N 180 (Fig. 5). The ®rst three directions delimit the cliff system; N 20/N 40 and N 170/N 180 prevail in the deepest and middle parts of the cliff, respectively. The fracture network becomes very dense in the upper part of the cliff where the two families N 50/N 70 and N 90/N 120 determine a succession of very sharp, vertical, low angle dihedrons 20 to 30 cm apart (Fig. 6). A general layering is observed with a steep dipping,

Fig. 5. Frequency diagram of fracturation observed during SAR 1 dive. Using MacCan software of J.P. Bouillin.

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Fig. 6. Exposures (sample SAR 1-07) of SAR 1 dive at 1824 m below sea level.

generally oriented westwards along the N 170/N 180 and N 20/N 40 directions or northward along the N 50/ N 70 and N 90/N 120 ones. Due of the presence of the encrusting ®lm on the fault planes, it was not possible to observe any slickensides, with an exception at a depth of 1824 m, where normal slickensides are probably present. The general geometry is that of a tensional system characterised by N±S and/or N 20/ N 40 normal faults and E±W to N 120 tear faults. A large number of tracks of various nature have been observed: bioturbations, feeding tracks either isolated (Crustacean) or converging toward a teepee structure, volcano-like mud mounds, a large trench attributed to some ®sh activities, and a dredging track. Large amounts of Pteropod and Argonaute shells occur near the cliffs. The presence of a dominant SSW striking current is evidenced by different facts such as: (i) large dissymetric ripple marks (amplitude 2±5 cm and wavelength 25±30 cm), (ii) presence of ®sh (Benthosaurus) facing the current on the sea¯oor), (iii) quick dilution of the sediments removed by the submersible (average

10 cm/min), and (iv) erosion at the base of the cliffs. The sedimentary charge of the sea water is important; the bottom temperature is a constant 13.48C. Very few living organisms were observed: ®shes belonging predominantly to three families (Benthosaurus, Macroridae, Gadidae), some shrimps and rare gambas, one small sea star (Solaster), rare Pagurids and swimming Holothurids. Fixed animals are extremely rare: some Cerianths on an amphora, two pedunculate Spungia, some living ahermatypic Polyps above the depth of 1750 m. Archeological remains are represented by two punic like amphoras at 2211 and 2157 m depth, respectively. There are traces of more recent human pollution (picture boxes, 200 l oil barrel, food and beverage tins, plates, plastic bags, coal blocks, ®shing line with lead, terracotta), too. 4. Analytical procedures Data were processed in the different institutes

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which participated in this program: Laboratoire de GeÂodynamique des ChaõÃnes Alpines CNRS/UJF Grenoble, GeÂosciences Azur CNRS/UPMC Paris (Centre de Villefranche sur Mer), Istituto di Petrogra®a e Mineralogia UniversitaÁ di Torino, Istituto di Geologia Marina UniversitaÁ di Parma, DeÂpartement de GeÂologie Ecole Nationale SupeÂrieure d 0 IngeÂnieurs de Sfax Universite du Sud-Tunisien aÁ Sfax with the help of IFREMER, GENAVIR, and CNRS-INSU. Mineral chemistry of the shoshonitic volcanic rock (SAR 1-05) and the lamprophyric xenoliths were analysed with a Cambridge SEM-EDS microprobe, operating at an accelerating potential of 15 kV and a counting time of 50 s, in the Dipartimento di Scienze Mineralogiche e Petrologiche of Torino University. Natural and synthetic minerals and oxide standards were employed. Contents of Fe 31 and structural formula of amphiboles were calculated according to Holland and Blundy (1994). Analyses of other minerals were processed by using the softwares of Petrakakis and Dietrich (1985) and of Ulmer (1986). Major elements were analysed at the Centre de Recherches PeÂtrographique et GeÂochimique (CRPG) of Nancy by ICP-OES. Trace elements were analysed by inductively coupled plasma mass spectrometry (ICP-MS) at the Laboratoire de GeÂochimie isotopique de l'Universite de Toulouse, using lithium borate fusion, and following the procedures of Valladon et al. (unpublished report). 100 mg of powdered rocks were weighed in a Pt crucible, with 320 mg lithium metaborate and 80 mg lithium borate (Fluka). After mixing the powders carefully, the crucible was heated for fusion at 10008C. After cooling, 8 ml doubledistilled HNO3 (12N) and HF were added for the dissolution of the glass. The ®nal dilution to 30 ml of a 15 ml aliquot, with MilliQTM water and addition of internal standards (In±Re), corresponds to a total dilution of 3000. Detection limits are: REE and Y ˆ 0:03 ppm; U, Pb and Th ˆ 0:5 ppm; Hf and Nb ˆ 0:1 ppm; Ta ˆ 0:03 ppm; and Zr ˆ 0:04 ppm: Sr and Nd isotopic compositions were determined on a Finnigan MAT261 multicollector mass spectrometer at the Laboratoire de GeÂochimie isotopique de l'Universite de Toulouse. Refer to Lapierre et al. (1997) for analytical techniques. Isotopic data on the rocks have been corrected for in situ decay with an age of 12 Ma (see Section 7). Complete analyses are listed in Table

3. Major element analyses were calculated on a free volatile basis. 5. Petrology and mineral chemistry of the Sarcya 1 volcanic rocks Six samples of plagioclase±biotite±phyric rocks were collected during the dive (Table 1). Carbonate concretions developed on the rock surface or in dikelets, within the groundmass. Most of the carbonates were removed during the sample preparation but some remained. Plagioclase and amphibole are partly replaced by zeolites and smectites, respectively. The glassy groundmass is replaced by microcrystalline quartz and feldspar. Biotite and clinopyroxene are always preserved. All the lavas exhibit porphyritic textures with a glass-rich groundmass. They differ on the modal percentage of phenocrysts, the presence of amphibole and/or clinopyroxene phenocrysts and xenoliths. Their petrographic features are listed in Table 2. Because all the studied rocks are either moderately (SAR 1-03) or strongly (the rest of the samples) altered, they have been classi®ed using a combined geochemical and petrographic approach. 5.1. Petrography of the SARCYA lavas Clots of plagioclase and biotite phenocrysts are ubiquitous in the Sarcya shoshonitic rocks. Biotite is generally included in the plagioclase. However, plagioclase crystals, which crystallise early, may occur as inclusions in large biotite isolated phenocrysts (SAR 1-07). Plagioclase phenocrysts show intense normal zoning with Ca-rich cores and oligoclase to albite rims. Biotite includes apatite and zircon, and less often Fe±Ti oxides. Amphibole and Ca-rich clinopyroxene are present in some rocks (Table 2). Amphibole is systematically zoned with pale-green or colourless cores and intense brown pleochroic rims. The glassrich groundmass includes feldspar microlites, which may be are ¯ow aligned, and amphibole and biotite microphenocrysts (Fig. 7). 5.2. Petrography of the xenoliths present in the SARCYA shoshonites Xenoliths found in the Sarcya shoshonitic rocks

