Sedimentary, stable isotope and micropaleontological records of paleoceanographic change in the Messinian Tripoli Formation (Sicily, Italy)

Palaeogeography, Palaeoclimatology, Palaeoecology 185 (2002) 255^286 www.elsevier.com/locate/palaeo

Sedimentary, stable isotope and micropaleontological records of paleoceanographic change in the Messinian Tripoli Formation (Sicily, Italy) M.-M. Blanc-Valleron a; , C. Pierre b , J.P. Caulet a , A. Caruso c , J.-M. Rouchy a , G. Cespuglio b , R. Sprovieri c , S. Pestrea d , E. Di Stefano c a

CNRS-FRE 2400, Lab. Ge¤ologie, Muse¤um National d’Histoire Naturelle, 43, rue Bu¡on, 75005 Paris, France b UMR 7617, CNRS-Universite¤-ORSTOM, Laboratoire d’Oce¤anographie Dynamique et de Climatologie, Universite¤ Pierre et Marie Curie, 4 Place Jussieu, 75252 Paris, Cedex 05, France c Dipartimento di Geologia and Geodesia, Universita' di Palermo, Corso Tukory 131, 90134 Palermo, Italy d Geological Institute, Lab. Paleontology, 3, Caransebes, 78344 Bucharest, Romania1 Received 21 June 2001; accepted 28 February 2002

Abstract The Tripoli Formation (6.96^5.98 Ma) of the Central Sicilian Basin provides a good record of the paleoceanographical changes that affected the Mediterranean during the transition from slightly restricted conditions to the onset of the Messinian Salinity Crisis. The Falconara/Gibliscemi section has been selected for an integrated approach at a high resolution scale using sedimentology, stable isotopes of the carbonates and microfossils. The sedimentary succession includes 46 precession-controlled cycles resulting from the periodical increase in biosiliceous productivity (diatomites) that followed the deposition of marls and pinkish laminites, which appear as sapropel-type deposits induced by the oceanic fertilization by terrestrial nutrients during wet periods. Higher scale environmental changes are superimposed to this precession forced rhythmicity. There is a general trend of increasing basin restriction from near marine conditions at the base of the Tripoli to semi-closed settings in its uppermost part, which are the prelude of the salinity crisis. This pattern reflects the hydrological response of the Mediterranean to the progressive decrease of the Atlantic inputs and an enhanced influence of the climate on depositional conditions. However, this evolution is not linear and shows successive phases of different duration. During the first period (until 6.71 Ma), open Atlantic^Mediterranean exchanges maintained relatively stable marine conditions. The second period (6.71^6.29 Ma), marks an important step in the basin restriction with a wider range of salinity fluctuations and an increased bottom stagnation. The 6.71-Ma event, which is correlated at a Mediterranean scale, may have resulted from shallowing of the Mediterranean gateway under a tectonic control. This shallowing reduced the oceanic inputs resulting in an increased climatic constraint of the Mediterranean hydrology. During the third period (6.29^6.03 Ma) an increase of the surface water salinity resulted in stressful conditions for the marine microfauna. The 6.29-Ma change is a major step in the restriction that may be correlated with the intensification of the glaciation recorded in the Atlantic, which could have enhanced the effects of the tectonic closure. The last two cycles (48 and 49), that underlie the ‘Calcare di

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Contribution to the Natenmar network funded by the European Union. * Corresponding author. Fax: +33-1-40-79-3739. E-mail address: [email protected] (M.-M. Blanc-Valleron).

0031-0182 / 02 / $ ^ see front matter > 2002 Elsevier Science B.V. All rights reserved. PII: S 0 0 3 1 - 0 1 8 2 ( 0 2 ) 0 0 3 0 2 - 4

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Base’, witnessed the rapid transition to a semi-closed Mediterranean setting characterized by large variations of salinity from diluted to hypersaline conditions, under a dominant climatic control, and by the nearly complete disappearance of the marine organisms. Long-trend environmental changes recognized within the Tripoli Formation resulted from a complex set of interfering factors controlling the water fluxes exchanged between the Mediterranean and the world ocean. Most of the rapid changes identified in Falconara/Gibliscemi at 7.16, 6.71 and 6.29 Ma, that occurred simultaneously in the western and eastern Mediterranean, were mainly controlled by the stepwise tectonic closure of the Atlantic connections, although a glacio-eustatic overprint cannot be completely excluded. > 2002 Elsevier Science B.V. All rights reserved. Keywords: Mediterranean; Messinian Salinity Crisis; Tripoli; stable isotopes; diatomites; micropaleontology

1. Introduction The salinity crisis that a¡ected the Mediterranean during the Messinian is one of the most signi¢cant evaporitic episodes of the stratigraphic record (Hsu« et al., 1973, 1978; Montadert et al., 1978; Rouchy, 1982; Cita and McKenzie, 1986). The Messinian Salinity Crisis (MSC), the paroxysmal phase of which is dated 5.7^5.96 to 5.33 Ma (Gautier et al., 1994; Krijgsman et al., 1999a), was preceded by sedimentation of cyclic bedded biosiliceous deposits (Tripoli Formation) in most Mediterranean peripheral basins from Algeria to Cyprus (Rouchy, 1982), as well as in some deepsea areas (ODP Site 372, Cita et al., 1978; ODP Site 654, Pierre and Rouchy, 1990). The Tripoli Formation records an important depositional change that occurred after the monotonous deposition of the middle Serravallian to lowermost Messinian marls in deep open marine conditions. Diatomites of the Tripoli Formation grade upward into the evaporitic succession, which is heralded in Sicily by the ‘Calcare di Base’ (Ogniben, 1957; Decima and Wezel, 1973). Evaporites are overlain by a detrital unit deposited in brackish waters, or in terrestrial environments, known usually as ‘Lago^Mare’ environments, which include the uppermost Messinian Arenazzolo Member in Sicily. Its sharp boundary with the overlying pelagic Early Pliocene Trubi Formation in most Mediterranean basins marks the rapid marine re-

£ooding of the Mediterranean by Atlantic waters when the Gibraltar gate opened (Cita and Ryan, 1973; Cita, 1975; Di Stefano et al., 1999; Sgarrella et al., 1999; Rouchy et al., 2001). The increased biosiliceous productivity that resulted in diatomites and preceded the salinity crisis in the Mediterranean, has been related to different factors, i.e. upwelling of marine deep waters (McKenzie et al., 1980; Rouchy, 1982, 1986; Mu«ller, 1985; Mu«ller and Hsu«, 1987) and increase of the terrestrial nutrients supply (Van der Zwaan, 1979; Hodell et al., 1994; Rouchy et al., 1998). These factors have been interpreted as related to either the increasing isolation of the Mediterranean (McKenzie et al., 1980) or the global eustatic control (Suc et al., 1995), as well as to the succession of these two factors (Rouchy, 1982; Rouchy and Saint-Martin, 1992). In the Caltanissetta Basin, Sicily, several sections provide a continuous record of the latest Tortonian^early Messinian environmental changes with a characteristically cyclic pattern of the Tripoli Formation. Most studies previously carried out on these sections were mainly focused on stratigraphy (Rouchy, 1982; Sprovieri et al., 1996a,b; Hilgen et al., 1995; Hilgen and Krijgsman, 1999). The purpose of this paper is to reconstruct the environmental changes leading to the MSC during the transition from the Tortonian open marine conditions to the Messinian evaporitic environ-

Fig. 1. (a) Cyclostratigraphic and biostratigraphic correlations of the Falconara and Gibliscemi sections used to construct the composite section of Fig. 2; correlations with the insolation curve following the age model of Hilgen and Krijgsman (1999) and Sierro et al. (2001). (b) Astronomical age of individual cycles of the composite section. Ages refer to the mid-points of pinky laminites without time lag relative to the correlative insolation maxima.

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ments, using a high resolution sampling of the whole Tripoli Formation in order to obtain a continuous record of the depositional evolution through the transition to evaporitic conditions. The Falconara/Gibliscemi section (Figs. 1a and 2) was selected due to its location in the deepest part of the Caltanissetta Basin, which could be compared to deeper Mediterranean areas and should not be a¡ected by nearshore environmental changes. To achieve this goal, a multidisciplinary approach combining micropaleontology (foraminifera, calcareous nannoplankton, radiolarians, diatoms), sedimentology and stable isotope analyses was developed.