Plagioclase isolated or included in biotite brown pleochroic amphibole glassy groundmass replaced by smectites 1 zeolites Glomeroporphyritic aggregates of plagioclase and biotite. Clots euhedral Ca-rich clinopyroxene. Glassrich groundmass partly replaced by smectites 1 zeolites Plagioclase 1 biotite clots ^ associated with abundant zoned amphiboles lamprophyric inclusion (with olivine pseudomorphs) Plagioclase partly replaced by zeolites with biotite inclusions Glass rich groundmass interstital inclusion Plagioclase associated or not with biotite, few amphibole 1 clinopyroxene microphenocrysts. Abundant glass replaced by spherulitic colourless highly birefringent smectites

Slightly porphyritic , 15% phenocrysts Highly phyric . 30% phenocrysts Porphyritic 1 ¯uidal Aphyric to slightly porphyritic # 15% phenocrysts Porphyritic 1 vesicular 20% phenocrysts

SAR1.07 SAR1.01

Porphyritic 1 ¯uidal

Plagioclase associated or not with early crystallizing biotite clots. Amphibole relics. Glass replaced by smectites 1 zeolites ¯ow alined feldspar microlites

Sample

Texture

Mineralogy

Table 2 Petrographic characteristics for the Miocene Sarcya lavas

SAR1.05 SAR1.04 SAR1.02

SAR1.06

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show intersertal (SAR 1-04) and porphyritic textures (SAR 1-05A and B). Contacts between the xenoliths and host lava (sample SAR 1.05B) are sharp with no evidence of reaction or mixture between the two rocks. The intersertal xenolith is composed of 90% plagioclase laths and microlites, locally cemented by pods of glass, recrystallised in smectites. This rock also contains tiny (,1 mm sized) needle shaped amphibole and biotite crystals. The lamprophyric xenoliths are formed of olivine pseudomorphs, zoned euhedral amphibole and biotite elongated phenocrysts, caught in a ®ne grained groundmass containing the same minerals (acicular amphibole and biotite) and anhedral plagioclase (Fig. 7). 5.3. Mineral chemistry of the shoshonitic rocks and the lamprophyric xenoliths 5.3.1. Plagioclase Plagioclase is the most abundant mineral in the trachy-andesitic host rock. It occurs as phenocrysts and microphenocrysts. Normal zoning is commonly observed. Phenocrysts present in the groundmass or in the glomeroporphyritic aggregates show labradorite (An54) to andesine (An38) cores with decreasing Ca towards oligoclase rims (An21 ± 24) (Fig. 8(1A)). The composition of the plagioclase remains constant, when this mineral occurs near the lamprophyric xenolith (Fig. 8(1A, B)). Microphenocrysts exhibit oligoclase compositions (An23 ± 28) while tiny crystals show anorthoclase compositions. Orthoclase content increases with increasing Albite (Ab) up to 10 moles (Fig. 8(1A±C)). In the lamprophyre, cores and rims of the phenocrysts are labradorite (An50 ± 55), whereas the microlites have andesine (An50 ± 51) and/or sanidine (Ab31Or67) compositions (Fig. 8(1B,C)). 5.3.2. Pyroxene Pyroxene is found solely as relics in the amphibole phenocrysts. It ranges from augitic to diopsidic compositions (Table 3; Fig. 8(2)) (Morimoto et al., 1988). 5.3.3. Amphibole Amphibole forms isolated large zoned phenocrysts or clusters with mica and plagioclase in glomeroporphyritic aggregates. In both occurrences, it ranges from Fe-rich pargasite to pargasite and in the zoned crystals it reveals a Si-enrichment from core to rim

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Fig. 7. Microphotographs of the lamprophyric xenolith and its shoshonitic host rock (sample SAR 1-05). (1) The lamprophyric xenolith is composed of hornblende phenocrysts and needle-shaped phologopite embedded in a matrix formed of plagioclase and smectites. (2) Visible boundary of the lamprophyric xenolith to the right and the shoshonitic host rock to the left. The shoshonite is formed of plagioclase phenocrysts and tiny amphibole and mica crystals; the ground mass includes feldspar microlites.

(Fig. 8(3A)). When developed around clinopyroxene, amphibole composition varies from edenite to ferroedenite. Microphenocrysts in the groundmass are ferro-pargasites and ferro-edenites. Amphibole shows a different composition when the host shoshonitic rock is in contact with the lamprophyric xenolith. All the crystals are zoned regardless of the occurrence, and show an evolution from pargasite or edenite cores to ferro-edenite rims (Fig. 8(3B)). In the lamprophyre, composition of amphibole clusters in the pargasite ®eld and rims does not depart signi®cantly from that of the cores (Fig. 8(3C)). 5.3.4. Mica Mica composition of the host shoshonitic rock depends heavily on its contact distance with the lamprophyric inclusion. Located far from the xenolith,

the mica falls in the biotite ®eld, whereas when the host rock and xenolith are in close contact, mica plots in the phlogopite domain (Fig. 8(4)). When plotted in the FeO±MgO±Al2O3 diagram (Fig. 8(5)), all the mica fall in the calc-alkaline ®eld, and show a nice trend, marked by a Fe-enrichment, correlated with a Mgdepletion at constant Al2O3 content. The most Mgrich crystals are present in the lamprophyre or nearby in the host rock while those which are the most Feenriched occur in the host rock far from the xenolith. 6. Major, trace element and Nd isotopic chemistry of the Sarcya volcanic rocks 6.1. Major and trace element compositions The Sarcya lavas in general are strongly altered

Fig. 8. SAR 1 shoshonitic rocks: mineralogical data. (1) Alibite±Anorthite±K±feldspar ternary diagram showing the feldspar major element composition (A) SAR 1-05 free of lamprophyric xenolith, (B) SAR 1-05B and (C) SAR 1-05A shoshonites with lamprophyric xenoliths. (2) Pyroxene composition of sample SAR 1-05. (3) Amphibole composition of sample SAR 1-05 and lamprophyric xenolith. (4) Mica composition of sample SAR 1-05 and lamprophyric xenolith. (5) Mica of sample SAR 1-05 and lamprophyric xenolith on MgO/FeO/Al2O3 diagram.

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Sample No.

SAR1.03

Total Cr Ni V Sc Y Ba Rb Sr Nb Ta Zr Hf U Th Pb La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er

SAR1.05A 63.03 0.32 18.75 2.26 1.13

2.57 5.48 5.66 0.13 3.47

2.68 5.72 5.91 0.14

99.12

99.94

SAR1.06

SAR1.07

41.86 0.41 18.57 6.75 7.13 0.04 1.84 3.94 3.35 0.27 15.51

49.77 0.48 22.08 8.02 8.47 0.04 218 4.68 3.98 0.32

30.08 0.64 17.62 7.93 12.43 0.06 3.07 2.38 2.46 0.39 22.61

39.03 0.83 22.86 10.29 16.13 0.08 3.98 3.09 3.19 0.51

39.34 0.3 17.91 2.31 10.15

49.8 0.38 22.52 2.9 12.77

0.98 4.08 4.32 0.11 20.24

1.23 5.13 5.43 0.14

100.02

99.67 5.34 130.16 112.5 10.12 19.19 825.14 132.09 811.01 10 0.66 149.64 4.6 1.68 21.85 45.61

99.99

99.74 1.11 66.6 13.61 2.29 11.42 554.24 188.77 155.96 19.4 1.51 237.57 6.47 7.79 24.8 31.06

99.98

12.17 8.5 20.24 7.47 412.44 100.11 155.96 13.36 1.08 127.4 3.88 3.48 17.98 27.23

99.67 12.32 233.11 57.56 6.35 10.14 816 63.62 435.07 18.9 1.73 240.54 7.83 1.32 37.17 51.17

35.19 59.78 6.08 19.73 2.79 0.56 2 0.26 1.56 0.31 0.88

55.31 122.74 10.67 36.24 5.2 1.4 3.74 0.44 2.37 0.41 1.13

Nd

60.29 0.31 17.94 2.17 1.09

SAR1.05B

54.08 109.08 12.5 46.5 7.48 1.57 5.69 0.7 3.67 0.7 1.96

40.9 92.49 7.69 26.07 3.84 0.65 2.67 0.38 2.36 0.47 1.37

38.71 0.27 16.83 2.85 9.16 0.02 4.16 4.09 4.5 0.15 18.95

47.94 0.33 20.84 3.52 11.34 0.02 5.15 5.06 5.57 0.18

99.69 0.8 49.67 14.86 2.48 14.92 536.94 204.66 238.94 17.7 1.42 225.61 6.23 8.19 29.77 27.32

99.96

47.21 96.57 8.81 29.38 4.43 0.81 3.52 0.53 2.87 0.58 1.83

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Rock type SiO2 TiO2 Al2O3 Fe2O3 MgO MnO CaO Na2O K2O P2O5 LOI

190

Table 3 Chemical analysis of the Miocene SAR1 lava

0.29 1.82 0.29 0:512063 ^ 6 0.51206 2 11.05 0.706894 ^ 10 3.3719 0.70629 0.21 1.36 0.18 0:512514 ^ 6 0.51251 2 2.25 0.707753 ^ 10 1.8572 0.70742 0.29 1.9 0.31 0:512650 ^ 5 0.51264 1 0.39 0.704733 ^ 18 0.4711 0.70465 0.12 0.76 0.13 0:512080 ^ 6 0.51207 2 10.71 Tm Yb Lu 143 Nd/ 144Nd 143 ( Nd/ 144Nd)1 eNd…T ˆ 12 Ma† 87 Sr/ 86Sr 87 Rb/ 86Sr 87 ( Sr/ 86Sr)1

0.15 1.07 0.15 0:512547 ^ 6 0.5124 2 1.60

SAR1.03 Sample No.

Table 3 (continued)

SAR1.05A

SAR1.05B

SAR1.06

SAR1.07

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191

rocks. They have been considered to belong to a shoshonitic suite based on their mineralogy with the exception of the lamprophyric inclusion (SAR1-05B). However, their SiO2 and MgO contents do not fall in the range of values commonly found for andesites. SiO2 is too low while MgO is too high. Moreover, these rocks have high loss of ignition (LOI # 20%). As mentioned earlier, the discrepancy between the SiO2 and MgO levels and the high LOI, re¯ect the presence of remaining carbonates and zeolites. The high Ni contents of the lamprophyric xenolith and host lava (SAR1-05B) re¯ect the presence of olivine pseudomorphs in the xenolith or as xenocrysts in the host rock. Because of the unreliable major element chemistry, we have plotted the Sarcya lavas on the Zr/Ti versus Nb/Y diagram (Fig. 9(1); Winchester and Floyd, 1977), based on trace elements considered to be relatively immobile during alteration and/or low grade metamorphism processes. The lamprophyric xenolith and host volcanic rock (SAR 1-05B) and samples SAR 1-01 and 1-07 cluster in the trachyandesite ®eld, while SAR 1-06 falls along the trachyte and trachy-andesite boundary. Only, SAR1-05B falls in the dacitic and rhyolitic ®eld. Similarly, because of the well known K2O mobility during hydrothermal or low grade metamorphism processes, diagrams based on trace elements considered as relatively immobile (Ce, Yb, Ta, Th) have been proposed to distinguish the shoshonitic suites from the calc-alkaline and tholeiitic ones (Muller et al., 1992). All the Sarcya lavas are plotted in the shoshonitic ®eld (Fig. 9(2)). The chondrite-normalized rare earth (REE) patterns of the Sarcya volcanic rocks are enriched in light (L)REE relative to heavy (H)REE ‰18:6 , …La=Yb†N , 37:10Š and have a marked Eu negative anomaly …0:62 # Eu=Eup # 0:73†; which is absent in the lamprophyric xenolith (SAR 1-05A; Eu=Eup ˆ 0:97). The primitive mantle-normalized (Sun and McDonough, 1989) multi-element patterns of the studied rocks (Fig. 10) have overall similarities with shoshonitic rocks of intermediate compositions developed in continental arc settings (Muller et al., 1992). All the samples have high Th/Nb (1.2 to 1.96) and Th/ Ta ratios (16.6 to 21.5) and the negative Ba and Sr

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Fig. 9. Classi®cation of SAR 1 shoshonitic rocks. (1) Using Zr/Ti versus Nb/Y diagram (Winchester and Floyd, 1977). (2) Using diagram Ce/ Yb versus Ta/Yb, Th/Yb versus Ta/Yb (Muller et al., 1992).

negative anomalies, coupled with the strong depletion in High Field Strengh elements (Nb, Ta and Ti), that characterise the volcanic potassic rocks from subduction-related settings (Foley et al., 1987; Muller et al., 1992). The highest Th/Nb (2.18) and Th/Ta (33.1) ratios of the lamprophyric xenolith, fall in the range

of Ti-poor lamproites. The Zr/Ce (2.13 to 2.56) and Ti/Nb (111.86 to 143.7) ratios of the lavas devoid of lamprophyric xenoliths or xenocrysts are higher than the ratios of the lamprophyric xenolith. Moreover, these ratios fall in the range of continental arcs. Thus, the Sarcya volcanic rocks and their

G.H. Mascle et al / Marine Geology 179 (2001) 179±202

lamprophyric xenolith belong to a shoshonitic suite emplaced in a continental arc or post-collisional settings. 6.2. Nd and Sr isotopic compositions The e Nd ratios of the Sarcya (Fig. 11) lavas are low and range from 10.39 to 210.71. The lava with the lamprophyric inclusion has the highest ratio (10.39)

193

which is close to the Bulk Silica Earth (BSE) composition. Samples SAR 1-03 and SAR 1-07 have the lowest e Nd ratios …, 2 11†: These e Nd ratios suggest that all these rocks are likely to be derived from the melting of a mantle source, which suffered contamination and/or assimilation of continental crust. The Sr isotopic compositions were determined for three samples only (Table 3). The ratio ( 87Sr/ 86Sr) ranges between 0.70465 and 0.70629 (Fig. 11).