2. Geological background Since the discovery of Messinian evaporites beneath the abyssal plain of the Mediterranean (Ryan et al., 1973), the Neogene deposits of central Sicily have been largely studied for analyzing the MSC and the depositional environment of the pre-evaporitic Tripoli Formation. The most important outcrops are found in the NE^SW trending Caltanissetta Basin (Fig. 2b), which is part of the Maghrebian thrust belt that runs through North Africa, Sicily and the Italian Apennines (Catalano et al., 1995). In the central part of the Caltanissetta Basin the evaporitic succession, which can be compared to the one known in the deep Mediterranean troughs, is intercalated between deep-water deposits of Tortonian^early Messinian and Zanclean age (Schreiber et al., 1976; Rouchy, 1982; Schreiber, 1988). It starts with the Lower Gypsum Unit and the massive Halite Unit, which are overlain by the Upper Gypsum Unit (Decima and Wezel, 1973; Rouchy, 1982). Many authors have interpreted the Caltanissetta Basin as a deep basin tectonically uplifted near the end of the Pliocene, due to major thrusting on the outer margin of the Maghrebian orogen (Kastens and Mascle, 1990). On the northern edge of the Caltanissetta Basin the Messinian sedimentation is thought to have occurred in perched sub-basins located in front of active thrusts (Grasso and Pedley, 1988). These sub-basins, which should be considered as synclines related

to underlying thrust structures of the frontal part of the Maghrebian chain, deepened following the tectonic subsidence caused by the lithosphere £exure under the load of the orogenic wedge developing towards the North (Grasso et al., 1990; Pedley and Grasso, 1993; Butler et al., 1995). The Tripoli Formation is 0^90 m thick, depending on its location within the basin structure (Monnier, 1978; Butler et al., 1995; Sprovieri et al., 1996b). The overlying ‘Calcare di Base’, up to 50 m in thickness, is generally rich in gypsum and halite pseudomorphs, indicating that these carbonates represent the initial stage of the evaporitic deposition (Ogniben, 1957; Decima and Wezel, 1973; Schreiber et al., 1976; McKenzie et al., 1980; Rouchy, 1982). The evaporitic minerals, usually observed in the ‘Calcare di Base’ as pseudomorphs, are mostly typical of interstitial growth processes from oversaturated brines trapped within carbonate muds, but they do not characterize a subaqueous precipitation from basin waters. Therefore, they represent a transitional stage between the beginning of hypersaline conditions observed in the upper part of the Tripoli (McKenzie et al., 1980; Bellanca et al., 2001), and the onset of the true MSC de¢ned by the massive deposition of subaqueous evaporites. However, the presence of primary laminated gypsum deposits, intercalated within and locally at the base of the ‘Calcare di Base’ in some parts of the Caltanissetta Basin, shows that the beginning of the subaqueous evaporite deposition occurred diachronously within the Caltanissetta Basin (Monnier, 1978; Rouchy, 1982; Bellanca and Neri, 1986; Bellanca et al., 2001).

3. The studied sections 3.1. Location and lithological description The studied sections are located in the southern part of the Caltanissetta Basin. The Gibliscemi section outcrops along the southern slope of Monte Gibliscemi (14‡16P05‘E, 37‡12P16PN). The Tripoli lies upon the Tortonian^lowermost Messinian marls (‘Argille di Licata’ or ‘Marne a Globigerine’) (Ogniben, 1966; Sprovieri et al., 1996a,b).

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Fig. 2. (a) Falconara/Gibliscemi composite section. The S column corresponds to the lithologic cycle numbers of Sprovieri et al., 1996a (levels numbered 122.1, 122.2, 122.3 and 124.1 were missing in the section originally sampled due to minor faulting and were sampled in a nearby gully at the end of 1999). The C column corresponds to the cycle numbers used in this paper and in Bellanca et al. (2001). Age refers to the astronomical tuning of Fig. 1b. (b) Map of the Caltanissetta Basin with location of the two sections studied. (c) Panoramic view of the Gibliscemi section showing the Tortonian marls at the base, the Tripoli and ‘Calcare di Base’ on top. (d) View of a typical cycle of the Tripoli Formation (Falconara cycle 15).

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As the upper Tripoli levels are tectonically disturbed at Gibliscemi, they were sampled about 23 km E^NE from the previous location, on the southern slope of Monte Cantigaglione (14‡34P08‘E, 37‡7P38PN), about 3.5 km NW of Castello di Falconara (Falconara section). The middle Tortonian/lower Messinian marls and the Tripoli Formation are characterized by 145 lithologic cycles controlled by the astronomical precession (Hilgen et al., 1995; Sprovieri et al., 1996a). The Tortonian^lowermost Messinian cycles are represented by couplets of grey homogeneous marls and more or less reddish laminites generally topped by a dark-brown manganiferous crust (Sprovieri et al., 1996a). One precursor white diatomite bed, 50 cm thick, is known within the Tortonian deposits at Gibliscemi, 29 cycles below the base of the Tripoli, with an age of 7.9 Ma. The total thickness of the Tripoli Formation measured in the Falconara/Gibliscemi composite section is about 39 m (27 m at Falconara) and corresponds to 46 cycles (4^49; Fig. 2a). The Tripoli Formation grades upward into the ‘Calcare di Base’ through three cycles with carbonate levels (McKenzie et al., 1980; Decima et al., 1988; Sprovieri et al., 1996a; Caruso, 1999; Bellanca et al., 2001), which display the ¢rst evidence of evaporite (Fig. 2a and c). Due to the poor development of the diatomites and the sandy nature of the laminites these three cycles were not studied in detail. The thicknesses of the Tripoli cycles vary between 10 and 170 cm at Falconara, and between 55 and 190 cm at Gibliscemi. Some cycles exhibit great variations in thickness between the two sections: cycle 16 is 70 cm thick at Falconara and 165 cm at Gibliscemi; cycles 17^23 total 333 cm at Falconara and 875 cm at Gibliscemi. The sedimentation rate of the Tripoli Formation is around 2.5 times higher at Gibliscemi than at Falconara, showing the signi¢cant e¡ects of local tectonics and/or varying sediment supply rates during the Tripoli Formation in the Caltanissetta Basin. In this paper, as in Bellanca et al. (2001), the numbering of the diatomite beds (Fig. 2a, numbers 4^49 at the right side of the lithological column) refers to the numbering proposed by Hilgen

and Krijgsman (1999). Our cycle de¢nition arbitrarily assumes that marlstones represent the base of the cycles, whereas in Hilgen and Krijgsman (1999, ¢gure 2) reddish-brown laminites (called sapropels) are considered as the base of the cycles and are correlated with maxima of the insolation curve and minima of the precession curve (Fig. 1b). The Gibliscemi and Falconara sections were correlated using the LO of G. nicolae and FCO of N. atlantica occurrences (Fig. 1a). 3.2. Stratigraphic framework The stratigraphic framework was provided by biostratigraphic data because the magnetostratigraphic analysis was not conclusive due to remagnetization of the sediments (Langereis and Dekkers, 1992). The stratigraphic position of the Tripoli Formation in the Caltanissetta Basin, calibrated to the paleomagnetic reversals record, signi¢cantly varies according to authors. The middle Tortonian to early Messinian Falconara/Gibliscemi series were recently the subject of many biostratigraphic and cyclostratigraphic studies (Hilgen et al., 1995; Krijgsman et al., 1995; Sprovieri et al., 1996a,b, 1999; Caruso, 1999; Hilgen and Krijgsman, 1999; Krijgsman et al., 1999a; Bellanca et al., 2001). It results that the age of the base of the Tripoli is 6.92 Ma for Sprovieri et al. (1999), and 7.01 Ma for Hilgen and Krijgsman (1999). Using the astronomical tuning of individual cycles of Hilgen and Krijgsman (1999), the base of the Tripoli is dated to 6.96 Ma in this paper (Figs. 1b and 2a; Table 1). Biostratigraphic events used for dating the Falconara reference section were ¢rst reported in Sprovieri et al. (1996a), and later re¢ned with a higher resolution sampling (Sprovieri et al., 1996b; Caruso, 1999; Hilgen and Krijgsman, 1999; Bellanca et al., 2001). The chronostratigraphic framework used here for the upper Tortonian^lower Messinian section is given in Table 1. The Tortonian/Messinian (T/M) boundary may be placed at the FO of Globorotalia conomiozea ss., whereas it is placed by other authors at the FO of the Globorotalia miotumida group (Sierro, 1985; Sierro et al., 1993; Hilgen et al., 2000),

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corresponding to the Globorotalia conomiozea group of Zachariasse (1979). This results in a difference of four precession cycles, i.e. around 80 ka, between the proposed ages for the T/M boundary, i.e. 7.16 Ma for Sprovieri et al. (1999) and 7.24 Ma for Hilgen and Krijgsman (1999). The integrated stratigraphy based on biostratigraphic and cyclostratigraphic data demonstrated that the cyclicity is constrained by astronomical precession (Krijgsman et al., 1995; Sprovieri et al., 1996a,b; Hilgen and Krijgsman, 1999). The average duration of one cycle is about 21 ka and, thus, the sampling interval (4^6 samples per cycle) can be estimated at 5250^3500 yr, which gives the average time-resolution of our sedimentology, micropaleontology and stable isotopes studies.

4. Methodology Sedimentological, geochemical and micropaleontological markers were analyzed on a compo-

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site section extending from cycles 4^23 in the Gibliscemi section, and cycles 24^49 in the Falconara section. In our interpretation, cycles 1^3 are not part of the Tripoli Formation, but cycle 3 was included in the study, as a basal reference. 4.1. Lithologic composition The bulk mineralogy and carbonate content of 293 samples were analyzed in 164 samples of the Falconara and 129 samples of the Gibliscemi sections. The carbonate content of each sample was measured on 100 mg of powdered sediment using a manocalcimeter (MCM). The mineral phases were identi¢ed by means of a Siemens D-500 X-ray di¡raction (XRD) instrument (CuKK, Ni ¢ltered, radiation). Calcite and dolomite percentages were calculated according to both MCM and XRD data. A semi-quantitative analysis of the other main well-crystallized minerals (mainly clay minerals, quartz, feldspars, gypsum, halite) was made using peak height above background of each mineral and quantitative coe⁄cient taking

Table 1 Chronostratigraphic framework of the lower part of the Messinian in the Caltanissetta Basin Events

Hilgen and Krijgsman, 1999

Sierro et al., 2001

Bellanca et al., 2001

This paper

Top of the Tripoli N. acostaensis sin. in£ux (up to 40%) N. acostaensis sin. in£ux (up to 90%) G. scitula group in£ux N. acostaensis S/D coiling change FCO T. multiloba LCO N. atlantica LO G. miotumida gr., LO Globorotalia conomiozea FCO N. atlantica LO G. nicolae FO G. nicolae Tripoli basis FO G. conomiozea s.s (Tortonian/Messinian boundary sensu Sprovieri et al., 1996a) In£ux of the high conical forms of the G. miotumida gr. sensu Sierro et al., 2001 FO G. miotumida group (Tortonian/Messinian boundary sensu Hilgen et al., 2000)

5.98 6.087 6.126 6.295 6.337 6.415

5.98 6.082 6.126 6.291 6.360 6.415

5.97 (50 T)

5.98 (49 T)

6.506

6.504

6.722

6.713 6.828

7.008

7.19

7.205

7.24

7.242

6.13 (43 T) 6.35 6.40 6.43 6.53 6.64 6.73

(32 (30 (28 (24 (18 (14

T) m) T) Tr) m) m)

6.34 (33 Tr) 6.415 (29 m) 6.425 (28 T) 6.51 (24 T) 6.65 (17 T) 6.72 (14 Tr) 6.82 (10 m) 6.96 (4 m) 7.16

Data reported from Hilgen and Krijgsman (1999); Bellanca et al. (2001); Sierro et al. (2001) and our work. Data within brackets correspond to the cycle number and facies within the composite Falconara/Gibliscemi section studied. Abbreviations: T, Tripoli; Tr, Tripoli rosato, i.e. pinkish laminites; m, marls).