Fig. 10. Chondrite normalized rate earth and mantle normalized multi-element diagram of the SAR 01 rocks. (Sun and McDonough, 1989).

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G.H. Mascle et al / Marine Geology 179 (2001) 179±202

Fig. 11. Neodymium geochemistry of SAR 1 rocks. Reported for comparison are the ®elds of lavas from Sardinia, Campania (after Beccaluva et al., 1994), Roccamon®na (after Hawkesworth and Vollmer, 1979, ref. in Beccaluva et al., 1994), Roman region (after Carter et al., 1978, ibid), Stromboli (after Ellam et al., 1988, ibid.) and Tuscany lamproõÈtes (after Peccerillo, 1999 and references therein).

In summary, the Sarcya lavas exhibit features of shoshonitic suites of intermediate composition. Some samples include amphibole and mica bearing lamprophyric xenoliths, which are geochemically similar to Ti-poor lamproites. Mica compositions re¯ect chemical exchanges between the lamprohyre and its andesitic host rock suggesting that both rocks were emplaced simultaneously. Nd compositions of these shoshonitic lavas and lamprophyric xenolith indicate that continental crust was largely involved in their genesis. The spatial association of lamprophyre and shoshonitic rocks is geochemically similar to K-rich and TiO2-poor igneous suites emplaced in post-collisional settings. 7. Geochronology Sample SAR 1-01 has been dated using the Ar/ 39Ar method with a laser microprobe on single grains. Analyses were carried out by progressive degassing of each grain by a series of distinct meltings 40

on a single mineral (Monie et al., 1997; MonieÂ, 1998). Ages are reported with 1s uncertainty. Fifteen dates were recorded from an automorph crystal of biotite, giving concordant values within the margin of uncertainty, and without variations between the core and the periphery of the crystal (Fig. 12, Table 4). The average age of 12:6 ^ 0:3 My appears representative of the volcanic activity. 8. Geodynamic setting of the 12.5 Ma Sarcya 1 shoshonitic suite and its place in the evolution of the Western Mediterranean In the Western Mediterranean region, volcanic rocks are widespread (Fig. 3) and characterise three main geodynamic settings: (1) ocean ¯oor basalts (OFB), (2) island-arc suites, (3) within-plate volcanism. 1. Ocean ¯oor magmatism began in the Algerian± Ligurian±ProvencËal basin at 20 Ma and lasted up to 12 Ma (Rehault et al., 1984, 1985). 6 Ma later, it occurred in the Tyrrhenian basin, ¯ooring the 6 to

G.H. Mascle et al / Marine Geology 179 (2001) 179±202

Fig. 12. Geochronology 39Ar/ 40Ar of SAR 1-01.

3 Ma old Vavilov±Magnaghi basin, and is still active in the Marsili basin (Sartori et al. 1987; Kastens et al., 1988; Bertrand et al., 1989; Beccaluva et al., 1990; Kastens and Mascle, 1990; Serri, 1990). This temporal evolution of OFB volcanism documents the progressive eastward migration of the opening of Western Mediterranean basins.

195

2. Subduction-related volcanic suites consist of three successive arcs, with the oldest one being more or less strongly fragmented by further geodynamic evolution (Beccaluva et al., 1985; Sartori et al., 1987, 1990; Sartori, 1990). The 20 to 12 Ma old (Montigny et al., 1981; Di Girolamo, 1984; Savelli, 1988) calc-alkaline sequences (Coulon, 1977) from western Sardinia and contemporaneous magmatic rocks from northern Tunisia (Galite archipelago) and Lesser Kabylia (Bellon, 1981; Rekhiss, 1985; Bouillin et al., 1986), represent the oldest arc. The development of this arc is linked to the subduction of Tethyan lithosphere below the European margin during Oligocene±Early Miocene times. This is coeval with the opening of the Algerian±Ligurian±ProvencËal back-arc basin (Le Pichon et al., 1971; Auzende et al., 1973; Rehault et al., 1984, 1985; Dercourt et al., 1986; Dewey et al., 1989). The second arc developed between 14 and 3 Ma (Civetta et al., 1978; Beccaluva et al., 1985, 1987; Savelli, 1988; Serri, 1990; Serri et al., 1993) and consists of the Late Miocene±Early Neogene Tuscan, Roman and Campanian lavas and the submerged remanent arc-suites, exposed between Ustica and Poseidone Seamount. The development of this arc is related to the eastward retreat of subduction, which provokes the opening of the Vavilov basin (Kastens et

Table 4 Argon geochronology of the Miocene SAR1 lava N8, SAR 1.01

40

Ar/ 39Ar

1 2 3 4 5 6 7 8 9 10 11 12 13 14 15

0.438 0.443 0.429 0.501 0.407 0.490 0.506 0.455 0.468 0.435 0.384 0.492 0.423 0.404 0.452

Total age ˆ 12.6 ^ 0.3 Ma

36 Ar/ 40Ar £ 1000, J ˆ 0:015585

39

Ar/ 40Ar

37

Ar/ 39Ar

0.255 0.135 0.116 1.627 1.006 0.029 0.387 0.470 0.051 0.028 0.480 0.647 0.111 0.301 0.030

2.1074 2.1665 2.2460 1.0357 1.7237 2.0186 1.7487 1.8901 2.1013 2.2777 2.2312 1.6414 2.2844 2.2516 2.1908

0 0.059 0.025 0 0.069 0.032 0 0.064 0 0 0 0.016 0 0.013 0.045

% Atm

Age ^ 1 sd

7.54 4.02 3.43 48.08 29.73 0.88 11.46 13.92 1.51 0.85 14.19 19.13 3.31 8.92 0.90

12:3 ^ 2:1 12:4 ^ 2:4 12:0 ^ 2:1 14:0 ^ 1:2 11.4 ^ 1.0 13:7 ^ 0:9 14:2 ^ 1:5 12:7 ^ 2:0 13:1 ^ 0:6 12:2 ^ 0:6 10:8 ^ 1:8 13:8 ^ 0:3 11:9 ^ 1:6 11:3 ^ 0:8 12:7 ^ 0:5

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Fig. 13. Geochemical comparison of the Sarcya shoshonitic lavas with neighbouring igneous suites of similar ages (data from Peccerillo, 1999 and references therein).

al., 1988; Kastens and Mascle, 1990; Sartori, 1990; Beccaluva et al., 1990, 1994). The Late Pliocene to Recent Roman, Campanian and Aeolian volcanic rocks, as well as the Arcestes, Anchise, the Albatros seamounts represent the third and youngest arc. Arc-volcanism began at 6 Ma and is still active (Francalanci and Manetti, 1994; Beccaluva et al., 1985). It is related to the new eastward retreat of the subduction, which led to the opening of the Marsili basin (Sartori et al., 1987; Kastens et al., 1988; Kastens and Mascle, 1990; Sartori, 1990). Retreat of subduction is interpreted as a selfcontrolled mechanism due to the slab buoyancy Kastens et al., 1988; Kastens and Mascle, 1990), or related to the slab detachment (Carminati et al., 1998; Van Der Meulen, 1999).

3. Pliocene-Quaternary within-plate alkaline volcanism took place in (i) central-eastern Sardinia and along its Tyrrhenian margin (Hole 654, Kastens and Mascle, 1990), M. Quirra, (ii) Ustica and (iii) Sicily Channel and mainland (Ambrosetti et al., 1987). The Cornacya shoshonitic lavas and their lamprophyric xenoliths emplaced after the 20 to 12 Ma old arc and before the Late Miocene±Early Neogene one. They show petrological and geochemical similarities with the K-rich rocks and lamproites from Tuscany (Peccerillo et al., 1988; Conticelli and Peccerillo, 1992; Serri et al., 1993; Peccerillo, 1999; Peccerillo and Panza, 1999). However, the rocks from Tuscany are younger than those from Cornacya and thus, cannot belong to the same magmatic episode and

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Fig. 14. The Sardinia Channel area at the time of Cornacya volcanism. 12 My is a key date for the Western Mediterranean. Before 2 My (between 20 and 12 My) the area underwent a distensive regime in its internal part, which corresponds to the gravitational collapse of the over thickened crust, the opening of the Algerian±ProvencËal Ligurian-basin and the rotation of Corsica±Sardinia. After 12 My (between 12 and 6 My) the distensive regime interests the future Western Tyrrhenian domain.

geodynamic setting. Lamproites contemporaneous with the Cornacya suite have been also described in Algeria, in the Constantine area (Raoult and Velde, 1971; Vila et al., 1974; Penven and Zimmerman,

1986), but are located rather far from the Tyrrhenian Sea. On the other hand, the Cornacya lavas display similar tectonic setting, ages and geochemical features of the lamproites from Sisco (Cape Corsica

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(Venturelli et al., 1984). Both suites are located on the western margin of the Tyrrhenian Sea and are Serravallian in age …< 12 Ma†: Moreover, the primitive mantle-normalized multi-elements plots of the Cornacya shoshonitic andesites and Sisco lamproite are similar (depleted in Nb, Ta, Ti, HREE and Y, enriched in large ion lithophile elements; Fig. 13). However, the Cornacya lavas differ from the lamprophyre by lower contents in incompatible trace elements, such as, Nb, Ta, Hf, Zr. According to Serri et al. (1993), the Sisco lamproites represent the very ®rst episode of postcollisional lithospheric extension of Northern Apennine. Similarly, the Cornacya K-rich andesites and lamprophyre were likely emplaced during the beginning of post-collisional lithospheric extension of the Corsica±Sardinia block, just after the end of its rotation and before the opening of the Tyrrhenian Sea, which is related to the ®rst slab retreat and the Late Miocene±Early Neogene arc development (Fig. 14). Drillings on the Sardinia margin (ODP Leg 107) have shown that the uppermost levels of the presentday margin (Hole 654) have suffered tectonic subsidence before the lowest part (Hole 652). This subsidence is interpreted to re¯ect an eastward migration of the extension during Late Miocene and Early Pliocene times (Kastens et al., 1988; Kastens and Mascle, 1990). Data presented in this paper are in good agreement with this model and provide chronological constraints for the beginning of the phenomenon. 9. Conclusion During the Sarcya cruise in the Sardinia Channel, SAR 1 dive was performed along the lower part of the Sardinian scarp which divides the Ichnusa and Cornaglia basins. The dive explored a previously unknown submerged volcano, named Cornacya. From 2385 m bsl and above, volcanic rocks are exposed and form a succession of subvertical cliffs separated by steep dipping talus. The intense fracturation corresponds to a N±S- to N 20/40-trending normal fault system, with E±W- to N 120 tear faults, which is consistent with a horst and graben, and tilted blocks structure, related to the opening of the deep Sardinia channel.

The volcanic rocks exhibit features of a shoshonitic suite of intermediate composition. Lamprophyres are often included in the K-rich andesites and are geochemically similar to Ti-poor lamproites. Mica compositions re¯ect chemical exchanges between the lamprophyre and its host rock suggesting that both rocks were emplaced simultaneously. The predominantly negative e Nd (10.39 to 210.71) of these rocks support involvement of continental crust in their genesis. The geochemical characteristics of this lamprophyre±shoshonitic suite are similar to those of K-rich and Ti-poor magmatic rocks emplaced in postcollisional settings. The Cornacya shoshonitic suite has been dated at 12:6 ^ 0:3 My and is contemporaneous with the Sisco lamproite. Although the igneous suites are 500 km apart they have similar structural location on the Western Tyrrhenian margin. They were emplaced during the post-collisional lithospheric extension of the Corsica±Sardinia block, just after its rotation and before the Tyrrhenian sea opening. This means that lithospheric thinning began in northern and southern Tyrrhenian at the same time, during the Serravallian. The differences between the geodynamic evolution of northern and southern Tyrrhenian basins occurred later due to the southward retreat of the Tyrrhenian subduction. Acknowledgements The authors acknowledge the European Community (European Program Human Capital and Mobility) for funding the SARCYA cruise, the IFREMER (SARTUCYA cruise), the staff and crew of N.O. Le Suroit, the staff of Cyana submersible, CNRS INSU programme Geosciences Marines (funding of data processing). Laboratory work of F.R. was ®nancially supported by M.U.R.S.T., grant ex-40%. L. Beccaluva, A. Peccerillo and an anonymous reviewer provided valuable suggestions to improve manuscript. References Ambrosetti, P., et al., 1987. Neotectonic map of Italy. Foglio 5. Bosi C., coordinatore. Amodio-Morelli, L., Bonardi, G., Colonna, V., Dietrich, D., Giunta, G., Ippolito, F., Liguori, V., Lorenzoni, S., Paglionico, A., Perrone, V., Piccarreta, G., Russo, M., Scandone, P., Zanettin