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in account the di¡racting abilities of each mineral; results were calibrated relative to the carbonate percentages. Traces of celestine and pyrite are also punctually present. The undosed fraction corresponds to amorphous opal (mainly within diatomites), and poorly crystallized components (e.g. Fe oxy-hydroxides). Opal cannot be directly estimated from the XRD results as it corresponds to more or less recristallized biosiliceous remains (from opal-A to opal CT) with no real di¡raction peak, but an elevation of the background around 22‡ 2a. The relative importance of opal was estimated through smear slide analysis, showing the importance of the biosiliceous remains, especially in the diatomite facies. In marls and pinkish laminites, signi¢cant amounts of amorphous components are also present. They do not correspond to organic matter (OM) as the total organic carbon (TOC) is generally low and partly corresponds to the terrestrial OM fragments observed in smear slides. TOC values are 6 0.2% in the central part of marls and diatomites of the Falconara section (Cortese and Bjorklund, 1999), and 6 0.15% in the pinkish laminites of the Falconara/Gibliscemi composite section (Rock^Eval analysis of 20 samples). These amorphous components, usually associated to darker color and high Fe2 O3 X-ray £uorescence (XRF) values, may be due to amorphous or poorly crystalline coatings of iron oxyhydroxides (hematite, goethite ?) on sediment particles. Some ¢ne-grained carbonates, biosiliceous deposits and diagenetic carbonates were also characterized by scanning electron microscopy. Smear slides were examined for each sample in normal light microscopy. They provided semiquantitative data, when calibrated with carbonate values, about the relative importance of the di¡erent biogenic carbonate and siliceous remains.

gen and carbon isotopic compositions (N18 O, N13 C) of the calcitic phase were, thus, determined on the bulk sediment of 201 samples. The isotopic composition of the dolomite fraction from 64 samples was analyzed in the dolomitic marl interval occurring at the base of many cycles and in the uppermost dolomitic interval underlying the ‘Calcare di Base’. The ‘calcite’ value is taken as corresponding to the fraction of CO2 produced after 20-min reaction at 25‡C with phosphoric acid. Dolomite from carbonate mixtures was isolated by selective attack with acetic acid 1N for 20 min and later reacted with phosphoric acid at 25‡C for four days. The CO2 gas was analyzed on a triple collector mass spectrometer (VGSIRA 9). The N values are expressed in x relative to the PDB reference, and the N18 O values of dolomites are corrected by 30.8x for the fractionation e¡ect during the phosphoric acid reaction (Sharma and Clayton, 1965). 4.3. Foraminifera Foraminiferal data already obtained by previous studies (Sprovieri et al., 1996a,b) were completed for the additional cycles. Only the fraction greater than 125 Wm was generally considered for the quantitative analysis of the foraminiferal assemblages (256 samples studied from cycles 1^48). Some samples, in cycles 48, 47, 46, 43, 35, 32+33, 30, 29, 25, 23, 21, 19, 18, 16, 14, 11, were not useful because either the foraminiferal tests were absent (smaller than 125 Wm) or badly preserved. The samples are generally rich in planktonic foraminifera. Rare samples, essentially from claystone layers, are barren, or with a foraminiferal assemblage including only benthic species such as Bulimina echinata, Bolivina dilatata. Some shallow water benthic foraminifera (Elphidium, Asterigerinata, Discorbis and Cibicides) may be present in the cycles of the upper part of the section.

4.2. Oxygen and carbon isotopic composition 4.4. Calcareous nannoplankton Because of the poor preservation of many foraminifera (i.e. recrystallization, carbonate cementation of the chambers, occasional presence of pyrite) it was decided to analyze the bulk carbonate rather than the individual foraminifera. The oxy-

Calcareous nannoplankton data were obtained from the 224 samples already studied by Sprovieri et al. (1996b). Nannofossils are generally well preserved. Semi-quantitative data (relative speci¢c

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percentages within the associations studied), as for the foraminiferal assemblages, were used as temperature or productivity proxies following the ecological data published in the literature (see synthesis about Messinian species in Caruso, 1999). 4.5. Radiolarians Occurrence of well-preserved radiolarians was observed in 126 samples and radiolarian debris were recognized in 44 other samples. Presence/absence of 135 morphotypes was recognized in the 126 radiolarian-rich samples prepared following the standard techniques of San¢lippo et al. (1985). All well-preserved morphotypes in each sample were identi¢ed at species level, on a smear slide containing a total of 6 000^10 000 specimens and radiolarian debris. For each sample, three radiolarian indices were calculated (Caulet et al., 1997): (1) diversity index (total number of morphotypes present), (2) ‘temperate North Atlantic’ index (number of temperate to subarctic Atlantic morphotypes present), and (3) ‘oceanic’ index (ratio of oceanic representatives of Stichocorys peregrina versus representatives of ‘local’ morphotypes of the same species). This last index was based on the total number of ‘local’ representatives of S. peregrina counted for 100 representatives of the ‘oceanic’ variety. 4.6. Diatoms Diatom assemblages were analyzed in 52 samples characteristic of 23 Tripoli cycles, with 300^

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500 specimens counted for each sample according to the method used by Schrader and Gersonde (1978). Occurrence of specimens pertaining to a total of 46 planktonic and 14 benthic species was counted in each sample. Chaetoceros were not taken into account as, ¢rstly, they are less abundant ( 6 5%) than in other sections such as Masseria il Salto, where Fourtanier (in: Gaudant et al., 1996) estimated a contribution up to 25%, and, secondly, they are resting spores often quoted in upwelling areas but di⁄cult to determine at speci¢c level and without temperature or bathymetry signi¢cance. Most diatoms are marine planktonic species classi¢ed as oceanic, oceanic^neritic, or littoral, using paleoecological data from Wornardt (1967), Barron (1975), Fourtanier et al. (1991), and Mansour et al. (1995). Some forms were also used as sea surface temperature indicators, i.e. ‘hot’ species of tropical and subtropical a⁄nities, ‘cold’ species of boreal a⁄nity and cosmopolite forms were distinguished using the classi¢cation proposed by Barron (1973) and data from Schrader (1973), Gardette (1979), Baldauf (1987), and Mansour et al. (1995).

5. Variations of mineralogy, micropaleontology and stable isotope compositions In the Falconara/Gibliscemi composite section, the Tripoli interval analyzed in this paper corresponds to the cycles 4^49, the three lower ones (1^3) reported in the Tripoli Formation by Hilgen and Krijgsman (1999) being without true diatomite layers. The section displays three di¡erent

Fig. 3. Variations of mineralogical and calcite stable isotope values versus depth scale within one cycle (Falconara cycle 16, 70 cm thick).

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Fig. 4. The Falconara/Gibliscemi composite section (Sicily): mineralogical data versus time scale.

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Fig. 5. The Falconara/Gibliscemi composite section (Sicily): percentages of the various biogenic components estimated from smear slides data versus time scale.

types of variations recorded by changes in mineralogy, fossil assemblages and stable isotope geochemistry of the carbonates (Figs. 3^9). The ¢rst type corresponds to the elementary sedimentary

cycles, which record a high frequency variation forced by the precession (Fig. 3). The second type of variation corresponds to groupings of precession-controlled cycles, which exhibit similar

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patterns. The third type is a general trend from the base to the top of the Tripoli Formation. Superposed to these changes, short events may be de¢ned by sharp negative excursions of the N13 C values of the calcite. 5.1. Elementary cycles 5.1.1. Mineralogy The cyclicity is mainly illustrated by wide variations of the mineralogy (Fig. 4), especially the carbonate (3^76%) and clay (5^48%) percentages which usually dominate the bulk mineralogy while the other terrigenous components (quartz and feldspars) remain in smaller proportions, generally below 12%. The other dominant component is the biogenic opal, which is mostly present in the diatomite levels. Carbonate fraction usually predominates in the basal marls of each cycle with an average value of 39% while its abundance decreases in the diatomites and pinkish laminites with mean values around 28%. The composition of the carbonate fraction is dominated by calcite, although dolomite is common, being sometimes the major component. Except in the uppermost part of the section (cycles 48 and 49) where it predominates, dolomite is mainly present in the basal marls with mean values of 17%, as compared to 3^4% in the laminites. The amount of siliciclastics (quartz, clays and feldspars) is generally higher in the lowermost part of the marl layers, increasing often in the pinkish laminites. Gypsum is a minor component ( 6 2%), except in some red clays and pinkish laminites where its proportion may be comprised between 10 and 70%. Halite content is generally low ( 6 3%) and must be of secondary origin, crystallizing from £uids trapped in interstitial pores during drying of the sediments. Celestine is irregularly distributed in the marls, or in the pinkish laminites. Pyrite has been identi¢ed in some samples, but its amount ( 6 1%) is probably underestimated by the XRD method. Iron oxy-hydroxides (hematite, goethite?), inferred from smear slide observations and high Fe2 O3 XRF results (up to 6.9%), correspond to badly crystallized products giving no obvious X-ray pattern ; they are more abundant

in the pinkish laminites and in the red clays immediately below. Opal of biogenic origin represents up to 80% of the diatomitic layer; it is made of di¡erent biosiliceous remains i.e. diatoms, radiolarians, silico£agellates, and sponge spicules. Mean values of biosiliceous fractions estimated after smear slide analysis decrease from around 45% in the diatomites to 10% in the pinkish laminites, and 5% in the marls. The diatoms, which are always dominant, may reach up to 80%, while both the radiolarian and silico£agellate fractions do not exceed 10%, and sponge spicules 5% (Fig. 5). 5.1.2. Stable isotope composition of calcite The N18 O values of calcites exhibit large variations within each cycle the range of which varies along the section around an average value comprised between 0 and 2x (Fig. 6). The range of variation is commonly comprised between 32x and 3x along most of the section, but it strongly increases in the two uppermost cycles preceding the ‘Calcare di Base’ with values comprised between 33.73x and 4.88x. The most striking feature is the strong decrease of the N18 O values in the pinkish laminites by more than 1 to 1.5x with respect to the N18 O values in the marls and diatomites (Figs. 3 and 6). This decrease also affects the oxygen isotopic composition of the carbonates in the basal sample of diatomitic layers immediately overlying the pinkish laminites. The N13 C values of calcites also display large variations from 321.19 to 30.73x. Most of the values commonly £uctuate from about 31x at the base of the Tripoli to 34x near the top, with very sharp excursions down to 39 to 321x in some intervals. The lowermost values are generally measured in the marls or in the transitional level between the Tripoli and the overlying marls. The stable isotope compositions of dolomites were measured in the cycles with signi¢cant dolomite percentages, more especially in the dolomiterich part at the base of the Tripoli (cycles 15^21) and in the upper part of the Tripoli. Corrected N18 O values of dolomites cover a wide range from 31.98 to +6.95x, whereas N13 C values exhibit a larger dispersion from 316.03 to 0.25x.

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5.1.3. Micropaleontology The percentages of the various biogenic carbonate components were roughly estimated after smear slide analysis: foraminifera (0^25%), coccoliths (0^40%) and discoasterids (0^25%). Foraminifera are usually more abundant in the laminites, as compared to the marls. In the calcareous nannoplankton the percentage of discoasterids versus nannoplankton is higher in the marls and in the pinkish laminites where mean values correspond respectively to 26 and 33%, as compared to 11% in the diatomites (Fig. 5). The diversity in the planktonic foraminifera assemblages (Sprovieri et al., 1996a,b; 1997) is lower in the marls than in the diatomites, which generally contain well-diversi¢ed assemblages with abundant warm-water planktonic species (Globigerinoides spp.) in their lower part. Cold-temperate water planktonic forms (Globigerina bulloides, Neogloboquadrina acostaensis, Turborotalita quinqueloba, Neogloboquadrina atlantica) are mainly present in the marls and in the topmost part of the diatomites (see cycles 4, 6, 11, 29, 37, 39; Figs. 7 and 8). Benthic foraminifera are mostly represented by a sparse, oligotypic assemblage essentially composed of Bulimina echinata and less frequently Bolivina dilatata. Oligotypic species assemblage, dominated by B. echinata and typical of low-oxygen conditions, is present in the grey claystones of cycles 16^47, increasing in abundance above cycle 28 (Sprovieri et al., 1996a). Shallow water benthic foraminiferal forms (Elphidium, Asterigerinata, Discorbis and Cibicides) are present above cycle 30, which may be indicative of shallower environments (Sprovieri et al., 1996b). Like foraminifera, calcareous nannoplankton assemblages are less diversi¢ed in the marls than in the reddish and white laminites. Some species considered as indicative of warm waters such as Discoaster pentaradiatus, Amaurolithus spp. and Sphaenolithus spp. (Bukry, 1981; Perch-Nielsen, 1985; Driever, 1988) are mostly present in the laminites (see cycles 13^48). Coccolithus pelagicus, considered as an indicator of cold waters with a low salinity (around 34x) (McIntyre and Be¤, 1967; Kennett, 1982) but also abundant in upwelling area (Cacha‹o and Moita, 2000) is mostly

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observed from the bottom up to cycle 22 with lowest percentages in the pinkish laminites (Fig. 8). Radiolarians are usually more frequent in the diatomites (Fig. 9). Diatom assemblages are mostly composed of marine planktonic species dominated by Thalassionema nitzschioides whereas benthic species are scarce (0^5%), poorly preserved, and often broken, suggesting an allochtonous origin. These assemblages are characterized by a wide geographical distribution, sometimes with a tendency to cosmopolitanism. Planktonic diatoms have been used as temperature or bathymetry indicators (Fig. 9). Their assemblages usually show, from bottom to top of the diatomitic layers, a decrease of the warm-water species and an increase of the cold-water species and littoral forms. Asterolampra acutiloba, A. grevillei, A. marylandica, Azpeitia nodulifer, Coscinodiscus asteromphallus, C. crenulatus, C. radiatus, C. stellaris, Hemidiscus cuneiformis, Nitzschia cylindrica, N. fossilis, N. marina, N. porterii, N. reinholdii, Thalassiosira convexa, T. eccentrica, T. cf. eccentrica, T. lineata, T. cf. lineata, T. miocaenica, T. oestrupii, T. praeconvexa, and T. symbolophora are considered as warm species of tropical and subtropical a⁄nities. Actinocyclus curvatulus, Coscinodiscus curvatulus, C. marginatus, C. oculus-iridis, Rhizosolenia hebetata, Thalassiothrix longissima are cold species of boreal a⁄nity. Occasionally high amounts (around 30%) of littoral forms are found in cycles 5^19 (Actinoptychus senarius-undulatus, Actinocyclus ehrenbergii), and in cycles 30^49 (A. cubitus). Some diatomite cycles (5 and 18), present high abundance of the marine/brackish species Paralia sulcata (ranging from 1 to 5.5%), but pure freshwater species are absent in diatom assemblages. 5.2. Variations involving groups of cycles The cycles of the Tripoli can be grouped into four intervals that present coeval variations of the mineralogical, microfossil, and geochemical markers. 5.2.1. First interval: cycles 4^14 (6.96^6.71 Ma) In this interval, the carbonate content varies

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Fig. 6. The Falconara/Gibliscemi composite section (Sicily): evolution of the stable isotope composition of the bulk calcite and dolomite fractions data versus time scale. Shaded areas correspond to intervals with very low values of N13 C calcite.

from 9 to 63% with the highest calcite contents usually located at the base of the marls. The dolomite content is low and represents 0^41% of the total carbonates.

The radiolarian assemblages are characterized by a high speci¢c diversity (up to 65 morphotypes), a relatively high ratio (up to 4%) of the oceanic forms versus Mediterranean forms of

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Fig. 7. The Falconara/Gibliscemi composite section (Sicily): relative frequencies of planktonic foraminifera species (calculated as percentages of the total planktonic foraminifera) data versus time scale. Planktonic foraminiferal events (see Fig. 8) have been numbered in stratigraphical order: (1) FO G. nicolae, (2) LO G. nicolae, (3) FCO N. atlantica, (4) LO G. miotumida, (5) LCO N. atlantica, (6) FCO T. multiloba, (7) N. acostaensis S/D coiling change, (8) N. acostaensis sinistral in£ux (up to 90%), (9) N. acostaensis sinistral in£ux (up to 40%).

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Fig. 8. The Falconara/Gibliscemi composite section (Sicily): relative frequencies of some planktonic foraminifera species, and selected species of nannofossils data versus time scale. The relative frequencies of the nannofossils species are calculated as percentages of the total nannofossils. For legend of main bioevents recognized, see Fig. 7.

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Fig. 9. The Falconara/Gibliscemi composite section (Sicily): radiolarian and diatoms data versus time scale. The ‘oceanic’ index corresponds to the ratio of oceanic representatives of Stichocorys peregrina versus representatives of ‘local’ morphotypes of the same species. The TNA index (Temperate North Atlantic) corresponds to the number of temperate to subarctic Atlantic radiolarian morphotypes present.