G.H. Mascle et al / Marine Geology 179 (2001) 179±202 Lorenzoni, E., Zuppetta, A., 1976. L'Arco Calabro-Peloritano nell'orogene apenninico-maghrebide. Mem. Soc. Geol. Ital. 17, 1±60. Auzende, J.M., Bonnin, J., Olivet, J.L., 1973. The origin of the western Mediterranean basin. J. Geol. Soc. London 129, 607± 620. Auzende, J.M., Olive, J.L., Bonnin, J., 1974. Le deÂtroit sardanotunisien et la zone de fracture nord-tunisienne. Tectonophysics 21, 357±374. Barbieri, F., Cornini, S., Marchetti, F., Morlotti, E., Terdich, P., Torelli, L., 1984. Geological investigations in the Sardinia Channel area: preliminary results. Ateneo Parm. 20, 177±207. Battici, G., Della Vedova, B., Pellis, G., Torelli, L., 1983. Lineamenti geologico-strutturali del canale di Sardegna e aree limitrofe. Atti Conv. Annuale Gruppo Nazionale Geo®sica della Terra Solida, Roma. Beccaluva, L., Gabbianelli, G., Lucchini, F., Rossi, P.L., Savelli, C., 1985. Petrology and K/Ar ages of volcanics dredged from the Eolian seamounts: implication for geodynamic evolution of the southern Tyrrhenian basin. Earth Planet. Sci. Lett. 74, 187±208. Beccaluva, L., Brotzu, P., Macciotta, G., Morbidelli, L., Serri, G., Traversa, G., 1987. Cainozoic tectonomagmatic evolution, inferred mantle in the Sardo-Tyrrhenian area. In: The Lithosphere in Italy. Adv. Earth Sci. Res. Acc. Naz. Lincei 80, 229±248. Beccaluva, L., Bonatti, E., Dupuy, C., Ferrara, G., Innocenti, F., Lucchini, F., Macera, P., Petrini, R., Rossi, P.L., Serri, G., Seyler, M., Siena, F., 1990. Geochemistry, mineralogy of volcanic rocks from ODP sites 650, 651, 655, 654 in the Tyrrhenian sea. Proc. ODP Sci. Res. 107, 49±74. Beccaluva, L., Coltorti, M., Galassi, B., Macciotta, G., Siena, F., 1994. The Cainozoic calcalkaline magmatism of the western Mediterranean, its geodynamic signi®cance. Boll Geof. Teor. Appl. 36, 294±308. Bellon, H., 1981. Chronologie radiomeÂtrique (K-AR) des manifestations magmatiques autour de la MeÂditerraneÂe occidentale entre 33 et 1 MA. In: Wezel, F.C. (Ed.), Sedimentary Basins of the Mediterranean Margins. CNR, Tecnoprint, Bologna, pp. 341±360. Bertrand, H., Boivin, P., Robin, C., 1989. Evolution geÂochimique d'un bassin d'arrieÁre arc: la mer TyrrheÂnienne au PlioceÁne. C.R. Soc. GeÂol. France, SeÂr. II 308, 1423±1429. Blundell, D., Freeman, R., Mueller, S., 1992. A Continent Revealed: The European Geotraverse. Cambridge University Press, Cambridge, 275pp. Bouillin, J.P., 1984. Nouvelle interpreÂtation de la liaison ApenninMaghreÂbides en Calabre; conseÂquences sur la paleÂogeÂographie teÂthysienne entre Gibraltar et les Alpes. Rev. GeÂol. Dyn. GeÂogr. Phys. 25 (5), 321±338. Bouillin, J.P., 1986. Le bassin maghreÂbin: une ancienne limite entre l'Europe et l'Afrique aÁ l'ouest des Alpes. Bull. Soc. GeÂol. Fr. II 4, 547±558. Bouillin, J.P., Durand-Delga, M., Olivier, P., 1986. Betic-Ri®an and Tyrrhenian arcs: distinctive features, genesis and development stages. In: Wezel, F.C. (Ed.), The Origin of Arcs. Elsevier, Amsterdam, pp. 281±304. L'eÂquipe scienti®que embarqueÂe, Bouillin, J.P., Poupeau, G.,

199

Tricart, P., Bigot-Cormie, F., Mascle, G., Torelli, L., 1998. PremieÁres donneÂes thermo-chronologiques sur les socles sarde et kabylo-peÂloritain submergeÂs dans le Canal de Sardaigne (MeÂditerraneÂe occidentale). C. R. Acad. Sci. Paris 326, 561± 566. L'eÂquipe sienti®que embarqueÂe, Bouillin, J.P., MonieÂ, P., Rolfo, F., Tricart, P., Mascle, G., Torelli, L., 1999. DonneÂes geÂochronologiques 40Ar/ 39Ar sur les socles sardes et kabylo-peÂloritain submergeÂs dans le Canal de Sardaigne (MeÂditerraneÂe occidentale). C. R. Acad. Sci. Paris 328, 529±534. Brocard, G., 1999. Le canal de Sardaigne au NeÂogeÁne. Analyse morphologique et structurale. Apports de la bathymeÂtrie multifaisceaux et des plongeÂes SARCYA et SARTUCYA. GeÂologie Alpine, 74, in press. Caire, A., 1973. Les liaisons alpines preÂcoces entre Afrique du Nord et Sicile, et la place de la Tunisie dans l'arc tyrrheÂnien. Ann. Mines GeÂol. Tunis 26, 87±110. Carminati, E., Wortel, M., Spakman, W., Sabadini, R., 1998. The role of slab detachment processes in the opening of the westerncentrel Mediterranean basins: some geological and geophysical evidence. Earth Planet. Sci. Lett. 160, 651±665. Castany, G., 1955. Le haut-bassin siculo-tunisien. Bull. Station OceÂanogr. Salammboà (Tunisie) 52, 3±17. Castany, G., 1959. La geÂologie profonde du territoire TunisieSicile. Topographie et GeÂologie des Profondeurs OceÂaniques. Coll. CNRS 83, 165±183. Catalano, R., D'Argenio, B., Montanari, L., Morlotti, E., Torelli, L., 1985. Marine geology of the NW Sicily offshore (Sardinia Channel) and its relationships with mainland structures. Boll. Soc. Geol. Ital. 104, 207±215. Civetta, L., Orsi, G.P.S., Pece, R., 1978. Eastwards migration of the Tuscan anatectic magmatism due to anticlockwise rotation of the Apennines. Nature 276, 604±606. Colantoni, P., Fabbri, A., Gallignani, P., Sartori, R., Rehault, J.P., 1981. Carta litologica e stratigra®ca dei mari italiani. Litogra®ca Artistica Cartogra®ca, 4. Compagnoni, R., Morlotti, E., Torelli, L., 1987. Lithostructural characters of the accoustic basement of the Sardinia Channel (Southwestern Tyrrhenian sea). Ann. Inst. Geol. Publ. Hung., Budapest 70, 227±233. Compagnoni, R., Morlotti, E., Torelli, L., 1989. Crystalline and sedimentary rocks from the scarps of the Sicily±Sardinia Trough and Cornaglia Terrace (southwestern Tyrrhenian Sea, Italy): paleogeographic and geodanamic implications. Chem. Geol. 77, 375±398. Conticelli, S., Peccerillo, A., 1992. Petrology and geochemistry of potassic and ultrapotassic volcanism from Central Italy: inferences on its genesis and on the mantle source evolution. Lithos 28, 221±240. Coulon, C., 1977. Le volcanisme calco-alcalin ceÂnozoõÈque de Sardaigne. PeÂtrologie, geÂochimie et geÂneÁse des laves andeÂsitiques et des ignimbrites. Signi®cation geÂodynamique. TheÁse Etat, Univ. Marseille, 385pp. Depardon, S., 1995. Le reÂamincissement crustal dans le Canal de Sardaigne. MeÂmoire de MaõÃtrise, Univ. Joseph Fourier Grenoble. Dercourt, J., Zoneshain, L.P., Ricou, L.E., Kazmin, V.G., Le