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S. peregrina, and a relative abundance of temperate North Atlantic forms reaching up to 20 species in cycle 14. Cold-water species predominate in diatom assemblages (mainly Coscinodiscus marginatus and Thalassiothrix longissima). Planktonic littoral diatoms (mainly Actinoptychus senariusundulatus) are relatively well represented (3^25%) especially in cycle 12. Planktonic foraminifera assemblages are characterized by high percentages of Globigerinoides representatives, especially Globigerinoides quadrilobatus in cycles 4^7 and 11^13, and G. obliquus in cycles 8^14. Globigerinoides are indicative of warm subtropical super¢cial waters (Sprovieri et al., 1999), G. quadrilobatus being considered as a good tracer of eutrophic (nutrient-rich) waters with normal salinity whereas G. obliquus s.l., extinct in the Lower Pleistocene but similar to G. ruber, is considered as a good tracer of oligotrophic (nutrient-poor) waters with salinity 6 34.5x or s 36x (Pujol and Vergnaud-Grazzini, 1995). So, occurrence of warm surface waters with varying salinity and nutrients content is inferred for the ¢rst interval. Slight decreases in the total Globigerinoides spp. and Globigerinoides quadrilobatus percentages coincide with higher abundances of temperate North Atlantic radiolarian species. Cycles 4^7 included the highest G. quadrilobatus percentages and the presence of G. miotumida and G. conomiozea which characterize intermediate temperate^cool waters incoming from the Atlantic (Kennett, 1982; Hilgen et al., 2000). G. nicolae is present in cycles 10^14; this species may be included in the ‘Globorotalia scitula’ group which is widespread in the Mediterranean in glacial times and is generally associated with cool waters (Be¤ and Hutson, 1977; Hemleben et al., 1989); its presence coincides with the highest percentages of temperate North Atlantic radiolarian species. Stable isotope compositions of the calcite show slight N18 O variations around an average value of 0.7x re£ecting relatively stable marine conditions, but the values lower than 0.4x measured in the pinkish laminites might indicate increased continental inputs during their deposition. N13 C values of the calcite exhibit a wide range (30.7 to 314.7x), with their minimum usually corresponding to the base of the marls and to highest

carbonate (calcite) contents. These minimum N13 C values decrease regularly from 31.7 to 314.7x between cycles 4 and 13. Five samples only, more dolomite-rich, were analyzed for the N18 O values of dolomite, which exhibit low values (32 to 2.2x). 5.2.2. Second interval: cycles 15^35 (6.71^6.29 Ma) The base of this interval is marked by one of the most important changes occurring along the section, that is de¢ned by three major modi¢cations in the global record of the chosen markers. The ¢rst modi¢cation is a sharp increase of the dolomite content, which reaches 49^100% of the total carbonate fraction of the marls whereas higher calcite values usually correspond to the diatomites (Figs. 4 and 6). The second modi¢cation is recorded by the assemblages of microfossils. The radiolarians exhibit a slight decrease of the diversity while the ‘oceanic index’ (ratio of oceanic versus Mediterranean forms of S. peregrina) falls to very low values ( 6 2%), except for a recurrence between cycles 21 and 26, and the number of temperate North Atlantic species falls to 1%. The diatoms show a signi¢cant decrease of the cold-water (2^35%) versus warm-water indicators (13^47%). Here the beginning of the predominance of the warm- versus cold-water planktonic diatoms is to be found. This still increases signi¢cantly in the upper part of the interval, between cycles 27 and 35. Planktonic littoral diatoms (mainly Actinoptychus senarius-undulatus, Actinocyclus ehrenbergii var. tenella, A. cubitus) are still present, but their number decreases signi¢cantly, except in cycles 16, 27, and 30. This interval is also characterized by the drastic reduction of the warm-water species within the planktonic foraminifera assemblage, specially G. quadrilobatus, whereas G. obliquus s.l. is still present in some diatomite levels. On the other hand, cold-water indicators such as Neogloboquadrina acostaensis and G. bulloides show a sharp increase whereas peaks of Turborotalita quinqueloba, N. atlantica, and temperate species (G. conomiozea ss. and G. miotumida) are also present. Neogloboquadrina atlantica, which is present in

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the interval between cycles 17 and 28, is now extinct, but it was exclusively present in the mid to high latitudes of the North Atlantic from late Miocene to late Pliocene (Poore and Berggren, 1975), suggesting that the areal distribution of this species was primarily controlled by cool surface water conditions. In the Mediterranean Pliocene, invasions of N. atlantica were associated with surface cooling events generally well correlated with glacial stages (Zachariasse et al., 1990). T. quinqueloba is classi¢ed as a super¢cially cool species common in nutrient-rich waters as its relative abundance increases from low to high latitudes (Parker, 1962; Tolderlund and Be¤, 1971). In the upper part of the interval (cycles 27^35), the warm-water species (Globigerinoides) became dominant with regard to the cold species (T. quinqueloba and G. bulloides). High percentages of G. quadrilobatus, a warm species characteristic of nutrient-rich surface waters with a normal salinity, are observed at the end of the interval (cycles 29^31). The calcareous nannoplankton shows a strong increase in the abundance of Sphenolithus from cycle 33 upwards. The third modi¢cation is clearly recorded by the oxygen isotope composition of the calcite (mean value 0.4x) which shows a greater range of variation (31.71x 6 N18 O 6 2.3x) than in the ¢rst interval. A signi¢cant shift towards more positive N18 O values of calcite (0.6x 6 N18 O 6 3x)

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occurs between cycles 22 and 26. In contrast, the N13 C values do not change signi¢cantly being always between 34 and 38x, except for few excursions towards very low values (down to 320.7x) and a diminution of the amplitude of variation from cycle 28 upwards where the values are close to 34x. The isotopic composition of the dolomite shows a clear evolution in this interval characterized by the drop of N13 C values from around 34x in cycles 15^21 to 316x and an increase of N18 O values from about 0.8x to about 6.8x. 5.2.3. Third interval: cycles 36^47 (6.29^6.03 Ma) This interval is mainly di¡erentiated by a new positive shift of the calcite N18 O values (0x 6 N18 O 6 3x), and the recurrence of very negative excursions of N13 C values down to 321x within an interval characterized by relatively homogeneous values close to 34x. It is also characterized by an increase of the N13 C values of the dolomite (311.1 to 32.8x) and by high, but variable, N18 O values (2.1^6.8x). The diversity of the planktonic foraminiferal assemblages continues to decrease. They are dominated by the association of T. quinqueloba and T. multiloba the positive £uctuations of which coincide with high N18 O values of calcite. If compared to the underlying cycles, this interval shows

Fig. 10. SEM microphotographs of the dolomite. (A) View of a dolomite-rich marl layer showing the euhedral to subhedral morphologies of the crystals. (B) Close-up of a dolomite crystal.

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the predominance of dextral forms of N. acostaensis versus senestral forms (especially in cycles 38^40, 44, 45 and 48). A recurrence of more abundant senestral forms exists in cycles 42 and 43, where T. multiloba is absent and N18 O values of calcite show a slight decrease indicating less saline or warmer waters. The abundance and diversity of radiolarians continue to decrease with less than 20 species from cycle 40 to 45, and the population becomes nearly monospeci¢c in cycles 46 and 47. Both the oceanic forms of S. peregrina and the temperate North Atlantic species nearly disappear in this interval. The diatom diversity also strongly decreases, with less than 15 species (instead of 25 species) from cycle 46 upwards. The proportion of warm-water indicators (2^63%) versus the temperate North Atlantic water indicators (0^40%) increases again except at the top of cycle 41, which presents a monospeci¢c population (99%) of cold-water species A. curvatulus. Percentages of planktonic littoral diatoms (mainly Actinocyclus ehrenbergii var. tenella) increase (0^30%) from bottom to top of the diatomitic layers. 5.2.4. Fourth interval: cycles 48 and 49 (6.03^5.98 Ma) This interval is characterized by major changes in the mineralogical, micropaleontological and isotopic records, which mark the transition between the true Tripoli unit and the ‘Calcare di Base’ where signi¢cant evaporitic conditions occurred. Dolomite becomes the dominant carbonate phase. Under SEM examination, it exhibits euhedral to subhedral shapes (average size generally around 5 Wm, Fig. 10), which are similar to those described in other Sicilian sections (McKenzie et al., 1980; Bellanca et al., 1986, 2001). These morphologies argue for an early diagenetic growth of the dolomite crystals. Both the abundance and the diversity of the microfossils lessen up to the disappearance of most of them in cycle 49. Only poorly diversi¢ed assemblages of diatoms persist, as well as scarce calcareous nannoplankton and low diversi¢ed assemblages of planktonic foraminifera. Warmwater diatoms (Coscinodiscus asteromphallus) generally predominate associated with littoral forms (Actinocyclus ehrenbergii var. tenella) in the

uppermost layer. The calcareous nannoplankton is mainly represented by nearly monospeci¢c assemblages of Sphaenolithus abies, which become very abundant in the pinkish laminites and reach maximum values just at the base of the white diatomite. The most striking feature is the wide range of variation in N18 O values of the calcite (from 33.7x to +4.9x), which indicates rapid changes from strongly diluted to highly evaporated solutions. The N13 C values display again negative excursions (315.3 6 N13 C 6 33.4x). This interval is also characterized by higher N13 C values (33.2 to +0.25x) and N18 O values (4^ 7x) of dolomite. 5.3. The general trend Although the di¡erent markers studied in these sections exhibit a clear stepwise evolution, marked by a succession of ¢ve major intervals, it appears that there is a general evolution from the base of the section to the two uppermost cycles. The mineralogy does not display a progressive change, but several modi¢cations that globally tend towards both a greater variability of the calcite/dolomite ratio and an increase of the dolomite proportion up to the last interval, where it constitutes the major carbonate mineral phase. From around 6.71 Ma upward, the dolomite becomes a major component in the marls although a decrease of its amount is observed in some cycles above. The last stage, which marks the dolomite dominance, corresponds to the uppermost transitional interval with the typical ‘Calcare di Base’. McKenzie (1985) and Bellanca et al. (2001) also reported the speci¢city of these last cycles. The microfossil assemblages display a general evolution towards the upward reduction of both abundance and species diversity, ending in a nearly complete disappearance of microfossils in the uppermost part of the section. Radiolarian assemblages show an upward decrease of diversity and a nearly complete disappearance in the four upper cycles. As mentioned above, this evolution occurs through two major changes recognized in the fossil assemblages successively at 6.71 Ma, which is marked by the decrease of the Atlantic

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and warm super¢cial and cold intermediate water forms. Around 6.29 Ma there is a sharp increase in T. quinqueloba and T. multiloba abundances, and Bulimina echinata FCO, which correspond to high salinity and cold super¢cial waters. The oxygen isotopic signal of the calcite exhibits a vertical evolution towards both heavier values up to the top of the third interval and higher amplitudes of variation. The range of variation signi¢cantly increases above 6.71 Ma, which consequently appears to represent one of the major steps in the evolution of local environmental conditions. The trend towards heavier isotopic compositions of the calcite appears well-marked above the second interval, where the lower values of N18 O are still positive and the variability of N18 O values increases abruptly in the uppermost interval. The N13 C values of the calcite display a complex behavior : a major set of relatively homogeneous values regularly decreases from ca 31x near the base of the section (cycle 4) to 33.5x at the top (cycle 48), and six intervals are characterized by sharp excursions towards negative values down to 321x. These excursions do not coincide with the intervals de¢ned using the N18 O values, suggesting they are dissociated from the processes controlling the oxygen isotopic composition. They may be interpreted as related to episodes of intense organic matter-induced diagenesis within bottom sediments. The N18 O values of the dolomites show a signi¢cant shift towards more positive values from near 0x in the ¢rst interval to 4^7x above the second interval, reaching 6^7x in cycle 49. Finally, a global evolutionary trend is observed along the section, but with two major changes, respectively at 6.71 and 6.29 Ma. At Falconara, another drastic event is recorded at the base of cycle 48 (at 6.03 Ma) when conditions became inhospitable for marine calcareous plankton and benthos. But, as shown by Bellanca et al. (2001), this change heralding the ‘Calcare di Base’ did not occur coevally in the di¡erent areas of the basin. It was recorded earlier in the marginal sections : around 6.13 Ma (cycle 43) at Marianopoli, 6.15 Ma (cycle 42) at Torrente Vaccarizzo and 6.32 Ma (cycle 34) at Serra Pirciata.

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6. Discussion The Falconara/Gibliscemi area (Caltanissetta Basin) is particularly important as it provides one of the most complete sedimentary succession available to study the main hydrological changes that occurred in the Mediterranean Basin at the beginning of the Messinian under the complex interaction of tectonic, eustatic and climatic constraints. The combined study of various paleoenvironmental markers shows that three types of environmental variations forced by diverse factors occurred on di¡erent time scales and interfered during the Tripoli deposition, i.e. (1) a general evolution towards a progressive reduction of the depth, more saline and highly unstable conditions and inhospitable environments for microfauna and micro/nanno£ora, (2) the successive steps in this restriction show no clear frequency, and (3) the high frequency changes, corresponding to individual cycles, are ascribed to the astronomical precession. In this section, we record the main hydrological changes in the basin and focus on ongoing, noncyclic trends. Then, we examine the cycle pattern in order to explain the high productivity of the Mediterranean during the formation of diatomites. 6.1. The depositional environments prior to the deposition of the Tripoli Formation The upper Tortonian^lower Messinian marls are not studied in this paper which is focused on the Tripoli Unit, as previous works have analyzed thoroughly the underlying succession both in the Falconara area (Sprovieri et al., 1999; Van der Zwaan, 1982) and the other basins i.e. Apennines (Kouwenhoven et al., 1999), Gavdos (Postma and ten Veen, 1999; Seidenkrantz et al., 2000), and Lorca (Rouchy et al., 1998). The results they provide on both the initial paleoenvironmental conditions of the basin and the precursor stages of the restriction are of a great importance to a better understanding of the changes that occurred during the deposition of the Tripoli until the onset of the MSC. These marls were deposited in a bathyal envi-

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ronment, at a water depth of at least 1000^1300 m and subtropical water conditions (Sprovieri et al., 1996a, 1999; Sierro et al., 2001). As in the overlying Tripoli Unit, the sedimentation displays a cyclic pattern forced by the precession, although the cycles di¡er from those of the Tripoli by the absence of diatomites, except for a precursor diatomitic layer astronomically calibrated around 7.9 Ma at Gibliscemi (Sprovieri et al., 1999). The ¢rst important event recognized in the Falconara section, around 7.50 Ma, is ascribed to the beginning of the late Miocene ‘Carbon shift’ (Sprovieri et al., 1999). In the Monte del Casino section (Apennines) this shift of the carbon isotopes towards lighter values is placed around 7.6 Ma and is correlated to a warming event, which may have led to a ¢rst slowing-down of the vertical circulation of the Mediterranean water masses (Kouwenhoven et al., 1999). After 7.6 Ma, the N13 C values in the Monte del Casino section remained stable until a new negative shift at 7.16 Ma (Kouwenhoven et al., 1999), while in the Falconara section the same interval exhibits a more progressive change. The 7.6^7.16-Ma interval corresponds globally, in the Mediterranean, to the late Miocene Carbon shift. Although less marked, the Carbon shift also exists in the Atlantic indicating a global event that received various interpretations (see references in Seidenkrantz et al., 2000). A new signi¢cant change is recorded at Falconara/Gibliscemi around 7.16 Ma by a drastic reduction and local disappearance of the benthic fauna (Van der Zwaan, 1982; Kouwenhoven et al., 1999). This event, related to bottom water conditions, is interpreted as being due to the increase of the water-mass strati¢cation related to a salinity increase of the bottom waters (Kouwenhoven et al., 1999). In somewhat shallower basins (Metochia and Monte del Casino sections) increasing environmental stress occurred between 7.16 and 6.8 Ma (Kouwenhoven et al., 1999), and a signi¢cant shallowing is inferred around 7.2 Ma in the Taza^Guercif Basin in the eastern part of the Ri¢an corridor (Krijgsman et al., 1999b). This decrease of benthic faunal diversity in the Mediterranean, from 7.16 Ma onward, has not been recorded on the Atlantic side of the Ri¢an corridor, so these changes are interpreted

as related to the relative uplift of the Atlantic^ Mediterranean waterway (Seidenkrantz et al., 2000). This tectonic control may have produced the constriction of the Atlantic connections through the Betic Strait that is recorded by an important change observed approximately at the Tortonian/Messinian boundary in some Betic Basins (Coppier et al., 1990; Montenat et al., 1990; Ott d’Estevou et al., 1990; Franseen et al., 1998; Rouchy et al., 1998; Soria et al., 1999). In contrast to bottom water changes, the surface waters did not undergo signi¢cant variation as indicated by the oxygen isotopic composition of the planktonic foraminifera at Falconara (Sprovieri et al., 1999) although Kouwenhoven et al. (1999) showed a general cooling beginning around 7.24 Ma and ending at 7.15 Ma in the Monte del Casino section. Prior to the beginning of the Tripoli sedimentation, the hydrology of the Mediterranean experienced thus several changes, already going towards a reduction of the vertical oceanic circulation due to the increasing strati¢cation of the water masses. This step-wise evolution was mostly controlled by the tectonics that increased the residence time of the Mediterranean bottom waters with no signi¢cant impact on surface marine conditions. 6.2. The progressive isolation of the basin during the Tripoli Formation The depositional environment became shallower from an average depth of 200^400 m at the bottom of the formation (Monnier, 1978; Broquet et al., 1984) to only a few metres at the top. The base of the Tripoli marks an important environmental change with a drastic increase in the biosiliceous productivity of Mediterranean waters, which occurred around 6.94 Ma. Thereafter, the shifts in faunal, sedimentological and geochemical parameters occurring at 6.71, 6.29 and 5.98 Ma, divide the Tripoli interval recorded in the Falconara/Gibliscemi sections into three major time slices, ending in the ¢nal closure of the evaporitic settings that occurred in the two uppermost cycles. They will be further discussed in chronological order.

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The N18 O values of calcite display a global evolution towards more positive values from around 0.7x near the base to values commonly comprised between 2 and 3x just below the uppermost cycles. This evolution can be interpreted as resulting from a global increase of water salinity. An ampli¢cation of the range of variation of N18 O values is superimposed to this global trend. It can be interpreted as the result of high frequency climatic variations the in£uence of which increased upwards as they became progressively less balanced by the oceanic inputs that were decreasing due to a worsening restriction of the Atlantic connections. The increasing in£uence of the carbon released by the oxidized organic matter is indicated by the regular decrease of the carbon isotopic composition from about 31x at the base of the Tripoli to 34x near the top. The sharply negative excursions of N13 C values down to 320x, which are superimposed to this global trend without any clear periodicity, could be due to periods of increased stagnation and processes of remineralization of organic matter through bacterial sulphate reduction. The lack of correlation with the N18 O curve does not permit, however, to clearly identify the factors that controlled these changes. All the assemblages of microfossils con¢rm this global trend towards worsening of the isolation as they exhibit a decrease in diversity to nearly complete disappearance near the top of the unit, except for diatoms. It seems that the diatoms are more tolerant to the environmental variations as diatomite layers still persist in the ‘Calcare di base’ as well as in the Upper Gypsum unit (Monnier, 1978; Rouchy, 1982; Bellanca et al., 1986, 2001). This evolution came to an end with the severe closure of the basin during the two uppermost cycles (6.03^5.98 Ma), which represent the transition to evaporitic conditions. The environment corresponded to semi-closed hydrological settings submitted to large salinity variations, from highly evaporated to diluted waters, recorded by wide variations of N18 O values, which range between 34 and +5x. The same evolutionary trend has been described in some sections located in more marginal settings, i.e. Serra Pirciata, Torrente