200

G.H. Mascle et al / Marine Geology 179 (2001) 179±202

Pichon, X., Knipper, A.L., Grandjaquet, C., Sbortshikov, I.M., Geyssant, J., Lepvrier, C., Pechersky, D.H., Boullin, J.P., Sibuet, J.C., Savostin, J.A., Biju-Duval, B., 1986. Geological evolution of the Tethys belt from Atlantic to Pamir since the Lias. Tectonophysics 45, 265±283. Dewey, J.F., Helman, M.L., Turco, E., Hutton, D.W.H., Knott, S.D., 1989. Kinematics of the western Mediterranean. In: Coward, M.P., Dietrich, D., Park, G. (Eds.), Alpine Tectonics on Spec. Publ., Geol. Soc. Lond., 45, 265±283. Di Girolamo, P., 1984. Magmatic character and geotectonic setting of some tertiary±quaternary Italian volcanic rocks: orogenic, anorogenic ant transitional association. A review. Bull. Volcanol. 47 (3), 421±432. Egger, A., 1992. Lithospheric structure along a transect from the northern Apennines to Tunisia derived from seismic refraction data. PhD Thesis. Univ. ETH ZuÈrich, 150pp. Fabretti, P., Sartori, R., Torelli, L., Zitellini, N., Brabcolini, G., 1995. La struttura profonda del margine orientale della Sadegna dall'interpretazione di sismica a ri¯essione e a rifrazione. Stud. Geol. Camerti 1995/2, 239±246. Foley, S.F., Venturelli, G., Green, D.H., Toscani, L., 1987. The ultrapotassic rocks: characterisitcs, classi®cation and constraints for petrogenetic models. Earth Sci. Rev. 24, 81±134. Francalanci, L., Manetti, P., 1994. Geodynamic models of the southern Tyrrhenian region: constraints from the petrology and geochemistry of the Aeolian volcanic rocks. Boll. Geof. Teor. Appl. 36, 283±292. Gennesseaux, M., Stanley, D., 1983. Neogene to recent displacement and contact of Sardinian and Tunisian margins, Central Mediterranean. Smithsonian Contr. Mar. Sci. 23, 21. Giese, P., Roeder, D., Scandone, P., 1992. Sardinia Channel and Atlas in Tunisia: extension and compression. In: Blundell, D., Freeman, R., Mueller, S. (Eds.), A Continent Revealed: The European Geotraverse. Press Syndicate of the University of Cambridge, Cambridge, pp. 199±202. Grandjacquet, C., Mascle, G., 1978. The structure of the ionian sea, Sicily and Calabria-Lucania. In: Nairn, A.E.M., Kanes, W.H., Stehli, F.G. (Eds.), The Oceans Basins and Margins, vol. 4B. Plenum Press, New York, pp. 257±329. Holland, T.J.B., Blundy, J.D., 1994. Non-ideal interactions in calci amphiboles and their bearing on amphibole-plagioclase thermometry. Contrib. Mineral. Petrol. 116, 433±447. Kastens, K., Mascle, J., 1990. The geological evolution of the Thyrrenian sea: an introduction to the scienti®c results of ODP leg 107. Proc. ODP Sci. Res. 107, 3±26. Kastens, K., Mascle, J., Auroux, C., Bonnati, E., Broglia, C., Channell, J., Curzi, P., Eemeis, K.C., GlacËon, G., Hasegawa, S., Hieke, W., Mascle, G., Mc Coy, F., Mc Kenzie, J., Mendelson, J., Muller, C., Rehault, J.P., Robertson, A., Sartori, R., Sprovieri, R., Torii, M., 1988. ODP Leg 107 in the Tyrrhenian Sea: insights into passive margin and back-arc basin evolution. Geol. Soc. Am. Bull. 100, 1140±1156. Lapierre, H., Dupuis, V., Mercier de LeÂpinay, B., Tardy, M., Ruiz, J., Maury, R.C., Hernandez, J., Loubet, M., 1997. Is the Lower Duarteigneous complex (Hispaniola) a remnant of the Caribbean plume-generated oceanic plateau? J. Geol. 105, 111±120. Le Pichon, X., Pautot, G., Auzende, J.M., Olivet, J.L., 1971. La