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Vaccarizzo and Marianopoli (Bellanca et al., 2001). The response of these marginal sub-basins to the increasing restriction was, indeed, ampli¢ed and they evolved earlier into semi-closed lagoonal conditions, as soon as 6.35 Ma at Serra Pirciata with a sharp decrease in calcareous plankton assemblages (Bellanca et al., 2001), that is to say around 350 kyr earlier than in the Falconara section. However, this global evolution was not linear, but went through several steps, as during the period prior to the deposition of the Tripoli (see 6.1. The depositional environments prior to the deposition of the Tripoli Formation). 6.2.1. The ¢rst period: large Atlantic^ Mediterranean exchanges until 6.71 Ma From 6.96 to 6.71 Ma (cycles 4^14), the marine conditions did not change signi¢cantly from those prevailing during the deposition of the underlying marl sequence. The surface waters were still characterized by normal salinity and subtropical temperatures as indicated by the relatively high diversity and great abundance of worldwide marine planktonic species, mostly the abundance of the warm-water Globigerinoides quadrilobatus. Except for the negative shift in the pinkish laminites, which is a common feature in all the individual cycles, the N18 O values around 0.7x attest to relatively stable marine conditions throughout the entire interval. Most of the N13 C values are close to 31x indicating also stable conditions except for sharp excursions towards very negative values down to 316x, which could result from the input of light carbon related to an enhanced oxidation of the organic matter. The high proportion of oceanic versus Mediterranean forms and temperate North Atlantic species in the radiolarian assemblages suggests that water exchanges with the Atlantic were active specially between 6.85 and 6.71 Ma. Due to the ecological requirements of some forms that are still living today, it may be inferred that some oceanic radiolarian species living near the thermocline entered the Mediterranean, suggesting that the oceanic sill was deep enough to permit the entrance of intermediate oceanic waters. Large exchanges with the Atlantic related to a high sea-level stand, or to the

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relative sinking of the sill between the Atlantic and the Mediterranean, may have been responsible for the presence of G. nicolae only in the upper part of this period (6.83^6.73 Ma), as cool and relatively deep waters £ew into the Mediterranean. In the entire Mediterranean, Rouchy (1982) and Rouchy and Saint-Martin (1992) correlated the ¢rst biosiliceous deposits accumulated basinward to a global sea-level highstand corresponding to the extensive development of the ¢rst Messinian carbonate platforms over most passive margins (for example at Melilla and Djebel Murdjadjo). 6.2.2. The second period (6.71^6.29 Ma): increased restriction The major change that is clearly recorded by all the studied markers towards 6.71 Ma in the Falconara/Gibliscemi section is particularly marked by a reduction of the diversity of radiolarian assemblages and the nearly complete disappearance of their Atlantic representatives. This lack of Atlantic forms is interpreted as corresponding to a reduction of the inputs of intermediate oceanic waters, or to increasing environmental stress. The presence of N. atlantica (6.65^6.425 Ma) indicates, however, that fairly large inputs of Atlantic water masses may have continued, while the benthic foraminifera assemblage (see 3 5.1) shows that bottom waters remained poorly ventilated. The increasing range of variation of the oxygen isotopic values also suggests that a wider range of salinity £uctuations took place in the surface waters. The composition of the Mediterranean water masses becoming less a¡ected by signi¢cant Atlantic inputs, there was an enhanced in£uence of the local climatic conditions marked by increased inputs of freshwater during wetter periods, and extension of more saline waters during drier intervals. From 6.71 to 6.55 Ma, the high N18 O values of the dolomite, commonly comprised between +4 and +6x, indicate precipitation from saline £uids trapped in bottom sediments. These values that signi¢cantly di¡er from those of the underlying interval (32 to +2x) provide also some evidence for hypersaline bottom conditions. These

were previously deduced, in several other areas, by the composition of the benthic foraminiferal assemblages (Kouwenhoven et al., 1999; Kouwenhoven, 2000; Seidenkrantz et al., 2000). Between 6.55 and 6.45 Ma the positive shift of the N18 O values of the calcite is interpreted as an overall increase of the salinity of the surface waters implying a signi¢cant reduction of the freshwater inputs during the deposition of the pinkish laminites. This shift, that is superimposed to the global trend, must have resulted from a climatic change tending towards drier conditions, which could have lasted around 100 ka. It was followed by a period (6.45^6.29 Ma) characterized by more negative and unstable N18 O values suggesting a stronger dilution by continental waters rather than an increase of oceanic inputs. From 6.5 Ma upwards the high values of the N18 O (around 6x) and the strong shift of the N13 C values (down to 316x) of the dolomite indicate a signi¢cant change in diagenetic conditions that may be interpreted as resulting from a bottom storage of more saline waters associated with processes of bacterial remineralization of organic matter in poorly ventilated conditions. Similar conditions were described in the upper part of the Tripoli in other Sicilian sections (Bellanca et al., 1986, 2001). This new step indicates a severe isolation of the basin the hydrological budget of which was no longer regulated by marine inputs, and had gone through the dominant control of climatic £uctuations, resulting in larger environmental changes. The environmental change recorded around 6.71 Ma may be correlated with events of the same age reported both in western and eastern Mediterranean basins. Kouwenhoven et al. (1999) observed a change at 6.7 Ma in the Monte del Casino section (Apennine) that they interpreted as a shallowing due to a regional uplift rather than to sea-level in£uence. Sapropel and diatomite deposition started at 6.79 Ma in the Sorbas Basin and around 6.7 Ma in the Melilla Basin (Morocco) in the westernmost part of the Mediterranean, and at the same time in Gavdos in the eastern Mediterranean (Sierro et al., 1999, 2001; Krijgsman et al., 1999a, 2001; Roger et al., 2000). In the Pissouri section of Cyprus, the

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¢rst sapropels appeared shortly after 6.7 Ma (Krijgsman et al., 2002). This was also the beginning of reef progradation in the Sorbas and Melilla basins (Martin and Braga, 1994; Roger et al., 2000). Therefore, the 6.7 Ma change must be considered as an important step in the Mediterranean restriction process under the major control of the tectonic closure of the connections with the Atlantic. This closure reduced signi¢cantly the thickness of the water column above the sill limiting the in£ux of oceanic waters entering the Mediterranean and preventing organisms living near the thermocline, as the oceanic radiolarians, to penetrate into the Mediterranean. The absence of signi¢cant change in the N18 O composition of the benthic foraminifera in the Atlantic (ODP site 982) during the same interval led Hodell et al. (2001) to conclude that this event was not in£uenced by a glacio-eustatic control. 6.2.3. The third period : further restriction at 6.29 Ma After 6.29 Ma, the reduction of the oceanic inputs in the Caltanisetta Basin must have been strong as suggested by the nearly complete disappearance of the Atlantic forms of radiolarians and the decrease of the planktonic foraminiferal assemblages. The abundance of T. multiloba and B. echinata, which are characteristic of stressed environments, probably indicates that hypersaline conditions developed. The global increase of salinity of the surface waters is also indicated by a new trend towards more positive values of N18 O values of calcite although the large variations of the values (0x 6 N18 O 6 3x) show that the surface conditions £uctuated due to periodic inputs of meteoric waters during the deposition of the pinkish laminites. The 6.29-Ma event recorded in the Caltanissetta Basin just predated the latest Miocene glaciation at 6.26 Ma, which has been identi¢ed in the ODP Site 982A in the North Atlantic (Hodell et al., 2001). There, the interval between 6.26 and 6 Ma, which corresponds to the onset of the MSC is characterized by several glacial^interglacial episodes forced by the astronomical obliquity. This would correspond to the intensi¢cation of the gla-

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ciation in the Antarctic and the development of glaciers in the Northern Hemisphere (Hodell et al., 2001). In contrast, the isotopic record of the foraminifera from the Bahamian transect rather indicates a warming during this period, the major cooling occurring later, near 6 Ma, that is to say more or less synchronously with the onset of the paroxysmal phase of salinity crisis (McKenzie et al., 1999). Zhang and Scott (1996) reported a global eustatic drop near 6.22 Ma in the North Atlantic Ocean (Sites 646B and 552A). Although problems of correlation exist between these di¡erent areas, it seems that during this period the world ocean experienced important glacio-eustatic variations. Sea-level lowstands may be considered as responsible for the aggravation of the e¡ects of the tectonic restriction of the Mediterranean and the subsequent increase of the water balance deficit. The 6.29-Ma event marked clearly the initiation of the rapid evolution towards hypersaline conditions that culminated in the next interval. 6.2.4. The isolation of the basin The two last cycles underlying the ‘Calcare di Base’ di¡er from the underlying succession by both the complete disappearance of the marine microfossils, except diatoms, and a wider range of variation of the N18 O values of calcites and dolomites. These data testify to the ¢nal restriction of the basin, which was submitted to rapid and large £uctuations of the environmental parameters under the dominant control of the climate. This interval can be considered as the true transition to evaporitic conditions which started during the ‘Calcare di base’ characterized by desiccation features and deposition of gypsum (Schreiber, 1974; Schreiber et al., 1976; Monnier, 1978; McKenzie et al., 1980; Rouchy, 1982; McKenzie, 1985; Decima et al., 1988; Bellanca et al., 2001). Similar variations of the depositional environment were reported at Serra Pirciata where the N18 O values of dolomite range between 34.7 to 7.1x (Bellanca et al., 2001). Such a transitional interval is well documented in other basins i.e. Sorbas with a carbonate^clay transitional interval (Sierro et al., 2001), Gavdos (Greece) with transitional carbonates (Krijgsman et al., 1999a), and Cyprus where a stromatolite-

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bearing interval displaying desiccation features and megabreccias just predates the onset of the gypsum deposition (Orszag-Sperber et al., 1980; Rouchy and Monty, 1981; Krijgsman et al., 2002). 6.3. The elementary cycles and the periodic increase of biosiliceous productivity The low frequency or aperiodical changes observed, although having an important impact on the biochemical parameters of the water masses, did not in£uence signi¢cantly the elementary cyclicity and the induced variations of productivity, which are under the precession control and continued up to near the ‘Calcare di Base’ and even inside it in several Sicilian sections. The formation of diatomites was related thus to a high frequency climatic control. The general trend of the calcareous faunal and £oral assemblages is the same in every cycle through the Tripoli Unit. The reddish and white laminites show a dominance of ‘warm’ planktonic assemblages (Globigerinoides spp., D. pentaradius, Sphaenolithus spp.) with often a sharp decrease in abundance in the topmost sample of each white level, in good agreement with the decrease of warm-water versus cold-water diatom species in the topmost part of the diatomites. It seems also that these assemblages indicate nutrientrich waters. Reciprocally ‘cold-water’ indicators (C. pelagicus, N. atlantica and T. quinqueloba) are mainly present in the marls and in the topmost part of the diatomites (Sprovieri et al., 1996b). The isotopic composition of oxygen displays a similar variation in each cycle characterized by a signi¢cant negative shift of N18 O values in the pinkish laminites (Fig. 3). The di¡erence, which rises from around 2x in the lower interval of the Tripoli up to 3x in the second interval, is too high to be explained by a temperature change (Fig. 6). It may have resulted from a periodic dilution by increased inputs of continental waters as indicated by a correlative increase of the amount of clastics (Fig. 3). The periodic climatic forcing was superimposed to the global evolution towards shallower and more saline water conditions. Regional climate oscillations, rather than

glacio-eustatically controlling in£uxes of Atlantic and/or Mediterranean waters, were responsible for this sedimentary cyclicity. Although the organic matter has been oxidized, the pinkish laminites are interpreted as sapropels, according to Krijgsman et al. (1995) and Hilgen and Krijgsman (1999), as indicated by the good preservation of the lamination, the presence of pyrite in some underlying Tortonian cycles and the high amounts of Fe-oxides that could result from the oxidation of sulphides. The formation of these sapropel horizons can be explained by a periodic strati¢cation of the water mass due to a surface dilution induced by large in£uxes of freshwater during episodes of warmer conditions. These interpretations rise the question of the origin of the periodical huge increases of biosiliceous productivity. They are not local events as diatomite-bearing units are known in many other Messinian basins i.e. Spain, Morocco, Algeria, Italy, Crete, and Cyprus (Rouchy, 1982, 1988; Howell et al., 1988; Fourtanier et al., 1991; Suc et al., 1995; Sierro et al., 1999; Krijgsman et al., 1999a; Triantaphyllou et al., 1999; Playa et al., 2000; Vazquez et al., 2000; Krijgsman et al., 2001; Sierro et al., 2001). The corresponding global increase of water fertility, which has contrasted with the more oligotrophic conditions that prevailed during deposition of the Tortonian^lower Messinian marls, has usually been ascribed to: (1) upwelling of deep waters (McKenzie et al., 1980; Rouchy, 1982, 1986; Mu«ller and Hsu«, 1987), (2) increased supply of continental waters (Van der Zwaan, 1979; Howell et al., 1988), and (3) water mass strati¢cation due to surface dilution of high concentrated waters by oceanic or continental in£uxes during sea-level £uctuations (Suc et al., 1995). Several of our data do not ¢t well with upwelling conditions as there are no typical radiolarian upwelling biomarkers in the biosiliceous levels, the foraminifer assemblages are rather dominated by warm-water forms, and representatives of G. bulloides, usually considered as an upwelling marker, are not abundant in the laminites. The fact that in each cycle the increase of biosiliceous productivity appears to have started rapidly, although progressively, after the deposition

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of the pinkish laminites rather favors a relationship with episodic inputs of continental waters enriched in nutrients of terrestrial origin, as previously proposed by Van der Zwaan (1979), Van der Zwaan and Gudjonsson (1986) and Howell et al. (1988). Therefore, the fertilization of marine waters was probably triggered by climatic £uctuations towards warmer and wetter conditions. It may be thought that these conditions of high productivity and rate of biogenic accumulation during the deposition of diatomites lasted a very short time compared to that of the underlying homogeneous marly intercalations. An important £ux of organic matter was rapidly transferred to the bottom water, increasing bottom-water anoxia. The low N13 C values of calcites in the marls (Fig. 6) indicate a large availability of 13 C-depleted biogenic CO2 derived from microbial oxidation of organic matter probably through processes of sulphate reduction.

7. Conclusions A combined approach of the sedimentology, micropaleontology, and oxygen and carbon isotope composition of the carbonates provides a high-resolution scenario of the successive hydrological changes through the transition from marine to hypersaline conditions in the Mediterranean. The environmental changes recognized along the section up to the base of the ‘Calcare di Base’ involved a complex set of interfering factors implying the tectonic closure of the Atlantic gateway connections the e¡ects of which on the water balance de¢cit of the Mediterranean was aggravated, in the upper part of the succession, by a glacio-eustatic forcing. High frequency climatic changes superimposed to these aperiodical events a prominent cyclicity characterized by variations of the biosiliceous productivity forced by the astronomical precession. Most of them have a global signi¢cance as they occurred simultaneously in several other basins. The most prominent feature is the global trend towards an increasing restriction of the Atlantic^Mediterranean water exchange that ended with the onset of the evaporitic conditions. It is not recorded as a linear

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trend, but as a step-wise evolution that is interpreted as a response to the tectonic closure of the Ri¢an gateway the e¡ects of which were probably ampli¢ed by glacio-eustatic sea-level changes. These evolutionary steps, as others observed in di¡erent regions, occurred at 7.16, 6.71 and 6.29 Ma, i.e. roughly with time lapses of about 400 kyr, suggesting that eccentricity forcing might have played a role in this evolution. The glacioeustatic variations of the world ocean level probably aggravated the e¡ects of the tectonic closure as the 6.29-Ma major change, which initiated the beginning of the evolution towards hypersaline conditions, roughly coincides with the beginning of the latest Miocene glaciation dated to 6.26 Ma in the North Atlantic. The reduction of the water depth above the sill limited the in£ux of less saline waters and the entrance of Atlantic microfaunal immigrants. The in£uence of the climatic constraint was ampli¢ed as the de¢cit of the water budget was less and less balanced by the entry of oceanic waters. The salinity and probably the temperature of the surface waters increased Mediterranean-wide whereas the reduction of the vertical circulation aggravated the bottom water stagnation. The evolution ended in the two uppermost cycles by the transition towards semi-closed hydrological settings submitted to large £uctuations of salinity from highly saline to diluted conditions, which marked in more marginal areas the prelude of the MSC. The overlying ‘Calcare di Base’ contains the ¢rst evidence of signi¢cant evaporitic precipitation and desiccation features. A regular cyclic sedimentary pattern forced by the astronomical precession is superimposed to this general evolution. It controlled periodical phases of enhanced biosiliceous productivity that occurred through the entire Tripoli interval until the top of the succession and locally continued into the overlying succession of the ‘Calcare di Base’, and even in the gypsum. These short-term climatic £uctuations induced strong changes in the surface water salinity especially within time intervals of increased continental inputs that might have corresponded to summer insolation maxima. They resulted in a dilution of the surface waters responsible for the strati¢cation of the water column at the origin of the deposition of

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Fig. 11. Main Messinian events recorded in the Mediterranean sections and in the North Atlantic, before the onset of the MSC.

the sapropel-like pinkish laminites of each cycle, and recorded by the N18 O curve. The fact that the periodical enhancement of the biosiliceous productivity occurred regularly and independently from both the restriction and global salinity trends, leads us to consider that the triggering factor was not so much the sporadic invasions of intermediate oceanic waters or upwelling waters, but rather the fertilization of surface waters by inputs of terrestrial nutrients during wetter and probably warmer periods. Although the 21-kyr orbital forcing is the only one that was demonstrated, a 100-kyr forcing is suspected in the interval between 6.71 and 6.29 Ma by the £uctuations of salinity recorded through the N18 O curve. These perturbations in an overall worsening of the restriction could be interpreted as related to variation in the volume of oceanic waters that periodically and more largely over£owed the barriers. This would need to be con¢rmed by more detailed studies. Some of the major steps as well as the shortterm orbital forcing recognized in the Falconara/ Gibliscemi section have also been reported from studies of sections in the Apennines, Gavdos, and Sorbas basins (Fig. 11). Their record and the age of some of them may nevertheless di¡er according

to the initial paleogeographical setting conditions of each basin, indicating that the regional constraint has to be considered when doing global correlations through the Mediterranean.

Acknowledgements This study was funded by the European Union, Human Capital and Mobility Programme, Contract ERBCH RXCT 930309 ^ Natenmar Network. Additional support was provided by the BQR Programme of the Museum National d’Histoire Naturelle (1996^1997), and MURST 60% to R. Sprovieri. This is a contribution to the project ‘La crise de salinite¤ messinienne’ of the French CNRS Eclipse Programme. We thank all people involved in the ¢eld work i.e. C. Taberner, J.J. Pueyo, J. Dinare's-Turell, and C. Taberner for helpful discussions concerning an earlier version of the manuscript. This paper was substantially improved by the thorough and constructive reviews of M.B. Cita and F. Hilgen. Thanks are due to P. Cle¤ment for XRD analysis, Ph. Blanc for SEM, M. Tamby and A.M. Brunet for sample processing, F. Savignac for Rock-Eval analyses, M. Destarac for the photography and A. Cambre-

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leng for the drawings. The help of S. Servant-Vildary during the study of the diatoms and of A. Foucault for the time series analysis are greatly acknowledged.

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