MeÂditerraneÂe occidentale depuis l'OligoceÁne. ScheÂma d'eÂvolution. Earth Planet. Sci. Lett. 13, 145±152. Mascle, G., Tricart, P., 1994. SARTUCYA. Proposition de campagne aÁ la mer IFREMER-I.F.R.T.P. Mascle, G., Tricart, P. (Eds.), 2001. Les escarpements sous-marins du canal de Sardaigne: reÂamincissement crustal et extension tardi-orogeÂnique, premier stade d'une ouverture arrieÁre-arc. ReÂsultats des campagnes de plongeÂes Cyana: SARCYA et SARTUCYA. MeÂm. GeÂol. Alpine 32, 202pp. Mauffret, A., Rehault, J.P., Gennesseaux, M., Bellaiche, G., Labarbarie, M., Lefebvre, D., 1981. Western Mediterranean basin evolution from a distensive to a compressive regime. In: Wezel, F.C. (Ed.), Sedimentary Basins of Mediterranean Margins. CNR Italian Project of Oceanogaphy, Bologna, pp. 67±81. MonieÂ, P., 1998. MeÂthodologie 40Ar/ 39Ar par sonde laser. Application aÁ l'eÂtude de l'orogeneÁse alpine en MeÂditerraneÂe occidentale. DHDR University, Montpellier. MonieÂ, P., Caby, R., Arthaud, M.H., 1997. The Neoproterozoic Brasiliano orogeny in northeast Brazil: 40Ar/ 39Ar and petrostructural data from Ceara. Precambr. Res. 81, 212±264. Montigny, R., Edel, J.B., Thuizat, R., 1981. Oligo-Miocene rotation of Sardinia: K±Ar ages and paleomagnetism data of Tertiary volcanics. Earth Planet. Sci. Lett. 54, 261±271. Morelli, C., Nicolich, R., 1990. A cross section of the lithosphere along the European Geotraverse Southern segment (from the Alps to Tunisia). Tectonophysics 176, 229±243. Morimoto, M., FabrieÁs, J., Ferguson, A.K., Ginzburg, I.V., Ross, M., Seifert, F.A., Zussman, J., 1988. Nomenclature of Clinopyroxenes. Mineral. Mag. 52, 535±550. Muller, D., Rock, N.M.S., Groves, D.I., 1992. Geochemical discrimination between shoshonitic and potassic volcanic rocks in different tectonic settings: a pilot study. Mineral. Petrol. 46, 259±289. Peccerillo, A., 1999. Multiple mantle metasomatism in centrasouthern Italy: geochemical effects, timing and geodynamic implications. Geology 27, 315±318. Peccerillo, A., Panza, G.F., 1999. Upper mantle domains beneath central-Southern italy: petrological, geochemical and geophysical constraints. Pure Appl. Geophys. 156, 421±443. Peccerillo, A., Poli, G., Serri, G., 1988. Petrogenesis of orenditic and kamafugitic rocks from Central Italy. Can. Mineral. 26, 45±65. Penven, M.J., Zimmerman, J.L., 1986. Mise en eÂvidence par la meÂthode potassium-argon d'un aÃge langhien pour le plutonisme calc-alcalin de la Kabylie de Collo (AlgeÂrie). C. R. Acad. Sci. 303, 403±406. Petrakakis, K., Dietrich, H., 1985. MINSORT: a program for the processing and archivation of microprobes analyses of silicate and oxide minerals. Neues Jahrb. Mineral. Monatsh. (1985), 379±384. Raoult, J.F., Velde, D., 1971. DeÂcouverte de trachytes potassiques aÁ olivine et d'andeÂsites en couleÂes dans le MioceÁne continental au Sud du Kef Hahouner (Nord du Constantinois, AlgeÂrie). C. R. Acad. Sci. 272, 1051±1054. Rehault, J.P., Boillot, G., Mauffret, A., 1984. The western Mediterranean Basin, geological evolution. Mar. Geol. 55, 447±477. Rehault, J.P., Mascle, J., Boillot, G., 1985. Evolution geÂodynamique

G.H. Mascle et al / Marine Geology 179 (2001) 179±202 de la MeÂditerraneÂe depuis l'OligoceÁne. Mem. Soc. Geol. Ital. 27, 85±96. Rekhiss, F., 1985. Les roches intrusives mioceÁnes de la Galite. Notes Serv. GeÂol. Tunisie 51, 177±196. Research Group for Lithospheric Structure in Tunisia, 1992. The EGT'85 seismic experiment in Tunisia: a reconnaissance of the deep structures. Tectonophysics 207, 245±267. Sartori, R., et al., 1990. The main results of ODP Leg 107 in the frame of Neogene to Recent geology of perityrrhenian areas. Proc. ODP Sci. Res. 107, 715±730. Sartori, R., Mascle, G., Amaudric du Chaffaut, S., et al., 1987. A review of circum-Tyrrhenian regional geology. Proc. ODP Init. Rep. 107, 37±63. Sartori, R., Mascle, G., Bouillin, J.P., Girault, J., Naud, G., Pasini, M., Piboule, M., 1990. Types and sources of large rock clasts and heavy minerals from ODP sites 652, 653, 654 and 656 in the Tyrrhenian sea. Implications about the nature of the East Sardinia Passive continental Margin, Proc. ODP Sci. Res. 107, 29± 35. Savelli, C., 1988. Late Oligocene to Recent episodes of magmatism in and around the Tyrrhenian Sea: implication for the process of opening in a young inter-arc basin of infra-orogenic (Mediterranean) type. Tectonophysics 146, 163±181. Serri, G., 1990. Neogene±Quaternary magmatism of the Tyrrhenian region: characterization of the magma sources and geodynamic implications. Mem. Soc. Geol. Ital. 41, 219±242. 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. 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, N.J. (Eds.), Magmatism in the Ocean Basins, Geol. Soc. London Spec. Publ. 42, 313±345.

201

Ulmer, P., 1986. NORM-program for cation and oxygen mineral norms. Computer Library, Institute fuÈr Minelalogie und Petrographie, ETH-Zentrum, ZuÈrich, Switzerland. Torelli, L., Tricart, P., Zitellini, N., Argnani, A., Bouhlel, H., Brancolini, G., De Cillia, C., De Santis, L., Peis, D., 1992. Une section sismique profonde de la chaõÃne MaghreÂbides-Apennins, du bassin tyrrheÂnien aÁ la plate-forme peÂlagienne (MeÂditerraneÂe centrale). C. R. Acad. Sci. Paris 315, 617±622. Tricart, P., Torelli, L., 1994. Extensionnal collapse related to compressionnal uplift in the Alpine Chain off northern Tunisia (Central Mediterranean). Tectonophysics 238, 317±329. Tricart, P., Torelli, L., Zitellini, N., Bouhlel, H., Creuzot, G., De Santis, L., Morlotti, E., Ouali, J., Peis, D., 1990. La tectonique d'inversion reÂcente dans le Canal de Sardaigne: reÂsultats de la campagne MATS 87. C. R. Acad. Sci. Paris 310, 1083±1088. Tricart, P., Torelli, L., Brancolini, G., Croce, M., De Santis, L., Peis, D., Zitellini, N., 1991. DeÂrive d'insulaires et dynamique meÂditerraneÂenne suivant le transect Sardaigne-Afrique. C. R. Acad. Sci. Paris 313, 801±806. Van Der Meulen, M., 1999. Slab detachment and the evolution of the Apenninic arc (Italy). Geol. Ultraiectina 170, 136. Venturelli, G., Thorpe, R.S., Dal Piaz, G.V., Del Moro, A., Potts, P.J., 1984. Petrogenesis of calc-alkaline, shoshonitic and associated volcanic rocks from the northwestern Alps, Italy. Contrib. Mineral. Petrol. 86, 209±220. Vila, J.M., Hernandez, J., Velde, D., 1974. Sur la preÂsence d'un ®lon de roche lamproõÈtique (trachyte potassique aÁ olivine) recoupant le ¯ysch de type Guerrouch entre Azzaba (exJemmapes) et Hammam-Meskoutine, dans l'Est du Constantinois (AlgeÂrie). C. R. Acad. Sci. 278, 2589±2592. Winchester, J.A., Floyd, P.A., 1977. Geochemical discrimination of different magma series and their differentiation products using immobile elements. Chem. Geol. 20, 325±343.